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
(c) 2011 Elsevier Inc. All Rights Reserved.
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
(c) 2011 Elsevier Inc. All Rights Reserved.
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.
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
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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
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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
(c) 2011 Elsevier Inc. All Rights Reserved.
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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
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Robert A. Duce Departments of Oceanography and Atmospheric Sciences Texas A&M University College Station TX, USA
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deceased
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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
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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
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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
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Stewart Turner Australian National University Research School of Earth Sciences Canberra Australia
James A. Yoder Woods Hole Oceanographic Institution Woods Hole MA, USA
deceased
(c) 2011 Elsevier Inc. All Rights Reserved.
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.
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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
(c) 2011 Elsevier Inc. All Rights Reserved.
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.
(c) 2011 Elsevier Inc. All Rights Reserved.
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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.
(c) 2011 Elsevier Inc. All Rights Reserved.
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
(c) 2011 Elsevier Inc. All Rights Reserved.
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Contributors
K G Foote Woods Hole Oceanographic Institution, Woods Hole, Massachusetts, USA L Franc¸ois 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
(c) 2011 Elsevier Inc. All Rights Reserved.
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
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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
(c) 2011 Elsevier Inc. All Rights Reserved.
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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
(c) 2011 Elsevier Inc. All Rights Reserved.
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
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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
(c) 2011 Elsevier Inc. All Rights Reserved.
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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 Muse´e Oce´anographique de Monaco, Monaco R A Dunn University of Hawaii at Manoa, Honolulu, HI, USA
(c) 2011 Elsevier Inc. All Rights Reserved.
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
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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
(c) 2011 Elsevier Inc. All Rights Reserved.
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
(c) 2011 Elsevier Inc. All Rights Reserved.
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
(c) 2011 Elsevier Inc. All Rights Reserved.
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.
(c) 2011 Elsevier Inc. All Rights Reserved.
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 Maran˜o´n 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.
(c) 2011 Elsevier Inc. All Rights Reserved.
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 Concepcio´n, Concepcio´n, 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
(c) 2011 Elsevier Inc. All Rights Reserved.
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.
(c) 2011 Elsevier Inc. All Rights Reserved.
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
(c) 2011 Elsevier Inc. All Rights Reserved.
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
(c) 2011 Elsevier Inc. All Rights Reserved.
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
(c) 2011 Elsevier Inc. All Rights Reserved.
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
(c) 2011 Elsevier Inc. All Rights Reserved.
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 Franc¸ois, 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
(c) 2011 Elsevier Inc. All Rights Reserved.
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
(c) 2011 Elsevier Inc. All Rights Reserved.
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 Nin˜o 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 Nin˜o 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.
(c) 2011 Elsevier Inc. All Rights Reserved.
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
(c) 2011 Elsevier Inc. All Rights Reserved.
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
(c) 2011 Elsevier Inc. All Rights Reserved.
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
(c) 2011 Elsevier Inc. All Rights Reserved.
xxxviii
Contents
Phytoplankton Blooms
D M Anderson
Phytoplankton Size Structure Plankton
M M Mullin
Plankton and Climate Plankton Viruses
432
E Maran˜o´n
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.
(c) 2011 Elsevier Inc. All Rights Reserved.
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.
(c) 2011 Elsevier Inc. All Rights Reserved.
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
(c) 2011 Elsevier Inc. All Rights Reserved.
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
(c) 2011 Elsevier Inc. All Rights Reserved.
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
(c) 2011 Elsevier Inc. All Rights Reserved.
ABRUPT CLIMATE CHANGE S. Rahmstorf, Potsdam Institute for Climate Impact Research, Potsdam, Germany & 2009 Elsevier Ltd. All rights reserved.
Introduction High-resolution paleoclimatic records from ice and sediment cores and other sources have revealed a number of dramatic climatic changes that occurred over surprisingly short times – a few decades or in some cases a few years. In Greenland, for example, temperature rose by 5–10 1C, snowfall rates doubled, and windblown dust decreased by an order of magnitude within 40 years at the end of the last glacial period. In the Sahara, an abrupt transition occurred around 5500 years ago from a relatively green shrubland supporting significant populations of animals and humans to the dry desert we know today. One could define an abrupt climate change simply as a large and rapid one – occurring faster than in a given time (say 30 years). The change from winter to summer, a very large change (in many places larger than the glacial–interglacial transition) occurring within 6 months, is, however, not an abrupt change in climate (or weather), it is rather a gradual transition following the solar forcing in its near-sinusoidal path. The term ‘abrupt’ implies not just rapidity but also reaching a breaking point, a threshold – it implies a change that does not smoothly follow the forcing but is rapid in comparison to it. This physical definition thus equates abrupt climate change with a strongly nonlinear response to the forcing. In this definition, the quaternary transitions from glacial to interglacial conditions and back, taking a few hundred or thousand years, are a prime example of abrupt climate change, as the underlying cause, the Earth’s orbital variations (Milankovich cycles), have timescales of tens of thousands of years. On the other hand, anthropogenic global warming occurring within a hundred years is not as such an abrupt climate change as long as it smoothly follows the increase in atmospheric carbon dioxide. Only if global warming triggered a nonlinear response, like a rapid ocean circulation change or decay of the West Antarctic Ice Sheet (WAIS), would one speak of an abrupt climate change.
Paleoclimatic Data A wealth of paleoclimatic data has been recovered from ice cores, sediment cores, corals, tree rings, and
other sources, and there have been significant advances in analysis and dating techniques. These advances allow a description of the characteristics of past climatic changes, including many abrupt ones, in terms of geographical patterns, timing, and affected climatic variables. For example, the ratio of oxygen isotopes in ice cores yields information about the temperature in the cloud from which the snow fell. Another way to determine temperature is to measure the isotopic composition of the nitrogen gas trapped in the ice, and it is also possible to directly measure the temperature in the borehole with a thermometer. Each method has advantages and drawbacks in terms of time resolution and reliability of the temperature calibration. Dust, carbon dioxide, and methane content of the prehistoric atmosphere can also be determined from ice cores. On long timescales, climatic variability throughout the past 2 My at least has been dominated by the Milankovich cycles in the Earth’s orbit around the sun – the cycles of precession, obliquity, and eccentricity with periods of roughly 23 ky, 41 ky, and 100 ky, respectively. Since the middle Pleistocene transition 1.2 Ma, the regular glaciations of our planet follow the 100-ky eccentricity cycle; even though this has only a rather weak direct influence on the solar radiation reaching the Earth, it modulates the much stronger other two cycles. The prevalence of the 100-ky cycle in climate is thus apparently a highly nonlinear response to the forcing that is likely linked to the nonlinear continental ice sheet and/or carbon cycle dynamics. The terminations of glaciations occur rather abruptly (Figure 1). Greenland ice cores show that the transition from the last Ice Age to the warm Holocene climate took about 1470 years, with much of the change occurring in only 40 years. The local Greenland response is not typical for the global response; however, since Greenland temperatures can be strongly affected by Atlantic Ocean circulation, which went through rapid changes during deglaciation. Globally, the transition from full Ice Age to Holocene conditions took around 5 ky. The ice ages were not just generally colder than the present climate but were also punctuated by abrupt climatic transitions. The best evidence for these transitions, known as Dansgaard–Oeschger (D/O) events, comes from the last ice age (Figure 2). D/O events typically start with an abrupt warming by up to 12 1C within a few decades or less, followed by gradual cooling over several hundred or thousand
(c) 2011 Elsevier Inc. All Rights Reserved.
1
2
ABRUPT CLIMATE CHANGE
18
O
LGM
I
II
III
IV
V
Eemian _ 500
_ 450
_ 400
_ 350
_ 300
_ 250
_ 200
_ 150
Holocene _ 50 0
_100
Age (1000 yr BP) Figure 1 Record of d18O from marine sediments (arbitrary units), reflecting mainly the changes in global ice volume during the past 50 ky. Note the rapid terminations (labeled with roman numbers) of glacial periods.
18
O (ppm)
_ 34
1 20 19 18
_ 38
17 161514
12 13
11 8 76 43 10 5 9
0
2
ΔT
_ 42
100
_ 20 80
60
40
20
0
Age (1000 yr BP) Figure 2 Record d18O from the GRIP ice core, a proxy for atmospheric temperature over Greenland (approximate temperature range DT (1C) is given on the right). Note the relatively stable Holocene climate during the past 10 ky and before that the much colder glacial climate punctuated by Dansgaard–Oeschger warm events (numbered). The timing of Heinrich events 1 to 6 is marked by black dots.
years. The cooling phase often ends with an abrupt final temperature drop back to cold (‘stadial’) conditions. Although first seen in the Greenland ice cores, the D/O events are not a local feature of Greenland climate. Figure 3 shows that subtropical sea surface temperatures in the Atlantic closely mirror the sequence of events in Greenland. Similar records have been found near Santa Barbara, California, in the Cariaco Basin off Venezuela, and off the coast of India. D/O climate change is centered on the North Atlantic and regions with strong atmospheric response to changes in the North Atlantic, with little response in the Southern Ocean or Antarctica. The ‘waiting time’ between successive D/O events is most often around 1470 years or, with decreasing probability, multiples of this period. This suggests the existence of an as yet unexplained 1470year cycle that often (but not always) triggers a D/O event. The second major type of abrupt event in glacial times is the Heinrich (H) event. H events involve surging of the Laurentide Ice Sheet through Hudson Strait, occurring in the cold stadial phase of some D/O cycles. They have a variable spacing of several thousand years. The icebergs released to the
North Atlantic during H events leave telltale dropstones in the ocean sediments when they melt, the socalled Heinrich layers. Sediment data suggest that H events shut down or at least drastically reduce the formation of North Atlantic Deep Water (NADW). Records from the South Atlantic and parts of Antarctica show that the cold H events in the North Atlantic were associated with unusual warming there (a fact sometimes referred to as ‘bipolar seesaw’). At the end of the last glacial, a particularly interesting abrupt climatic change took place, the so-called Younger Dryas event (12 800–11 500 years ago). Conditions had already warmed to near-interglacial conditions and continental ice sheets were retreating, when within decades the climate in the North Atlantic region switched back to glacial conditions for more than a thousand years. It has been speculated that the cooling resulted from a sudden influx of fresh water into the North Atlantic through St. Lawrence River, when an ice barrier holding back a huge meltwater lake on the North American continent broke. This could have shut down the Atlantic thermohaline circulation (i.e., the circulation driven by temperature and salinity differences), but evidence is controversial.
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3
22 20 18 _ 35
16
12 13
11 10
8 9
_ 40
GISP2 18O (ppm)
Temperature (°C)
ABRUPT CLIMATE CHANGE
_ 45 60
50
40
30
Age (1000 yr BP) Figure 3 Sea-surface temperatures derived from alkenones in marine sediments from the subtropical Atlantic (Bermuda Rise, upper curve) compared to d18O values from the GISP2 ice core in Greenland (lower curve).
Alternatively, the Younger Dryas may simply have been the last cold stadial period of the glacial following a temporary D/O warming event. Does abrupt climate change occur only during glacial times? Early evidence for the last interglacial, the Eemian, suggested abrupt changes there, but has since been refuted. During the present interglacial, the Holocene, climate was much more stable than during the last glacial. However, two abrupt events stand out. One is the 8200-year event that shows up as a cold spike in Arctic ice cores and affected the North Atlantic region. The second major change is the abrupt desertification of the Sahara 5500 years ago. There is much evidence from cave paintings, fire remains, bones, ancient lake sediments, and the like that the Sahara was a partly swampy savannah before this time. The best evidence for the abrupt ending of this benign climate comes from Atlantic sediments off northeastern Africa, which show a sudden and dramatic step-function increase in windblown dust, witnessing a drying of the adjacent continent.
Mechanisms of Abrupt Climate Change The increased spatial coverage, quality, and time resolution of paleoclimatic data as well as advances in computer modeling have led to a greater understanding of the mechanisms of abrupt climate change, although many aspects are still in dispute and not fully understood. The simplest concept for a mechanism causing abrupt climatic change is that of a threshold. A gradual change in external forcing (e.g., the change in insolation due to the Milankovich cycles) or in an internal climatic parameter (e.g., the slow buildup or melting of continental ice) continues until a specific
threshold value is reached where some qualitative change in climate is triggered. Various such critical thresholds are known to exist in the climate system. Continental ice sheets may have a stability threshold where they start to surge; the thermohaline ocean circulation has thresholds where deep-water formation shuts down or shifts location; methane hydrates in the seafloor have a temperature threshold where they change into the gas phase and bubble up into the atmosphere; and the atmosphere itself may have thresholds where large-scale circulation regimes (such as the monsoon) switch. For the D/O events, H events, and the Younger Dryas event discussed above, the paleoclimatic data clearly point to a crucial role of Atlantic Ocean circulation changes. Modeling and analytical studies of the Atlantic thermohaline circulation (sometimes called the ‘conveyor belt’) show that there are two positive feedback mechanisms leading to threshold behavior. The first, called advective feedback, is caused by the large-scale northward transport of salt by the Atlantic currents, which in turn strengthens the circulation by increasing density in the northern latitudes. The second, called convective feedback, is caused by the fact that oceanic convection creates conditions favorable for further convection. These (interconnected) feedbacks make convection and the large-scale thermohaline circulation self-sustaining within certain limits, with well-defined thresholds where the circulation changes to a qualitatively different mode. Three main circulation modes have been identified both in sediment data and in models (Figure 4): (1) a warm or interglacial mode with deep-water formation in the Nordic Seas and large oceanic heat transport to northern high latitudes (Figure 4(a)); (2) a cold or stadial mode with deep-water formation south of the shallow sill between Greenland, Iceland, and Scotland and with greatly reduced heat transport
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4
ABRUPT CLIMATE CHANGE
(a) 0
Depth (km)
1 2 3 4
‘Warm’
5 (b) 0
Depth (km)
1 2 3 4
‘Cold’
5 (c) 0
Depth (km)
1 2 3 4 5 30° S
‘Off’ 0°
30° N
60° N
90° N
Figure 4 Schematic of three major modes of Atlantic Ocean circulation. (a) ‘Warm’ or interglacial mode; (b) ‘cold’ or stadial mode; (c) ‘off’ or Heinrich mode. In the warm mode the Atlantic thermohaline circulation reaches north over the Greenland– Iceland–Scotland ridge into the Nordic Seas, while in the cold mode it stops south of Iceland. Switches between circulation modes at certain thresholds can pace and amplify climatic changes.
to high latitudes (Figure 4(b)); and (3) a ‘switched off’ or ‘Heinrich’ mode with practically no deepwater formation in the North Atlantic (Figure 4(c)). In the last mode, the Atlantic deep circulation is dominated by inflow of Antarctic Bottom Water (AABW) from the south. Many features of abrupt glacial climate can be explained by switches between these three circulation modes. Model simulations suggest that the cold stadial mode is the only stable mode in a glacial climate; it prevails during the cold stadial periods of the last glacial. D/O events can be interpreted as temporary incursions of warm Atlantic waters into the Nordic Seas and deep-water formation there, that is, a switch to the warm mode causing abrupt climatic warming in the North Atlantic region. As this mode is not stable in glacial conditions, the
circulation starts to gradually weaken and temperatures start to decline again immediately after the incursion, until the threshold is reached where convection in the Nordic Seas stops and the system reverts to the stable stadial mode. H events can be interpreted as a switch from the stadial mode to the H mode, that is, a shutdown of North Atlantic deepwater formation. As this mode is probably also unstable in glacial conditions, the system spontaneously reverts to the stadial or to the warm mode after a waiting time of centuries, the timescale being determined partly by slow oceanic mixing processes. This interpretation is consistent with the observed patterns of surface temperature change. The warming during D/O events is centered on the North Atlantic because this is where the change in oceanic heat transport occurs; the warm mode delivers heat to much higher latitudes than does the cold mode. A switch to the H mode, on the other hand, strongly reduces the interhemispheric heat transport from the South Atlantic to the North Atlantic. This cools the Northern Hemisphere while warming the Southern Hemisphere, explaining the ‘bipolar seesaw’ response in climate. It should also be noted that the initial transient response can differ from the equilibrium response as the oceanic heat storage capacity is large. The patterns of these abrupt changes differ from the longer-timescale (many thousands of years) response to the Milankovich cycles because, for the latter, the slow changes in atmospheric greenhouse gases (e.g., CO2) and continental ice cover act to globally synchronize and amplify climatic change. While the threshold behavior of the Atlantic ocean can dramatically shape and amplify climatic change, the question remains what triggers the mode switches. As mentioned above, D/O switches appear to be paced by an underlying 1470-year cyclicity that is as yet unexplained. This could either be an external (astronomical or solar) cycle or an internal oscillation of the climate system, perhaps also involving the Atlantic thermohaline circulation. A superposition of two major shorter solar cycles can, in climate model simulations, trigger events spaced 1470 years apart. The irregularity in D/O event timing is probably the result of the presence of stochastic variability in the climate system as well as the presence of longer-term trends such as the slow buildup of large continental ice sheets. The ocean circulation change during H events, on the other hand, can be explained by the large amounts of fresh water entering the North Atlantic at these times in the form of icebergs. Simulations show that the observed amounts of fresh water are sufficient to shut down deep-water formation in the North Atlantic. The nonlinear dynamics of ice sheets
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ABRUPT CLIMATE CHANGE
provide a plausible trigger mechanism. Ice sheets may grow for many thousands of years until their base melts owing to geothermal heating, when the ice sheet becomes unstable and surges. Thresholds in ocean circulation and continental or sea ice dynamics are not the only mechanisms that can cause abrupt climatic changes. In the desertification of the Sahara in the mid-Holocene, probably neither of these mechanisms were involved. Rather, an unstable positive feedback between vegetation cover (affecting albedo and evapotranspiration) and monsoon circulation in the atmosphere appears to have been responsible. It is almost certain that there are further nonlinearities in the climate system that could have caused abrupt climatic changes in the past or may do so in the future. We are only beginning to understand abrupt climate change, and the interpretations presented here – while consistent with data and model results – are not the only possible interpretations. Reflecting the state of this science, they are current working hypotheses rather than established and welltested theory.
Risk of Future Abrupt Changes The prevalence of abrupt nonlinear (rather than smooth and gradual) climatic change in the past naturally leads to the question whether such changes can be expected in the future, either by natural causes or by human interference. The main outside driving forces of past climatic changes are the Milankovich cycles. Close inspection of these cycles as well as modeling results indicate that we are presently enjoying an unusually quiet period in the climatic effect of these cycles, owing to the present minimum in eccentricity of the Earth’s orbit. The next large change in solar radiation that could trigger a new ice age is probably at least 30 ky away. If this is correct, it makes the Holocene an unusually long interglacial, comparable to the Holstein interglacial that occurred around 400 ka when the Earth’s orbit went through a similar pattern. This stable orbital situation leaves unpredictable events (such as meteorite impacts or a series of extremely large volcanic eruptions) and anthropogenic interference as possible causes for abrupt climatic changes in the lifetime of the next few generations of humans. Significant anthropogenic warming of the lower atmosphere and ocean surface will almost certainly occur in this century, raising concerns that nonlinear thresholds in the climate system could be exceeded and abrupt changes could be triggered at some point. Processes that have been (rather speculatively)
5
mentioned in this context include a collapse of the WAIS, a strongly enhanced greenhouse effect due to melting of permafrost or triggering of methane hydrate deposits at the seafloor, a large-scale wilting of forests when drought-tolerance thresholds are exceeded, nonlinear changes in monsoon regimes, and abrupt changes in ocean circulation. Of those possibilities, the risk of a change in ocean circulation is probably the best understood and perhaps also the least unlikely. Two factors could weaken the circulation and bring it closer to a threshold: the warming of the surface and a dilution of high-latitude waters with fresh water. The latter could result from an enhanced atmospheric water cycle and precipitation as well as meltwater runoff from Greenland and other glaciers. Both warming and fresh water input reduce surface density and thereby inhibit deep-water formation. Model simulations of global-warming scenarios so far suggest three possible responses: a shutdown of convection in the Labrador Sea, one of the two main NADW formation sites; a complete shutdown of NADW formation (i.e., similar to a switch to the H mode); and a shutdown of AABW formation. A transition to the stadial circulation mode has so far not been simulated, perhaps because convection in the Nordic Seas is strongly wind driven and is more effectively switched off by increased sea ice cover than by warming. A shutdown of Labrador Sea convection would be a significant qualitative change in the Atlantic Ocean circulation, but would probably affect only the surface climate of a smaller region surrounding the Labrador Sea. Effects on ecosystems and fisheries have not been investigated but could be severe. A complete shutdown of NADW formation would have wider climatic repercussions. Temperatures in northwestern Europe could initially rise several degrees in step with global warming, then abruptly drop back to near present values (the competing effects of raised atmospheric CO2 and reduced oceanic heat transport almost balancing). If CO2 levels decline again in future centuries as expected, European temperatures could remain several degrees below present as the Atlantic thermohaline circulation is not expected to recover perhaps for millennia. Further effects of a shutdown of deep-water renewal include reduced oceanic uptake of CO2 (enhancing the greenhouse effect), shifts in tropical rainfall belts, accelerated global sea level rise (due to a faster warming of the deep oceans), and rapid regional sea level rise in the northern Atlantic. The probability of major climatic thresholds being crossed in the coming centuries is difficult to establish and largely unknown. Currently, this possibility
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ABRUPT CLIMATE CHANGE
lies within the (still rather large) uncertainty range for future climate projections, so the risk cannot be ruled out. The IPCC 4th assessment report assigns a probability of up to 10% to a shutdown of the Atlantic overturning circulation within this century.
See also Authigenic Deposits. Deep-Sea Drilling Results. Past Climate from Corals. Ocean Circulation: Meridional Overturning Circulation.
Further Reading Abrantes F and Mix A (eds.) (1999) Reconstructing Ocean History – A Window Into the Future. New York: Plenum.
Broecker W (1987) Unpleasant surprises in the greenhouse? Nature 328: 123. Clark PU, Webb RS, and Keigwin LD (eds.) (1999) Mechanisms of Global Climate Change at Millennial Time Scales. Washington, DC: American Geophysical Union. Clark PU, Alley RB, and Pollard D (1999) Northern Hemisphere ice sheet influences on global climate change. Science 286: 1104--1111. Houghton JT, Meira Filho LG, and Callander BA (1995) Climate Change 1995. Cambridge, UK: Cambridge University Press. Sachs JP and Lehman SJ (1999) Subtropical North Atlantic temperatures 60,000 to 30,000 years ago. Science 286: 756--759. Stocker T (2000) Past and future reorganisations in the climate system. Quarterly Science Review (PAGES Special Issue) 19: 301--319. Taylor K (1999) Rapid climate change. American Scientist 87: 320.
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ABSORBANCE SPECTROSCOPY FOR CHEMICAL SENSORS R. Narayanaswamy, The University of Manchester, Manchester, UK F. Sevilla, III, University of Santo Tomas, Manila, The Philippines & 2009 Elsevier Ltd. All rights reserved.
the optical fiber to the phototransducer. The phototransducer generates an electrical signal that is related to the concentration of the analyte. This article focuses on chemical sensors that are based on absorbance spectroscopy. The various principles involved are reviewed, and some applications are presented.
Introduction
Absorbance and Reflectance Chemical sensors have introduced an alternative Spectroscopy technology for chemical measurements. These analytical devices generate an electrical signal in response to the presence of a specific substance, the signal being related to the concentration of the analyte. The application of these sensors has simplified chemical analysis, since the need for obtaining a laboratory sample is eliminated and a real-time and on-site measurement can be carried out. Furthermore, the configuration of these devices enables the detection with great sensitivity of low concentrations of chemical species. A number of analytical principles have been exploited in the development of chemical sensors. Among them are the optical methods of chemical analysis, which rely on the interaction of electromagnetic radiation with matter for the quantitation of a large number of substances. These methods provide a rapid and nondestructive tool for the measurement of chemical species. Optical methods have indeed played an important role, and continue to do so, in various field of chemical analysis. One group of chemical sensors that are based on the optical methods of chemical analysis is called optodes (or optrodes). Other types of optical sensing techniques that utilize waveguides, for example, surface plasmon resonance, also exist but are not reviewed here. The basic concept of optodes is depicted diagrammatically in Figure 1. An optode consists of an optical fiber, an optoelectronic instrumentation that incorporates a light source and a phototransducer, and a solid-phase molecular recognition element. Light from a suitable source is launched into the optical fiber and directed to the sensing zone which contains the molecular recognition element. The molecular recognition element reacts with the analyte, resulting in a modification of its optical property. This change is probed by the supplied radiation which is subsequently guided via
Ultraviolet (UV)/visible absorbance spectroscopy has been employed extensively in analytical chemistry. The basis for analysis here is the ability of the analyte or its derivative to absorb radiation that impinges into the measuring system. The radiation can have a wavelength occurring in the UV (200–400 nm), visible (400–780 nm), near-infrared (780–3000 nm), or infrared (3–50 mm) region. The absorption of radiation causes a reduction in the intensity of the radiation after it has passed through the system (Figure 2). This phenomenon is mathematically described by the Beer–Lambert law, as expressed by the following equation: I0 A ¼ log ¼ ecl It
½1
where A is the absorbance, c is the concentration of the absorbing species, I0 and It are the intensity of incident and transmitted light, respectively, e is the
Optoelectronic instrumentation Light source
Phototransducer
Optical fiber system
Molecular recognition element Figure 1 Basic design of optical fiber chemical sensor.
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8
ABSORBANCE SPECTROSCOPY FOR CHEMICAL SENSORS
Cladding
(a)
(b)
Transmitted radiation (It)
Incident radiation (Io)
Core
Analyte solution
Figure 3 Structure of an optical fiber. (a) Longitudinal cross-section view, and (b) latitudinal cross-section view.
Figure 2 Absorption of radiation.
molar absorptivity of the species, and l is the optical path length of the absorbing species. Absorbance spectroscopy is applicable only when the measurand system containing the analyte is transparent. However, if the medium is optically dense or even opaque, in which case the absorbance measurements would produce a high background, a technique that is complementary to absorptiometry (viz., reflectometry) can be employed for analytical measurements. In this case, the radiation infringes on the boundary interface of two media having different dielectric constants and reflection occurs. Two distinct types of reflection are possible, namely (1) specular (or mirror-type) reflection and (2) diffuse reflection. Specular reflection occurs at the interface of a medium with no transmission through it, and reflection is at the same angle as the incident light; whereas, in diffuse reflection, the light penetrates the medium and subsequently reappears at the surface after partial absorption and multiple scattering within the medium. Of these two processes, diffuse reflection has found to be useful in chemical measurements. Specular reflection is minimized or eliminated through appropriate sample preparation and optical engineering. The distribution of diffusely reflected light is rather homogeneous and largely independent of size and shape of the particles. The optical characteristics of diffuse reflectance are dependent on the composition of the system. Among several theoretical models that have been proposed for diffuse reflectance, the most widely used is the Kubelka–Munk theory. Here, the scattering layer is assumed to be infinitely thick, which may be effectively the case with molecular recognition element utilized in optical sensors, and the reflectance (R) is related to the absorption coefficient (K) and the scattering coefficient (S), as follows: FðRÞ ¼
ð1 RÞ2 K ¼ 2R S
½2
where F(R) is known as the Kubelka–Munk function. The absorption coefficient K can be expressed in
terms of the molar absorptivity (e) and the concentration (c) of the absorbing species. Thus, eqn [2] can be rewritten as FðRÞ ¼
ð1 RÞ2 ec ¼ ¼ kc 2R S
½3
where k ¼ e/S, and S is assumed to be independent of concentration. Equation [3] is analogous to the Beer– Lambert relationship (eqn [1]) and holds true within a range of concentrations for solid solutions in which the absorber is adsorbed onto the surface of a scattering particle. The reflectance values (R) are generally evaluated relative to the reflectance of standard reference materials such as barium sulfate.
Optical Fibers Almost all optical sensors employ optical fibers to transmit light to and from the molecular recognition element. The fiber couples the optoelectronic instrumentation to the molecular recognition element, resulting in an integrated analytical system. This integration has contributed to the simplification of the chemical measurement process. Optical fibers consist of a cylinder (known as the ‘core’) of transparent dielectric with a certain refractive index (n1), surrounded by a thin film of another dielectric (called the ‘cladding’) of a lower refractive index (n2). Most optical fibers are then covered with a protective jacket that has no influence on the wave-guiding properties of the optical fiber. The basic structure of an optical fiber is shown in Figure 3. The common materials used in optical fibers include plastic (poly(methyl methacrylate)), glass, and quartz. Incident light is transmitted through the fiber when it impinges the core–cladding interface at an angle greater than the critical angle, so that there is total internal reflection at the core–cladding interface. A series of total internal reflection takes place until the light reaches the other end of the fiber (Figure 4). The optical fiber light transmission characteristics are described by its numerical aperture
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ABSORBANCE SPECTROSCOPY FOR CHEMICAL SENSORS
Figure 4 Transmission of a light beam inside an optical fiber through total internal reflection.
Evanescent wave
Figure 5 Evanescent wave due to a light beam being transmitted inside an optical fiber.
(NA), which is directly proportional to the sine of the half angle (a) of the acceptance cone of light entering it which, in turn, is related to n1 and n2, as in eqn [4]: NA ¼ sin a ¼ ðn21 n22 Þ1=2
½4
where n0 is the refractive index of the surrounding medium, for example, air. Basically there are three kinds of optical fibers in use for sensing purposes. These are the multimode step index, the multimode graded index, and the single-mode step index fibers. These fibers differ in the number of light beams that travel through the length of the waveguide. The light rays accepted into the fiber interact, and only those which undergo constructive propagation traverse through the fiber. A small portion of light that is transmitted through an optical fiber or a waveguide by total internal reflection extends outside core and is referred to as the ‘evanescent wave’. In an optical fiber, the evanescent wave penetrates the cladding material (Figure 5). The intensity (I) of the evanescent wave decreases exponentially with increasing distance (d) from the surface according to eqn [5]: I ¼ I0 ed=dp
½5
where I0 is the electric field intensity at the core– cladding interface and dp, the depth of penetration, is the distance from the interface of the point at which the electric field intensity has reduced to 1/e of its value at the interface. This characteristic depth of penetration is dependent on wavelength of the propagating light, refractive indices of the core and the cladding, and the acceptance angle of the light
9
beam entering the fiber core. The distance dp is typically of the order of a fraction of a wavelength of light. The incorporation of optical fibers in chemical sensors imparts a number of advantages to optical sensors over the conventional devices in many areas of application. Optical sensors are electrically passive and immune to electromagnetic disturbances. They are geometrically flexible, corrosion-resistant, and capable of being miniaturized. They are compatible with telemetry and capable of operation in remote and hostile environments. They can be of low-cost, of rugged construction, and intrinsically safe. The optical fibers used in these sensors are capable of transmission of optical signals over great distances with low attenuation of optical power. Thus optical sensors are capable of measurements of samples in their dynamic environment, no matter how distant, difficult to reach, or harsh that environment is. Intrinsic safety aspects are imparted to these sensing devices by the low optical power utilized in them. Furthermore, the chemical sensing process itself is nonelectrical. With these sensors only very small sample volumes are needed for analysis, which has the advantages of nonperturbation of samples in real-time monitoring applications. However, optical sensing devices possess certain limitations, such as interference from ambient light, limited dynamic range, long response times, limited specificity, and nonreversibility. Many of these limitations can be eliminated or reduced by the use of appropriate instrumentation and sensing phases, and thus the sensor devices can be used advantageously in specific applications.
Optoelectronic Instrumentation Optical fiber sensors involve similar instrumentation as those involved in spectrophotometers. These sensors require both optical and electrical components, including a light source; a wavelength selector; a photodetector; and a readout device. A block diagram of the basic instrumentation associated with optical sensors is presented in Figure 6. The radiation supplied by the light source is first made monochromatic by a wavelength selector before it enters into optical fiber sensor system. Then, the radiation emanating from the sensor is directed to the photodetector which subsequently generates an electrical signal. This magnitude of the final signal is displayed in the readout device. Several types of sources, including incandescent lamps (tungsten and quartz–halogen), lasers, lightemitting diodes (LEDs), and laser diodes, have been
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ABSORBANCE SPECTROSCOPY FOR CHEMICAL SENSORS
Light source
Signal modulator
Monochromator
Photodetector
Lock-in amplifier
Optical fiber Readout device
Measurand system
Figure 6 Schematic diagram of a typical instrumentation system associated with optical fiber chemical sensors.
used. Each of these types of source has its own advantages and disadvantages. The detection of light is carried out using a photocounting device, which converts optical signals into electrical signals that can be amplified electronically. The photodetectors used in optical sensors include photomultiplier tubes, p–i–n photodiodes, avalanche photodiodes, and photodiode arrays. A wavelength selector, such as a filter or a monochromator, isolates the desired wavelength for the measurement. Optical couplers and lenses are used to focus the beam of light to the optical fiber and to direct light to the detector. It should be emphasized that efficient coupling of light ensures the attainment of high sensitivity of sensor signals. It is also common practice to exclude extraneous light reaching the detector by suitable modulation of the light source and by synchronizing the detector to this modulation frequency so as to detect only those source signals. Instrumental drift may be eliminated or reduced by the use of a suitable referencing system. Drifts in sensor response due to aging of the sensing phase and variable light source may be encountered. But these contributions may not be significant due to short periods of measurement periods used with the sensing devices.
Molecular Recognition Element The molecular recognition element constitutes the primary sensing unit of a chemical sensor. It interfaces the chemical sensor with the measurand system, generating an optical signal in response to the presence of the analyte. It transforms the analyte into another substance with distinct optical properties, so that it can be considered as a chemical transduction unit. The transformation can be detected and quantified by the instrumentation system via the optical fiber. However, if the analyte itself possesses a detectable optical characteristic, then that property can be directly measured using optical fibers and quantified to the concentrations of the analyte.
For absorbance- and reflectance-based chemical sensors, a number of reagents employed in colorimetry can serve as a molecular recognition element. Chromogenic reagents, such as pH colorimetric indicators and chelating reagents, have been used as chemical transduction elements in optical sensors. Novel materials, such as conducting polymers, molecularly imprinted polymers, and nanoparticles, have also been exploited for molecular recognition in optical fiber chemical sensors. The molecular recognition element is often employed in the solid state. The reagent(s) is usually immobilized onto an inert and stable solid material by physical methods (adsorption, entrapment, and electrostatic attraction) or by chemical means (covalent bond formation). Among the solid supports that have been used to immobilize the molecular recognition element are glass, silica gel, or organic polymers. The physical methods are simple and economical to carry out, but they do not necessarily produce stable reagent matrices. On the other hand, the chemical means of immobilization produce a strongly bound reagent, but this is often achieved after several reaction steps in the synthesis or modification of the reagent and/or the support material in order to realize stable chemical bond(s) between them. The sensor response characteristics in chemical transducer-based sensors depend on the manner in which the analyte–reagent interaction takes place. In a simple system, where direct indicators are employed, the analyte concentrations can be correlated to the optical changes that occur in the reagent phase or in the product or both. The correlated sensor signal may also be dependent on the equilibrium constant of the analyte–reagent reaction. For example, pH can be measured by monitoring the changes in optical property of acid-base indicators. In many sensors, reversible reactions are preferred in the chemical transduction process because they can be used in continuous monitoring applications. In this case, the response time of these sensors (i.e., the
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ABSORBANCE SPECTROSCOPY FOR CHEMICAL SENSORS
time to reach the reaction equilibrium) is dependent on mass transfer processes. Irreversible reactions may also be employed in sensors, which can result in ‘one-shot’ devices. Although such sensors would be of limited merit, measurements with high sensitivity can be attained here. In certain cases, the reagent phases can be regenerated by the use of another chemical reaction and the sensor reused. Indirect chemical reactions involving two or more reagents and/or reactions can also be adapted as optical chemical transducers. Many enzyme-based reactions fall into this category.
Sensor Design A variety of configurations have been utilized in sensors based on absorbance and other spectroscopic measurements. The sensor designs can be classified into two categories: extrinsic and intrinsic sensors. In extrinsic sensors, the optical fiber acts only as a light guide between the light source and chemical sensor and between the sensor and the detector. In intrinsic sensors, the optical fiber becomes a part of the transducer. In absorbance-based sensors, light is usually fed at one end of an optical fiber, guided to a sensing cell, collected through a second optical fiber, and detected at the other end of this fiber. If no chemical
(a)
From light source
11
transduction is employed (Figure 7(a)), the light interacts directly with the measurand system. This type of sensors has been described as ‘spectroscopic sensors’, since they are based on the spectral properties of the analyte. In sensors involving a molecular recognition element, the reagent phase is often translucent and is set at one end of the feed fiber and of the collector fiber (Figure7(b)). The analyte reacts with the immobilized reagent, and the product absorbs the light that passes through the reagent phase. In reflectance-based sensors, chemical transduction is often employed and the molecular recognition element is placed at one end of an optical fiber system. A single optical fiber or optical fiber bundles can be used in this type of optical sensors. In single-fiber sensors (Figure 8(a)), the source light and detected light travel through the same optical fiber, and are discriminated either temporally or by wavelength with the aid of a beam splitter. Optical fiber bundles are often configured as a bifurcated system (Figure 8(b)), wherein the incident radiation travels through one branch and the reflected light is directed to the photodetector through the other branch. Evanescent wave interactions have been exploited in optical fiber chemical sensors. In these devices, the cladding material of the optical fiber is removed and replaced with the analyte system itself (Figure 9(a))
Flow cell
To photodetector Optical fiber
Optical fiber
Analyte
(b)
From light source
Reagent film
Optical fiber
Optical fiber
To photodetector
Analyte inlet Figure 7 Typical configuration of absorbance-based chemical sensors: (a) employing no chemical transduction, and (b) with chemical transduction.
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ABSORBANCE SPECTROSCOPY FOR CHEMICAL SENSORS
(a)
(b)
From light source
From light source
To photodetector
Beam splitter
To photodetector
Optical fiber
Optical fiber
Molecular recognition element
Membrane
Molecular recognition element
Membrane
Figure 8 Typical configuration of reflectance-based chemical sensors: (a) employing a single fiber, and (b) employing a bifurcated optical fiber bundle.
(a)
Cladding
Optical fiber
From light source
To photodetector
Flow cell Analyte inlet (b)
From light source
To photodetector
Cladding
Reagent phase
Optical fiber
Figure 9 Typical configuration of evanescent-wave chemical sensors: (a) not employing chemical transduction, and (b) with chemical transduction.
or with a thin layer of the molecular recognition element (Figure 9(b)). The presence of the analyte affects the evanescent wave interaction and generates changes in the optical signal. In optical fiber absorptiometry or reflectometry, reference signals are employed for correcting
variations in the background caused by source, detector, and also intrinsic absorption of light by optical fibers. This can be done by subtracting the blank signal from the sensor signal electronically, or by the use of two identical optical fibers, one of which is used as a reference.
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ABSORBANCE SPECTROSCOPY FOR CHEMICAL SENSORS
Sensor Applications A number of optical fiber pH sensors have been developed based on the absorption characteristics of a colorimetric acid-base indicator. The reagent is immobilized on a finely powdered solid support or on a membrane that is deposited on the tip of the optical fiber and held in place by an enclosing membrane. One of the earliest reported optical fiber sensor for pH was based on the measurement of the absorbance of phenol red covalently bound to polyacrylamide mixed with polystyrene microspheres to scatter the light within the reagent phase. The sensor was configured as a probe employing a bifurcated optical fiber, with one arm for the probing radiation and the other arm for the scattered radiation from the reagent phase. The sensor measures pH in the range 7.0–7.4 to 0.01 pH units. This example illustrates the feasibility of the approach though the sensors may not be directly applicable to pH measurements in seawater. Different indicators may be employed for such applications. Absorbance-based optical sensing has been applied for the measurement of metal ions. These ion sensors involve the use of immobilized metal-ion-selective reagents interfaced to optical fibers. These sensors rely on the fact that the metal ion (M) reacts with the immobilized reagent (R) to form a metal complex (MR), accompanied by either an enhancement or change of color of the immobilized reagent. This change can be correlated to metal ion concentrations. Thus, an optical sensor for micromolar amounts of cobalt was developed employing a reagent phase consisting of pyrogallol red held on a cellulose acetate film. Likewise, a fiber optic sensor for copper ions occurring in ppm levels was devised based on a-benzoinoxime, a highly selective colorimetric reagent for copper ion, immobilized on hydrophobic reagent Amberlite XAD-2 microspheres. Optical sensors based on absorbance measurement have been developed for organic compounds, such as pesticides, occurring in environmental water systems. In the case of pesticide sensors, more than one type of transduction system is involved – an enzyme reagent and a pH indicator co-immobilized on suitable polymeric material. In the absence of pesticides, the reaction of the immobilized enzyme, such as acetylcholine esterase, with its specific substrate will be accompanied by a change in pH that can then be measured by the pH sensor using absorbance or reflectance. However, in the presence of the pesticides, the substrate/immobilized enzyme reaction is inhibited and the degree of inhibition, transduced in the pH sensor, can be correlated to the pesticide
13
concentration. Organophosphate and carbamate pesticides, and also many toxic metal ions, have been quantified using this principle. Absorbance measurement has also been used in gas-phase sensing. A nitrogen dioxide sensor has been constructed using optical fibers and measurement of absorbance of the gas at 0.5 mm using an argon-ion laser source. Real-time measurement in the lower ppm concentration range has been conducted at remote locations that are 20 km away. A similar device has been used for the measurement of methane at low concentrations by recording absorbance at 1.33 mm. Many sensors for gases and vapors such as ammonia, carbon dioxide, humidity, hydrogen cyanide, hydrogen sulfide, etc., have developed based on the use of analyte-specific immobilized reagents and measurements of absorbance or reflectance. For instance, the optical sensing of ammonia can be carried out through the change in color of polyaniline deposited on a polystyrene substrate. An optical sensor for gaseous hydrogen sulfide was based on the development of a grayish color on a cellulose membrane impregnated with lead acetate.
Conclusions As described above, many types of optical sensor designs have been studied for a variety of analytes using different types of transduction reactions in the development of absorbance- and reflectance-based sensors. A current trend in this development is to construct multianalyte sensing systems based on the use of single or a few reagent phases, together with the employment of appropriate signal-processing techniques such as pattern recognition and artificial neural networks. Most of the optical sensors described above can be designed for use in oceanographic measurements in the analysis of heavy metal ions, dissolved gases, and other species. Applications of absorbance-based sensors to ocean sciences have yet to demonstrate their potential. Most of the applications published in the literature describe ‘proof-of-concept’ studies with the sensors and with very little or no practical demonstration in such areas. Some of the recent studies in this area have been focused on the monitoring of dissolved CO2 in seawater using absorbance and fluorescence measurements with transducers that incorporate pH-sensitive indicators. The use of infrared absorption measured through the evanescent waves in optical fibers for subsea monitoring of organic compounds (at ppm levels) has also been demonstrated. Parameters such as sensitivity, specificity,
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14
ABSORBANCE SPECTROSCOPY FOR CHEMICAL SENSORS
lifetime, aging, etc., need to be investigated and addressed before the sensors can be used for seawater measurements. These studies would present great challenges in order to realize practical sensors. However, there is substantial interest in new sensors for oceanographic applications including monitoring of nutrients and pollutants, and optical sensors have great potential here. The devices are clearly attractive in concept and require expertise from several scientific disciplines including analytical chemistry, polymer chemistry, environmental chemistry, fiber optics, and opto-electronics.
See also Fluorometry for Chemical Sensing. Inherent Optical Properties and Irradiance. Wet Chemical Analysers.
Further Reading Andres R, Kuswandi B, and Narayanaswamy R (2001) Optical fiber biosensors based on immobilized enzymes – a review. The Analyst 126: 1469--1491. Eggins BR (2002) Chemical Sensors and Biosensors. London: Wiley. Hales B, Burgess L, and Emerson S (1997) An absorbancebased fiber-optic sensor for CO2(aq) measurement in pure waters of sea floor sediments. Marine Chemistry 59: 51--62. Harmer AL and Narayanaswamy R (1988) Spectroscopic and fibre-optic transducers. In: Edmunds TE (ed.) Chemical Sensors, ch. 13. Glasgow: Blackie. Lieberzeit PA and Dickert FL (2007) Sensor technology and its applications in environmental analysis. Analytical and Bioanalytical Chemistry 387: 237--247. Mizaikoff B (1999) Mid-infrared evanescent wave sensor – a novel approach for sub-sea monitoring. Measurement Science and Technology 10: 1185--1194.
Narayanaswamy R (1991) Current developments in optical biochemical sensors. Biosensors and Bioelectronics 6: 467--475. Narayanaswamy R (1993) Chemical transducers based on fibre optics for environmental monitoring. The Science of the Total Environment 135: 103--113. Narayanaswamy R (1993) Optical chemical sensors: Transduction and signal processing. Analyst 118: 317--322. Narayanaswamy R and Sevilla FS, III (1988) Optical fibre sensors for chemical species. Journal of Physics E: Scientific Instruments 21: 10--17. Narayanaswamy R and Wolfbeis OS (2004) Springer Series on Chemical Sensors and Biosensors, Vol. 1: Optical Sensors – Industrial Environmental and Diagnostic Applications. Berlin: Springer. Oehme I and Wolfbeis OS (1997) Optical sensors for determination of heavy metal ions. Microchimica Acta 126: 177--192. Orellana G and Moreno-Bondi MC (2005) Springer Series on Chemical Sensors and Biosensors, Vol. 3: Frontiers in Chemical Sensors – Novel Principles and Techniques. Berlin: Springer. Rogers KR and Poziomek EJ (1996) Fiber optic sensors for environmental monitoring. Chemosphere 33: 1151--1174. Seitz WR (1988) Chemical sensors based on immobilised indicators and fiber optics. CRC Critical Reviews in Analytical Chemistry 19: 135--171. Sevilla F, III and Narayanaswamy R (2003) Optical chemical sensors and biosensors. In: Alegret S (ed.) Integrated Analytical Systems, ch. 9. Amsterdam: Elsevier. Tokar JM and Dickey TDN (2000) Chemical sensor technology. Current and future applications. Ocean Science and Technology 1: 303--329. Varney MS (2000) Chemical Sensors in Oceanography. London: Taylor and Francis. Wise DL and Wingard LB, Jr. (1991) Biosensors with Fiberoptics. Clifton, NJ: Humana. Wolfbeis OS (1991) Fiber Optic Chemical Sensors and Biosensors, vols. I and II. Boca Raton, FL: CRC Press.
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ABYSSAL CURRENTS W. Zenk, Universita¨t Kiel, Kiel, Germany Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 1, pp 12–28, & 2001, Elsevier Ltd.
Introduction Historically the term ‘abyss’ characterizes the dark, apparently bottomless ocean under extreme static pressure far beyond coastal and shelf areas. Today this ancient definition remains still rather unfocused in earth sciences. Geographers, marine biologists, and geologists use abyss for deep-sea regions with water depths exceeding 1000 or 4000 m. In physical oceanography a widely accepted definition of the abyss denotes the water column that ranges from the base of the main thermocline down to the seabed. The main thermocline itself – occasionally also called warm-water sphere – extends laterally between the polar frontal zones of both hemispheres. In contrast to deep strata the surface exposition of the thermocline allows the direct exchange of heat, substances, and kinetic energy with the atmosphere. This
wind-driven part of the water column and its fluctuations are consequently an immediate subject of weather and climatic conditions. The base of the thermocline at about 1000–1200 m represents the lower boundary of the warm-water sphere with temperatures well 451C. The abyss or cold-water sphere below, is clearly colder. Below 2000 m potential temperatures o41C are found virtually everywhere. Below 4000 m values of 0–21C are more characteristic. Until the advent of modern self-recording instrumentation abyssal currents were believed to be very slow (o2 cm s1) and negligible in comparison with rather vigorous and variable surface currents (sometimes 4100 cm s1). Only subsurface passages in rises and ridges that subdivide ocean basins (Figure 1) seemed to allow for more energetic deep interior currents funneling through gaps and channels. Until a few decades ago practically all knowledge about abyssal currents had to be inferred from the perspective of property fields like potential temperature, salinity, or potential density. Since these fields seem to be much more stable than velocity fields this method still yields reasonable general
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Figure 1 Major subsurface fracture zones, passages, and sills of the world’s ocean. (A) Atlantic, (B) Pacific, (C) Indian Ocean. These topographic features represent important constraints for deep and bottom water spreading paths (after http:// www.soc.soton.ac.uk/JRD/OCCAM/sills.html).
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15
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large-scale flow patterns even with scarce and nonsynoptic data sets. However, for any resolution beyond the low-frequency band of the current spectrum direct current observations by floating objects, moored instruments, or remotely measuring methods are essential. In low latitudes of all oceans the permanent presence of the enormous, almost invariant temperature gradient between the surface at 4251C and the floor at o21C requires a deep advection of cold water. The latter circulates freshly ventilated cold water from polar latitudes towards the Equator. For reasons of continuity a compensating poleward flow of water heated from the surface through the thermocline must occur above the abyssal layer of the tropics and the subtropics. Hence, the cold water drift propelled by sinking of cooled water masses (convection in selected areas) accelerates an endless global circulation cell known as the meridional overturning circulation (MOC), occasionally called the global conveyor belt (Figure 2). Its bottomnearest limb, and sometimes also one or two layers in motion above it are characterized by the abyssal circulation.
The Stommel-Arons Concept and Diffusivity in the Interior The concept that rising water from a flat deep-sea bottom without density stratification has to be replaced by convectively formed water is a key element of the modern Stommel-Arons theory of abyssal circulation. While the freshly ventilated water sinks in only a few selected semi-enclosed polar regions, the rising process itself occurs over broad lateral scales everywhere at lower latitudes. As a consequence it is possible to distinguish between two different dynamical regimes of the abyssal circulation (Figure 3). (1) A small (typically only 100– 200 km wide) corridor at the continental rises and near the slopes on the western margins of the ocean’s basins is occupied by deep western boundary currents (DWBCs). (2) The huge remaining interiors of the basins, fed laterally by DWBCs, are ruled by a uniform broad-scale upwelling regime. The inherent vertical velocities imply vortex stretching on top of the abyss. Conservation of potential vorticity then requires the poleward return drift in the interior. Not long after the concept of DWBC was hypothesized in 1958, the first observations confirming its existence were made beneath the Gulf Stream in the north-western Atlantic using neutrally buoyant Swallow floats. Shortly afterwards, the newly detected
deep equatorward drift led to a supplementary experiment to find the slow poleward countercurrent of the interior flow. The test failed at the time (1962) in so far as no flow with a slow, persistently northward velocity component could be proved. Instead, the mesoscale eddy phenomenon at great depths was discovered. In steady state the vertical flux of heat at the base of the thermocline in lower latitudes can be formulated as the balance of temperature advection and its vertical diffusion: @T @2T ! ¼k 2 u rT þ w @z @z
½1
-
where u is the lateral current vector, T temperature, w vertical velocity component, z vertical coordinate, and k eddy diffusivity. This temperature equation implies for the interface between the base of the thermocline and the top of the abyss the generation of a vertical velocity to balance the upwelled cold water and the downward diffusion of heat. An assumed sinking rate of 20 106 m3 s1 in high latitudes and an active lower thermocline interface of 3 108 km6 yields a global integral upwelling speed of O(0.1 dm d1). An adjoined downward diffusion with a diffusivity of O(1 cm2 s1) can be estimated under the assumption of a vertical scale of 1 km within the abyssal upwelling regime. Figure 4 depicts a sketch of the integrated form of [1] under the assumption of negligible lateral currents along isotherms. w T ¼ kd
@T @z
½2
where kd is the cross-isothermal diffusivity. Observations of kd are difficult to conduct. On a global scale numerical values of kd fluctuate in a wide range (1–500 cm2 s1). A summary of the available sparse estimates from the Atlantic is reproduced in Figure 5. Figure 6 gives a rare example of the heterogeneity of cross-isopycnal diffusivity in the interior Brazil Basin. The diffusivity is quantified from uniform microstructure observations at all depths above the smooth abyssal plains and west of the Mid-Atlantic Ridge. Estimates range from very weak (o0.1 cm2 s1) to moderate (45 cm2 s1) turbulent signals. The latter enhancement is observed above the rough flanks of the ridge. The response is especially prominent within 500 m above the bottom. The abyss contains recirculation cells. Significant amounts of cold water do not participate directly in the global meridional overturning cell. They circulate horizontally in response to the upwelling process.
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Figure 2 The Atlantic thermocline circulation as a key element of the global oceanic circulation. (After Broecker (1991), modified by Meier-Reimer.)
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Figure 3 The Stommel-Arons (1958) concept of the abyssal circulation in the Atlantic. A system of western boundary currents feeds a slow broad-scale upwelling regime in the remaining interior of the basins. (Reprinted from Deep-Sea Research 5, Stommel H, The abyssal circulation, pp. 80–82, Copyright (1958), with permission from Elsevier Science.)
Such basin-wide recirculation cells (Figure 7) are distinctly influenced by the local bottom topography. Their persistence remains widely unexplored. In summary, it is of considerable interest to quantify rates of sinking waters in high latitudes because the compensating abyssal upwelling is believed to drive the internal horizontal circulation of the oceans.
are interconnected by the Southern Ocean as the circum-Antarctic water ring is called. In this strongly simplified pattern of the global water mass circulation the abyss ranges from the lower third of the displayed water column down to the seafloor.
3
Sources of Abyssal Waters Figure 8 represents a refined global update of the highly schematized deep Atlantic circulation pattern in Figure 3. Antarctica lies in the center. The North Pole cap is split; it is situated towards the outer edges of the Pacific and Atlantic blocks. All three oceans
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Figure 4 Schematic representation of eqn [2], the balance of downward diffusivity and upwelling across the 11C isotherm level of the western South Atlantic. (After Hogg et al., 1994 ^ Springer Verlag.)
Figure 5 Some estimates of cross-isotherm diffusivity in cm2 s1 in the abyss of the world’s oceans. (After Hogg, 2001, ^ Academic Press.)
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Figure 6 Directly observed distribution of diffusivity in the South Atlantic. The section covers the range between the continental rise off Brazil and the western flanks of the Middle Atlantic Ridge. Highest diffusivity values are correlated with the rough topography on the slopes of the ridge. (After Polzin et al., 1997 ^ Science.)
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Figure 7 Example of abyssal recirculation cells in the tropical North Atlantic Ocean. Numbers indicate volume transports of lower North Atlantic Deep Water in 106 m3 s1. (After Friedrichs and Hall, 1993. Courtesy of the Journal of Marine Research.)
Regions of sinking waters are symbolized in Figure 8 by near-surface downward arrows. They are unevenly distributed in both hemispheres. In fact, they are limited to the Antarctic Circumpolar
Current System (ACCS) and to the high latitudes of the North Atlantic. In addition to bottom and deep water sources there are contributions from outflows of marginal seas (not shown in Figure 8). Although
ic
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Figure 8 The complex global circulation. Colors: red and purple, thermocline waters; green and blue, abyssal waters. For details see text. Abbreviations: SAMW, SubAntarctic Mode Water; LOIW, Lower Intermediate Water; SLW, Surface Layer Water; UPIW, Upper Intermediate Water; NIIW, North-west Indian Intermediate Water; BIW, Banda Intermediate Water; RSW, Red Sea Water; CDW, Circumpolar Deep Water; NADW, North Atlantic Deep Water; NPDW, North Pacific Deep Water; IODW, Indian Ocean Deep Water; AABW, Antarctic Bottom Water. (Reproduced with permission from the Woods Hole Oceanographic Institution, Schmitz 1996a, 1996b.)
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ABYSSAL CURRENTS
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Figure 9 Global near-bottom potential temperature distribution and inferred flow paths of abyssal waters at 4000 m depth. (Reproduced with permission from an article by Charnock, The Atmosphere and the Ocean; in Summerhayes CP and Thorpe SA, Oceanography, An Illustrated Guide ^ 1996 Manson Publishing.)
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Figure 10 Examples of deep entrainment from the subpolar North Atlantic. Numbers represent volume transports in 106 m3 s1. Note the observed significant increase of transport along the east Greenland side which is caused by Denmark Strait Overflow Water. The Denmark Strait between Iceland and Greenland is also called Greenland Strait. (Reproduced with permission from Dickson RR and Brown J, Journal of Geophysical Research 99, 12319–12341 (1994) ^ American Geophysical Union.)
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Figure 11 Demonstration of interannual variability (1938–97) of potential temperature in 1C (colors and white isotherms) and potential density d1.5 (kg m3) 1.5 (dashed black lines) referenced to 1500 dbar pressure level. (After I Yashayaev et al., www.mar.dfo-mpo.gc.ca/science/ocean/woce/labsea/labsea_poster.html.)
Pressure (db)
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Figure 12 Influence of the Mediterranean Water on the salinity distribution at 1100 m of the North Atlantic. Isohalines represent positive anomalies relative to 35.01. Circles depict observed high salinity eddies of Mediterranean Water (Meddies). (Reprinted from Progress in Oceanography 4, Richardson PL, Bower AS and Zenk W, A census of meddies tracked by floats, pp. 209–250 ^ (2000) with permission from Elsevier Science and the American Meteorological Society.)
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Figure 13 Salinity distribution (PSU) of Antarctic and North Pacific Intermediate Waters in the tropical western Pacific along 1501E. (After Holfort and Zenk, 1999.)
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ABYSSAL CURRENTS
these flows reach only intermediate depth levels they affect characteristic water masses of the abyss beneath their own spreading level. Several discrete locations for water sinks have been identified around Antarctica. Completely homogenizing wintertime convection down to the bottom of this semi-enclosed basin was observed only in Bransfield Strait south of the South Shetland Islands. However, in all other locations around Antarctica cooled and freshly ventilated water sinks to the bottom where it behaves like a contour current on the rotating earth. These densest waters in thin layers on the bottom mix with surrounding slightly warmer water masses and finally are transformed into Antarctic Bottom Water (AABW) spreading equatorward in all oceans. Known sinking areas are the Weddell Sea, Ross Sea, Wilkes-Ade´lie Coast, and some locations off Enderby Land and the Prydz Bay. Because the generation of AABW is characterized by highest densities (due to the freezing temperatures
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Figure 14 Salinity distribution (PSU) on a section across the Rio Grande Rise, South Atlantic, at nominally 301S between the slope off Brazil and the Middle Atlantic Ridge. The Vema and Hunter Channels allow an active equatorward flux of Antarctic Bottom Water (AABW). More saline North Atlantic Deep Water (NADW) is advected poleward above the bottom water. On top of the abyssal layers lower saline Antarctic Intermediate Water (AAIW) again spreads equatorward. (Reproduced with permission of the American Meteorological Society from Hogg et al., 1999.)
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in the source regions), AABW can be detected by low (potential) temperature signals close to the bottom everywhere in the Southern Hemisphere and in lower latitudes of the Northern Hemisphere. Figure 9 demonstrates the global deep flow at 4000 m inferred from temperature distribution in the world ocean. The main production regions of sinking waters in the Weddell Sea (and near Iceland in the Northern Hemisphere) are labeled by arrows. The northern counterpart of AABW is called North Atlantic Deep Water (NADW). Since its potential density is slightly different than that of AABW both water masses create an abyssal stratification where they encounter. Formation areas of NADW are located in the Nordic seas, i.e. the Norwegian and Greenland Seas. Convectively formed deep water from polar regions spills through outflow channels over sills to the north west and south east of Iceland. The outflow products are called Denmark Strait and Iceland Scotland Overflow Waters. On their way into the subpolar North Atlantic, the swift overflow plumes with speeds exceeding 50 cm s1 are subject to strong topographic control along the slopes of Greenland and the Mid-Atlantic Ridge, respectively. The high propagation speed of overflow waters favors further transformations by entrainment from intermediate waters increasing its transport rate in selected regions by a factor of 2 or more (Figure 10).
Depth (m)
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Figure 15 Eulerian long-term observation of the southwardflowing North Atlantic Deep Water at nominally 181S. Isoclines of current speeds in cm s1. The prime deep western boundary current of Antarctic Bottom Water is pressed against the continental slope off Brazil. Dashed lines denote the conventional upper and lower boundaries of NADW. (Reproduced with permission of the American Meteorological Society from Weatherley et al. 2000.)
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ABYSSAL CURRENTS
Additional mixing occurs with seasonally generated Labrador Sea Water. The latter also is formed by deep wintertime convection; its seasonal generation rate fluctuates substantially on interannual scales (Figure 11). Local formation regions are not restricted to the western and central Labrador Sea itself. The western Irminger Sea seems to contribute to NADW formation as well. Formation and sinking rates due to deep convection of AABW around Antarctica and of NADW in the North Atlantic are about equal in volume. Each hemisphere contributes B10 106 m3 s1. In the Southern Ocean generation sites are less focused compared with the Nordic Seas of the Atlantic. In the Pacific and Indian Oceans northern sources of
25
abyssal waters are unknown (or at least are insignificant in the Pacific). Overlying waters from two additional sources influence the abyssal circulation. (1) The excess evaporation over precipitation in the Mediterranean and Red Seas increases salinities and densities of these marginal seas substantially in comparison with the adjacent oceans. These density differences force exchange currents in the connecting straits (Strait of Gibraltar and Bab-el-Mandeb). The outflows sink to their equilibrium levels where they form pronounced intermediate high salinity layers in the North Atlantic (Figure 12) and northern Indic, respectively. (2) In selected regions of the southern polar frontal zone notable amounts of low salinity Antarctic Intermediate Water (AAIW) downwell and become
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7 Figure 16 Spreading paths of Antarctic Bottom Water in the Atlantic after miscellaneous observations. Numbers indicate volume fluxes in 106 m3 s1. Originally the AABW spreading path is restricted to the western side of the South Atlantic. It enters the subtropical region through the Vema and Hunter Channels at 301S. The eastern basin is only accessible via the Vema Fracture Zone (111S) and through the Romanche and Chain Fracture Zones near the equator. (Reproduced with permission of the American Meteorological Society from Stephens and Marshall, 2000.)
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part of the intermediate circulation. AAIW is found northward of subpolar regions in all three oceans. In the Pacific there is also a northern mode of intermediate water which is called North Pacific Intermediate Water (Figure 13).
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Deep Western Boundary Currents (DWBC) In contrast to the slow interior drift in ocean basins DWBC are much easier to access by direct observations. They can be revealed by property characteristics of parameters like temperature, salinity, nutrients or other chemical tracers, or be quantified by moored instrumentation. Following the schematic circulation diagram (Figure 8) we expect three stacked opposite DWBC cores at intermediate and abyssal levels in the South Atlantic: northward propagating AABW with mixing components of Circumpolar Deep Water and of the Weddell Sea Deep Water, southward-flowing NADW at least reaching the southern rim of the subtropics, and AAIW with a northward drift again. Water mass distributions and their dynamical imprints are exemplarily displayed in form of a hydrographic section across the Rio Grande Rise at B301S. This ridge separates the abyssal Argentine from the Brazil Basins (Figure 14). Long-term current observations with moored instruments farther north at 181S are reproduced in Figure 15. The core of the southward flowing NADW with averaged speeds of 410 cm s1 is centered at 2200 m about 200 km offshore. The integrated transport of the NADW plume amounts to 39 106 m3 s1. The deepest northward AABW component along the continental rise with long-slope speeds of o3 cm s1 is less developed. The most recent transport estimates within the AABW from the Atlantic are shown in Figure 16. The northern part of this graph is shown in more detail in a schematic diagram jointly with lower NADW invading the Sargasso Sea and adjacent regions from the northern end of the Americas (Figure 17). Mixing of both abyssal water masses is symbolized by stars in Figure 17. The NADW drift grows from B13 10 m6 s1 south of the southern tip of Greenland (Cape Farewell) to B40 106 m6 s1 off the Caribbean island arc. In contrast to the Atlantic the basin-wide spreading of abyssal waters in the Pacific is exclusively controlled by the Antarctic Circumpolar Current System (ACCS). The lowest stratum is filled with Circumpolar Deep Water (CDW), a mixture of converted AABW and transformed NADW, possibly homogenized by repeat circulation in the Southern
45°N 40°N 35°N 30°N 25°N 20°N 15°N 10°N 5°N 0° 80°W
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Figure 17 The complex abyssal circulation in the western North Atlantic. Colors: red and brown, flow of North Atlantic Deep Water; blue, Antarctic Bottom Water. Dashed lines indicate conceivable recirculation branches. Stars denote regions of abyssal entrainment. (Reproduced with permission from McCartney and Curry, 2001.)
Ocean. Potential temperatures of CDW at 321S range from 0.6 to 1.21C. It enters the South Pacific as a DWBC and returns to the ACCS as Pacific Deep Water (PDW) after internal mixing and upwelling decades or centuries later. The region north east of northern New Zealand delineates the gateway for CDW into the southeastPacific Basin (Figure 18). Its boundary current system was found from 2-year long observations to be B700 km wide. The maximum mean velocity of the applied current meter array was 9.6 cm s1 on the eastern flank of the Tonga-Kermadec Ridge. The time-averaged transport amounts to 15.8 106 m6 s1 with a standard error of 9.2 106 m6 s1 for the focused northward advection of the CDW core. A striking aspect of the overall observations at 321S are the total transport (CDW plus PDW) fluctuations ranging from ( 17 to þ 51) 106 m6 s1, typically oscillating over periods of a few months with amplitudes of 1–2 times the mean. About 201 farther north the northward flow of CDW into the main basins of the Pacific is topographically controlled by the Samoan Passage. Recent observations with moored current meters have
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ABYSSAL CURRENTS
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30˚S
2500 m 35˚S 30˚S
4000 m 35˚S 30˚S
_1
5 cm s
Bottom 35˚S 180˚
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Figure 18 Long-term observations of the deep western boundary current system at 321300 S at three instrumented abyssal layers (2500, 4000 m, near-bottom) of the Pacific. Mooring locations at dots are shown by the origin of the presented mean velocity vectors. Ellipses inform about the current’s stability in form of the root mean square amplitudes. Small italics in the lower panel indicate depths in km. Bolder numbers are mooring identifiers. Note the clear bottom-intensified flow of Antarctic Bottom Water. The returning North Pacific Deep Water flow is less confined and highly variable. (Reprinted from Progress in Oceanography 43, Whitworth et al., on the deep western-boundary current in the Southwest Pacific Basin, 1–54 ^ (1999) Elsevier Science and the American Meteorological Society.)
yielded northward transports of 10.671.7 106 m6 s1 from an 18 month measuring program. In accordance with the Stommel-Arons concept further decreases of DWBC transports were estimated at 101N (9.6 106 m6 s1) and 241N ((4.9 and 9.1) 106 m6 s1) from snapshot hydrographic surveys (Figure 19).
To date estimates of the returning PDW transports are rare and quantitatively inconsistent. The abyssal conditions of the Indic resemble those of the Pacific: no deep convective sources in the north are available (Figure 8). Instead, the deep Indian Ocean is controlled by deep and bottom waters from the ACCS (CDW) with a particular influence from
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ABYSSAL CURRENTS 55°N
40°N
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0° 10.6 ± 9.1
20°S 15.8 ± 1.4
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Figure 19 Inflow of Antarctic Bottom Water into the western basins of the subtropical/tropical Pacific. Heavy arrows were inferred from long-term current meter observations, open arrows represent single realizations from hydrography. Numbers indicate volumes of the deep western boundary current transports in 106 m3 s1. (After Hogg, 2001 ^ Academic Press.)
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Figure 20 The distribution of potential density s4 in kg m3 referenced to 4000 dbar is used as a tracer for the spreading of Antarctic Bottom Water in the Indian Ocean. Note the three distinct tongues of dense water invading the Indian Ocean from the Southern Ocean. (Adapted from Mantyla AW and Reid JL, Journal of Geophysical Research 100, pp. 2417–2439 (1995) ^ by the American Geophysical Union and from Schmitz, 1996b.)
the southern South Atlantic (NADW). The Indonesian passages are too shallow to affect the Indic abyss immediately. The topography of the Indic is different to that of the other oceans. Several meridionally aligned ridges divide this ocean into sub-basins. The pathways of deep flows are therefore more complex then assumed in the simple approach by the Stommel-Arons concept. Figure 20 displays the distribution of water density, referenced to 4000 dbar at the bottom of the Indic. The graph enables a view of large-scale bottom water spreading on the base of hydrographic observations. Three major inlets for deep water (CDW or modified NADW) are obvious at the northern tips of the Mozambique (381E) and the Crozet (601E) Basins. The spreading into the Australian Basin (1251E) is constrained by the Australian–Antarctic Discordance at 501S, 1231E. The near-Equator Somali and Australian Basins are only accessible for deep flows through the Amirante and the Diamartina
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ABYSSAL CURRENTS
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Figure 21 Displacement vectors from 800 days of the drift of North Atlantic Deep (A) and Antarctic Bottom Waters (B) in the Brazil Basin. Nominal observation levels were 2500 m (A) and 4000 m (B). Note the clear signal of boundary currents and the predominant zonal structure of the flow in the central basins. Numbers þ 1900 indicate the years in which the Lagrangian current observations by neutrally buoyant floats were started. (From Hogg, 2001 ^ Academic Press.)
Passages, respectively. Long-term transport estimates of CDW spreading are rare. They lie clearly below the 10 106 m6 s1 range.
Conclusions During the past experimental period of the World Ocean Circulation Experiment (1990–99) significant progress in understanding of the abyssal circulation has been achieved. Deep Western Boundary Currents have been quantified on virtually all rises along continents and parallel to meridionally aligned ocean ridges. Pathways and transports of Deep Western Boundary Currents (DWBCs) were identified and quantified mostly with moored current meter arrays. Today the concept of a quasi-steady mean deep circulation has been outdated. In view of the omnipresent eddy kinetic energy with surprisingly large fluctuations more observations on a wide scale of temporal and spatial scales from the whole water column are essential. The need for long-term timeseries in ‘ocean observatories’ has been recognized in international programs like the Climate Variability and Predictability Programme (CLIVAR) or the Global Ocean Observation System (GOOS). The Stommel-Arons concept seems to be confirmed in respect to the balance of abyssal upwelling and downward heat flux through the base of the
main thermocline. However, although the first Lagrangian vector time-series from the abyss of the Brazil Basin are now available, according to the theory, the slow poleward return drift remains inaccessible to direct observations, as was found some 30 years ago in the deep Sargasso Sea. In contrast to expectations, net flow displacements in the Brazil Basin over 800 days indicate preferred zonal advection of bottom water partly in opposing directions (Figure 21). Besides current meter moorings improved instrumental approaches with autonomous arrays of density recorders (‘moored geostrophy’) or acoustic ocean tomography at a basin-scale lie at the frontier of new technologies for future integral observations. Such methods attenuate eddy noise before data are recorded, and hence reveal motion characteristics that may be closer to robust circulation patterns. Progress has also been made with respect to the experimental determination of abyssal diffusivities. Heat transports through deep passages appears to be significantly higher and steady compared with unconstrained areas. This implies enhanced mixing above passages converting them into important stirring agents of the abyss. Finally, recirculation cells delineate a significant limb between the swift DWBC regime and the interior slow drift regions. These deep powerful current loops on a sub-basin-wide scale may explain the
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enormously high speed of the DWBC cores in lower latitudes, particularly in the Atlantic. Their variability strongly affects the meridional overturning cell and questions the role of the conveyor belt (Figure 2) in climate variability.
See also Antarctic Circumpolar Current. Atlantic Ocean Equatorial Currents. Benguela Current; Brazil and Falklands (Malvinas) Currents. California and Alaska Currents. Canary and Portugal Currents. Current Systems in the Atlantic Ocean. Current Systems in the Southern Ocean. Deep Convection. Dispersion and Diffusion in the Deep Ocean. Double-Diffusive Convection. East Australian Current. Florida Current, Gulf Stream and Labrador Current. Heat Transport and Climate. Kuroshio and Oyashio Currents. Non-Rotating Gravity Currents. Open Ocean Convection. Overflows and Cascades. Pacific Ocean Equatorial Currents. Rotating Gravity Currents. Water Types and Water Masses
Further Reading Broecker W (1991) The great ocean conveyor. Oceanography 4: 79--89. Dickson RR and Brown J (1994) The production of North Atlantic Deep Water: Sources, rates and pathwalks. Journal of Geophysical Research 99: 12319--12341. Friedrichs MAM and Hall MM (1993) Deep circulation in the tropical North Atlantic. Journal of Marine Research 51: 697--736. Hogg N (2001) Deep circulation. In: Siedler G, Church J, and Gould JW (eds.) Ocean Circulation and Climate. London: Academic Press. Hogg N, Siedler G, and Zenk W (1999) Circulation and variability at the southern boundary of the Brazil Basin. Journal of Physical Oceanography 29: 145--157. McCartney M and Curry R (2001) Abyssal potential vorticity in the western North Atlantic and the formation
of Lower North Atlantic Deep Water. Progress in Oceanography (in preparation). Mantyla AW and Reid JL (1995) On the origin of deep and bottom waters of the Indian Ocean. Journal of Geophysical Research 100: 2417--2439. Pedlosky J (1979) Geophysical Fluid Dynamics. New York: Springer. Polzin KL, Toole JM, Ledwell JR, and Schmitt RW (1997) Spatial variability of turbulent mixing in the abyssal ocean. Science 276: 93--96. Richardson PL, Bower AS, and Zenk W (2000) A census of Meddies tracked by floats. Progress in Oceanography 4: 209--250. Schmitz WJ Jr (1996a) On the World Ocean Circulation: Some Global Features/North Atlantic Circulation. 1, Woods Hole Oceanographic Institution, Technical Report, WHOI-96-03. Schmitz WJ Jr (1996b) On the World Ocean Circulation: The Pacific and Indian Oceans/A Global Update. 2, Woods Hole Oceanographic Institution, Technical Report, WHOI-96-08. Siedler G, Church J, and Gould JW (eds.) (2001) Ocean Circulation and Climate. London: Academic Press. Stephens JC and Marshall DP (2000) Dynamical pathways of Antarctic Bottom Water in the Atlantic. Journal of Physical Oceanography 30: 622--640. Stommel H (1958) The abyssal circulation. Deep-Sea Research 5: 80--82. Summerhayes CP and Thorpe SA (1996) Oceanography. An Illustrated Guide. London: Manson Publishing. Warren BA and Wunsch C (1981) Evolution in Physical Oceanography. Weatherly GL, Kim YY, and Kontar EA (2000) Eulerian measurements of the North Atlantic Deep Water Western Boundary Current at 181S. Journal of Physical Oceanography 30: 971--986. Wefer G, Berger WH, Siedler G, and Webb DJ (eds.) (1996) The South Atlantic: Present and Past Circulation. Berlin: Springer-Verlag. Whitworth T III, Warren BA, Nowlin WD Jr, et al. (1999) On the deep western-boundary current in the Southwest Pacific Basin. Progress in Oceanography 43: 1--54.
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ACCRETIONARY PRISMS J. C. Moore, University of California at Santa Cruz, Santa Cruz, CA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 1, pp 28–34, & 2001, Elsevier Ltd.
Introduction Subduction of oceanic lithosphere along a convergent plate boundary transfers sediments and rocks from the underthrust lithosphere to the overriding plate, producing an accretionary prism. Accretionary prisms develop beneath the inner slopes of the deep ocean trenches that typically mark convergent plate boundaries. The subduction process destabilizes the mantle after about 100 km of underthrusting beneath the upper plate to produce magmas of the volcanic arcs that virtually always occur along convergent plate boundaries (Figure 1). As accretionary prisms grow through addition of oceanic material, they become coastal mountain ranges. When a continent collides with a subduction zone, the intervening accretionary prism becomes incorporated into the resultant great mountain belts. Thus, rocks in accretionary prisms sometimes are the only record of ancient vanished ocean basins. Accretionary prisms typically form on the upper plate of subduction zone thrust faults, which host the world’s largest earthquakes. Because accretionary prisms incorporate soft sediments at high rates of deformation, they produce
Figure 1 The setting of an accretionary prism in a generalized cross-section of a convergent plate boundary. Although the accretionary prism builds up primarily by scraping off material riding on the oceanic crust, some portions are deeply underthrust and flow back to the surface, while other portions are deeply subducted and participate in the formation of the igneous rocks of the volcanic arc.
some of the world’s most complexly deformed rocks, commonly called melanges. Sediments offscraped to form accretionary prisms are like sponges that yield fluids as they are squeezed and deformed during prism growth. The fluids affect the mechanics of faults; chemically dissolve, transfer, and deposit material; and support chemosynthetic biological communities. The shape of the accretionary prism is mechanically controlled by the strength of the material comprising the accretionary prism and its internal fluid pressure. At slightly fewer than half of modern convergent plate boundaries, accretionary prisms are not currently forming and the incoming sediment and rock is deeply underthrust. Sometimes the accretionary process is reversed and the upper plates of subduction zones are mechanically abraded or eroded by the underthrust plate, causing subsidence and contraction of the overriding plate.
Origin and Variation of Materials Incorporated in Accretionary Prisms Accretionary prisms vary in composition depending on the type of material on the subducting oceanic plate. The ideal sequence incoming to a subduction zone consists of oceanic basaltic igneous rocks covered by oceanic or pelagic sediments that change upsection to more rapidly deposited, continentally derived sandstones, shales, and even conglomerates. At a subduction zone starved for sediments, material available for accretionary prism construction may be the igneous rocks of the oceanic plate with thin overlying sedimentary deposits. Alternatively, incoming plates may be sediment-dominated and covered with a kilometers-thick sequence of deposits that are available for accretion. The resulting accretionary prisms may consist of slices of oceanic igneous rocks with minor amounts of interspersed sediments to thick thrust sheets of continentally derived clastic rocks. The Marianas subduction zone of the western Pacific is an example of the former, whereas the Cascadia subduction zone off the northwestern United States and Canada is an example of the latter. Because the sediment-dominated accretionary prisms (Figure 2) are more voluminous, they tend to be well recognized in the stratigraphic record, for example, parts of the Franciscan Complex of California. Sediment-starved accretionary prisms are thinner and typically dominated by basaltic and ultramafic igneous rocks. They are harder
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Figure 2 Cross-section of an accretionary prism showing the principal structural elements. Paths of fluid flow from deep sources to surface utilize high-permeability conduits, whether along sedimentary layers or faults.
to recognize in the ancient stratigraphic record. Factors controlling the amount of sediment and type of sediment available for accretion include the age of the oceanic plate, the types and rates of sediments being deposited along the transport path of the oceanic plate to the subduction zone, and rate of travel or residence time of any plate in a particular sedimentary environment.
Solid Material Transfer in Accretionary Prisms Sediments and rocks incoming to a subduction zone may be (1) offscraped as a series of thrust sheets at the frontal edge of the accretionary prism; (2) underplated or emplaced at depth along the base of the upper lithospheric plate; or (3) underthrust to great depths to participate in production of volcanic arc magmas or ultimately to be carried into the earth’s mantle (Figures 1 and 2). In the zone of offscraping, incoming materials are typically accreted as a series of imbricate thrust sheets extending from the surface to a basal detachment fault or decollement, beneath which all other material is underthrust. The decollement forms in a layer weaker than the adjacent sediment. With continued underthrusting, the weak layer may become stronger because of mineralogical changes or decreasing fluid pressure and step or migrate down through the underthrust plate. This down-stepping process underplates fault-bounded rock packages to the overlying plate (Figure 2).
The surfaces of some accretionary prisms and nonaccretionary convergent margins (see below) are marked by volcanoes of fluidized mud or serpentine (Figure 2). The fluidized mud and serpentine rise through the upper plate of the convergent margin because these materials are of lower density than the surrounding sediments and rocks. Deep underthrusting of mud and the associated production of natural gas or oil produces mud volcanoes. Serpentine, a low-density rock, is formed by addition of water from underthrust sediments or rocks to mantle rocks. In both cases the low-density rock occurs beneath higher-density rock and buoyancy forces drive the low-density material to the surface. Incoming sediment that is not offscraped at the front of or underplated beneath the main part of the accretionary prism continues to be underthrust. This underthrust sediment may be underplated beneath volcanic arc basement rocks at any point until it reaches the melting zone (Figure 1). Short-lived radioactive isotopes and other chemical tracers in volcanic rocks indicate that sediments less than several million years old are underthrust to the depths of melting (75–100 km) beneath the volcanic arc. Residual material from the melting zone may be deeply underthrust into the earth’s mantle. Continuing sediment accumulation occurs on top of the growing accretionary prism, forming an apron of slope deposits. Locally erosion may remove material from the accretionary prism, redepositing it on the oceanic plate for recycling back into the accretionary prism. In addition to the thrust faults and the
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decollement associated with accretionary processes, accretionary prisms are cut by thrust, normal, and strike-slip faults that form as the prism adjusts its shape in response to its continuing growth (see Mechanics below). Accretionary prisms include some of the most complexly deformed and puzzling rocks on the earth. Melanges or ‘mixed rocks’ and stratally disrupted rocks are included in this category. These rocks are marked by not only stratal discontinuity (Figure 3A) but by mixing of incompatible sedimentary and metamorphic environments (Figure 3B). These intricately deformed rocks form from hard igneous rocks and sediments that are partially consolidated and lithified in a high-strain and high-strain-rate environment. The extreme variation in strength between the hard rocks and soft sediments and the high
(A)
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strain deformation results in a heterogeneously deformed rock mass. Various return flow processes at depth (Figure 1), faulting, and erosion and redeposition of previously accreted rocks contribute to the mixing of rocks derived from differing sedimentary and metamorphic environments (Figure 3B). Accretionary prisms include metamorphic rocks formed under high-pressure, low-temperature conditions. These rocks, called blueschists (Figure 4), are diagnostic of this subduction zone metamorphic environment. The high-pressure–low-temperature condition is caused by the rapid underthrusting of old, cold oceanic plates. Burial rates for underthrust rocks at subduction zones exceed 20 km My 1. The material is buried and returned to shallow depths more rapidly than it can be warmed by conduction from adjacent warmer parts of the earth. Thus, the low-temperature–high-pressure conditions of the subduction system are imprinted on the rocks and preserved by rapid uplift. The lowest average geothermal gradients through accretionary prisms are less than 101C km 1, or about a third of the typical gradient through continents. Thermal gradients beneath accretionary prisms may be much higher where young, hot oceanic crust is being subducted. In these cases blueschists are not formed but are replaced by metamorphic rocks characteristic of higher temperature regimes (greenschist and amphibolite). Overall, the material transfer in accretionary prisms is similar to that in thrust belts in overall form, fault geometry, and mechanics. Thrust belts on land tend to deform more consolidated and lithified sedimentary rocks at slower rates than occur at oceanic convergence zones. Therefore, in thrust belts, both the structural complexity and the effects
(B) Figure 3 Melanges. (A) Dismembered layers of light-colored sandstone in shale matrix. This type of deformation is typically developed along thrust faults with substantial displacement or along the decollement. (B) Melange consisting of block of basalt with overlying pelagic or open oceanic limestone included in shale matrix. Limestone and shale are incompatible depositional environments that are mixed together in the deformational environment of the accretionary prism. Both photographs are from the 60–65 My old accretionary prism of the Kodiak Islands, Alaska.
Figure 4 Blueschist metamorphic rock from an accretionary prism, indicating high pressures and low temperatures and showing ductile or plastic deformation (fold). This rock was metamorphosed about 200 My ago and occurs in the accretionary prism of the Kodiak Islands of Alaska.
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due to fluid expulsion are less pronounced than in accretionary prisms. A good example of a thrust belt is the Canadian Rocky Mountains just west of Calgary Alberta.
Fluids In and Out of Accretionary Prisms Accretionary prisms are like sponges, saturated as they begin to form but squeezed virtually dry as their rocks reach maximum depths of burial. About 40% of the sediment section entering the world’s subduction zones is composed of water in pore spaces. Additional water resides in pores in the igneous rocks of the crust and is bound to minerals both in the sediments and the oceanic crust. The rapid burial of incoming sediments and rocks due to incorporation into and underthrusting beneath the accretionary prism increases stress on the rock framework and raises fluid pressure (the sponge is squeezed). Burial also increases temperature, which releases water bound in minerals and converts sedimentary organic matter to oil and natural gas. Sediment microrganisms may also produce natural gas from organic matter. By 1501C, the minerals are substantially dehydrated and most of the organic matter has been converted to hydrocarbons. By about 4 km of burial, the pore volume of sedimentary rocks in the accretionary prism is reduced by about 90%. The high fluid pressure drives fluids out of the accretionary prism through sediment pores and fractures and along faults (Figure 2). Fluid migration out of the accretionary prism affects everything from surface biology, to large-scale structural features, to the fabric of prism rocks. The high fluid pressures that result from the fluid generation process weaken the rocks by reducing the contact stresses on mineral grains and fault surfaces, thereby reducing friction. This pressure-related weakening facilitates long-distance lateral transport on thrust faults (Figure 2), such as the decollement, and also controls the shape of the accretionary prism (see Mechanics). At temperatures from 1001C and up, rocks dissolve and re-precipitate (undergo pressure solution) with formation of preferred orientations of minerals and solution seams (fabrics) that record stress orientations. Rocks such as slates commonly form through this process. The constituents of the dissolved carbonate and silicate minerals may be locally precipitated or transported with the fluids and precipitated elsewhere as veins and cements that are common in prism rocks. The fluidmediated processes of dissolution, transport, and precipitation significantly change the physical
properties of the accretionary prism and ultimately affect how it deforms. Fluids expulsed from the surface of accretionary prisms contain dissolved methane and hydrogen sulfide, which are utilized by chemosynthetic organisms that are the basis for cold seep biological communities on the seafloor. In the sub-surface, microorganisms both produce and consume various fluid-borne chemical constituents, modifying the physical properties and composition of the host sediments or rocks. Because fluids alter the physical properties of sediment and rock, they affect the seismic reflection images of the prism interior. Fluid-enriched zones along faults reduce velocity and density and produce strong reflections in the seismic images. Methane near the surface of the accretionary prism freezes to form methane hydrate. Progressively deeper in the prism, the methane hydrate is unstable and is transformed back to free methane. Minor accumulations of free methane gas below the hydrate produce a large change in rock physical properties that is seen as a prominent bottomsimulating reflection in seismic images.
Seismogenesis and Accretionary Prisms Subduction zone thrust faults produce the largest earthquakes on the Earth because the plate-boundary thrust is in a zone of high strain rate and is inclined shallowly. The shallow inclination provides a large surface area, or seismogenic zone, subject to brittle failure and the production of earthquakes (Figure 1). In contrast, more steeply inclined faults, such as the San Andreas Fault, transition down-dip from the region of brittle or seismogenic deformation into the realm of ductile deformation over a shorter distance, therefore limiting the area capable of producing an earthquake. The low thermal gradients characteristic of many subduction zones also extends the brittle–ductile transition to greater depths than in other plate boundary settings and increases the area subject to catastrophic seismogenic failure. Commonly, the seismogenic zone earthquakes occur beneath accretionary prisms, such as the great Alaskan earthquake of 1964 (magnitude 9.2). The ability of accretionary prisms to build up enough elastic strain to be released in a sudden earthquake event testifies to strength developed during the evolution from soft sediments to hard rocks. In addition to becoming strong and rigid, the materials of the accretionary prism must evolve to fail in a ‘velocity weakening’ manner, such that there is an acceleration of slip along the fault surface. This accelerating slip produces a discrete seismic event, as opposed to a
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decelerating creep event that would not produce an earthquake.
Mechanics In 1983 Davis, Dahlen, and Suppe articulated and formalized the mechanics of accretionary prisms in the widely accepted ‘critical Coulomb wedge theory’ (Figure 5). Virtually all accretionary prisms approximate a wedge in shape, being thinner on the oceanic side and thicker toward the associated volcanic arc. According to Davis et al. accretionary wedges resemble piles of dirt (prism materials) being pushed forward by bulldozer (the volcanic arc basement). The stresses driving the prism seaward are that of the arc basement pushing the wedge from the rear and a seaward-directed lateral stress due to unequal gravitational stress resulting from the wedge shape. The latter is similar to the stress causing a pile of sand to fail if it is oversteepened. These driving stresses are resisted by a shear stress along the base of the wedge controlled by the frictional strength of the material, which varies with overburden stress, the fluid pressure along the base, and the material properties. Additionally, the component of the overburden stress acting parallel to the decollement must be overcome to, in essence, lift the prism up the decollement. The wedge is just as thick as it can be in order to move forward; that is, it is at its threshold of failure determined by its frictional strength and internal fluid pressure. If the wedge grows at its leading edge, the area of the decollement resisting motion increases and the prism must thicken at the rear in order to increase the area to which the driving stresses in the arc basement can be applied. Wedge theory has been largely successful in explaining the shape of accretionary prisms and the observations of fluid pressure, though the latter are not numerous. Because fluid pressure can counteract normal stresses along the decollement, it is a prime
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parameter controlling the mechanics of accretionary prisms. The generally high fluid pressure along the decollement sharply decreases the frictional resistance there, allowing narrowly tapered prisms to be mechanically stable. The necessity to thicken the landward portion of the prism to keep a stable wedge taper can explain much of the faulting observed in the landward parts of prisms (Figure 2). Other mechanical conceptualizations of wedges utilize differing material properties; however, the Coulomb wedge theory, dependent on basic frictional behavior, is most successful at depths up to 10–20 km.
Non-accretionary Convergent Plate Boundaries According to a compilation by Von Huene and Scholl, somewhat more than half of the world’s convergent plate boundaries are forming accretionary prisms now. The remainder of convergent plate boundaries have inner trench slopes underlain by older accretionary prisms or igneous and metamorphic rocks of the continental crustal or volcanic arc origin (Figure 6). Some are underlain by igneous and metamorphic rocks of uncertain origin that may represent accreted pieces of seamounts or normal oceanic crust. Recent Ocean Drilling Program results off Costa Rica unequivocally demonstrate that virtually all of the sediment riding on the oceanic plate (Cocos Plate) is underthrust beneath the upper plate (Caribbean Plate) of the subduction zone. This process must also occur at many other convergent margins without accretionary prisms. At a number of other convergent margins, the presence of continental or volcanic arc rocks close to the trench suggests that portions of the forearc may have been tectonically eroded by underthrusting. Moreover, Ocean Drilling Program holes show that rocks currently located in deep water on inner slopes of trenches have subsided from much shallower depths.
Figure 5 Diagram showing stresses that control motion of accretionary prisms (and thrust belts). These stresses, integrated over the areas where they act, form a force balance. Forces driving the prism seaward (a push from the rear and internal gravitational forces) are offset by resisting forces (primarily frictional forces and gravitational forces acting along the base of the decollement).
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Figure 6 A convergent margin showing a range of features seen worldwide supporting nonaccretion or tectonic erosion. The absence of offscraping indicates no accretion at the front of the margin. At deeper levels, sediment may be being underplated. Locally, sediments and contained fluids are underthrust to great depths and participate in the melting process beneath the volcanic arc (see Figure 1). A previous accretionary prism much older than the age of incoming sediment may be being underthrust. Arc basement may be anomalously close to the trench, suggesting erosion of the upper plate. Slope deposits may be of much shallower water origin. These shallow-water deposits suggest substantial subsidence of the margin, which is commonly explained by tectonic erosion of the upper plate.
Presumably this subsidence is due to tectonic erosion of the trench inner slope by the underthrusting plate. So, although there are accretionary prisms forming at many convergent margins, the offscraping and underplating at the front of the margin is not a constant process. Accretionary prism formation may be episodic even with continuous subduction. Tectonic erosion may occur. Gaps in the record of accretion are to be expected. What determines whether convergent plate boundaries form accretionary prisms or tectonically erode the upper plate? The primary control is sediment supply. Where thick sequences of sediment enter the subduction zones, accretionary prisms form. Where the sedimentary sequences are thin, there is less material available to accrete and irregularities in the underthrusting oceanic plate may interact with the overthrusting plate to cause erosion. Protruding fault blocks or seamounts may tectonically abrade the lower surface of the upper plate. Also, as these high areas on the lower plate are underthrust, they may oversteepen and otherwise disturb the upper plate, causing it to fail at the surface by landsliding. The landslides accumulate in the trench, where they are underthrust. Accordingly, the front of the convergent margin is decimated and cycled to deeper levels in the earth.
Conclusions Accretionary prisms form at the leading edge of convergent plate boundaries by skimming-off sediments and rocks of the lower plate. In detail, the accretion process involves offscraping of rocks and
sediments at the front of the prism or underplating (emplacement beneath the prism). This deformational process stacks the sediments into thick vertical piles and shortens them horizontally. Consequently, fluids are expulsed and the sediments are progressively transformed to rocks. Because the deformation in accretionary prisms is large and fast, rocks emplaced therein are often severely disrupted and mixed, forming melanges. The rapid rate of underthrusting of the lower plate may carry rocks to great depths before they can heat up, forming a characteristic type of metamorphic rock called a blueschist. The fault surface bounding the base of the accretionary prism, the decollement, is the plate boundary thrust. Because it is shallowly inclined, this fault has a large area undergoing brittle deformation and produces the largest earthquakes on the planet. Mechanically, accretionary prisms resemble a pile of dirt being pushed by a bulldozer. Prisms are pushed from the rear by the volcanic arc basement, with forward motion being resisted by frictional and gravitational forces. Accretion does not occur at the leading edge of all convergent plate boundaries. In these cases, sediments and rocks are underthrust beneath the crustal framework of the upper plate. In some cases accreted rocks have been removed from the upper plate, apparently by tectonic erosion by the lower plate. Thus, an accretionary prism may be a very discontinuous recorder of the incoming sediments and rocks at a convergent plate boundary. Sediments, rocks, and fluids not emplaced in the accretionary prism are carried to great depths and either catalyze subduction zone volcanism or are transported into the mantle.
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ACCRETIONARY PRISMS
Acknowledgments I thank the National Science Foundation for grants supporting my studies of accretionary prisms since 1974 (most recently grant # OCE 9802264). The Ocean Drilling Program provided the opportunity and support to participate in many cruises investigating accretionary prisms. Eli Silver’s insight into convergent margin tectonics substantially improved this review. Hilde Schwartz thoughtfully reviewed the final manuscript.
See also Mid-Ocean Ridge Geochemistry and Petrology. Mid-Ocean Ridge Seismic Structure. Mid-Ocean Ridge Tectonics, Volcanism and Geomorphology. Seismic Structure.
Further Reading Bebout GE, Scholl DW, Kirby SH, and Platt JP (1996) Subduction Top to Bottom, American Geophysical Union Monograph 96. Washington, DC: American Geophysical Union. Davis DJ, Suppe J, and Dahlen FA (1983) Mechanics of fold-and-thrust belts and accretionary wedges. Journal of Geophysical Research 88: 1153--1172. Fisher DM (1996) Fabrics and veins in the forearc: a record of cyclic fluid flow at depths of o15 km. In: Bebout GE, Scholl DW, Kirby SH, and Platt JP (eds.) Subduction Top to Bottom, American Geophysical Union Monograph 96, pp. 75--89. Washington, DC: American Geophysical Union. Fryer P, Mottl M, Johnson LE, et al. (1995) Serpentine bodies in the forearcs of Western Pacific convergent margins. In: Taylor B and Natland J (eds.) Active Margins and Marginal Basins of the Western Pacific,
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American Geophysical Union Monograph 88, pp. 259--279. Washington, DC: American Geophysical Union. Hyndman RD (1997) Seismogenic zone of subduction thrust faults. The Island Arc 6: 244--260. 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 for oceanic chemical fluxes. Philosphical Transactions of the Royal Society of London, Series A 335: 243--259. Meschede M, Zweigel P, and Kiefer E (1999) Subsidence and extension at a convergent plate margin: evidence for subduction erosion off Costa Rica. Terra Nova 11: 112--117. Moore JC and Vrolijk P (1992) Fluids in accretionary prisms. Reviews in Geophysics 30: 113--135. Moores EM and Twiss RJ (1995) Tectonics. New York: WH Freeman. Morris JD, Leeman WP, and Tera F (1990) The subducted component in island arc lavas; constraints from B-Be isotopes and Be systematics. Nature 344: 31--36. Silver EA (2000) Leg 170 Synthesis of fluid-structural relationships of the Pacific Margin of Costa Rica. In: Silver EA, Kimura G, Blum P and Shipley TH (eds) Proceedings of the Ocean Drilling Program, Scientific Results 170 [Online at http://www.odp.tamu.edu/ publications/170_SR/VOLUME/CHAPTERS/ SR170_04.PDF] Taira A, Byrne T, and Ashi J (1992) Photographic Atlas of an Accretionary: Geologic Structures of the Shimanto Belt. Japan, Tokyo: University of Tokyo Press. Tarney J, Pickering KT, Knipe RJ, and Dewey JF (1991) The Behaviour and Influence of Fluid in Subduction Zones. Philosophical Transactions of the The Royal Society of London, Series AE 335: 225--418. von Huene R and Scholl DW (1991) Observations at convergent margins concerning sediment subduction, subduction erosion, and the growth of continental crust. Reviews in Geophysics 29: 279--316.
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ACOUSTIC MEASUREMENT OF NEAR-BED SEDIMENT TRANSPORT PROCESSES P. D. Thorne and P. S. Bell, Proudman Oceanographic Laboratory, Liverpool, UK & 2009 Elsevier Ltd. All rights reserved.
Introduction: Sediments and Why Sound Is Used Marine sediment systems are complex, frequently comprising mixtures of different particles with noncohesive (sands) and cohesive (clays and muds) properties. The movement of sediments in coastal waters impacts on many marine processes. Through the actions of accretion, erosion, and transport, sediments define most of our coastline. Their deposition and resuspension by waves and tidal currents in estuarine and nearshore environments control seabed morphology. Fine sediments, which act as reservoirs for nutrients and contaminants and as regulators of light transmission through the water column, have significant impact on water chemistry and on primary production. Therefore, an improved understanding of sediment dynamics in coastal waters has relevance to a broad spectrum of marine science ranging from physical and chemical processes, to the complex biological and ecological structures supported by sedimentary environments. However, it is commonly acknowledged that our capability to describe the coupled system of the bed, the hydrodynamics, and the sediments themselves is still relatively primitive and often based on empiricism. Recent advances in observational technologies now allow sediment processes to be investigated with greater detail and precision than has previously been the case. The combined use of acoustics, laser and radar, both at large and small scales, is facilitating exciting measurement opportunities. The enhancement of computing capabilities also allows us to make use of more complex coupled sediment–hydrodynamic models, which, linked with the emerging observations, provide new openings for model development. It is readily acknowledged by sedimentologists that the presently available commercial instrumentation does not satisfy a number of requirements for nearbed sediment transport processes studies. Here we focus on the development of acoustics to fulfill some of these needs. The question could be asked: Why use acoustics for such studies? Sediment transport can be
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thought of as dynamic interactions between: (1) the seabed morphology, (2) the sediment field, and (3) the hydrodynamics. These three components interrelate with each other in complex ways, being mutually interactive and interdependent as illustrated schematically as a triad in Figure 1. The objective therefore is to measure this interacting triad with sufficient resolution to study the dominate mechanisms. Acoustics uniquely offers the prospect of being able to nonintrusively provide profiles of the flow, the suspension field, and the bed topography. This exclusive combination of being able to measure all three components of the sediment triad, co-located and simultaneously, has been and is the driving force for applying acoustics to sediment transport processes. The idea of using sound to study fundamental sediment processes in the marine environment is attractive, and, in concept, straightforward. A pulse of high-frequency sound, typically in the range 0.5–5 MHz in frequency, and centimetric in length, is transmitted downward from a directional sound source usually mounted a meter or two above the bed. As the pulse propagates down toward the bed, sediment in suspension backscatters a proportion of the sound and the bed generally returns a strong echo. The amplitude of the signal backscattered from the suspended sediments can be used to obtain vertical profiles of the suspended concentration and particle size. Utilizing the rate of change of phase of the backscattered signal provides profiles of the three orthogonal components of flow. The strong echo from the bed can be used to measure the bed forms. Sediments Bed load−suspended load
Sediment triad
Flow Waves−current
Bed forms Flat−rippled
Figure 1 Illustration of the sediment processes triad and their interactions.
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ACOUSTIC MEASUREMENT OF NEAR-BED SEDIMENT TRANSPORT PROCESSES
Acoustics therefore has the potentiality to provide profile measurements of near-bed sediment processes, with sufficient spatial and temporal resolution to allow turbulence and intrawave processes to be probed; this coupled with the bedform morphology observations provides sedimentologists and coastal engineers with an extremely powerful tool to advance understanding of sediment entrainment and transport. All of this is delivered with almost no influence on the processes being measured, because sound is the instrument of measurement.
Some Historical Background For over two decades the vision of a number of people involved in studying small-scale sediment processes in the coastal zone has been to attempt to utilize the potential of acoustics to simultaneously and nonintrusively measure seabed morphology, suspended sediment particle size and concentration profiles, and profiles of the three components of flow, with the required resolution to observe the perceived dominant sediment processes. The capabilities to measure the three components have developed at different rates, and it is only in the past few years that the potentiality of an integrated acoustic approach for measuring the triad has become realizable. A schematic of the vision is shown in Figure 2. The figure shows a visualization of the application of acoustics to sediment transport processes. ‘A’ is a
A
B
C
39
multifrequency acoustic backscatter system (ABS), consisting, in this case, of three downward-looking narrow beam transceivers. The differential scattering characteristic of the suspended particles with frequency is used to obtain profiles of suspended sediment particle size and concentration profiles. ‘B’ is a three-axis coherent Doppler velocity profiler (CDVP) for measuring co-located profiles of the three orthogonal components of flow velocity; two horizontal and one vertical. It consists of an active narrow beam transceiver pointing vertically downward and two passive receivers having a wide beam width in the vertical and a narrow beam width in the horizontal. This system uses the rate of change of phase from the backscattered signal to obtain the three velocity components. ‘C’ is a pencil beam transceiver which rotates about a horizontal axis and functions as an acoustic ripple profiler (ARP). This is used to extract the bed echo and provide profiles of the bed morphology along a transect. These measurements are used to obtain, for example, ripple height and wavelength, and assess bed roughness. ‘D’ is a highresolution acoustic bed (ripple) sector scanner (ARS) for imaging the local bed features. Although the ARS does not provide quantitative measurements of bed form height, it does provide the spatial distribution and this can be very useful when used in conjunction with the ARP. ‘E’ is a rapid backscatter ripple profiler system (BSARP) for measuring the instantaneous relationship between bed forms and the suspended sediments above it.
D
E
B
Waves
B Current Suspension
Bed forms
Figure 2 A vision of the application of acoustics to sediment transport processes. A, multifrequency acoustic backscatter for measuring suspended sediment particle size and concentration profiles. B, coherent Doppler velocity profiler for measuring the three orthogonal components of flow velocity. C, bed ripple profiler for measuring the bed morphology along a transect. D, high-resolution sector scanner for imaging the local bed features. E, backscatter scanning system for measuring the relationship between bed form morphology and suspended sediments.
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ACOUSTIC MEASUREMENT OF NEAR-BED SEDIMENT TRANSPORT PROCESSES
Acoustic ripple scanners, ARSs, are based on sector-scanning technology, which has been specifically adapted for high-resolution images of bed form morphology. They typically have a frequency of around 1–2 MHz with beam widths of about 11 in the horizontal and 301 in the vertical. As the pulse is backscattered from the bed, the envelope of the signal is measured and usually displayed as image intensity. An example of the data collected by an ARS is shown in Figure 4. As can be seen, this provides an aerial image of the bed, clearly showing the main bed features. The advantage of the ARS is the area coverage that is obtained, as opposed to a single line profile with the ARP; however, direct information on the height of features within the image cannot readily be extracted. Ideally one would like to combine the two instruments and recently such systems have become available. This is essentially an ARP which also rotates horizontally through 1801 and therefore allows a three-dimensional (3-D) measurements of the bed. The sonar gathers a single swath of data in the vertical plane and then rotates the transducer around the vertical axis and repeats the process until a circular area underneath the sonar has been scanned in a sequence of radial spokes. An example of data collected by a 3-D ARP operating at a frequency of 1.1 MHz is shown in Figure 5.
What Can Be Measured? The Bed
Whether the bed is rippled or flat has a profound influence on the mechanism of sediment entrainment into the water column. Steep ripples are associated with vortices lifting sediment well away from the bed, while for flat beds sediments primarily move in a confined thin layer within several grain diameters of the bed. Therefore knowing the form of the bed is a central component in understanding sediment transport processes. The development of the ARP and the ARS has had a significant impact on how we interpret sediment transport observations. These specifically designed systems typically either provide quantitative measurements of the evolution of a bed profile with time, the ARP, or generate an image of the local bed features over an area, the ARS. Figure 3 shows data collected with a 2-MHz narrow pencil beam ARP in a marine setting. The figure shows the variability of a bed form profile, over nominally a 3-m transect, covering a 24-h period. Over this period the bed was subject to both tidal currents and waves and the figure shows the complex evolution of the bed with periods of ripples and less regular bed forms. The figure clearly shows the detailed quantitative measurements of the bed that can be obtained with the ARP.
Height (m)
0.2
0
−0.2 1.5 1
25 0.5 Di
20
sta
0
nc
e(
m)
15 −0.5
10 −1
5 −1.5
)
e (h
Tim
0
Figure 3 ARP measurements of the temporal evolution of a 3-m profile on a rippled bed, covering a 24-h period.
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ACOUSTIC MEASUREMENT OF NEAR-BED SEDIMENT TRANSPORT PROCESSES
2
y (m)
1
0
−1
−2
−2
−1
0
1
2
x (m)
z (m)
Figure 4 Image of ripples on the bed in a large-scale flume facility, the Delta Flume, collected using an acoustic ripple sector scanner, ARS. The 0.5-m-diameter circle at the center right of the image is one of the feet of the instrumented tripod used to collect the measurements, while the dark area at the center of the image is a blind spot directly beneath the sonar.
−0.25 −0.5 −2
−1
0 )
m
x(
2 1
1 0 2
−1
y (m)
−2
Figure 5 Three-dimensional measurements of a rippled bed collected in the marine environment using a 3-D ripple scanner, 3-D ARP. The artifacts in the image are associated with reflections from the instrumented tripod used to collect the measurements.
The data shown in the figure were collected during a recent marine deployment. The bed surface shown is based on a 3-D scan, with a 100 vertical swaths at 1.81 intervals, each of which comprised 200 acoustic
samples at 0.91 intervals and spanning 1801. The plot clearly shows a rippled bed. There are one or two artifacts in the plot; these are associated with reflection from the rig on which the 3-D ARP was
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ACOUSTIC MEASUREMENT OF NEAR-BED SEDIMENT TRANSPORT PROCESSES
mounted. However, the figure plainly contains information on both the horizontal and vertical dimensions of the ripples and can therefore be used to precisely define quantitatively the features of the bed over an area. The 3-D ARP is a substantial advance on both the ARP and the ARS.
The Flow
The success of acoustic Doppler current profilers (ADCPs), which typically provide mean current profiles with decimeter spatial resolution, and more recently the acoustic Doppler velocimeter (ADV), which measures, subcentimetric, subsecond, three velocity components at a single height, has stimulated interest in using acoustics to measure near-bed velocity profiles. The objective is to use the same backscattered signal as used by the ABSs, but process the rate of change of phase of the signal (rather than the amplitude as used by ABSs) to obtain velocity profiles with comparable spatial and temporal resolution to ABSs. The phase technique is utilized in CDVPs and the phase approach has been the preferred method for obtaining high spatial and temporal resolution velocity profiles. An illustration of a three-axis CDVP is shown in Figure 6. A narrowbeam, downwardly pointing transceiver, Tz, transmits a pulse of sound. The scattered signal is picked up by Tz, and two passive receivers Rx and Ry which are orthogonal to each other and have a wide beam in the vertical and narrow in the horizontal. Receiver Rx
Receiver Ry z Transmit transceiver
Tz
To examine the capability of such a system, measurements from a three-axis CDVP have been compared with a commercially available ADV. The system had a spatial resolution of 0.04 m, operated over a range of 1.28 m, and provided 16-Hz velocity measurements of the vertical and two horizontal components of the flow. Figure 7 shows a typical example of CDVP and ADV time series, power spectral density, and probability density function plots for u, the streamwise flow. The velocities presented in Figure 7 show CDVP results that compare very favorably with the ADV measurements; having time series, spectra, and probability distributions in general agreement. There are differences in the spectrum; the CDVP spectra begins to depart from the ADV above about 4 Hz, with the CDVP measuring larger spectral components at the higher frequencies. This trend was common to all the records and is a limitation of the CDVP system used to obtain the data, rather than an intrinsic limitation to the technique. Figure 8 illustrates the capability of the CDVP for flow visualization in the marine environment. The figure shows mean zeroed velocity vectors, u–w, v–w, and u–v, plotted over a 5-s time period, between 0.05 and 0.7 m above the bed. The length of the velocity vectors is indicated in the figure. A single-point measurement instrument such as an ADV can provide the time-varying velocity vectors at a single height above the bed; however, the spatial profiling which is achievable with the threeaxis CDVP provides a capability to visualize structures in the flow. The structures seen in Figure 8 are associated with combined turbulent and wave flows. This type of plot exemplifies the value of developing a three-axis CDVP with co-located measurement volumes, since it clearly illustrates the fine-scale temporal and spatial flow structures which can be measured in the near-bed flow regime. Linking such measurements with ABS profiles of particle size and concentration will provide a very powerful tool for studying near-bed fluxes and sediment transport processes.
y
The Suspended Sediments
Range bin
x
Figure 6 Schematic of the transducer arrangement for a threeaxis CDVP.
Multifrequency acoustic backscattering (ABS) can be used to obtain profiles of mean particle size and concentration. The ABS is the only system available that profiles both parameters rapidly and simultaneously. Also the bed echo references the profile to the local bed position. This is important because all sediment transport formulas use the bed as the reference point and predict profiles of suspension parameters relative to the bed location. Examples of the results that can be obtained are shown
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ACOUSTIC MEASUREMENT OF NEAR-BED SEDIMENT TRANSPORT PROCESSES
u (m s−1)
(a)
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1.5 1 0.5 0 −0.5 0
Time (s)
(b)
100
(c) 3 2.5 2 10−2
PDFu
PSDu (m2 s−2 Hz−1)
10−1
10−3
1.5 1 0.5
10−4 10−1
100
0 −1
101
−0.5
0
0.5
1
Velocity (m s−1)
Frequency (Hz)
Figure 7 Comparison of the streamwise flow, u, measured by an ADV (red) and a CDVP (black). (a) The velocities measured at 16 Hz over 100 s. (b) The power spectra of the zero-mean velocities. (c) The probability density functions of the zero-mean velocities for the ADV (open circles) and the CDVP (crosses).
Height above bed (m)
(a) 0.8 0.6 0.4 0.2 0 155
155.5
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157
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Height above bed (m)
(b)
158
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159
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160
0.8 0.6 0.4 0.2 0 155
Time (s)
(c) Height above bed (m)
157.5 Time (s)
160 −1
1ms
0.8 0.6 0.4 0.2 0 155
155.5
156
156.5
157
157.5 Time (s)
158
158.5
159
159.5
160
Figure 8 Demonstration of the capability of the triple-axis CDVP to provide visualizations of intrawave and turbulent flow. Plots in (a)–(c) show a time series over a 5 s period of the zero-mean velocities displayed as vectors u–w, v–w, and u–v, respectively.
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ACOUSTIC MEASUREMENT OF NEAR-BED SEDIMENT TRANSPORT PROCESSES
(b)
0.8
0.8
0.7
0.7
0.6
0.6
0.5
0.5
z (m)
z (m)
(a)
0.4
0.4
0.3
0.3
0.2
0.2
0.1
0.1
0
0 0
100
200
300
10−2
d (µm)
10−1 C (kg
100
m−3)
Figure 9 Profiles of suspended sediments: (a) particle size and (b) concentration.
in Figures 9 and 10. These observations were collected off Santa Cruz Pier, California, as a seasonal storm passed through the area over the period of a couple of days. Figure 9 shows profiles of the suspended sediment particle size and concentration. The profiles of particle size are seen to be relatively consistent with a mean diameter of c. 180 mm near the bed and reducing to c. 120 mm at 0.7 m above the bed. The variability in size is relatively limited and due to changing bed and hydrodynamic conditions as the storm passed by. In Figure 9(b) the suspended concentrations are comparable in their form, although they have absolute values that vary by greater than an order of magnitude. This variation in suspended concentration is associated with the changing conditions as the storm passed through the observational area. It is interesting to note that over the period even though there is a large variation in concentration, the particle size remains nominally consistent. Figure 10 shows the temporal variation in particle size and
concentration with height above the bed. It can be clearly seen that some of the periods of increased particle size are associated with substantial suspended sediment events, as one might expect, however, there are one or two events where the correlation is not as clear. Figures 9 and 10 clearly illustrate the capability of ABS to simultaneously measure profiles of concentration and particle size and the combination of both significantly adds to the assessment and development of sediment transport formulas.
A Case Study of Waves over a Rippled Bed Here the use of acoustics is illustrated by application to a specific experimental study. Over large areas of the continental shelf outside the surf zone, sandy seabeds are covered with wave-formed ripples. If the ripples are steep, the entrainment of sediments into the water column, due to the waves, is considered
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ACOUSTIC MEASUREMENT OF NEAR-BED SEDIMENT TRANSPORT PROCESSES
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(a)
0.6
180
µm
z (m)
160 0.4 140 0.2 120
0
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15
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100
(b)
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z (m)
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log(kg m−3)
−1
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−3 0 14.8
15
15.2
15.4
15.6
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16
16.2
Date Mar. 2003 Figure 10 Measurements of the temporal variations with height above the bed, of: (a) particle size with the color bar scaled in microns and (b) logarithmic concentration with the color bar scaled relative to 1.0 kg m 3.
(a)
(b)
(c)
(d)
(e) v1
v1
v1
v1
(f) v1
v1 v2
v2
v2
Figure 11 A schematic of vortex sediment entrainment by waves over a steeply rippled bed. The arrows show the direction and relative magnitude of the near-bed wave velocity. v1 and v2 are the lee slope-generated vortices.
to be primarily associated with the generation of vortices. This process is illustrated in Figure 11. As shown in Figures 11(a) and 11(b), a spinning parcel of sediment-laden water, v1, is formed on the leeside of the ripple at the peak positive velocity in the wave
cycle. This sediment-rich vortex is then thrown up into the water column at around flow reversal (Figures 11(c) and 11(d)), carrying sediment well away from the bed and allowing it to be transported by the flow. At the same time, a sediment-rich vortex,
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ACOUSTIC MEASUREMENT OF NEAR-BED SEDIMENT TRANSPORT PROCESSES
v2, is being formed on the opposite side of the ripple due to the reversed flow. As shown in Figures 11(d)– 11(f), v2 grows, entrains sediment, becomes detached, and moves over the crest at the next flow reversal carrying sediments into suspension. The main feature of the vortex mechanism is that sediment is carried up into the water column twice per wave cycle at flow reversal. This mechanism is completely different to the flat bed case where maximum near-bed concentration is at about the time of maximum flow velocity. To study this fundamental process of sediment entrainment, experiments were conducted in one of the world’s largest man-made channels, specifically constructed for such sediment transport studies, the Delta Flume; this is located at the De Voorst Laboratory of the Delft Hydraulics in the north of the Netherlands. The flume is shown in Figure 12(a); it is 230 m in length, 5 m in width, and 7 m in depth, and it allows waves and sediment transport to be studied at full scale. A large wave generator at one end of the flume produced waves that propagated along the flume, over a sandy bed, and dissipated on a beach at the opposite end. The bed was comprised of coarse sand, with a mean grain diameter by weight of
(a)
(b)
330 mm, and this was located approximately halfway along the flume in a layer of thickness 0.5 m and length 30 m. In order to make the acoustic and other auxiliary measurements, an instrumented tripod platform was developed and is shown in Figure 12(b). The tripod, STABLE II (Sediment Transport and Boundary Layer Equipment II), used an ABS to measure profiles of particle size and concentration, a pencil beam ARP to measure the bed forms, and, in this case, electromagnetic current meters (ECMs) to measure the horizontal and vertical flow components. Figure 12(c) shows a wave propagating along the flume with STABLE II submerged in water of depth 4.5 m, typical of coastal zone conditions. To investigate and then model the vortex entrainment process it was necessary to establish at the outset whether or not the surface waves were generating ripples on the bed in the Delta Flume. Using an ARP a 3-m transect of the bed was measured over time. The results of the observations over a 90-min recording period are shown in Figure 13. Clearly ripples were formed on the bed and the ripples were mobile. To obtain flow separation and hence vortex formation requires a ripple steepness (ripple height/
(c)
Figure 12 (a) Photograph of the Delta Flume showing the sand bed at approximately midway along the flume and the wave generator at the far end of the flume. (b) The instrumented tripod platform, STABLE II, used to make the measurements. (c) A surface wave propagating along the flume.
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ACOUSTIC MEASUREMENT OF NEAR-BED SEDIMENT TRANSPORT PROCESSES
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−0.4
Height (m)
−0.6 −0.8 −1 −1.2 0
20
e
m Ti
40
)
in
(m
60 −0.5
80
0.5 1 1.5
−1
−1.5
0
ce (m)
Distan
Figure 13 Acoustic ripple profiler measurements of a transect of the bed, over time, in the Delta Flume.
ripple wavelength) of the order of 0.1 or greater; an analysis of the observations showed that this was indeed the case. Using the ABS, some of the most detailed full scale measurements of sediment transport over a rippled bed under waves were captured. These measurements from the Delta Flume are shown in Figure 14. The images shown were constructed over a 20-min period as a ripple passed beneath the ABS. The suspended concentrations over a ripple, at the same velocity instants during the wave cycle, were combined to generate a sequence of images of the concentration over the ripple with the phase of the wave. Four images from the sequence have been shown to illustrate the measured vortex entrainment. The length and direction of the arrows in the figure give the magnitude and direction of the wave velocity, respectively. Comparison of Figure 14 with Figure 11 shows substantial similarities. In Figure 14(a), there can be observed the development of a high-concentration event at high flow velocity above the lee slope of the ripple, v1. In Figure 14(b), as the flow reduced in strength, the near-bed sediment-laden parcel of fluid travels up the leeside of the ripple toward the crest. As the flow reverses, this sediment-laden fluid parcel, v1, travels over the crest and expands. As the reverse flow increases in strength (Figure 14(d)), the parcel v1 begins to lift
away from the bed and a new sediment-laden lee vortex, v2, is initiated on the lee slope of the ripple. In order to capture the essential features of these data within a relatively simple, and hence practical, 1-DV (one-dimensional in the vertical) model, the data has first been horizontally averaged over one ripple wavelength at each phase instant during the wave cycle. The resulting pattern of sediment suspension contours is shown in the central panel of Figure 15, while the upper panel shows the oscillating velocity field measured at a height of 0.3 m above the bed. The concentration contours shown here are relative to the ripple crest level, the mean (undisturbed) bed level being at height z ¼ 0. The measured concentration contours presented in Figure 15 show two high concentration peaks near the bed that propagate rapidly upward through a layer of thickness corresponding to several ripple heights. The first, and the strongest, of these peaks occurs slightly ahead of flow reversal, while the second, weaker and more dispersed peak, is centered on flow reversal. The difference in the strengths of the two peaks reflects the greater positive velocity that can be seen to occur beneath the wave crest (time ¼ 0 s) than beneath the wave trough (time ¼ 2.5 s). Between the two concentration peaks the sediment settles rapidly to the bed. Maybe rather unexpectedly this settling effect occurs at the times of strong forward and backward velocity at measurement
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48
ACOUSTIC MEASUREMENT OF NEAR-BED SEDIMENT TRANSPORT PROCESSES
Concentration (kg m−3)
1
6 cm
v1 v1 0 42 cm
v1
v2
Figure 14 Acoustic imaging of suspended sand entrainment over a rippled bed due to waves, at four phases of the wave velocity. The length of the white arrow in each plot gives the magnitude and direction of the near-bed wave velocity.
levels well above the bed. The underlying mechanism of sediment entrainment by vortices shed at or near flow reversal is clearly evident in the spatially averaged measurements shown in Figure 15. Any conventional ‘flat rough bed’ model that attempts to represent the above sequence of events in the suspension layer runs into immediate and severe difficulties, since such models predict maximum near-bed concentration at about the time of maximum flow velocity, and not at flow reversal. Here therefore, for the first time in a 1-DV model, it has been attempted to capture these effects realistically through the use of a strongly time-varying eddy viscosity that represents the timing and strength of the upward mixing events due to vortex shedding. The model initially predicts the size of the wave-induced ripples and the size of the grains found in suspension, and then goes on to solve numerically the equations governing the upward diffusion and downward settling of the suspended sediment. The resulting concentration contours in the present case are shown on the lower panel of Figure 15. The essential twopeak structure of the eddy shedding process can be seen to be represented rather well, with the initial
concentration peak being dominant. The decay rate of the concentration peaks as they go upward is also represented quite well, though a phase lag develops with height that is not seen to the same extent in the data. Essentially, the detailed acoustic observations of sediment entrainment under waves over ripples of moderate steepness have begun to establish a new type of 1-DV modeling, thereby allowing the model to go on to be used for practical prediction purposes in the rippled regime, which is the bed form regime of most importance over wide offshore areas in the coastal seas.
Discussion and Conclusions The aim of this article has been to illustrate the application of acoustics in the study of near-bed sediment processes. It was not to detail the theoretical aspects of the work, which can be found elsewhere. To this end, measurements of bed forms, the hydrodynamics, and the movement of sediments have been used. These results show that acoustics is progressively approaching the stage where it
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ACOUSTIC MEASUREMENT OF NEAR-BED SEDIMENT TRANSPORT PROCESSES
Velocity (m s−1)
0.5
0
−0.5 Observed
0.25
z (m)
0.2
0.15
0.1
0.05
Modeled 0.25
z (m)
0.2
0.15
0.1
0.05 0
1
2
3
4
5
Time (s)
−0.5 0.25 log10[Concentration] (re: kg m−3)
1
Figure 15 Measurement and modeling of suspended sediments with height z above a rippled bed under a 5-s period wave.
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49
50
ACOUSTIC MEASUREMENT OF NEAR-BED SEDIMENT TRANSPORT PROCESSES
can measure nonintrusively, co-located and simultaneously, with high temporal and spatial resolution, all three components of the interacting sediment triad. Figure 16 shows the instrumentation used in a further recent deployment in the Delta Flume. This shows the convergence of the instrumentation and also its use in conjunctions with instruments such as laser in situ scattering transmissometry (LISST). Although substantial advances have been made in the past two decades, the application of acoustics to sediment transport processes is still in an ongoing developmental phase and there are limitations and shortcomings that need to be overcome, and further applications explored. Although there have been few reports to date on data collected using 3-D ARP, such systems are now becoming available. The 3-D ARP is a substantial development over the single-line ARP and the ARS, and should make a considerable contribution to the measurement and understanding of the formation and development of bed forms. There have been a number of reports on single-axis CDVP; however, it is the three-axis CDVP which is the way ahead. Again, these instruments are coming online, though they are still very much a research tool. However, the concept has now essentially been
proven and its use with a vengeance in sediment studies will begin to make an important impact in the next couple of years. It was with the concept of using acoustics to measure suspended sediment concentration that its application to the other components of the small-scale sedimentation processes triad followed. To date the use of sound to measure suspended sediment concentration and particle size has been successful when systems have been deployed over nominally homogeneous sandy beds. However, all who use acoustics recognize that the marine sedimentary environment is frequently much more complicated, and suspensions of cohesive sediments and combined cohesive and noncohesive sediments are common. To employ acoustics quantitatively in mixed and cohesive environments requires the development of a description of the scattering properties of suspensions of cohesive sediments and sediment mixtures. This would be interesting and very valuable work, and should significantly extend the deployment regime over which acoustic backscattering can be employed quantitatively. In conclusion, the objective of this article has been to describe the role of acoustics, in near-bed sediment transport studies. It is clearly acknowledged
ARP
LISST
ARS
PT
PS
ADV ABS
Three-axis CDVP
Figure 16 The instrumentation used in a recent Delta Flume experiment. LISST, laser in situ scattering transmissometry; PS, pumped samples; PT, pressure transducer.
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ACOUSTIC MEASUREMENT OF NEAR-BED SEDIMENT TRANSPORT PROCESSES
that acoustics is one of a number of technologies advancing our capabilities to probe sediment processes. However, its nonintrusive profiling ability, coupled with its capability to measure all three components of the sediment dynamics triad, make it a unique and very powerful tool for studying the fundamental mechanisms of sediment transport.
See also Acoustic Scattering by Marine Organisms. Acoustics in Marine Sediments. Breaking Waves and Near-Surface Turbulence. Estuarine Circulation. Offshore Sand and Gravel Mining.
51
Hay AE and Mudge T (2005) Principal bed states during SandyDuck97: Occurrence, spectral anisotropy, and the bed storm cycle. Journal of Geophysical Research 110: C03013 (doi:10.1029/2004JC002451). Thorne PD and Hanes DM (2002) A review of acoustic measurements of small-scale sediment processes. Continental Shelf Research 22: 603--632. Vincent CE and Hanes DM (2002) The accumulation and decay of near-bed suspended sand concentration due to waves and wave groups. Continental Shelf Research 22: 1987--2000. Zedel L and Hay AE (1999) A coherent Doppler profiler for high resolution particle velocimetry in the ocean: Laboratory measurements of turbulence and particle flux. Journal of Atmosphere and Ocean Technology 16: 1102--1117.
Further Reading Crawford AM and Hay AE (1993) Determining suspended sand size and concentration from multifrequency acoustic backscatter. Journal of the Acoustical Society of America 94(6): 3312--3324. Davies AG and Thorne PD (2005) Modelling and measurement of sediment transport by waves in the vortex ripple regime. Journal of Geophysical Research 110: C05017 (doi:1029/2004JC002468).
Relevant Websites http://www.aquatecgroup.com – Aquatec Group Ltd. http://www.marine-electronics.co.uk – Marine Electronics Ltd. http://www.pol.ac.uk – POL Research, Proudman Oceanographic Laboratory.
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ACOUSTIC NOISE I. Dyer, Marblehead, MA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 1, pp 35–44, & 2001, Elsevier Ltd.
Introduction Some ocean scientists consider ambient noise to be a fairly simple and well-behaved property of the ocean. Ambient noise, after all, is often reported and summarized in highly averaged form, its naturally large variance mostly unstated. Other ocean scientists consider the variational complexity of ambient noise a richly colored portrait carrying images of basic ocean processes, including the physics of various noise sources and the acoustics of multiple noise propagation paths. Space and incomplete knowledge precludes a description here that can fully satisfy all ocean scientists or technologists. Instead, the objective is to summarize those aspects of ocean ambient noise that convey the more important recent research results and the more significant remaining research questions. A 1962 summary of ambient noise measurements in the ocean (see Figure 1) is still useful today, at least to classify the various noise sources and their average levels and smooth frequency spectra. Prevailing noises (those observed almost always) are caused by wave–wave interactions at the sea surface, by distributed seismic activity in the earth, by atmospheric or oceanic turbulence, by distant shipping, by windinduced sea surface agitation, and by thermally induced molecular agitation. According to Wenz, wave–wave interaction effects, seismic background, and/or turbulence dominates the noise at VLF (very low frequency band: 1ofo20 Hz), with power spectral density of the pressure field Sðf Þpf 4 . Distant shipping noise dominates at LF (low frequency band: 20ofo200 Hz), has a broad spectral peak around 50 Hz, and falls off sharply for f > 200 Hz as f 6 . At MF (midfrequency band: 200 Hzo fo50 kHz), noise caused by sea surface agitation typically dominates, with a broad peak within 200 Hzofo2 kHz and, beyond f E2 kHz, with Sðf Þpf 1:7 . Finally, molecular agitation typically dominates the noise at HF (high frequency band: f4100 kHz), with Sðf Þpf 2 . Other noise sources are classified as temporally intermittent or spatially discrete, rather than prevailing, and can often dominate. These include
52
sounds from marine earthquakes, from marine animals, from nearby ships or other nearby commercial activities in the ocean, from rain/hail/snow striking the sea surface, and from fractures of ice in the north or south polar oceans. With such a large number of prevailing and other noise sources, the band designations given in the previous paragraph are unlikely to be associated unequivocally with just one noise source or, for that matter, adopted fully by most ambient noise researchers or practitioners. They are of use, however, to help present the material to follow. The spectral summaries used in this Introduction are based on Wenz, and although still useful, modifications and additions are needed in the light of new knowledge. Urick published an excellent summary of ambient noise data acquired in various measurement programs through about 1980. Practitioners commonly use these data, plus the Wenz results, for prediction. Nevertheless, basic understanding of many ambient noise mechanisms through about 1980 was meager and, indeed, some suggested mechanisms were considered speculative. Fortunately, mechanisms for prevailing ambient noises have received considerable research attention since then, particularly from 1985 or so. The other noises have also been researched, in general to a lesser degree. Two volumes edited by Kerman and one by Buckingham and Potter are conference proceedings of recent ambient noise research, and are extraordinary seminal contributions to the understanding of ambient noise mechanisms in the ocean. The continuing flow of research results in archival journals and books, and the aforementioned volumes, provide important modifications and additions to the classical summary of ambient noise by Wenz. In what follows the more important new knowledge, or lack thereof, is summarized.
ULF Band: Wave–Wave Interaction Noise Measurements within 0.1ofo2 Hz, which has come to be called the ultralow frequency band (ULF), extended the Wenzian picture one decade lower in frequency1, and showed ULF noise to be a function of wind speed. The data (Figure 2), have a strong peak f0 located between about 0.2 and 0.7 Hz, with
1 To accommodate the ULF band, the band scheme described in the Introduction is redefined to be VLF2–20 Hz.
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ACOUSTIC NOISE
53
120 Intermittent and local effects Earthquakes and explosions
Biologics
Precipitation Ships, industrial activity Sea ice
26
KEY Limits of prevailing noise Wind-dependent bubble and spray noise Low-frequency very-shallow-water wind dependence Heavy precipitation Heavy traffic noise Usual traffic noise _ shallow water Usual traffic noise _ deep water Thermal noise General pattern of noise from earthquakes and explosions Extrapolations
_
Ln ′ Spectrum level, dB re 0.0002 dyn cm 2 and 1 Hz
80
60
6
_14
Wind force (Beaufort) _ 34
_
40
Ln ′ dB re 1 dyn cm 2 and 1Hz (add 100 dB for ref. of 1 μPa and 1 Hz)
100
8 5
_ 54
20 3 2
Prevailing noises Turbulent-pressure fluctuations
0
1
_ 74
Oceanic traffic Bubbles and spray (Surface agitation)
Surface waves _ Second-order pressure effects (Seismic background)
_ 20 1
10
10
2
Molecular agitation 3
10
4
10
5
10
Frequency (Hz) Figure 1 Ambient noise spectra summarized by Wenz. (The ordinates are Ln ¼ 10 log10 Sðf Þ, with respect to the reference value. Add 100 dB to the right-hand scale to obtain Ln in dB re 1 mPa and 1 Hz.) The Beaufort Force translates to wind speeds, in ms 1, as follows: 1, 0.5–1.5; 2, 2–3; 3, 3.5–5; 5, 8.5–10.5; 8, 17–20. (Reproduced from Wenz, 1962.)
Sðf Þ at higher f proportional to about f 3 2f 5 , dependent upon wind speed. A long history of measurements, as well as theoretical surface wave interaction studies, presaged this result. The appearance of systematic data such as in Figure 2 apparently sparked even more research efforts that ultimately confirmed the basic aspects of ULF noise.
Pressure Spectral Density
Wave–wave interaction noise is caused by opposing wind-driven surface gravity waves, each at frequency fw , that to second-order create a pressure field in the water at f ¼ 2fw . (Orders higher than the second have been shown to be negligible.) In its simplest
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54
ACOUSTIC NOISE
The noise spectral density Sðf Þ is proportional to [Sw ðf =2Þ2 , the surface elevation density squared and shifted in frequency. Since the frequency dependence of the last three terms in eqn [1] is relatively weak around the peak frequency fwo of Sw , the peak of S is essentially fo E2fwo. Both fwo and Sw are functions of wind speed U, as affected by other sea conditions (fetch, sea age, etc.), and similarly lead to the dependence of fo and S on U.
160
Ambient noise spectrum level (dB re μPa2/Hz)
150
140
Microseism Spectral Density 130
120
110 _
Time (z)
100
90 0.05
Wind speed (m s 1)
17 OCTOBER (0340) 17 OCTOBER (1940)
30 25
18 OCTOBER (1940) 19 OCTOBER (1940)
15 7.5
20 OCTOBER (1140) 20 OCTOBER (1540)
5 <5
20 OCTOBER (1940) 24 OCTOBER (0340)
<5 35
0.5 Frequency (Hz)
0.1
1.0
2
Figure 2 Ambient noise pressure derived from vertical displacement (seismometer) data at the water/bottom interface (depth about 100 m). (Reproduced from Kibblewhite and Evans, 1985.)
Displacement, rather than pressure, is often measured on the seafloor. Termed microseism noise, it can be obtained from eqn [1] with the substitution of a modified transfer function Td ðf ; zÞ for Tp ðf ; zÞ, each in general given for all z, to yield the displacement spectral density Sd ðf ; zÞ rather than Sðf ; zÞ. These transfer functions incorporate the acoustics of the medium, including the seismoacoustics of the bottom. It is no surprise, therefore, that noise pressure or displacement data may show subsidiary peaks in the frequency domain, additional to the major peak at fo , associated with acoustic modes of the oceanic waveguide. With modifications as outlined in the foregoing and as detailed in the literature, and with site-specific seismic–acoustic bottom properties and time- and site-specific sea surface elevations, data such as in Figure 2 are remarkably well predicted with use of the theory symbolized by eqn [1]. The theory may also be extended to cover interaction of swells with wind waves, with use of Sw Ss instead of S2w , where Ss is the swell spectral density, and with F similarly including the swell azimuthal directivity.
form, the spectral density S of the pressure field is
VLF Band: Atmospheric Turbulence Forcing as a Noise Source Sðf ; zÞ ¼ p r g =2c ½Sw ðf =2Þ f Fðf =2ÞTp ðf ; zÞ ½1 3 2 2
2
2 3
where r is water density, c is water sound speed, g is acceleration of gravity, Sw is the spectral density of the surface wave elevation, F is an integral over all azimuth angles f of the product of the normalized azimuthal directivity of the opposing surface waves (for f and f þ p), and Tp is a function normalized by its value at depth z ¼ 0 that relates the pressure spectral density at the surface to that at z, which importantly includes acoustic coupling to the sea bottom. For simplicity, eqn [1] includes contributions to the noise field only for horizontal phase speeds given by cx c. Because of its exponential decay in z, the remaining regime cx oc is important mostly for depths approaching z ¼ 0, especially for lower f, and is included in fuller analyses.
The heading of this section is in reality a question. In the VLF band (2–20 Hz), and to somewhat higher frequencies in seas with very low shipping density, noise measurements do not match direct extrapolation of wave–wave interaction data as in Figure 2. Other mechanisms may become important in this band, and one has attracted considerable attention, namely turbulent pressure fluctuations in the atmosphere that drive the ocean surface. Is the latter the responsible mechanism? And if not, what is? A robust answer is not yet available. In one of the more extensive data summaries dating from the 1960s, Crouch and Burt showed that in the absence of shipping noise, and at low wind speeds, S is p f 1 U0 between about 10–50 Hz. This holds until a crossover wind speed Uc is reached
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ACOUSTIC NOISE
(E16 m s1 at 11 Hz), beyond which S is roughly pf 3 U4 . The crossover speed Uc tends to increase with increasing frequency up to about 50 Hz. Nichols reported some data that also included frequencies lower than those of Crouch/Burt. These data show that for a small spread of wind speeds around UE9 m s1 , Spf 22:5 from 3 to 8 Hz, and Spf 0 from 8 to 20 Hz. These results compare well with those of Crouch/Burt at the overlapping frequencies, and add a functional form below 8 Hz not covered by Crouch/Burt. Nichols also summarized unpublished data for U varying from about 3 to 9 m s1, that show Spf 25 U5 from 2 to 5 Hz, and Spf 0 U5 from 5 to 10 Hz. (The f 0 U5 functional form could be argued from the data shown to be f 0 U0 at lower U, and f 0 U>5 at higher U.) This form at the lowest frequencies has a much sharper falloff with frequency than the Nichols’ form, but preference should be given to it because the frequency resolution was much finer. Accordingly, the overall result, labeled ‘Nichols’, is stated here as Spf 25 U5 from 2 to 5 Hz, and f 0 U5 from 5 to 20 Hz. More recently, as part of LF noise measurement programs, data in the upper part of the VLF band were published. Deep ocean noise versus U at f ¼ 13 and 50 Hz has been reported. For 5rUr15 m s1, these data have a functional form approximately Spf 0:8 U1:3 (perhaps more like f 0 U1:6 at the highest U). A report of shallow water noise versus U in third-octave bands for 10rf r20 Hz, with form Spf 3:2 U3:4 for U ¼ 3 and 5 m s1 in late fall, had data that differed so widely for U ¼ 5 and 9 m s1 in late spring that a functional form could not be stated. One report of deep ocean noise at f ¼ 15 and 25 Hz, for 2rUr12 m s1, suggests Spf 0 U0:5 for Uo Uc, and Spf 0 U3 for U > Uc, with the crossover speed Uc E8 m s1. What noise mechanism could account for all the foregoing observations? An extrapolation of wave– wave interaction noise to the VLF band, from data such as in Figure 2, suggests that Spf 4 U1=2 , give or take one integer in the exponent of f , and one-half integer in the exponent of U. However, this is unacceptably far from the data. The Crouch/Burt data suggest f 1 U0 and f 3 U4 for low and high U, respectively. The overall Nichols result is f 5 U5 from 2 to 5 Hz, and f 0 U5 from 5 to 20 Hz. Other data give f 0:8 U1:3 , f 3:2 U3:4 or an indefinite form, and f 0 U0:5 and f 0 U3 for low and high U, respectively. Without significant modifications applicable to the VLF band, it seems that the wave–wave interaction possibility must be set aside. Next, consider the atmospheric turbulence model. It has evolved as most theories do, but is contentious. It predicts Spf 0 U4 . This, too, is mostly far from the
55
functional form of the foregoing data, but does come close to Nichols and others (at the higher wind speeds) for the 5–25 Hz range. It seems inappropriate, however, to choose among available data sets for the ones that confirm a model. The difference between the data sets may well be caused by some mechanism that we are collectively ignorant of. Finally, it is possible that available data are at least partially contaminated by hydrophone flow noise, whose functional form goes as f 4 . None of the data sets matches this form. Thus, it can be concluded that the hydrophone flow noise mechanism is an unlikely cause of VLF noise. The identification of the mechanism responsible for VLF noise can thus not be made with confidence. In searching for candidate VLF noise mechanisms, one is inclined to look toward appropriate extensions or modifications of mechanisms in the adjacent ULF and LF bands, mainly because wave–wave interactions and distant shipping, respectively, are well established. Nonprevailing mechanisms should also be considered. For example, whale vocalizations are observed for 15of o35 Hz and can affect the VLF band. Noise data sets beyond those referred to here, supported by environmental data as suggested by candidate mechanisms, may well be needed. The Crouch/Burt data set incorporated a plausible but convoluted data analysis path to extract the VLF noise. The reported database of Nichols is not large. The VLF data of other workers could have been affected, as the authors acknowledged, by distant shipping noise. Perhaps because many of these research efforts were aimed at other objectives, environmental data provided with the noise data are generally too fragmentary to aid the search for VLF mechanisms.
LF Band: Distant Shipping Noise Evolving technology has altered the view of distant shipping noise. Increasing use of large aperture acoustic arrays, with attendant high-resolution beamwidths and focused scanning in range and use of highresolution frequency filters, blurs the distinction once sharp between distant and local ships. That is, ambient noise at LF can be observed with high-resolution technology as a countable number of discrete ship noise sources, rather than as a sum of noise from a very large number of widely distributed ships. Frequency Spectra
Figure 3 shows the noise radiated by a contemporary cargo ship. The radiation is largely tonal, as has long
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ACOUSTIC NOISE
0.5 Hz Bandwidth levels (dB re. 1 μPa at 1 m)
190 B 4
B 3 G F 4 2
180 B F B 1 1 2
170
160
B 5 F 3
G 5
G 6
G 2
140
B 8
B 7
B 9 F 6
B 10
B 11
F 5
G G 8 9
140 rpm G G 10 11
G 3
150
B 6 F 4
G 7
G G G 15 G 12 G 14 16 G G 13 17 18
G 1
G 20
G 30
22
58 rpm 32
25
G 19
27 21
23
24
26
31 28
33
29
130
20
40
60
120 80 100 Frequency (Hz)
140
160
180
200
Figure 3 On-axis source level spectra of a cargo ship at 8 and 16 knots (4 and 8 m s1) measured directly below the ship. Noise levels at distances beyond 1 m may be obtained by subtracting the transmission loss. Levels for bandwidths other than 0.5 Hz can not be determined from this figure because the bandwidths of the tones are not given. (B, F, and G in the figure identify, respectively, the harmonics of the (propeller) blade rate, the (diesel engine) firing rate, and the (ship’s service) generator rate.) (Reproduced from Arveson and Vendittis, 2000.)
been known. However, this is not obvious from the spectra shown in Figure 1, because they entail sums over many ships. The tonal envelopes in Figure 3 maximize between about 20 and 80 Hz, in good agreement with the summation spectra shown in Figure 1. Acoustic propagation losses in the ocean change the shape of the source spectrum shown in Figure 3; above about 80 Hz, the spectrum observed distantly is increasingly reduced with increasing f and with increasing distance from the ship. Directional Spectra
Noise radiated by a ship is a function of azimuth fs , and vertical angle ys , in a cylindrical coordinate system attached to the ship. The azimuthal spectral shape Sf ðf; f ; zÞ observed for a single ship at longer ranges is close to that measured near the ship. However, propagation of the noise to large ranges fundamentally affects the vertical directional spectrum Sy ðy; f ; zÞ. In these spectra, f and y are in a coordinate system attached to the observer (y ¼ 0 is the local horizontal plane). Figure 4 shows Sy as measured in deep water by a vertical line array. It sums over the directional spectrum in azimuth Sf , and therefore over the areal distribution of ships. A prominent feature of Sy is a pedestal of high noise around the horizontal, which is weakly dependent on f and z, and varies in half-width yw from about 715 to 201. These values are consistent with
cosyw Ecz =cb , where cz and cb are the sound speed at the observation depth and at the bottom, respectively. Distant shipping noise in deep water thus arrives mostly from source radiation at the surface near 7ys 7 ¼ 0, and then propagates to the observer via refraction and surface reflection paths in the water. Bottom reflection or transmission losses are relatively high, so that these paths are less important. Other effects influence the pedestal including, but not limited to, surface waves that modulate the source amplitude, scattering rather than specular reflection from the rough sea surface, and scattering from a seamount or continental margin. Because of these oceanographic and topographical complexities, the shape of the noise pedestal in Figure 4 is not general but instead suggestive of the main features of Sy in the deep ocean. In shallow waters, if cz =cb > 1 (downward refracting profile), then Sy is governed by path or mode losses, including those attributable to the bottom. If cz =cb o1 (upward refracting profile), then Sy would have a pedestal, but with yw typically an order of magnitude smaller than that observed in the deep ocean. Summation Issues
With several thousand ships underway in each of the heavily traveled oceans, taking account of all at the same time to determine distant shipping noise is
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ACOUSTIC NOISE
Wind speed High Mid Low Low
70 60 50
Power (dB / μPa/√Hz deg)
40 (A)
57
from limits that may be imposed by the focusing, filtering, or beam-forming processor, the spectral valleys are set by the ship’s radiation of continuous rather than tonal noise, and the spatial valleys by the sum of noise from more distant or less powerful ships. In addition, the valleys can be influenced by nonprevailing sources, the most prominent of which at LF is whale vocalization from about 15 to 35 Hz.
High 70
VLF Implications
Mid
60
With respect to the VLF band, the tonal envelope of a ship for f o20 Hz, at all of its higher speeds is about f 0 . Because very low frequency sounds can be detected at transoceanic distances, distant ships could cause the measured result to be Spf 0 U0 found in some VLF experiments, in the apparent but not real absence of distant ships.
Low
50 40 (B)
High 70
Mid
60
40 (C)
MF Band: Wind-driven Sea Surface Noise
Low
50 _
80
_
60
_
40 _ 20 0 20 Angle (deg)
40
60
80
Figure 4 Vertical angle directional spectrum SY summed in azimuth in deep water for f ¼ 75 Hz and U ¼ 3; 7, and 11 m s1. Positive angles are upward looking. (A) Shallow, (B) mid, and (C) deep depth refers to the vertical line array (VLA) centered at z ¼ 850, 1750, and 2650 m. The half-power beamwidth of the VLA at this frequency is about 1.31 when steered to y ¼ 0. (Reproduced from Sotrin and Hodgkiss, 1990.)
neither feasible nor necessary. At one extreme, lowresolution data in the LF band (20–200 Hz) are relatively insensitive to the detailed noise source characteristics of individual ships. Because summation from a large number of ships merges the details, only the broad and slowly evolving trends in shipping lane location, in shipping density, and in shipping composition will affect the level and horizontal directionality of the noise. At the other extreme, high-resolution data resolve the frequency and spatial spectra associated with distant ships, thus giving the experimenter the in situ noise field in relevant detail. The experimenter can avoid such noise in the spectral valleys between tones or the spatial valleys between high-noise beams2. Aside
2
Many ocean processes broaden tonal bandwidths and spatial beamwidths as the sound propagates from source to receiver. Such broadening is typically large enough to be a candidate ocean monitoring tool, but not so much that it completely fills the valleys.
Bubbles created by wind-driven surface waves have long been thought to be the dominant source of prevailing noise in the MF band (0.2–50 kHz). Many basic physical details, however, have only recently become better understood, and some relevant additional questions only recently posed. Various wavebreaking processes of wind-driven surface waves entrain air in the upper part of the ocean. Air-filled bubbles in the water are pinched off from the entrained air, which in turn oscillate and radiate noise as acoustic monopoles. Such noise thus entails wavebreaking and bubble hydrodynamics, both of which are addressed elsewhere in this encyclopedia. The acoustical aspects are addressed here. Vertical Directional Spectrum
In its simplest form, the theory for noise generated by a uniform distribution of sources on the surface is, from ray acoustics, 1 Sþ y ðy; f ; zÞ ¼ ðcs =cz ÞSDfsinys ð1 Rb Rs Þg yw ryrp=2
½2
þ S y ðy; f ; zÞ ¼ Rb Sy ; p=2ryr yw
½3
where cosyw ¼ cz =cs or yw ¼ 0
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for for
cz =cs o1; cz =cs 41
½4
ACOUSTIC NOISE
and where the directional spectrum in vertical angle Sy , per unit solid angle, is a function of f (at least through S, the pressure spectral density of the source per unit surface area) and of observation depth z (through sound speed cz at z). The superscripts þ and – refer to positive (upward looking) and negative (downward looking) y obtained from cos cosy ¼ ðcz =cs Þcos ys (the grazing angle ys and sound speed cs pertain to z ¼ 0). Dðys ; f ) is the directivity of the elemental bubble noise sources distributed just below the surface. For present purposes, the surface source distribution S, and the acoustic waveguide, is taken as uniform and independent of range and azimuth, contributions from propagation in the bottom are neglected, and volumetric absorption in the seawater is neglected. (These simplifications are adopted to keep the main ideas clear but, for more precise needs, can readily be replaced by assumptions that are more realistic.) For a downward refracting sound speed profile to depth zðcz =cs o1), the ray theory of [3][4] predicts a refractive shadow zone or notch of width 2yw around y ¼ 0, the horizontal plane. Wave theory must be used to properly predict the field in the notch, which also can be partially filled by scattering of the noise from midwater depths by fish schools and by ocean inhomogeneities. For an upward refracting profile (cz =cs 41), the field around y ¼ 0 is directly due to the surface-generated noise, plus possible scattering contributions. In eqns [2] and [3], Rb and Rs are, respectively, the coefficients of bottom and surface specular reflection. Terms involving these parameters can be important in the directional spectrum (but since perfect reflection is not likely for an acoustic waveguide in the ocean, they do not lead to singularities as eqns [2]–[4] might appear to suggest). For example, consider that Rb and Rs approach unity (but do not reach it) as the grazing angles at the bottom and the surface, respectively, approach zero. Then, for y within about 7p=4; Sy can be increased in typical situations by about 10 dB. In addition, the bottom propagation paths neglected here can actually contribute, especially at the lower MF frequencies. Thus details of the acoustic waveguide affect Sþ y and Sy and, along with the sound speed profile cðzÞ, could account for the plethora of somewhat dissimilar measured MF vertical directional spectra in the literature. Eqn [2] contains the bubble source directivity Dðys ; f ) that, unfortunately, is not known with confidence. At least two models for directive radiation from aggregated bubbles have been considered. One assumes an exponential decrease of uncorrelated monopoles below a horizontal perfectly reflecting
surface, and the other assumes a similarly situated monopole distribution concentrated on a submerged plane. Then, respectively, D ¼ 2f1 sincð2ks dsinys Þg
½5
2
½6
D ¼ 4sin ðks dsinys Þ
where ks is the acoustic wavenumber at the surface, d is the effective depth (the e-folding depth and the d-function depth, respectively), and sinc (x)(sin xÞ=x. In the limit ks d sin ys 51, these functions have the same shape, and close to the same magnitude (E1 dB different). Data, however, show that the two are distinct. For eqn [5], the data suggest ks dEp, whereas for eqn [6] ks dEp In either case, the idealized states assumed in eqn [6] might not represent the relevant complexity of the radiating bubbles beneath a breaking wave. For example, the exponential decay of bubble density with depth may well be a good model for horizontally isotropic bubbles quasistatically present as a result of previous wave breaking events, but a poor model for radiating bubbles immediately caused by a new event. Integration of eqns [2]–[4] over y to obtain the noise spectral density Sðf ; z) also depends sensitively on Rb and Rs (and on possible bottom propagation paths). This emphasizes the need to compare experimentally derived values of Sðf ; z) with appropriate knowledge of the acoustic waveguide. Alternatively, with use of eqns [2]–[4],S may be extracted from vertical line array (VLA) data. When a VLA is steered to y ¼ p=2, the specular reflection and the bottom propagation paths will contribute at most
65
Noise level (dB// 1 μPa2/sr-Hz)
58
60
55
50
45
5
10 25 15 20 20 Log (Wind speed) (kn)
30
Figure 5 Level of the source spectral density S, in dB re mPa2 per m2 Hz, for f ¼ 110 Hz. The 10 m wind speed U is in kn (1 knE1/2 m s1). (Reproduced from Chapman and Cornish, 1993.)
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ACOUSTIC NOISE
Source Spectral Density
Chapman and Cornish measured S in deep water with an upward-looking VLA. They apparently assumed eqn [6] for D, with ks dEp=2 . Their data at f ¼ 110 Hz, and for the wind speed interval 2oUo15 m s1, are reproduced in Figure 5, and show a crossover wind speed Uc E4.3 m s1. The frequency interval for their measurements is 13of o300 Hz, within which they found that Uc is about 4.3 m s1 for f r110 Hz, and that Uc is somewhat smaller (E3.5 m s1) for f 4110 Hz. Chapman/Cornish attribute the crossover speed to a transition in source mechanism physics. Furthermore, by regression analyses, their data show that for UoUc ; Spf 2:1 U0:6 , and for U > Uc ; Spf 2:1 U2:7. These results hold on average within the speed and frequency intervals measured. Kewley, Browning, and Carey reviewed and compared several data sets, mostly deep water VLA measurements, to extract S. They also used eqn [6] and ks dEp=2, and concluded that for 30of o 1300 Hz and 1oUo15 m s1, Uc E6 m s1 with SpU1 for UoUc and Sp U3 for U4Uc. The Kewley et al. wind speed exponents of 1 and 3 are not too different from those of Chapman/Cornish. When one considers that the former tilted their exponent choices somewhat to agree with extant physical models proposed for the below and above Uc regimes, the agreement can be considered quite satisfactory.3 What is more relevant, however, is that a universal spectral shape is not evident for either regime in the Kewley et al. comparisons. More likely than not S, f , and D need to be scaled by hydrodynamic parameters other than or additional to U, as shown below. Basic Wave-Breaking Correlates
Research results on hydrodynamically based scaling of noise from breaking waves have been reported. Kerman has proposed that at u =uc E1, where u is the friction velocity and uc is the minimum phase speed of gravity/capillary surface waves, the wavebreaking process transitions from one that has an aerodynamically smooth sea surface to another that
3
When compared at the same U and f , S is about 3 dB higher in the Chapman/Cornish data set than in the data reviewed by Kewley et al.
is rough. Kennedy analyzed VLA data in a deep, acoustically isolated bay (40of o4000 Hz, 2oUo15 m s1), with unlimited wind fetch but limited wave fetch. It was found that u =uc 0.9 defined a rough surface regime. (It may therefore be presumed that the crossover speed discussed in the foregoing section is Uc E0:9uc ) Figure 6 shows that the spectral data for u =uc 40.9 aggregate to an almost universal scalable spectrum. What garners the caveat of ‘almost’ is that frequency is scaled by fp, the observed peak frequency. Both Kerman and Kennedy point out that fp does not vary strongly. It ranges from about 300 to 800 Hz in the Kennedy data, and is not unlike that sketched by Wenz (Figure 1). But experimental interest does not always include measurement of fp , in which case a user of Figure 6 must slide the frequency scale without benefit of Kennedy’s judgement. Neither, however, can properly be accused of intellectual sloth. Kerman provides a model for fp, which contains wave-breaking parameters that unfortunately are poorly known. Kennedy’s collapsed spectral spread although acceptably small, is large enough, and the frequency dependence for 1/3of =fp o10 is weak enough, to forego fine attention to fp . Although apparently not used, fp is related semi-empirically to breaking-wave whitecap size fp (in Hz)E1400/(OLW), where L and W are, respectively, the whitecap along-crest length and crosscrest width, both in meters]. The source spectral density S in Figure 6 is obtained by Kennedy from a dipole directivity model. In effect, eqn [6] was used with ks d51, in this limit known as a compact dipole. With this assumption, the integral of Dðf ; ys Þ over a hemispherical surface _7
10
u* > 0.9 uc _8
10
ADSD (f)B pawcu *3
weakly. Such a measurement is thus dominated by local surface sources, so that S may be compared among measurements with less concern for waveguide properties.
59
_9
10
_ 10
10
_ 11
10
0.01
1
0.10
10
_f fp
Figure 6 Source spectral density S versus f , both nondimensional as described in the text. The source directivity model used is presumably eqn [6] with ks d 51 (the compact dipole model). Data are for the aerodynamically rough regime (u 4uc ). (Reproduced from Kennedy, 1992.)
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60
ACOUSTIC NOISE
yields AD ¼ 2p=3 and this appears as one of the S scaling terms in Figure 6.4 Another term is B ¼ 1000 Hz, the nominally observed bandwidth of the noise; its role is simply to create an integral measure SB for dimensional clarity. In the term E ¼ 5ra u3 ; ra is air density, and E is the major scaling variable, the average rate of energy dissipated per unit surface area by the breaking waves. Finally, AD SB=ðrw cs ), where rw is water density, is the average rate of acoustic energy radiated per unit surface area. Virtually simultaneously and independently, others have researched in greater depth some concepts that are related to the Kerman/Kennedy result. Noise spectral density Sðf ; zÞ has been correlated with dissipation E, in deep water under steady wind and wave conditions, for the intervals 4.3of o14 kHz and 2oUo12 m s1. The data on average show SpSpf 0:4 E0:74 , with the exponent of E varying from 0.86 to 0.67 from the low to the high end of the frequency interval. At constant E, the frequency dependence agrees reasonably with an extrapolation of Figure 6. But the dissipation dependence can not be compared without scaling the peak frequency fp, which was not observed. Thus, for a range of E one can seek the range of fp to satisfy linear scaling in E. The peak frequency fp would then need to decrease about a factor of 4 from low to high E, a factor so large as to suggest that a major change in noise physics occurs at these higher frequencies. Does the quasistatic bubble layer below the sea surface increasingly attenuate the noise, or increasingly inhibit its generation, at these frequencies? Bubbles are known to attenuate sound as a function of frequency and size distribution, but data analyses do not consider this. With use of the Fresnel field of an array of hydrophones, sound radiated by individual breaking waves has been measured in deep water (0.35of o4 kHz, 4oUo15 m s1). The on-axis source levels of individual breaking events, were obtained and modeled as spatially and temporally discrete compact dipoles eqn ([6] with ks d51). The individual source levels were correlated with U and cb , the latter being the speed of a breaking wave event, a measure closely connected to breaking wave dissipation E. The correlation with cb was found to
4 A spherical surface for the integral would seem more appropriate, since the only way a monopole can become a dipole is by including the negative image above the free surface, in which case AD ¼ 4p=3. Had a noncompact dipole been assumed with ks d ¼ p=2 (eqn [6]), then AD ¼ 2p. There is as much as 5 dB difference in these values compared to the one used by Kennedy.
be significantly better than that with U and, via physical arguments it was concluded that the source levels are well correlated with E. This measurement technique is also important as it determined the probability density of the dipole source levels, and the spatial density of discrete breaking wave events. It was also concluded, again via physical arguments, that the source spectral density for the frequencies measured are on average pE0.83, which in view of the lower frequencies observed might be taken as reasonably consistent with the E0.74 obtained by other workers. Thus the question remains on a possible frequency-dependent bubble layer effect. The foregoing results clearly have not answered all questions on MF noise caused by breaking waves. They do, however, provide more general predictive tools than those previously available, and identify at least some of the more important physical attributes of noise from breaking waves.
HF Band: Molecular Noise Molecules impinging on the surface of a pressure sensor cause noise, as estimated from physical principles and as plotted in Figure 1. Molecular motion, and thus momentum reversal on the sensor (i. e., force per unit area) is a function of molecular kinetic energy, and thus seawater temperature. On an absolute temperature scale, all oceans may be considered at a constant temperature. Hence, one line in Figure 1 is sufficient to estimate the noise.
See also Acoustics, Arctic. Acoustics, Acoustics, Shallow Water. Ships.
Deep
Ocean.
Further Reading Arveson PT and Vendittis DJ (2000) Radiated noise characteristics of a modern cargo ship. Journal of the Acoustical Society of America 107: 118--129. Buckingham MJ and Potter JR (eds.) (1995) Sea Surface Sound ’94, vol. 3. Singapore: World Scientific, 494pp. Chapman NR and Cornish JW (1993) Wind dependence of deep ocean ambient noise at low frequencies. Journal of the Acoustical Society of America 93: 782--789. Crouch WW and Burt PJ (1972) The logarithmic dependence of surface-generated ambient-sea-noise spectrum level on wind speed. Journal of the Acoustical Society of America 51: 1066--1072. Kennedy RM (1992) Sea surface sound source dependence on wave-breaking variables. Journal of the Acoustical Society of America 91: 1974--1982.
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ACOUSTIC NOISE
Kerman BR (1984) Underwater sound generation by breaking wind waves. Journal of the Acoustical Society of America 75: 149--165. Kerman BR (ed.) (1988) Sea Surface Sound, vol. 1. Dordrecht: Kluwer Academic Publishers, 639pp. Kerman BR (ed.) (1993) Sea Surface Sound, vol. 2. Dordrecht: Kluwer Academic Publishers, 750pp. Kewley DJ, Browning DG, and Carey WM (1990) Lowfrequency wind-generated ambient noise source levels. Journal of the Acoustical Society of America 88: 1894--1902. Kibblewhite AC and Evans KC (1985) Wave–wave interactions, microseisms, and infrasonic ambient noise
61
in the ocean. Journal of the Acoustical Society of America 78: 981--994. Nichols RH (1981) Infrasonic ambient ocean noise measurements: Eleuthera. Journal of the Acoustical Society of America 69: 974--981. Sotrin BJ and Hodgkiss WS (1990) Fine-scale measurements of the vertical ambient noise field. Journal of the Acoustical Society of America 87: 2052--2063. Urick RJ (1986) Ambient Noise in the Sea. Los Altos, CA, Peninsula Publishing. Wenz GM (1962) Acoustic ambient noise in the ocean: spectra and sources. Journal of the Acoustical Society of America 34: 1936--1955.
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ACOUSTIC SCATTERING BY MARINE ORGANISMS K. G. Foote, 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 1, pp 44–53, & 2001, Elsevier Ltd.
Historical Overview Development of underwater sonar as a tool for navigation and military operations, following sinking of the Titanic in 1912, led inevitably to applications to marine organisms. By the 1930s, echoes from fish schools had been detected. In the 1940s, the deep sound-scattering layer was observed. Its biological origin in mesopelagic fish was identified in the 1950s. At the same time, applications to commercial fish were pursued with vigor, and both scientific echo sounders and fishery echo sounders began to be manufactured. Steady improvements in transduction enabled individual fish of certain species and sizes to be detected at ranges of hundreds of meters. The ultrasonic frequency of 38 kHz was becoming a standard at this time; it was subsequently shown to be near the optimum for achieving detection of commercially important fish in the presence of attenuation due to spherical spreading and absorption. Parallel to studies of single-fish scattering at ultrasonic frequencies were studies of scattering at sonic frequencies, especially to determine the resonance frequency in swimbladder-bearing fish, which is a measure of size. Echo integration was introduced in 1965 as a tool for quantifying fish aggregations at essentially arbitrary conditions of numerical density. This was rapidly developed, and it has been used routinely in surveys of fish stock abundance since about 1975. Introduction of standard-target calibration in the early 1980s served the cause of quantification by providing a rapid, high-accuracy method of enabling the results of echo integration to be expressed in absolute physical units. With few exceptions, standard-target calibration has become the method of choice. Sonar, with one or more obliquely oriented or steerable beams, began to find common application in the 1970s for counting fish schools that might be missed by a vertical echo sounder beam. This was a significant development for acknowledging the narrowness of the sampling volume of vertically
62
oriented directional echo sounder beams and the possibility of fish avoidance reactions to the transducer platform, typically a research vessel. In another parallel development, the Doppler principle was exploited to measure the rate of approach or recession of fish targets. Both horizontally oriented echo sounder beams and sonar beams were used. Early applications determined the swimming speeds of schools of small pelagic fish and individual salmon in rivers. Applications of acoustics to fish in the 1970s were accompanied by notable applications to zooplankton, if pursued less intensively owing to differences in commercial importance. Because of the enormous diversity of zooplankton species in size, shape, and composition, it was recognized early that insonification over a band of frequencies is required, even for routine observation. This has usually been achieved by the use of multiple resonant transducers, but genuinely broadband sonars are also proving successful in yielding spectra of individual euphausiids and copepods. Recognition of the importance of bandwidth in scattering by zooplankton was accompanied by appreciation of the role of interpretive models. Acoustic scattering models have been developed and applied to fish since the 1950s and to zooplankton since the 1970s. The transition from analog to digital technologies in the 1970s facilitated processing of echo data. This has become steadily more automated and sophisticated, but always with operator control of important decisions through the man–machine interface. Other developments in technology since the 1970s have extended the range of applications of acoustic scattering by marine organisms. Multiple-element transducers have been used to determine the threedimensional locations and movement, as well as the target strength, of individual animals. Compact, high-frequency sonars have been mounted on fish capture gear to observe the behavior of fish during catching operations. Steerable high-frequency sonars have been used to track fish schools during capture and to map their three-dimensional shapes.
Physical Basis for Scattering Acoustic scattering by a marine organism is, in principle, no different from that of any other kind of scattering. Differences in the physical properties of the causative bodies with respect to the surrounding
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ACOUSTIC SCATTERING BY MARINE ORGANISMS
medium are accompanied by reflection and refraction, or more generally diffraction, of incident waves. Organisms, with contrasts in mass density or elasticity relative to sea water, are thus sources of scattering. The processes of reflection, refraction, and diffraction occur at surfaces, both external and internal, marking discrete changes in physical properties and throughout the volume or inside embedded inhomogeneities, as characterized by continuous changes in properties. The net result of the individual processes is a redistribution in space of the incident energy field. Changes in direction and amplitude characterize the scattering.
Classification of Marine Organisms as Scatterers Marine organisms are conveniently divided into groups based on considerations of taxonomy and anatomy. Two major groups are those of fish and zooplankton, but others are also treated. Fish may be distinguished as cartilaginous or bony. Bony fish may be acoustically distinguished because the fish possesses or lacks a gas-filled swimbladder. Swimbladders may be closed, with gas exchange effected by the rete mirabile, or open, with gas exchange effected by gulping air at the surface or by releasing a sphincter muscle on a duct leading to the exterior. The respective swimbladder types are called physoclists and physostomes. They are illustrated by cod (Gadus morhua) and herring (Clupea harengus), respectively. Some mesopelagic fish possess gas-filled swimbladders, including a number of myctophid species. Some other myctophids, as well as the deepwater fish orange roughy (Hoplostethus atlanticus), possess swimbladders that are invested with wax esters. The whiptail (Coryphaenoides subserrulatus), a macrurid, possesses a swimbladder that contains gas in a spongy matrix of tissue. Swimbladderless fish are illustrated by mackerel (Scomber scombrus). Cartilaginous fish lack a swimbladder, but their liver is large and presents a marked density contrast with the surrounding fish flesh. Zooplankton come in many shapes and sizes, but acoustically their variable physical composition admits of a severe reduction. Three prominent classes have been identified: the liquidlike, the hard-shelled, and the gas-bearing. These are illustrated by, respectively, euphausiids, pteropods, and siphonophores. Other marine organisms have also been detected by scattering. These include squid, gelatinous zooplankton, algae, benthos, marine mammals, and
63
even diving birds. The first five groups are considered in a separate section in the following.
Dependences of Scattering In general, scattering by marine organisms is affected by a number of factors. Some are listed here.
Intrinsic factors. Intrinsic to the scatterers are size, shape, internal composition, and condition. Condition may be affected by the stage of development, presence of reproductive products, and degree of stomach filling. Behavior is another intrinsic factor, if often directly affected or determined by the external environment. It is typically quantified through the attitude, or orientation, of the organism and its velocity of movement. Extrinsic factors. Scattering is affected by the insonification signal, hence by its spectral composition. For impulsive signals, the spectrum may be broadly continuous. For a typical pulsed sinusoid containing many wavelengths, the spectrum will be narrow, and the signal can be characterized by the center frequency, pulse duration, and amplitude. Depth and history of depth excursion may also influence the scattering, as in the case of rapid depth changes for physoclists. For swimbladdered fish lacking rete mirabile, depth excursions will necessarily affect the swimbladder form, with the volume changing in accordance with Boyle’s law, thus inversely with the ambient pressure.
Quantification of Scattering Nomenclature
Scattering properties of organisms are distinguished as belonging to individual organisms or to aggregations of organisms. The fundamental scattering property of a single organism is the scattering amplitude. This is described through the idealization of a plane harmonic wave incident on a finite scattering body. At a great distance r from the body, the scattered pressure field or amplitude psc is related to the incident pressure amplitude pinc by eqn [1]. psc ¼ pinc f expðikrÞ=r
½1
In eqn [1] f is the far-field scattering amplitude, r is the distance from the scatterer, k is the wavenumber 2p/l, and l is the acoustic wavelength. The scattering
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ACOUSTIC SCATTERING BY MARINE ORGANISMS
amplitude f describes the angular characteristics of the scattered field. The differential or bistatic scattering cross-section is jf j2 . In the backscattered direction f ¼ fb, and the backscattering cross-section is given by eqn [2], where the dual convention of using both sbs and s is shown. sbs ¼ jfb j2 ¼
s 4p
½2
The target strength TS is a logarithmic measure (eqn [3]; where r0 is the reference distance, typically 1 m). TS ¼ 10 log
sbs r20
½3
When many scatterers are concentrated in a volume in which individual scatterers cannot be distinguished by their echoes, a collective standard measure of scattering is used. This is the volume scattering coefficient. In the backscattered direction, the volume backscattering coefficient sv is given by eqn [4], where fb,i is the backscattering amplitude for the ith scatterer of N, and V is the volume. sv ¼ V 1
N X fb;i 2
½4
i¼1
The volume backscattering strength is given by eqn [5]. Sv ¼ 10 logðr0 sv Þ
½5
A quantity useful in echo integration is the area or column backscattering coefficient sa, (eqn [6]), where the integration is performed over the range interval [r1, r2].
sa ¼
Zr2
sv dr
½6
r1
In scattering by fish, a numerically more convenient measure of sa is eqn [7], which refers the backscattering to the reference area of one square nautical mile. sA ¼ 4p18522 sa
characteristic or mean backscattering cross section. sA ¼ r A s
Another measure of scattering is the extinction cross-section. This measures the relative loss of energy due to scattering and internal absorption. It may be defined for an individual scatterer, but is generally applied to aggregations of organisms if they are sufficiently numerous. With few exceptions, the issue of calibration must be addressed when making measurements. Standard methods are available for this, the aim being to define the system characteristics so that the result of a measurement, a voltage signal for instance, can be expressed as a pressure-wave amplitude in the water medium. Measurement
There are dozens of techniques for measuring the scattering properties of individual organisms and aggregations of organisms. These are commonly distinguished as being in situ, without constraint in the natural environment of the organisms, or ex situ, hence constrained in some way, wherever this might be. Target strength is a key quantity in many investigations. It may be determined with a single-beam echo sounder; for example, by repeated measurement of similar organisms that are acoustically resolved and by appropriate statistical reduction of these measurements. Alternatively, it may be measured directly with a dual- or split-beam echo sounder, in which the beam pattern can be determined in the direction of the organism, enabling the backscattering cross-section to be extracted from each individual echo. Similar measurements can be performed on single organisms ex situ with greater control and hence knowledge of their state during measurement. Measurements on tethered organisms, constrained to maintain a given orientation during insonification, are popular. Aggregations of organisms are frequently quantified acoustically through the volume backscattering coefficient. If the number and occupied volume of the organisms are known, then the characteristic target strength can be inferred through eqn [9].
½7
This form is particularly useful, for the fundamental equation of echo integration is simply eqn [8], where rA is the numerical density of fish referred to the same area of one square nautical mile, and s is the
½8
Sv ¼ 10 logn þ TS
½9
Here n is the numerical density of organisms, and TS is the so-called mean target strength corresponding to a single organism, but derived as the logarithmic measure of the mean backscattering cross-section.
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ACOUSTIC SCATTERING BY MARINE ORGANISMS
Cages are often employed to confine a known or knowable number of organisms to a fixed volume. Measurement of Sv can then yield a value for TS. Modeling
The importance of target strength in many studies involving scattering by marine organisms is so great that recourse is frequently made to theoretical models. On the basis of assumptions about the shape and internal composition of subject organisms, mathematical expressions may be derived that can be evaluated for particular conditions of concentration or frequency that might not be realistically explored through measurement. Ultimately, measurements may be used to refine models, and models to interpret measurements.
where g is the ratio of specific heats at constant pressure and volume, P is the ambient pressure at depth, r is the mass density of fish flesh, and a is the equivalent spherical radius. For elongated bubbles or swimbladder shapes, the resonance frequency is modified. The backscattering cross-section s at frequency v is given by eqn [11]. s¼
4pa2 h i2 ½n0 =ðnH Þ2 þ ðn0 =nÞ2 1
Fish as Scatterers The swimbladder shape varies with species and with condition of the individual specimen. An example of a swimbladder in corpus is shown in Figure 1. Here the swimbladder of an Atlantic herring (Clupea harengus) has been exposed by careful dissection. Low frequencies At low frequencies, with acoustic wavelengths much greater than characteristic swimbladder dimensions, the effect of a pressure wave on the swimbladder is essentially that of uniform compression and rarefaction. Consequently, a spherical model can be used. In fact, some swimbladder-bearing fishes have been modelled successfully as a spherical gas bubble surrounded by a finite layer of fish flesh that acts as a viscous fluid medium supporting surface tension on the interface between the shell and fish flesh. The volume of a bubble of radius a is equivalent to that of the swimbladder. Equation [10] gives the resonance frequency v0 of an immersed spherical gas bubble, 1 3gP 1=2 n0 ¼ 2p ra2
2pan x þ 2 ; n0 c pa n0 r
Figure 1 Drawing of a specimen of Atlantic herring (Clupea harengus), female, 36.0 cm long, 453 g, with exposed swimbladder. (Drawing by H. T. Kinacigil, used with permission.)
½12
Some numerical values for the various parameters are r ¼ 1050 kg m3 and x ¼ 50 Pa s. The speed of sound in sea water varies over the range 1450– 1550 m s1, depending on temperature, salinity, and pressure. For gadoids and clupeoids in the size range 8– 30 cm, n0 varies over 2.2–0.3 kHz. Given the inverse relationship of resonance frequency and size in eqn [10], smaller fish will have higher resonance frequencies. Thus, mesopelagic fish with partially waxinvested swimbladders may have resonance frequencies in the low ultrasonic range. Very large swimbladdered fish, say with a total fish length exceeding 1 m, will have resonance frequencies of the order of hundreds of hertz. The corresponding backscattering cross-section, hence target strength, can be computed from eqs [10]–[12]. It is important to note that the quality factor of the resonance condition, eqn [13], where Dn describes the range in frequency over which s decreases to one-half its maximum value, may be of the order of 1.5–3. Q ¼ n0 =Dn
½10
½11
H is the damping factor given by eqn [12], where c is the speed of sound in water, and x is the viscosity of fish flesh. H 1 ¼
Swimbladder-bearing Fish
65
½13
Implicit in the low-frequency condition of the model is that s is independent of orientation. Averages of s with respect to arbitrary orientation distributions will be identical to s itself. When computing average values of s for aggregations of swimbladdered fish of varying size, s must be averaged with respect to the size distribution. The characteristic target strength is determined from the definition in eqn [3]. Intermediate frequencies As the acoustic wavelength decreases toward characteristic
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ACOUSTIC SCATTERING BY MARINE ORGANISMS
swimbladder dimensions, the scattering becomes markedly directional, and the backscattering begins to depend sensitively on the orientation of the fish. From measurements made both in situ and ex situ, the empirical relationship of eqn [14] between mean target strength TS at 38 kHz and total fish length l in centimeters has been derived for a number of gadoids. TS ¼ 20 log l 67:5
½14
Equation [15] applies for clupeoids TS ¼ 20 log l 71:9
½15
The average backscattering cross-section s may be determined immediately from eqn [3]. For a cod of length l ¼ 50 cm, TS ¼ 33.5 dB and s ¼ 56 cm2. For a herring of length l ¼ 30 cm, TS ¼ 42.4 dB and s ¼ 7.2 cm2. Blue whiting is an important commercial stock in both hemispheres, and it is routinely surveyed by acoustics. To convert measurements of acoustic density at 38 kHz to numerical density in accordance with the echo integration equation [8], eqn [16], where l is the fork length in centimeters. is used for the northern-hemisphere blue whiting (Micromesistius poutassou): TS ¼ 21:7 log l 72:8
½16
Equation [17] applies for the southern-hemisphere southern blue whiting (Micromesistius australis), where l is again the fork length in centimeters. TS ¼ 25:0 log l 81:4
½17
Coincidentally, perhaps, the target strength of yellowfin tuna (Thunnus albacares) at 38 kHz is nearly identical to that of Micromesistius australis and is given by eqn [18]. TS ¼ 25:3 log l 80:6
is the total fish length in centimeters). TS ¼ 20 log l 72:7
Some stocks of orange roughy (Hoplostethus atlanticus) are being surveyed about their seamount habitats. Determination of the target strength of this deepwater fish with fat-invested swimbladder is admittedly problematical. Some work suggests convergence of the mean target strength of a 35 cm long orange roughy at 38 kHz to about 48 dB. If the standard equation for mean target strength–length were used, namely eqn [21], TS ¼ 20 log l þ b
½21
the coefficient b would be 79 dB. For modeling scattering by swimbladdered fish at these frequencies, the Kirchhoff approximation model can be used. This assumes that the fish is represented by the swimbladder, which acts as a pressure-release surface where it is directly insonified, and as a surface without response otherwise. A more general scattering model is that of the boundary-element method. The swimbladder is represented by a mesh of points, called nodes, spanning the surface, illustrated in Figure 2. The harmonic wave equation is solved numerically, assuming continuity of pressure and normal component of velocity at each node. It is thus possible to model the effects of internal gas density and pressure. To convert modeled values for s as a function of orientation to an average value, an orientation distribution is required. Ideally, this is done on the basis of in situ observations, but often such data are lacking and an orientation distribution must be assumed. Some orientation distributions are described in the literature. In some special circumstances it has
½18
The target strength of bigeye tuna (Thunnus obesus) under similar conditions is given by eqn [19]. TS ¼ 24:3 log l 73:3
½20
z
½19 x
These relations were established from specimens in the approximate size range 50–130 cm and 3–50 kg. The whiptail (Coryphaenoides subserrulatus), with a swimbladder containing gas-filled spongy tissue, seems to have a mean in situ target strength at 38 kHz that is consistent with the equation developed for another macrurid, the blue grenadier or hoki (Macruronus novaezelandie) (eqn [20], where l
y Figure 2 Boundary element model of the swimbladder of a specimen of pollack (Pollachius pollachius), 34.5 cm in length, with anterior end to the lower right (y direction). (Model by D. T. I. Francis, used with permission.)
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ACOUSTIC SCATTERING BY MARINE ORGANISMS
been possible to infer the orientation distribution by a combination of acoustic measurement and modeling. The relationship of maximum and average measures of s is given approximately by eqn [22]. smax E7save
½22
Alternatively, eqn [23] can be used. TSmax ETSave þ 5 dB
½23
Measures of the extinction cross-section are relatively rare, there being few occasions when it is necessary to compensate for scattering losses. However, measurement or inference suggest that the extinction cross-section is very roughly 1–3 times the backscattering cross-section at intermediate frequencies. Ultimately, the cross-sections and their ratio must depend on the behavior of the organism, as quantified through the orientation distribution. High frequencies When the acoustic wavelength becomes very small compared to the swimbladder size, scattering by other tissues may become important. The contributions of head structure, vertebrae, and even scales at very high frequencies have been established through ex situ measurement. Modeling of scattering by such structures can be computationally excessive, suggesting the advantages of stochastic modeling if direct measurement is not possible or convenient. Swimbladderless Fish
The mackerel is a prominent example of a swimbladderless fish. Its target strength must be attributed to the non-swimbladder structures and hence is intrinsically complicated at nearly all frequencies. At intermediate frequencies, the mean target strength is roughly 10 dB less than that of a gadoid of comparable size (eqn [24]). TSmackerel ETSgadoid 10 dB
½24
For cartilaginous fish, such as sharks, the liver may be very large. In pelagic sharks, this may be of the order of 7–23% by weight; in demersal sharks, 3–6%. The specific gravity of lipids is of the order of 0.87–0.92 in pelagic sharks and 0.93–0.94 in demersal sharks, further suggesting the role of the liver in buoyancy and its significance in acoustic scattering. At least for the pelagic sharks, the size and difference in mass density may explain much of the target strength. Were a model to be constructed,
67
a pelagic shark might be represented by a body with the size, shape, and physical properties of the liver.
Zooplankton as Scatterers Liquid-like Bodies
A number of prominent and abundant zooplankton can be classified as liquidlike in their acoustic properties. Extensive modeling and measurement have demonstrated that internal shear waves have negligible influence in scattering by such organisms. The animals are thus generally fluidlike in their properties. If the same animals lack sizable organs or other tissue presenting large contrasts in mass density or compressibility relative to the sea water immersion medium, then the acoustic properties of the organisms are more particularly liquidlike, and their acoustic scattering is consequently relatively weak. Two examples of zooplankton with liquidlike properties are euphausiids and copepods. These are also representative of homogeneous and inhomogeneous scatterers, respectively. Homogeneous liquidlike bodies The expectation of relatively weak scattering by euphausiids has been confirmed by measurement. For example, the target strength of Antarctic krill (Euphausia superba) of mean lengths 30–39 mm is in the range from 88 to 83 dB at 38 kHz and from 81 to 74 dB at 120 kHz. The respective acoustic wavelengths are 39 and 12.5 mm. For a scattering body that is relatively long compared to the wavelength, the scattering will be inherently directional. Laboratory measurement has demonstrated strong effects of orientation on scattering by euphausiids in the size range 30–42 mm at frequencies of 120 kHz and higher. In modeling scattering by homogeneous liquidlike zooplankton, there are just two significant material properties, the mass density and compressibility, or longitudinal-wave sound speed. A variety of models can be used to represent shape. At low frequencies, a single euphausiid can be represented by a finite circular cylinder or even a sphere, with volume equal to that of the animal. At higher frequencies, the same animal might be represented as a finite, bent, tapered cylinder or, better, by the actual shape of the exoskeleton. Scattering models for euphausiids have demonstrated the sensitive dependence of target strength on both the material properties and orientation of the organism. Given the rarity of measurements of material properties, their seasonal and individual
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variability, and the generally unknown orientation, there has been little systematization of measured values of target strength. Theoretical understanding of scattering by euphausiids has succeeded in associating large lobes with the echo spectrum at rather short acoustic wavelengths. When these are combined with knowledge of the target strength to within about an order of magnitude, it is possible to classify euphausiids by their acoustic signature. Inhomogeneous liquidlike bodies Copepods, like euphausiids, also display relatively weak acoustic scattering. Unlike euphausiids, however, their internal structure is acoustically distinct, being composed of two dominant scatterers, a prosome and an embedded oil sac. Because of the low density of lipids in the oil sac, of the order of 900 kg m3, the prosome must be correspondingly more massive. Because the copepod body as a whole is close to neutral buoyancy in sea water, the target strength is due to the internal contrast in mass density and compressibility, or longitudinal-wave sound speed, between the prosome and oil sac. Measurement has shown that the target strength of a 2 mm long copepod, Calanus finmarchicus, is in the approximate range from 95 to 90 dB over the frequency range 1600–2400 kHz. Copepods have been modeled as composite twoliquid-body structures. Numerical values for the mass density and longitudinal-wave sound speed have been derived from measurements or have been assumed. The shapes of embedded oil sac and encompassing prosome, illustrated in Figure 3, have been determined from videomicroscopic cross-sections in dorsal and lateral views. Results of modeling of copepods have shown the expected weak dependence on orientation at low or moderate frequencies, and an overall mean target strength that is in line with measured values. Hard-shelled Bodies
An example of a hard-shelled zooplankton is the pteropod Limacina retroversa, a marine snail with a spiral shell, opercular opening, and wings that propel it through the mid-water column. The target strength of specimens of shell length 2 mm has been measured over the approximate frequency range from 350 to 750 kHz. The target strength varies between 80 and 60 dB, depending on both frequency and orientation. The pteropod has been modeled as a rough spherical shell with a circular opening. Predictions of scattering have been in reasonable agreement with
z
y x z
y x Figure 3 Boundary element models of the prosome and oil sac of a specimen of Calanus finmarchicus, stage 6 female, 2.74 mm in length, with anterior end to the lower left (x direction). (Models by D. T. I. Francis, used with permission.)
measurements at wavelengths roughly comparable to the maximum shell dimension. Gas-bearing Bodies
Siphonophores are representatives of gas-bearing zooplankton, with gas inclusions in the pneumatophores. These are generally small compared to overall dimensions of specimens, and the target strength varies widely over the frequency range 350– 750 kHz. In particular, the target strength varies over the range from 90 to 60 dB, but with no apparent systematic dependence on frequency. This wide range is suggestive of interference between echoes from the gas inclusions and the nongaseous tissue, the basis of an acoustic model.
Other Organisms as Scatterers Squid
A number of specimens of squid have been observed by acoustics. These include Todarodes pacifica, Loligo opalescens, and Loligo vulgaris reynaudii. In a survey of the second species, performed at
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ACOUSTIC SCATTERING BY MARINE ORGANISMS
120 kHz, the target strength of specimens of mean dorsal mantle length 11.6 cm and mean mass 23.7 g was about 59 dB. Thus in the standard target strength–length equation [21], but with l representing the mean dorsal mantle length, b is about 80 dB. For Todarodes pacifica of mean dorsal mantle length 16 cm and mean mass 95 g, the target strength is about 51 dB at 28.5 kHz and 55 dB at 96.2 kHz, corresponding to values of b of 75 and 79 dB, respectively. For Todarodes pacifica of mean dorsal mantle length 23.7 cm and mean mass 340 g, the respective mean target strengths at 28.5, 50, 96.2, and 200 kHz are about 45.7, 46.5, 48.0, and 47.6 dB, with respective values of b of 75, 74, 76, and 76 dB. For Loligo vulgaris reynaudii, the target strength was measured at 38 kHz for sufficiently dispersed animals of mean mass 300 g. The target strength when referred to 1 kg was 42.5 dB. This compares favorably with the measurements on Loligo opalescens at 120 kHz and Todarodes pacifica at 28.5 kHz. When expressed relative to 1 kg, the respective target strengths are 42.3 and 41.1 dB.
69
and the masses 0.33–0.8 kg. Target strengths expressed relative to 1 kg of biomass vary from 35 to 28 dB at 50 kHz, from 33 to 24 dB at 70 kHz, and from 29 to 22 dB at 200 kHz. Smaller algae, the phytoplankton Prorocentrum micans, Peridinium triquetrum, Olistodiscus luteus, Dunaliella salina, Platimous viridis, and Phaeodactilum tricornutum, are also being studied by acoustics. Measurements of reverberation, in particular, are being used in attempts to quantify the volume of gas vacuoles. Clams
Both the razor clam (Tagelus dombeii) and the surf clam (Mesodesma donacium) have been surveyed by acoustics. Beds of the razor clam have been surveyed in shallow water over a flat bottom. Echograms that show the bottom–surface–bottom reflection in addition to the first bottom reflection show an enhanced registration above the so-called second bottom echo. Counting of its characteristic serrations provides a quantitative measure of clam density. Marine Mammals
Common Jellyfish
In anticipation of acoustic surveying of the ctenophore Mnemiopsis leidyi and other gelatinous zooplankton, namely Aurelia aurita and Pleurobrachia pileus, in the Black Sea, measurements have been made of the target strength of the common jellyfish Aurelia aurita. Functional regression equations have related the mean target strength in decibels to the disk diameter d in centimeters. At 120 kHz, the relation is eqn [25]. TS ¼ 14:7 log d 74:6
½25
At 200 kHz it is eqn [26]. TS ¼ 39:6 log d 104:4
½26
Thus for a specimen with mean diameter 10 cm, TS ¼ 59.9 and 64.8 dB at the respective frequencies. Algae
Algae, such as kelp, are being surveyed by acoustics. For purposes of quantification, the acoustic properties of the plants themselves are being studied, both by experiment and by theoretical modeling. Measurements have been performed on leaves of Laminaria saccarina and L. digitata at three ultrasonic frequencies. The lengths of these span the range 0.7– 2 m; the widths 0.4–0.9 m; the thicknesses 1–5 mm;
A few measurements have been reported on the target strength of the sperm whale (Physeter catodon) and the humpback whale (Megaptera novaeangliae) in situ. Measurements have been made of the Atlantic bottlenose dolphin (Tursiops truncatus) in captivity. Measurements made on a 2.2 m long 126 kg female dolphin in broadside aspect at the surface revealed a mean target strength that decreased from about 10 dB at the lowest measurement frequency of 23 kHz to about 24 dB at 45 kHz, rising to about 20 dB at 65 kHz, then falling to 25 dB at 80 kHz. The observed degree of variability about these nominal values due to repeated insonification was 4–11 dB to within the first standard deviation to either side.
Challenges For all of the instances and applications of acoustic scattering by marine organisms, there is an enormous demand for enhanced imaging capability and more quantitative understanding, including both improved measurement methods and models. In addition to refinement of current measurement methods, including those for quantifying concentrations of marine organisms, instruments are being developed or adapted for application. These include high-frequency sonars, multibeam sonars, and continuously broadband echo sounders, operating at both low and high frequencies.
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In general, the addition of bandwidth to acoustic devices, whether achieved by multiple frequencies or a continuous spectrum, is a firm objective of many development efforts. Its usefulness in classification is appreciated from certain studies in zooplankton scattering, but it would aid studies of nekton scattering if successful. Recognition of the importance of understanding the acoustic properties of individual organisms is similarly influential in promoting developments and applications. Determining the properties of single organisms when found en masse remains a challenge, as does quantifying avoidance reactions or avoiding inducing them. While there are many techniques for determining target strength, their application requires ingenuity to elucidate some of the principal dependences. The general lack of information on the depth dependence of target strength for gas-bearing organisms is a particular, prominent example. Modeling of scattering by marine organisms offers much potential for resolving physically intractable problems, such as those involving separation of echoes from individual organisms in the midst of their social aggregations or inferring the acoustic properties of organisms that are very fragile or that occur in extreme environments. Both analytical and numerical models, however, require knowledge of the physical properties, shape, and behavioral characteristics, such as the orientation distribution, of the subject organisms. Acoustic inference of the in situ properties of organisms, by special measurement techniques and aided by models, appears very attractive if generally difficult. An enhanced imaging capability based on acoustic scattering is also valuable. If realized in a compact device, this could aid fishing practice, as in providing fishers with information on the species and size of organisms present in the water column or on the bottom without actually having to capture the organism to make the determination. For the researcher, being able to distinguish different organisms with overlapping distributions would be invaluable in aiding the study of relationships, ultimately to advance the goals of ecosystem analysis and understanding.
Acknowledgments This is Woods Hole Oceanographic Institution contribution number 10271.
See also Acoustics, Deep Ocean. Bioacoustics. Fish Locomotion. Mesopelagic Fishes. Pelagic Fishes. Plankton. Sonar Systems.
Further Reading Craig RE (ed.) (1984) Fisheries Acoustics A symposium held in Bergen, 21–24 June 1982, Rapports et ProcesVerbaux des Reunions, vol. 184. Copenhagen: International Council for the Exploration of the Sea. Foote KG (1997) Target strength of fish. In: Crocker MJ (ed.) Encyclopedia of Acoustics, vol. 1, pp. 493--500. New York: Wiley. Foote KG and Stanton TK (2000) Acoustical methods. In: Harris RP, et al. (ed.) ICES Zooplankton Methodology Manual, pp. 223--258. London: Academic Press. Freon P and Misund OA (1999) Dynamics of Pelagic Fish Distribution and Behaviour: Effects on Fisheries and Stock Assessment. Oxford: Fishing New Books. Karp WA (ed.) (1990) Developments in Fisheries Acoustics, A symposium held in Seattle, 22–26 June 1987, Rapports et Proces-Verbaux des Reunions, vol. 189. Copenhagen: International Council for the Exploration of the Sea. Margetts AR (ed.) (1977) Hydro-acoustics in Fisheries Research, A symposium held in Bergen, 19–22 June 1973, Rapports et Proces-Verbaux des Reunions, vol. 170. Copenhagen: International Council for the Exploration of the Sea. Medwin H and Clay CS (1998) Fundamentals of Acoustical Oceanography. San Diego, CA: Academic Press. Nakken O and Venema SC (eds.) (1983) Symposium on Fisheries Acoustics, Selected papers of the ICES/FAO Symposium on Fisheries Acoustics, Bergen, Norway, 21–24 June 1982, FAO Fisheries Report no. 300. Rome: Food and Agriculture Organization of the United Nations. Ona E (1990) Physiological factors causing natural variations in acoustic target strength of fish. Journal of the Marine Biological Association of the United Kingdom 70: 107--127. Physics of Sound in the Sea (1969) Reprint of the 1946 edition. Washington, DC: Department of the Navy. Progress in Fisheries Acoustics (1989) Proceedings of the Institute of Acoustics, vol. 11, no. 3. St. Albans: Institute of Acoustics. Simmonds EJ and MacLennan DN (eds.) (1996) Fisheries and Plankton Acoustics, Proceedings of an ICES international symposium held in Aberdeen, Scotland, 12–16 June 1995. ICES Journal of Marine Science, vol. 53, no. 2. London: Academic Press. Urick RJ (1983) Principles of Underwater Sound, 3rd edn. New York: McGraw-Hill.
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ACOUSTIC SCINTILLATION THERMOGRAPHY P. A. Rona, Rutgers University, New Brunswick, NJ, USA C. D. Jones, University of Washington, Seattle, WA, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction: Mapping Diffuse Flow at Seafloor Hydrothermal Sites Acoustic scintillation thermography or AST is a method to detect and map diffuse flow at seafloor hydrothermal sites. Diffuse flow is the discharge of low-temperature hydrothermal solutions (up to tens of degrees Celsius) as seepage through areas of the seafloor. It is considered widespread at low intensity (temperature less than 1 1C; flow rate less than 1 cm s 1) in ocean basins and at higher intensity (temperature less than 100 1C; flow rate less than 1 m s 1) in seafloor hydrothermal fields. Quantitative assessment of diffuse flow is important because the cumulative thermal flux of diffuse flow through areas of the seafloor may equal or exceed that of focused flow from associated high-temperature (up to 405 1C) higher-flow-velocity (flow rate tens to hundreds of centimeters per second) black smoker vents. Chemical fluxes in diffuse flow may be selectively significant. Measuring diffuse flow is difficult. Occurrence is patchy and temperatures and flow velocities are low. Because it is clear, lacking the suspended particulate matter of black smoker plumes, diffuse flow cannot be detected by measuring attenuation or backscatter of light and sound.
Method The AST method uses the phase-coherent correlation of acoustic backscatter from consecutive sonar
scans of the seafloor to detect weak fluctuation in the index of refraction of the water near the seafloor. Random index of refraction changes result from temporal variations in the water temperature caused by turbulent mixing, which produce detectable changes in travel time of an acoustic ray as the ray propagates from a stationary acoustic transducer through the diffuse flow to the seafloor and is scattered back through the diffuse flow. These changes in travel time increase from the nearer to the farther portion of the sonar footprint, causing the associated echo waveform to change in shape with each transmission and resulting in decorrelation of successive sound pulses. In essence, the scintillation of the acoustic wave as it passes through a turbulent flow field and scatters off the underlying seafloor makes the seafloor appear to shimmer, much as the eye would detect hot water or hot air shimmering against a stationary backdrop. For certain types of seafloor hydrothermal flow (such as the case with diffuse flow where buoyant turbulent microplumes are concentrated near the seafloor), the AST method can provide a means of remotely detecting areas of flow and a potential method of measuring the scales of temperature and velocity fluctuations in the near-bottom boundary layer. Consider the scattering geometry illustrated in Figure 1, where a turbulent boundary layer overlays a rough seafloor and the incident acoustic field propagates forward and back between the transducer and the backscattering seafloor through the boundary layer. Assume that the seafloor is not changing in time and there is no motion of the acoustic transducer, but the turbulent flow is evolving temporally. Between consecutive and rapid sonar returns from the same spot on the seafloor, turbulent mixing will cause slight spatial and temporal fluctuations of the index of refraction of the water near the seafloor. These fluctuations will cause weak random forward scattering of the acoustic field as it propagates
z Stationary source receiver
Vertical beam pattern
Incident field Turbulent flow x
Seafloor Figure1 Propagation of an acoustic pulse through a turbulent boundary layer near the seafloor produced by diffuse hydrothermal flow and backscatter of the pulse from the seafloor.
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through the turbulent flow, resulting in temporal fluctuations in the amplitude and phase of the acoustic field measured at the receiver. The temporal changes in backscatter can be measured by estimating the phase-coherent temporal decorrelation of backscatter between successive echoes, whereas the backscatter intensity from the turbulence itself is too weak to measure. In this case, the rate and magnitude of temporal decorrelation can be used to detect and potentially measure scales of mixing in the boundary layer.
Application The method was first applied to detect and map diffuse flow in seafloor hydrothermal fields using a human occupied vehicle (HOV). Several later experiments have shown that deep ocean diffuse flow can be detected over relatively large areas of the seafloor using a remotely operated vehicle (ROV) platform. In these experiments, backscatter was recorded using a 200 kHz multibeam sonar mounted on an ROV as it hovered above the seafloor at intervals along a track line. Figure 2 shows an area of detected diffuse flow (as verified by video) and the corresponding backscattered intensity image. The two images are derived from the same sonar scans of the seafloor taken at 0.1 s intervals. Future applications will extend the AST method to seafloor observatories to record changes in area of diffuse flow at a seafloor site on timescales up to years. An informative representation is to drape areas of diffuse flow detected by the AST method over a corresponding bathymetric map to show the relation between diffuse flow and seafloor morphology. Backscatter intensity (dB)
Volume and heat fluxes from areas of diffuse flow can be estimated when in situ measurements of water temperature and vertical flow velocity are made in the areas of diffuse flow. For example, in situ measurements in diffuse flow of temperature using thermistors and of vertical flow velocity using video to record rise rate against a calibrated rod in an acoustically imaged area of diffuse flow and black smokers in the Main Endeavour Field indicated that heat flux of the diffuse flow was a multiple of that of the associated black smokers. Figure 3 shows areas of decorrelation intensity (draped onto bathymetry) for a relatively large area of seafloor surveyed at the Clam Bed hydrothermal field on the Juan de Fuca Ridge (near the Main Endeavour Field site). The AST method was applied using a sonar mounted on a hovering ROV to map the diffuse flow over a 3500 m 900 m area of the axial valley of the Endeavour segment of the Juan de Fuca Ridge. Using in situ sensors to simultaneously measure temperature and flow velocity, a diffuse heat flux of 150 MW was integrated over the areas of the AST decorrelation anomalies, indicating that the heat transferred by diffuse flow is a significant component of total heat flux in the study area. The temporal correlation of the backscattered field is found by correlating collocated returns from the seafloor as a function of transmission time. Consider the monostatic geometry for backscatter as illustrated in Figure 1, where a source/receiver system is fixed in space and measures backscatter at discrete pulse transmission times tn. The temporal correlation of backscatter signals between a transmission at time tm and a later transmission at tn is defined as Cðr; tm ; tn Þ ¼ huðr; tm Þu ðr; tn Þi
50
50
40
40
30
30
20
20
10
10
0
20
0 Meters
20
40
0
½1
Decorrelation intensity (dB)
20
0 Meters
20
40
Figure 2 An AST image of diffuse flow near Hulk vent in the Main Endeavour Field on the northern Juan de Fuca Ridge is shown in the right panel. The areas of diffuse flow (light patches) are detected as an increase in the decorrelation intensity. A conventional sonar image of the same area is shown in the left panel. Colors indicate level of backscatter intensity (red is highest intensity) related to bottom roughness.
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ACOUSTIC SCINTILLATION THERMOGRAPHY
73
30 20 10
50 100 150
Me
150
ter
s
200 100 250 50 300
Figure 3 Areas of diffuse hydrothermal flow from the seafloor (Clam Bed site on the northern Juan de Fuca Ridge) detected by backscatter decorrelation (area in meters; Jones et al., 2000). Colors indicate level of decorrelation intensity (yellow is higher intensity) corresponding to intensity of diffuse flow.
where u(r; tn) ¼ A(r; tn) exp[if(r; tn)] is the complex envelope of the backscattered field as a function of range r along the seafloor. Range is determined by round-trip acoustic time-of-flight as is usual in sonar. The complex envelope, also known as the baseband signal, incorporates the amplitude (A) and phase (f) of the echo signal, and the asterisk in eqn [1] denotes complex conjugation. For m ¼ n, the correlation function is real and proportional to the signal intensity. With man, the correlation function will be complex due to signal fluctuations in amplitude and phase, with phase dominating. If the area of seafloor observed has not changed because of motion of the source/receiver, fluctuations in the signal will be due to changes in the water above the seafloor as a function of time tn. With m ¼ 0 as the reference time and n ¼ 1, 2, 3,y increasing monotonically, eqn [1] is the cumulative temporal correlation function. For a single realization and finite number of discrete data points, an estimate of the ideal correlation function is found by windowing the backscattered signals in range and summing over range bins, M X ˆ i ; t0 ; t n Þ ¼ 1 uðri ; t0 Þu ðri ; tn Þ Cðr M i¼1
½2
where ri is a set of M discrete range values that correspond to a finite area (or patch) of the seafloor centered at range r.
The correlation function is a measure of the change in the scattering medium, but noise and motion will also cause signal decorrelation. To correct for the effects of low signal-to-noise levels when the return from the seafloor is weak, it is convenient to define the decorrelation intensity as Ir ðr; t0 ; tn Þ ¼ Iðr; t0 ; tn Þ½1 rðr; ˆ t0 ; tn Þ
½3
where the normalized temporal correlation coefficient is .qffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi ˆ ˆ t n ; tn Þ ˆ t0 ; t0 Þ Cðr; t0 ; tn Þ rðr; Cðr; ˆ t0 ; tn Þ ¼ Cðr; ½4 The scalar intensity I is the geometric mean of the intensity in the range window, defined as
Iðr; t0 ; tn Þ ¼
qffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi ˆ t0 ; t0 Þ Cðr; ˆ tn ; t n Þ Cðr;
½5
The decorrelation intensity [3] is useful as a relative measure of change when the scattering strength over an area of the seafloor is nearly uniform and helps avoid the comparison of areas where the scattered signal level is relatively low. Motion of the
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source/receiver is more difficult to compensate and will depend on the randomness in the seafloor. Angular motion of the source/receiver between recordings should be much smaller than the angular correlation scale of the scattering from the seafloor. If there is motion the rate of signal decorrelation due to motion must be less than the rate of signal decorrelation due to the diffuse flow. The size of the area of seafloor (or patch) used to estimate the correlation is defined by the length of the transmitted pulse and the number of samples needed to form an accurate estimate. For a bandwidth-limited Gaussian signal with bandwidth B, it is well known that the normalized mean-square error of the correlation estimate is given as h i h i ˆ t0 ; tn Þ ¼ 1 1 þ r2 ðr; t0 ; tn Þð1 þ N=SÞ2 e2 Cðr; 2BT ½6 where r is the desired true correlation coefficient and T is the length of the record used to form the estimate (corresponding to the number of points M or patch size). The signal-to-noise ratio (S/N) is assumed constant and uncorrelated between transmission times t0 and tn.
Scattering by Marine Organisms. Acoustics, Deep Ocean. Acoustics in Marine Sediments. Acoustics, Shallow Water. Bioacoustics. Dispersion from Hydrothermal Vents. Hydrothermal Vent Biota. Hydrothermal Vent Deposits. Hydrothermal Vent Ecology. Hydrothermal Vent Fluids, Chemistry of.
Further Reading Bendat JS and Piersol AG (2000) Random Data: Analysis and Measurement Procedures, 3rd edn., pp. 291–296. New York: Wiley-Interscience. Johnson HP, Hautala SL, Tivey MA, et al. (2002) Survey studies hydrothermal circulation on the northern Juan de Fuca Ridge. EOS, Transactions, American Geophysical Union 83(73): 78--79. Jones CD, Jackson DR, Rona PA, and Bemis KG (2000) Observations of hydrothermal flow (abstract). Journal of the Acoustical Society of America 108(5): 2544-2545. Rona PA, Jackson DR, Wen T, Jones CD, Mitsuzawa K, and Bemis KG (1997) Acoustic mapping of diffuse flow at a seafloor hydrothermal site: Monolith Vent, Juan de Fuca Ridge. Geophysical Research Letters 24(19): 2351--2354.
See also Acoustic Measurement of Near-Bed Sediment Transport Processes. Acoustic Noise. Acoustic
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ACOUSTICS IN MARINE SEDIMENTS T. Akal, NATO SACLANT Undersea Research Centre, La Spezia, Italy Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 1, pp 72–88, & 2001, Elsevier Ltd.
Introduction Because of the ease with which sound can be transmitted in sea water, acoustic techniques have provided a very powerful means for accumulating knowledge of the environment below the ocean surface. Consequently, the fields of underwater acoustics and marine seismology have both used sound (seismo-acoustic) waves for research purposes. The ocean and its boundaries form a composite medium, which support the propagation of acoustic energy. In the course of this propagation there is often interaction with ocean bottom. As the lower boundary, the ocean bottom is a multilayered structure composed of sediments, where acoustic energy can be reflected from the interface formed by the bottom and subbottom layers or transmitted and absorbed. At low grazing angles, wave guide phenomena become significant and the ocean bottom, covered with sediments of different physical characteristics, becomes effectively part of the wave guide. Depending on the frequency of the acoustic energy, there is a need to know the acoustically relevant physical properties of the sediments from a few centimeters to hundreds of meters below the water/ sediment interface. Underwater acousticians and civil engineers are continuously searching for practical and economical means of determining the physical parameters of the marine sediments for applications in environmental and geological research, engineering, and underwater acoustics. Over the past three decades much effort has been put into this field both theoretically and experimentally, to determine the physical properties of the marine sediments. Experimental and forward/ inverse modeling techniques indicate that the acoustic wave field in the water column and seismoacoustic wave field in the seafloor can be utilized for remote sensing of the physical characteristics of the marine sediments. Sediment Structure as an Acoustic medium
Much of the floor of the oceans is covered with a mixture of particles of sediments range in size from
boulder, gravels, coarse and fine sand to silt and clay, including materials deposited from chemical and biological products of the ocean, all being saturated with sea water. Marine sediments are generally a combination of several components, most of them coming from the particles eroded from the land and the biological and chemical processes taking place in sea water. Most of the mineral particles found in shallow and deep-water areas, have been transported by runoff, wind, and ice and subsequently distributed by waves and currents. After these particles have been formed, transported, and transferred, they are deposited to form the marine sediments where the physical factors such as currents, dimensions and shapes of particles and deposition rate influence the spatial arrangements and especially sediment layering. Particles settle to the ocean floor and remain in place when physical forces are not sufficiently strong to move them. In areas with strong physical forces (tidal and ocean currents, surf zones etc.) large particles dominate, whereas in low motion energy areas (ocean basins, enclosed bays) small particles dominate. During the sedimentation process these particles, based on the physical and chemical interparticle forces between them, form the sedimentary acoustic medium: larger particles (e.g., sands) by direct contact forces; small particles (e.g., clays and fine silts) by attractive electrochemical forces; and silts, remaining between sands and clays are formed by the combination of these two forces. The amount of the space between these particles is the result of different factors, mainly size, shape, mineral content and the packing of the particles determined by currents and the overburden pressure present on the ocean bottom. Figure 1 is an example of a core taken very carefully by divers, to ensure an undisturbed internal structure of the sediment sample. Sediment structures have been quantified by using X-ray computed tomography to obtain values of density with a millimeter resolution for the full three-dimensional volume. The image shown in Figure 1 is a false color 3D reconstruction of a core sample at a site where sediments consist of sandy silt (75% sand, 15% silt, and 10% clay) and shell pieces. The results of the X-ray tomography of the same core, can also be shown on an X-ray cross-section slice along the center of the core (Figure 2A) and the corresponding two-dimensional spectral density levels for that cross-section (Figure 2B). The complex
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300 250 200 150 100 110
50
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350 300 250
60 200 150 100
10
50
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Figure 1 Three-dimensional reconstruction of a sediment core sample containing 75% sand, 15% silt and 10% clay. Scale in mm.
_ 100
CT number 2250
X-ray CT Section
dB re m
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_
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_
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Figure 2 (A) X-ray cross-section slice along the center of the core shown in Figure 1. (B) 2-D spectral density levels for the crosssection shown in (A).
structure of the sediments can be with seen strong heterogeneity (local density fluctuation) of the medium that controls the interaction of the seismoacoustic energy. In addition to the complex fine structure described above, the seafloor can also show complex layering (Figure 3A) or a simpler structure (Figure 3B). These structures result from the lowering of sea levels during the glaciation of the Pleistocene epoch during which sand was deposited over wide areas of the continental shelves. Unconsolidated
sediments subsequently covered the shelves as the sea level are in postglacial times.
Biot–Stoll Model Various theories have been developed to describe the geoacoustic response of marine sediments. The most comprehensive theory is based on the Biot model as elaborated by Stoll. This model takes into account various loss mechanisms that affect the response of
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ACOUSTICS IN MARINE SEDIMENTS
N 42 56.536 E10 13.836
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Figure 3 Seismic reflection profiles showing (A) complex and (B) simple structures of sediment layers.
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porous sediments that are saturated with fluid. The Biot–Stoll theory shows that acoustic wave velocity and attenuation in porous, fluid-saturated sediments depend on a number of parameters including porosity, mean grain size, permeability, and the properties of the skeletal frame. According to the Biot–Stoll theory, in an unbounded, fluid-saturated porous medium, there are three types of body wave. Two of these are dilatational (compressional) and one is rotational (shear). One of the compressional waves (‘first kind’) and the shear wave are similar to body waves in an elastic medium. In a compressional wave of the ‘first kind’, the skeletal frame and the sea water filling the pore space move nearly in phase so that the attenuation due to viscous losses becomes relatively small. In contrast, due to out-of-phase movement of the frame and the pore fluid, the compressional wave of the ‘second kind’ becomes highly attenuated. The Biot theory and its extensions by Stoll have been used by many researchers for detailed description of the acoustic wave–sediment interaction when basic input
parameters such as those shown in Table 1 are available.
Seismo-acoustic Waves in the Vicinity of the Water–Sediment Interface As mentioned above, when acoustic energy interacts with the seafloor, the energy creates two basic types of deformation: translational (compressional) and rotational (shear). Solution of the equations of the wave motion shows that each of these types of deformation travels outward from the source with its own velocity. Wave type, velocity, and propagation direction vary in accordance with the physical properties and dimensions of the medium. The ability of seafloor sediments to support the seismo-acoustic energy depends on the elastic properties of the sediment, mainly the bulk modulus (incompressibility, K) and the shear modulus (rigidity, G). These parameters are related to the compressional and shear velocities CP and CS respectively, by: CP ¼ ½ðK þ 4G=3Þr1=2
Table 1
Basic input parameters to Biot–Stoll model
Frequency-independent Variables Porosity (%) P Mass density of grains rr Mass density of pore fluid rf Bulk modulus of sediment grains Kr Bulk modulus of pore fluid Kf Variables affecting global fluid motion Permeability k Viscosity of pore fluid Z Pore-size parameter a Structure factor a Variables controlling frequency-dependent response of frame Shear modulus of skeletal frame m¯ ¼ mr ðoÞ þ imi ðoÞ Bulk modulus of skeletal frame K¯ b ¼ Kbr ðoÞ þ iKbi ðoÞ
Table 2
CS ¼ ðG=rÞ1=2 where, r is the bulk density. Table 2 shows basic seismic–acoustic wave types and their velocities related to elastic parameters These two types of deformations (compressional and shear) belong to a group of waves (body waves) that propagate in an unbounded homogeneous medium. However, in nature the seafloor is bounded and stratified with layers of different physical properties. Under these conditions, propagating energy undergoes characteristic conversions every time it interacts with an interface: propagation velocity,
Basic seismo-acoustic wave types and elastic parameters
Basic wave type Body wave Ducted wave
Wave velocity Compressional
CP ¼ ½ðK þ 4G=3Þ=r1=2
Shear
CS ¼ ðG=rÞ1=2
Love
CL ¼ ðG=rÞ1=2
CSCH ¼ ðG=rÞ1=2 Elastic parameters in terms of wave velocities (C) and bulk density (q) Bulk modulus (incompressibility) K ¼ rðCP2 4CS2 =3Þ Compressibility b ¼ 1=K Young’s modulus E ¼ 2CS2 rð1 þ sÞ Poisson’s ratio (transv./long. strain) s ¼ ð3E rCP2 Þ=ð3E þ 2rCP2 Þ Shear modulus (rigidity) G ¼ rCS2 Lame’s constant l ¼ rðCP2 2CS2 Þ Surface wave
Scholte
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ACOUSTICS IN MARINE SEDIMENTS
energy content, and spectral structure and propagation direction changes. In addition, other types of waves, i.e., ducted waves and surface waves, may be generated. These basic types of waves and their characteristics together with their arrival structure as synthetic seismograms for orthogonal directions are illustrated in Figure 4.
Body Waves
These waves propagate within the body of the material, as opposed to surface waves. External forces can distort solids in two different ways. The first involves the compression of the material without changing its shape; the second implies a change in shape without changing its volume (distortion). From earthquake seismology, these compressional and distortion waves are called primus (P) and secoundus (S), respectively, for their arrival sequence on earthquake records. However, the distortional waves are very often called shear.
79
Shear waves Shear waves are those in which the particle motion is perpendicular to the direction of propagation. These waves can be generated at a layer interface by the incidence of compressional waves at other than normal incidence. Shear wave energy is polarized in the vertical or horizontal planes, resulting in vertically polarized shear waves (SV) and horizontally polarized shear waves (SH). However, if the interfaces are close (relative to a wavelength), one cannot distinguish between body and surface waves. Ducted Waves
Love waves Love waves are seismic surface waves associated with layering; they are characterized by horizontal motion perpendicular to the direction of propagation (SH wave). These waves can be considered as ducted shear waves traveling within the duct of the upper sedimentary layer where total reflection occurs at the boundaries; thus the waves represent energy traveling by multiple reflections. Interface Waves
Compressional waves Compressional waves involve compression of the material in such a manner that the particles move in the direction of propagation.
Propagation direction
Basic wave types z A. Body waves
Seismic interface waves travel along or near an interface. The existence of these waves demands the combined action of compressional and shear waves. Thus at least one of the media must be solid, whereas
Particle motion
y
x x
Compressional
Shear
Synthetic seismogram
Sv Sh
B. Ducted waves
Love C. Surface waves
Scholte
Figure 4 Basic seismo-acoustic waves in the vicinity of a water–sediment interface.
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y
z
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ACOUSTICS IN MARINE SEDIMENTS
the other may be a solid (Stoneley wave), a liquid (Scholte wave), or a vacuum (Rayleigh wave). For two homogeneous half-spaces, interface waves are characterized by elliptical particle motion confined to the radial/vertical plane and by a constant velocity that is always smaller than the shear wave. When different types of seismic waves propagate and interact with the layered sediments they are partly converted into each other and their coupling may create mixed wave types in the vicinity of the interface. Figure 4 shows the basic seismo-acoustic waves in the vicinity of a water/sediment interface. Basic characteristics of these waves together with their particle motion and synthetic seismograms at three orthogonal directions (x, y and z) are also illustrated in the same Figure. Under realistic conditions, in which the seafloor cannot be considered to be homogeneous, isotropic, nor a half-space, some of these waves become highly attenuated or travel together, making identification very difficult. In fact, the interface waves shown in Figure 4 are for a layered seafloor, where the dispersion of the signal is evident (homogeneous half-space would not give any dispersion). Under realistic conditions, i.e., for an inhomogeneous, bounded and anisotropic seafloor, some of these waves convert from one to another. The different wave types may travel with different speeds or together, and they generally have different attenuation. As an example, Figure 5 shows signals from an explosive source (0.5 kg trinitrotoluene (TNT))
Hydrophone
a
c
b
Vertical
d
received by a hydrophone and three orthogonal geophones placed on the seafloor, at a distance of 1.5 km from the source. The broadband signal of a few milliseconds duration generated by the explosive source is dispersed over nearly 18 s demonstrating the arrival structure of different types of waves. The characteristics of these waves are indicated in the figure for four different sensors. They can be identified in order of their arrival time as: (a) head wave; (b) water arrival (compressional wave); (c) interface wave; (d) Love waves.
Seafloor Roughness The roughness of the water–sediment interface and layers below is another important parameter that needs to be considered in sediment acoustics. The seafloor contains a wide spectrum of topographic roughness, from features of the order of tens of kilometers, to those of the order of millimeters. The shape of the seafloor and its scattering effect on acoustic signals is be covered here.
Techniques to Measure Geoacoustic Parameters of Marine Sediments The geoacoustic properties of the seafloor defined by the compressional and shear wave velocities, their attenuation, together with the knowledge of the material bulk density, and their variation as a function of depth, are the main parameters needed to solve the acoustic wave equation. To be able to determine these properties of the seabed, different techniques have been developed using samples taken from the seafloor, instruments and divers conducting measurements in situ, and remote techniques measuring seismo-acoustic waves and inverting this information with realistic models into sediment properties. Some of the current methods of obtaining geoacoustic parameters of the marine sediments are briefly described here.
Transversal
Laboratory Measurements on Sediment Core Samples c Radial
0
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Time (s) Figure 5 Signals from an explosive source of 0.5 kg trinitrotoluene received by a hydrophone and three orthogonal geophones at 1.5 km distance (a, Head wave; b, Compressional wave; c, Interface wave; d, Love wave).
Most of our knowledge of the physical properties of sediments is acquired through core sampling. A large number of measurements on marine sediments have been made in the past. In undisturbed sediment core samples, under laboratory conditions, density and compressional velocity can be measured with accuracy, and having measured values of density and compressional velocity, the bulk modulus can be selected as the third parameter, where it is can be calculated (Table 1).
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ACOUSTICS IN MARINE SEDIMENTS
There are several laboratory techniques available to measure some of the sediment properties. However, the reliability of such measurements can be degraded by sample disturbance and temperature and pressure changes. In particular, the acoustic properties are highly affected by the deterioration of the chemical and mechanical bindings caused by the differences in temperature and pressure between the sampling and the laboratory measurements. Controlling the relationships between various physical parameters can be used to check the accuracy of measured parameters of sediment properties. It has been shown that the density and porosity of sediments have a relationship with compressional velocity, and different empirical equations between them have been established. Over the past three decades, at the NATO Undersea Research Centre (SACLANTCEN), a large number of laboratory and in situ measurements have been made of the physical properties of the seafloor sediments. These measurements have been conducted on the samples with the same techniques as when the same laboratory methods were applied. This data set with a great consistency of the hardware and measurement technique, has been used to demonstrate the physical characteristics of the sediment that affect acoustic waves. The relationship between measured physical parameters From 300 available cores, 20 000 measured data samples for the density and porosity, and 10 000 samples for the compressional velocity were obtained. To be able to handle this large data set taken from different oceans at different water depths, all bulk density and compressional velocity data were converted into relative density ðrÞ and relative velocity ðCÞ with respect to the in situ water values:
than used here and shown to have a strong linear correlation. Theoretically this linearity only exists if the dry densities of the mineral particles are the same for all marine sediments. The density of the sediment would then be the same as the density of the solid material at zero porosity, and the same as the density of the water at 100% porosity. Figure 6 shows this relationship. Porosity and relative compressional wave velocity The relationship between porosity and compressional wave velocity has received much attention in the literature because porosity can be measured easily and accurately. Data from the SACLANTCEN sediment cores giving the relationship between porosity and compressional wave velocity are shown in Figure 7. As shown in Figure 6 due to the linear relation between density and porosity, the relationship between density and compressional wave velocity is similar to the porosity compressional wave velocity relation. In situ Techniques
There are several in situ techniques available that use instruments lowered on to the seafloor mainly by means of submersibles, remotely operated vehicles (ROVs), and autonomous underwater vehicles AUVs and divers. The first deep-water, in situ measurements of sediment properties were made from the bathyscaph Trieste in 1962. These measurements provided accurate results due to the minimum disturbance of temperature and pressure changes compared to bringing the sample to the surface and for 2.6
= 2.635_ 0.01675 P Number of data points = 13885
2.4
where, rS is sediment bulk density, rW is water density, CS is sediment compressional velocity and CW is water compressional velocity. The data are not only from the water–sediment interface but cover sedimentary layers of up to 10 m deep. Relative density and porosity The density and porosity of the marine sediment are least affected during coring and laboratory handling of the samples. Porosity is given by the percentage volume of the porous space and sediment bulk density by the weight of the sample per unit volume. The relationship between porosity and bulk density has been investigated by many authors with fewer data
Relative density ( = s /w)
r ¼ rS =rW C ¼ CS =CW
81
2.2 2.0 1.8 1.6 1.4 1.2 1.0 0
20
40 60 Porosity (%)
80
100
Figure 6 Relationship between relative density and porosity.
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Figure 9 Examples of received compressional wave signals from fine sand sediment.
Figure 7 Relationship between relative compressional wave velocity and porosity.
Figure 10 Examples of signals recorded from two shear wave receivers in hard-packed fine sand sediments. Figure 8 In Situ Sediment Acoustic Measurement System (ISSAMS) for near-surface in situ geoacoustic measurements.
analysis in the laboratory. The most reliable direct geoacoustic measurement techniques for marine sediments are in situ techniques. Some of the approaches used in recent years are discussed below. Near-surface method A system has been developed to measure sediment geoacoustic parameters, including compressional and shear wave velocities and their attenuation at tens of centimeters below the sediment–water interface. Figure 8 shows the main features of the In situ Sediment Acoustic Measurement System (ISSAMS). Shear and compressional wave probes are attached to a triangular frame that uses weights to force the probes into the sediment. In very shallow water, divers can be used to insert the probes into the sediment, whereas in deeper water, a sleeve system (not shown) allows the ISSAMS to penetrate into the seafloor.
Using the compressional and shear wave transducers measurements are made with a continuous wave (cw) pulse technique where the ratio of measured transducer separation and pulse arrival time yields the wave velocity. Samples of compressional and shear wave data are shown in Figure 9 and 10 respectively. Table 3 gives a comparison of laboratory and in situ values of compressional and shear wave velocities from two different types of Adriatic Sea sediment. Cross-hole method Measurements as a function of depth in sediments can be made with boreholes, using either single or cross-hole techniques. Boreholes are made by divers using water–air jets to penetrate thin-walled plastic tubes for cross-hole measurements. Figure 11 shows the experimental set up for the cross-hole measurements. The source is in the form of an electromagnetic mallet securely coupled to the inner wall of one of the plastic tubes with a hydraulic clamping device.
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ACOUSTICS IN MARINE SEDIMENTS
Table 3
83
Comparison of laboratory and in situ measurements Wave velocity (ms1)
Sediment type
Porosity (%)
Mean grain size (f)
In situ (CP)
Laboratory (CP)
In situ (CS)
Laboratory (CS)
37 68
3.5 8.6
1557–1568 1467–1488
1580–1604 1468–1487
78–82 27–31
50 15
Sand Mud
Pump
Projector Clamp
Receiver
S
Geophones Clamp
Electromagnetic mallet
Figure 11 The experimental setup for cross-hole measurements.
0
4
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_4
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Relative velocity amplitude (V)
8
0.1 s
_ 20
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Figure 12 Cross-hole shear wave signal (A) and its spectrum (B) for a silty-clay bottom in the Ligurian Sea.
With a separation range of 2.9 m, moving-coil geophone receivers are also coupled to the inner wall of the second plastic tube with a hydraulic clamping device. The electromagnetic mallet generates a point source on the thin plastic tube wall, which results in a multipolarized transient signal to the sediment. Depending on the orientation of receiving sensors, vector components of the propagating compressional and shear waves are received and
analyzed for velocity and attenuation parameters. Examples of a time series and a frequency spectrum for a shear wave signal received by a geophone with a natural frequency of 4.5 Hz are shown in Figure 12. Figure 13 shows shear wave velocity as a function of depth obtained in the Ligurian Sea. It can be seen that the shear wave velocities are around 60 m1 s at the sediment interface and increase with depth.
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Remote Sensing Techniques
Even though in situ techniques provide the most reliable data, they are usually more time consuming and expensive to make and they are limited to small areas. Remote sensing and inversion to obtain geoacoustic parameters can cover larger areas in less time and provide reliable information. These techniques are based on the use of a seismo-acoustic 0
signal received by sensors on the seafloor and/or in the water column. Figure 14 illustrates a characteristic shallow water signal from an explosion received by a hydrophone close to the sea bottom. Three different techniques to extract information relative to bottom parameters from these signals are described briefly below. Reflected waves The half-space seafloor. When the seafloor consists of soft unconsolidated sediments, due to its very low shear modulus it can be treated as fluid.
1
Depth (m)
2 w 3
r
Water
Cw
Sediment
s
t
Cs 4
5 0
75 100 50 _ Shear wave velocity (m s 1)
125
Figure 13 Shear wave velocity profile obtained from cross-hole measurements.
Figure 15 Geometry and notations for a simple half-space water–sediment interface.
Refracted wave
Interface wave
Waterborne wave 0.2
0.4
0.6 Time (ms)
0.8
1.0
Figure 14 A characteristic shallow-water signal.
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ACOUSTICS IN MARINE SEDIMENTS
Reflection may occur whenever an acoustic wave is incident on an interface between two media. The amount of energy reflected, and its phase relative to the incident wave, depend on the ratio between the physical properties on the opposite sides of the interface. It is possible to calculate frequency
independent reflection loss as the ratio between the amplitude of the shock pulse and the peak of the first reflection (after correcting for phase shift, absorption in the water column, and differences in spreading loss). If the relative density r ¼ rS =rW, and the relative compressional wave velocity
0°
0°
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angle Grazing
85
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3 4 5m
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(B)
Figure 16 Acoustic signals and reflection loss as a function of grazing angle and sediment core properties: (A) for critical angle; (B) for angle of incidence case.
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However, the phase of the reflected wave is then shifted relative to the phase of the incidence wave by an angle varying from 01 to 1801 and is given as:
C ¼ CS =CW , are used to present the contrast between the two media, reflection coefficient for such a simple environmental condition can be written as:
p
pffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi rsinY ð1=C2 cos2 YÞ pffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi R¼ rsinY ð1=C2 cos2 YÞ
F ¼ 2 arctan
ðcos2 Y 1=C2 Þ rsinY
Figure 16 shows measured and calculated reflection losses (20 log R) for this simple condition together with the basic physical properties of the core sample taken in the same area for explosive signals at different grazing angles.
Where Y is the grazing angle with respect to the interface as shown in Figure 15. For an incident path normal to a reflecting horizon, i.e., Y ¼ 901, the reflection coefficient is:
Angle of intromission case. Especially in deepwater sediments the sound velocity in the top layer of the bottom is generally less than in the water above (Co1). In such conditions there is an angle of incidence at which all of the incident energy is transmitted into the sedimentary layer and the reflection coefficient becomes zero:
R ¼ rC 1=rC þ 1 Critical angle case. When the velocity of compressional wave velocity is greater in the sediment layer (C>1), as the grazing angle is decreased, a unique value is reached at which the acoustic energy totally reflects back to the water column. This is known as the critical angle and is given by:
cosYi ¼
p r
2
1=C2 r2 1
The phase shift is 01 when the ray angle is greater than the intromission angle and 1801 when it is smaller. Thus, acoustical characteristics of the bottom, such as the critical angle, the angle of incidence, the phase shift,
Ycr ¼ arccosð1=CÞ When the grazing angle is less than this critical angle, all the incident acoustic energy is reflected. Direct signal Reflected signal D(t )
R(t )
F
F
D(f )
R(f ) R(f ) D(f ) H (f ) Transfer func.
F
Impulse response
_1
20 Time (ms)
0 A (f ) e Reflection coefficient A (f )
i(f )
Phase shift (f ) Reflection loss _ 20 log A dB
+180° 0° _ 180° 0 10 20 30 dB 0
1
2
3 kHz
Figure 17 Acoustic reflection data and analysis technique for a layered seafloor.
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4
5
40
Loss dB
Grazing angle
ACOUSTICS IN MARINE SEDIMENTS
Computed
87
Measured
0
36° 10 20 10 54° 20 30 10 72° 20 30
0
2 kHz
4
2 kHz
0
4
Impulse responses Measured
Computed 18° 36° 54° 72° 90° 0
10 0 Time (ms) For layered bottom
10 Time (ms)
Figure 18 Measured and calculated reflection losses for a layered seafloor.
and the reflection coefficient, are primarily influenced by the relative density and relative sound velocity of the environment as a function of the ray angle. The layered seafloor. Since the seafloor is generally layered, a simple peak amplitude approximation cannot be implemented because of the frequency dependence. In this case one calculates the transfer function (or reflection coefficient) from the convolution of the direct reference signal with the reflected signal. Examples of the phase shift and reflection loss as a function of frequency are shown in Figure 17. The reflectivity can also be described in the time domain by the impulse response, which is the inverse Fourier transform of the transfer function as shown in the same figure. This type of data can be utilized for inverse modeling to obtain the unknown parameters of the sediments. Figure 18 is an example of measured and calculated reflection losses and impulse responses for a layered seafloor.
ray
one ar
Geoph
Explosive source
Figure 19 Experimental setup to measure seismo-acoustic waves on the seafloor.
Refracted waves Techniques developed for remote sensing of the uppermost sediments (25–50 m below the seafloor) utilize broad-band sources (small
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0 0
25
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75
100
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250
500
Time (ms)
1000 1500 2000 2500 3000
8.39 8.50 4.87 8.84 7.22 8.94 8.40 10.26 10.00
6.81 6.96 3.64 7.26 7.40 7.54 7.68 7.82 7.95 8.08 8.21 4.24 8.46 8.59 8.71 8.83 8.95 9.07
4000
0.19 0.23 0.06 0.27 0.57 0.68 0.37 0.77 0.47 3.52 3.75 3.76 3.92 5.02 3.09 3.86 3.44 4.88 6.49
3500
Figure 20 Signals received by geophone array.
The slope of the travel-time curve (fitted parabola) gives the rate of change of the range with respect to time that is also the velocity of propagation of the diving compressional wave at the level of its deepest penetration (turning point) into the sediment. At each range D, the depth corresponding to the deepest penetration is then calculated using the following integral zðVÞ ¼ 1=p
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explosives) and an array of geophones deployed on the seafloor. To obtain estimates of the bottom properties as a function of depth, both refracted compressional and shear waves as well as interface waves are analyzed. Inversion of the data is carried out using modified versions of techniques developed by earthquake seismologists to study dispersed Rayleigh-waves and refracted waves. Figure 19 illustrates the basic experimental setup and Figure 20 the signals received by an array of 24 geophones that permit studies of both interface and refracted waves. These data are analyzed and inverted to obtain both compressional and shear wave velocities as a function of depth in the seafloor. Studies of attenuation and lateral variability are also possible using the same data set. Figure 21 shows the expanded early portion of the geophone data shown in Figure 20, where, the first arriving energy out to a range of about 250 m has been fitted with a curve showing that compressional waves refracted through the sediments just beneath the seafloor travel faster as they penetrate more deeply into the sediment. At zero offset, the slope indicates a velocity of 1505 m s1 whereas at a range offset at 250 m the slope corresponds to a velocity of 1573 m s1. At ranges over 250 m, a strong head wave becomes the first arrival and, the interpretation would be that there is an underlying rock layer with compressional wave velocity of about 4577 m s1. A compressional wave velocity–depth curve for the upper part of the seafloor can be derived from the first arrivals shown in Figure 21 using the classical Herglotz–Bateman–Wieckert integration method.
100 150
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Figure 21 Expanded early portion of the data shown in Figure 21.
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ACOUSTICS IN MARINE SEDIMENTS
The result is the solid velocity–depth curve shown in Figure 22.
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Interface waves In order to obtain a shear wave velocity–depth profile from the data, later arrivals corresponding to dispersed interface waves may be utilized (Figure 20). The portion of each individual signal corresponding to the interface-wave arrival can be processed using multiple filter analysis to create a group velocity dispersion diagram (Gabor diagram). The result of applying this technique to a dispersed signal is a filtered time signal whose envelope reaches a maximum at the group velocity arrival time for a selected frequency. The envelope is computed by taking the quadrature components of the inverse Fourier transform of the filtered signal. Filtering is carried out at many discrete frequencies over selected frequency bands. Once the arrival times are converted in to velocity, the envelopes are arranged in a matrix and contoured and dispersion curves are obtained by connecting the maximum values of the contour diagram (Figure 23). Having obtained the dispersion characteristics of the interface waves, the geoacoustic model, made of a stack of homogeneous layers with different compressional and shear wave velocities for each layer that predict the measured dispersion curve is determined. Figure 24 illustrates a number of examples from the Mediterranean sea covering data from soft clays to hard sands.
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Figure 22 Compressional wave velocity versus depth curves derived from data and predicted (dashed line) by the model.
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Frequency (Hz) Figure 23 Dispersion diagram and measured and predicted dispersion curves.
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Velocity (m s 1) Figure 24 Summary of shear wave velocity profiles from the Mediterranean Sea.
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Figure 25 Comparison of transmission loss (TL) data and SAFARI prediction. Effects of changing bottom parameters from sand to clay is also shown together with the input profiles utilized for SAFARI predictions.
Transmission Loss Technique
The seafloor is known to be the controlling factor in low-frequency shallow water acoustic propagation. Forward modeling is performed with models giving exact solutions to the wave equation, i.e., SAFARI where, compressional and shear wave velocities, the attenuation factors associated with these waves, and the sediment density as a function of depth are the main input parameters. Acoustic energy propagating through a shallow water channel interacts with the seafloor causing partitioning of waterborne energy into different types of seismic and acoustic waves. The propagation and attenuation of these waves observed in such an environment are strongly dependent on the physical characteristics of the sea bottom. Transmission loss (TL), representing the
amount of energy lost along an acoustic propagation path, carries the information relative to the environment through which the wave is propagating. Figure 25 shows a comparison of TL data and model predictions together with the input parameters used at 400 Hz. The effects of changes in bottom parameters are also shown in the figure. This technique becomes extremely useful when seafloor information is sparse.
Conclusions Acoustic/seismic characteristics of the marine sediments have been of interest to a wide range of activities covering commercial operations involving trenching and cable lying, construction of offshore
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ACOUSTICS IN MARINE SEDIMENTS
foundations, studies of slope stability, dredging and military applications like mine and submarine detection. Experimental and theoretical work over the years has shown that it is possible to determine the geoacoustic properties of sediments by different techniques. The characteristics of the marine sediments and techniques to obtain information about these characteristics have been briefly described. The studies so far conducted indicate that direct and indirect methods developed over the last four decades may give sufficient information to deduce some of the fundamental characteristics of the marine sediments. It is evident that still more research needs to be done to develop these techniques for fast and reliable results.
See also Acoustics, Deep Ocean. Acoustics, Shallow Water. Benthic Boundary Layer Effects. Calcium Carbonates. Clay Mineralogy. Ocean Margin Sediments. Pore Water Chemistry. Seismic Structure.
Further Reading Akal T and Berkson JM (eds.) (1986) Ocean SeismoAcoustics. New York and London: Plenum press. Biot MA (1962) Generalized theory of acoustic propagation in porous dissipative media. Journal of Acoustical Society of America 34: 1254--1264.
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Brekhovskikh LM (1980) Waves in Layered Media. New York: Academic Press. Cagniard L (1962) Reflection and Refraction of Progressive Seismic Waves. New York: McGraw-Hill. Grant FS and West GF (1965) Interpretation and Theory in Applied Geophysics. New York: McGraw Hill. Hampton L (ed.) (1974) Physics of Sound in Marine Sediments. New York and London: Plenum Press. Hovem JM, Richardson MD, and Stoll RD (eds.) (1992) Shear Waves in Marine Sediments. Dordrecht: Kluwer Academic. Jensen FB, Kuperman WA, Porter MB, and Schmidt H (1999) Computational Ocean Acoustics. New York: Springer-Verlag. Kuperman WA and Jensen FB (eds.) (1980) Bottom Interacting Ocean Acoustics. New York and London: Plenum Press. Lara-Saenz A, Ranz-Guerra C and Carbo-Fite C (eds) (1987) Acoustics and Ocean Bottom. II F.A.S.E. Specialized Conference. Inst. De Acustica, Madrid. Pace NG (ed.) (1983) Acoustics and the Sea-Bed. Bath: Bath University Press. Pouliquen E, Lyons AP, Pace NG et al. (2000) Backscattering from unconsolidated sediments above 100 kHz. In: Chevret P and Zakhario ME (ed.) Proceedings of the fifth European Conference on Under water Acoustics, ECUA 2000 Lyon, France. Stoll RD (1989) Lecture Notes in Earth Sciences. New York: Springer-Verlag.
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ACOUSTICS, ARCTIC P. N. Mikhalevsky, Science Applications International Corporation, McLean, VA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 1, pp 53–61, & 2001, Elsevier Ltd.
Introduction The Arctic Ocean is an isolated mediterranean basin with only limited communication with the world’s oceans, principally the Atlantic Ocean via the Fram Strait and the Barents Sea, and the Pacific Ocean via the Bering Strait. The ubiquitous feature of the Arctic Ocean is the sea ice that covers the entire Arctic basin during the winter months and only retreats off the shallow water shelf areas in the summer months, creating a permanent cap over most of the central Arctic basin (Figure 1). The presence of the yearround sea ice cover determines the unique character of acoustic propagation and ambient noise in the Arctic Ocean. The sea ice nsulates the Arctic Ocean from solar heating in the summer months, creating a yearround upward refracting sound speed profile with the sound speed minimum at the water–ice interface. Sound, therefore, is refracted upward and is continuously reflected from the ice as it propagates, causing attenuation by scattering, mode conversion, and absorption that increases rapidly with frequency. The lack of solar forcing and the Arctic Ocean’s restricted communication with the other oceans of the world creates a very stable acoustic channel with significantly reduced fluctuations of acoustic signals in comparison with the temperate oceans. In contrast to the central basin, acoustic propagation on the Arctic shelves and in the marginal ice zones (MIZs, those areas between the average ice minimum and maximum) (Figure 1), is quite complex and variable owing to the seasonal retreat of the sea ice, river run-off, and bottom interaction (see Acoustics, Shallow Water, Acoustics in Marine Sediments). Over the last half-century Arctic acoustics research and development has largely supported submarine operations. The importance of the Northern Sea Route to the Soviet Union, and the prospect of Soviet nuclear ballistic missile submarines exploiting the unique Arctic acoustic environment to remain undetected provided the need for this research. Since the end of the Cold War and the beginning of concern about ‘global warming’ there has been a new focus for Arctic acoustics on acoustic thermometry
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and acoustic remote sensing (see Tomography). The Arctic Ocean is the world’s ‘air-conditioner’, maintaining the surface heat balance, and it provides fresh water to the world’s oceans, principally in the form of sea ice discharged from the Fram Strait. The latter regulates convective overturning in the Greenland and Norwegian Seas that in turn drives the global thermohaline circulation with significant impact on climate. Monitoring changes in the temperature and stratification of the Arctic Ocean and sea ice thickness using acoustics is an important capability that will improve our understanding of the Arctic Ocean and its role in global climate change.
History The first large program dedicated to acoustics research in the Arctic Ocean was undertaken by the US Navy Underwater Sound Laboratory, as it was then known, in 1958 in connection with the International Geophysical Year. Also in that year the US nuclear submarine Nautilus (SSN 571) made its historic voyage to the North Pole, marking the beginning of regular submarine operations in the Arctic Ocean. From 1958 to 1975 the USA conducted much of its acoustics research from manned ice islands (thick tabular sections of land-fast ice that occasionally break away from the Ellesmere ice shelf) that remained within the polar pack ice (see Cryosphere: Sea Ice). The most famous of these was Fletcher’s Ice Island, also known as T3, which was discovered in June 1950. T3, originally 14.5 km long, 6.4 km wide and 52 m thick, provided an ideal platform for Arctic research including acoustics. The first experiments in 1958 were aimed at investigating the feasibility of using the RAFOS (Ranging and Fixing of Sound) for submarine navigation in Arctic waters. Explosive signals were deployed from T3 and other ice stations while the USS Skate (SSN 578) attempted to receive the signals for acoustic crossfixing. While RAFOS was never operationally deployed, a significant amount was learned about the upward refracting sound speed profile, propagation, and resulting scattering losses, as well as ambient noise and reverberation. The Soviet Union operated at least two manned year-round ice stations in the Arctic continuously from 1937 through 1992. In 1971 they began POLEX (Polar Experiment), which was an intense set of oceanographic, ice, and atmospheric measurements over the entire Arctic Ocean basin that
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Figure 1 Map of the Arctic Ocean showing the average minimum and maximum sea ice extent, as well as the major bathymetric features, and other significant geographic locations.
continued for a decade. From 1971 through 1994 the US Office of Naval Research sponsored or co-sponsored many international acoustic-related research science programs based from seasonal ice camps in the Arctic, staged out of facilities in Alaska, Canada, Greenland, and Svalbard. These included the Arctic Ice Dynamics Experiments (AIDJEX), 1971–76; the FRAM and MIZEX series in the central Arctic basin and marginal ice zones respectively, 1979–87; the Coordinated East Arctic Experiment (CEAREX), 1988–89; the Sea Ice Mechanics Initiative (SIMI), 1993–94; and the Transarctic Acoustic Propagation Experiment (TAP) in 1994. These programs
quantified sea ice properties and ice scattering processes, bottom and surface reverberation, Arctic plate tectonics with seismic reflection and refraction, identified the mechanisms of ice generated ambient noise, discovered the exceptional phase stability of low frequency propagation (B20–30 Hz), and demonstrated the feasibility of basin scale acoustic thermometry in the Arctic Ocean. The US, UK, and Soviet Union (now Russia) have operated submarines in the Arctic Ocean and conducted a significant amount of acoustics research. In particular, the US Navy’s Submarine Ice Exercises (SUBICEXs) were conducted over almost three
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decades by the Arctic Submarine Laboratory (ASL) under the leadership of Dr Waldo Lyon, often in conjunction with seasonal ice camps. Dr Lyon was one of the first to investigate propagation in polar stratified waters, beginning shortly after the end of World War II, working with Canadian researchers. From the Naval Electronics Laboratory (NEL) in San Diego, Dr Lyon led the development of upwardlooking ice profiling sonar, and the forward-looking ice avoidance sonar, that proved critical for safe under-ice submarine operations. SUBICEXs were designed to test all aspects of submarine operations including surfacing through ice, and weapons effectiveness (see Under Ice in Further Reading). In 1993 the US Navy supported a dedicated 60 day science cruise in the Arctic with the USS Pargo (SSN 650). This started the Submarine Science Ice Expeditions (SCICEX) that were conducted each year from 1995 to 1999. These multidisciplinary cruises have confirmed the large changes in the Arctic Ocean thermohaline structure, including the warming of the Arctic Intermediate Water (AIW, see below) that was measured acoustically during the TAP experiment in 1994. Other data and modeling have shown strong evidence of decadal Arctic oscillations, but longer observational time series are needed. A trans-basin acoustic section was started in October 1998 from a source located in the Franz Victoria Strait to a receiving array in the Lincoln Sea. This 3-year US/ Russian collaborative effort is expected to be followed by more permanent long-term monitoring of the Arctic Ocean using acoustics.
Arctic Ocean Sound Speed Structure The structure of the typical sound speed profile in the Arctic consists of components that correspond closely to the primary stratified water masses of the Arctic Ocean (Figure 2). The uppermost layer is Polar Water (PW) defined by temperatures o01C and salinities o34.5 ppt and extends from the surface to depths of 100–200 m. There is often a mixed layer of Polar Water 30–60 m thick just below the ice with a nearly constant temperature near the freezing point and nearly constant salinity at around 33.6 ppt. Below the mixed layer the temperature and salinity increase uniformly, attaining 01C and 34.5 ppt respectively, at the base of the Polar Water layer. Arctic Intermediate Water (AIW) – also known as the Atlantic Layer reflecting its origin (and sometimes referred to as Atlantic Intermediate Water) – is defined by temperatures 401C and salinities 434.5 ppt and extends from the base of the Polar Water to a depth of 1000 m. Temperature increases with depth through the AIW layer to a maximum at approximately 200–500 m depth and then decreases with depth to 01C at about 1000 m. AIW enters the Arctic Ocean via the West Spitzbergen Current from the Fram Strait and via the Barents Sea east of Franz Josef Land (see Ocean Currents: Arctic Basin Circulation). The shallower depths (200 m) of the temperature maximum of the AIW occur in the eastern Arctic and approaches 21C. This temperature maximum deepens to 500 m the western Arctic in the Canada Basin and is approximately 0.41C. Salinity increases with depth in the AIW layer
_
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Figure 2 The Arctic sound speed profile shown on the right was computed from the measured temperature and salinity, also shown, from an ice camp in the eastern Arctic Ocean in April 1994. A ray trace for this sound speed profile is plotted for a source at a depth of 100 m. The mode shapes computed at 20 Hz using this profile are shown on the left. The major Arctic water masses (see text) are indicated.
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from 34.5 ppt to 34.9 ppt at about the same depth of the temperature maximum and remains constant at 34.9 ppt below this depth. Deep Arctic Water (DW) is a relatively homogeneous water mass from a depth of 1000 m to the bottom, with temperatures o01C and a nearly constant salinity of 34.9 ppt. The resulting sound speed profile (Figure 2) for these typical deep Arctic basin conditions has a minimum at the ice–water interface of 1435–1440 m s1 and increases with depth to the bottom. In the nearsurface mixed layer of the PW the sound speed gradient is þ 0.016 s1, due entirely to the increase of pressure with depth, as temperature and salinity are constant. Below the mixed layer the sound speed increases rapidly with the increasing temperature and salinity (temperature having the far dominant effect) to the depth of the maximum temperature in the AIW layer with gradients of þ 0.1 s1 or more, reaching a sound speed between 1455 and 1465 m s1. Below the depth of the AIW temperature maximum the sound speed continues to increase, but with reduced gradients of þ 0.01 s1 or less as the temperature decreases, but the pressure increases and the salinity is constant. Below 1000 m in the DW both temperature and salinity are nearly constant and the sound speed continues to increase with the þ 0.016 s1 gradient to the bottom associated with the increasing pressure. In general the Arctic or polar profile can be characterized (and is often approximated) roughly as a bilinear upward refracting profile with the large positive gradient at the near surface creating a strong near-surface duct and a smaller positive gradient below this ‘knee’ to the bottom. This structure is important to understanding the resulting propagation effects described below. There are both regional and temporal variations of this typical sound speed structure associated with the corresponding variations in the water masses described above. The strong positive gradient associated with the upper AIW layer weakens as the ‘knee’ in the sound speed profile (Figure 2) deepens from 200 m to about 500 m as one moves from the eastern Arctic Ocean near Svalbard (where the profile in Figure 2 was measured) into the western Arctic Ocean and the Beaufort and Chukchi Seas. In the Beaufort and north Chukchi Seas the influx of warmer Bering Sea water can create a sound speed maximum just below the mixed layer of the PW at depths between 50 and 80 m, creating a double duct. Because of the year-round ice cover, seasonal, interannual, and recently identified decadal time scales dominate the temporal variations of the sound speed profile. The seasonal and interannual variability is confined almost entirely to the PW layer. The decadal variability is seen in the PW and AIW layers. Mesoscale and internal wave
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Figure 3 A photograph of the pack ice in the central Arctic Ocean taken at the FRAM II Ice Camp in May 1980. A weathered pressure ridge is shown on the left bounding a relatively flat multi-year ice floe on the right that is typical of the central arctic. The blocks of ice in the foreground are 3–4 m high. (Photo by P. Mikhalevsky.).
dynamics are approximately one order of magnitude less energetic in the Arctic Ocean than in the temperate seas and so have a small influence on the sound speed profile and this results in the exceptional signal stability at low frequencies (discussed below).
Propagation Acoustic propagation in the Arctic is dominated by the repeated reflection of the sound from the sea ice as illustrated in Figure 2. The ray trace shown was computed using the sound speed profile of Figure 2. The morphology of sea ice is quite complicated in general. Figure 3 is a photograph taken in the pack ice at an ice camp in the eastern Arctic. It shows a typical pressure ridge on the left that forms between two ice floes when winds and currents compress the pack. Pressure ridges can extend as much as several tens of meters below the ice, and consist of unconsolidated blocks in newly formed ridges as well as refrozen consolidated structures in older ridges. In general the sea ice in the Arctic consists of relatively smooth floes (the right part of Figure 3) laced with pressure ridges and/or open leads as the pack works between convergent and divergent conditions. Average ice roughness and thickness measurements derived from upward-looking submarine sonars have historically (1958–76) been in the range of 1–5 m rms and 2–4 m respectively.1 Recent analysis of SCICEX data (1993–97), however, has shown that
1 LeSchack LA (1980) Arctic Ocean Sea-ice Statistics Derived from Upward-Looking Sonar Data Recorded during Five Nuclear Submarine Cruises. Technical Report to ONR under Contract N0001476-C-0757/NR 307-374. Maryland, LeSchack Associates, Ltd.
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_ 50
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Figure 4 The transmission loss plotted as a function of range for 20, 50, 100, and 200 Hz. The source and receiver were located at 60 m depth. The ice was modeled as a rough elastic plate including statistical parameters for the upper and lower side of the ice representative of the ice floe and ridge structure typical of the central Arctic.
since the late 1970s the ice has thinned by as much as 40%,2 which would also imply a reduction in ice roughness as well. The thicker ice and correspondingly greater roughness is typically found in the eastern Arctic and just north of the eastern Canadian Archipelago and Greenland, where the transpolar drift pushes the ice. The acoustic energy is reflected and scattered from the rough ice and converted to both shear waves and compressional waves within the ice, resulting in significant frequency-dependent attenuation loss. Because of this, trans-basin propagation in the Arctic is limited to frequencies typically below 30 Hz or wavelengths exceeding about 50 m. Figure 4 shows the propagation loss plotted as a function of range for several frequencies. The curves in Figure 4 were generated using a model that includes a rough elastic plate representing the ice cover, with a two scale roughness spectrum that models pressure ridges and the smoother intervening floes.3 The model includes the statistics for both the underside and surface of the ice since at lower frequencies as the wavelength becomes much greater than the ice thickness the reflection occurs at the ice–air interface. Submarine ice draft statistics from the eastern Arctic over the
2 Rothrock DA, Yu Y and Maykut GA (1999) Thinning of the arctic sea-ice cover. Geophysical Research Letters 26(23): 3469– 3472. 3 Kudryashov VM (1996) Calculation of the acoustic field in an arctic waveguide. Physical Acoustics. 42: 386–389.
Nansen Basin (Figure 1) were used for Figure 4. As Figure 4 shows the Arctic acoustic waveguide is a low pass filter. So much so in fact that in comparison with propagation in the temperate oceans which scatter from surface waves that have a similar roughness spectrum, at a given range, equivalent propagation loss in the Arctic requires transmitting at a frequency as much a factor of ten less. On the left side of Figure 2 the acoustic mode shapes for modes 1–4 at 20 Hz are plotted for the profile shown. The acoustic modes can be thought of as the interference pattern of upward and downward going acoustic rays (Figure 2) whose turning depths (the depth where the rays are horizontal) correspond with the depth of the deepest peak of the mode amplitude. At 20 Hz, mode 1 consists of those paths that are trapped in the strong near-surface duct created by the thermocline in the transition from the PW to AIW. The higher modes correspond to rays with greater launch angles that turn at successively deeper depths in the Arctic Ocean. Rough surface scattering theory tells us that the loss per bounce or reflection increases with grazing angle and frequency (i.e., the steeper rays, or higher modes have greater loss per bounce). However, even though the lower modes (lower grazing angles) have lower per bounce loss, those that are trapped in the upper duct experience many more interactions with the ice, as clearly seen in Figure 2, and consequently their net loss per kilometer is higher. At shorter ranges for a source and receiver in the upper duct (Figure 4) the propagation is dominated by the lower modes (or equivalently the lower grazing angle rays that are trapped), but these get stripped away as the range increases and at longer ranges (>100–200 km) the higher modes (or equivalently the higher grazing angle rays that are not trapped) dominate. Because the source and receiver are shallow (at 60 m depth) for the model run in Figure 4, there is a less efficient excitation of mode 1 at 20 Hz (i.e., destructive interference with the surface reflected energy) than at the higher frequencies. Thus less of the energy is trapped in the near-surface duct. This results in lower energy being received on the shallow receiver at closer ranges and hence the higher initial propagation loss at 20 Hz. As the frequency increases the mode shapes tend to compress toward the surface (because the wavelength is getting smaller). For a shallow source more of the energy is trapped in this upper duct at higher frequencies (see discussion of frequency dispersion and Figure 6 below). Given the higher per bounce losses at the higher frequencies, and the greater number of bounces experienced by these trapped modes, it can be seen why attenuation in the Arctic increases rapidly with frequency. For
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Figure 6 A composite of various measurements of Arctic ambient noise. (1) Central Arctic pack ice Nansen Basin, firstand multi-year floes, 2–3 m thick, with ridging, April 1982. (Reproduced with permission from Dyer, 1984.) (2) Beaufort Sea, summer and fall conditions: (a) highest levels observed, (b) typical cold weather situation following rapid temperature drop and thermal-induced cracking, (c) quietest conditions during warmer stable temperature periods with low winds September– October, 1961 and May–September 1962. (Reproduced with permission from DiNapoli et al., 1978, in Von Winkle, 1984.) (3) Shore-fast ice in the Canadian Archipelago, winter conditions, thermal cracking, February 1963. (Reproduced with permission from Milne AR and Ganton JH (1964) Ambient noise under ArcticSea ice. Journal of the Acoustic Society of America 36(5): 855– 863.) (4) Shore-fast ice in the Canadian Archipelago, spring conditions: (a) noisiest conditions observed during diurnal cooling, (b) quietest conditions observed during diurnal warming, April 1961 (Milne and Ganton 1964). (5) Chukchi Sea, spring conditions, cold stable temperatures, low winds, April 1999 (Mikhalevsky, APLIS Ice Camp).
deeper sources the near-surface duct is not as important and higher grazing angle (higher mode) propagation and loss are dominant at all ranges; however, the higher loss per bounce at higher frequencies still results in more rapid attenuation as the frequency increases. There is a significant body of research on the scattering and reverberation of acoustic energy from the sea ice. The earliest models relied on rough free surface scattering theory with empirical fits to important parameters such as rms ice roughness and average ridge spacing. The free surface scattering models that included ridge-like morphology performed reasonably well at frequencies from 250 Hz to 2000 Hz. At lower frequencies they tended to underpredict the loss. When elastic coupling and scattering into the ice (particularly conversion to shear waves and the attenuation in the ice) are included, better agreement at the low frequencies has been achieved. In particular
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the ice keels and their spacing are important in the conversion of acoustic compressional waves into flexural modes in the ice. Le Page and Schmidt (see Further Reading) have shown that the intensity of acoustic scattering into the flexural modes of a rough ice plate strongly depends on the width of the roughness spatial spectra. In order to achieve consistent agreement with data the propagation loss models have become more complex, demanding additional input data, particularly about the structure of the sea ice and its constitutive properties. Bottom interaction must also be included, for those paths which cross major features like the Lomonosov Ridge, and propagation on the arctic shelves (Figure 1). The Lomonosov Ridge can rise to 1500 m below the surface. Figure 2 shows that modes 5 and higher (rays 411–121) will start to interact with this feature and will begin to get stripped away for longrange propagation. As the frequency decreases below 20 Hz the modes tend to expand away from the surface and finally all modes will begin to interact with the bottom except in the very deepest parts of the Arctic basin. This leads to a lower frequency bound at 5–10 Hz for efficient long-range propagation. However, for high source levels at these frequencies there is significant backscatter reverberation from these bathymetric features that can be detected and mapped. Excellent correlation with known Arctic bathymetry as well as the discovery of uncharted features at basin scale ranges (1000 km) has been achieved using large explosive sources deployed from the ice. Over the shelves the shallow water propagation dictates bottom and surface/ice interaction at all frequencies; however, there are optimal propagation frequencies similar to deep water that depend upon the depth and boundary properties. The stable upward refracting Arctic sound speed profile causes a very predictable modal and frequency dispersion. As can be seen in Figure 2, successively higher modes are propagating at higher sound speeds corresponding to higher group velocities. As a consequence at a given range, mode 1 arrives last, preceded in order by modes 2, 3, 4, etc. This can be observed by transmitting a wideband waveform such as an impulsive or explosive source and plotting the signal intensity as a function of frequency and time. Figure 5 shows the arrival time from the computed group velocities of modes 1–3 for an explosive shot at a range of approximately 580 km in the Beaufort Sea. Although not plotted, this calculation agreed very closely with the measured data. Note also from Figure 5 that as the frequency increases (as described above), the modes compress towards the surface and first mode 1, then mode 2 and finally mode 3 become trapped in the slower near-surface duct, as evidenced
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These frequencies are used for short-range applications including the upward-looking ice profiling sonar, forward-looking ice avoidance sonar, and downward-looking depth sounders and bottommapping sonar. Torpedoes also operate at these frequencies and ice capture can be a problem, but Doppler processing can distinguish moving targets from the stationary ice (see Sonar Systems).
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4
Figure 6 The arrival time of a broadband acoustic pulse plotted as a function of frequency showing the modal dispersion with mode 1 arriving last, preceded by modes 2 and 3 respectively. The absolute travel time can be obtained by adding 396.5 s to the time shown on the vertical axis. This modeled result agreed very closely with the received data. (Adapted with permission from DiNapoli et al., 1978, in Von Winkle, 1984.)
by the longer travel times. This modal dispersion and arrival pattern is a classic Arctic acoustic result. The data upon which Figure 5 is based were taken in 1962; 32 years later in 1994 the TAP experiment measured the same arrival pattern, attesting to the very long-term stable water mass structure in the Arctic Ocean. Observations of the changes in the arrival times of these modes are directly related to the temperature changes of these water masses (since sound speed increases approximately 5 m/s per 11C). In particular, at 20 Hz mode 2 is most sensitive to temperature changes in the AIW, as can be seen from Figure 2. A decrease in the travel time of mode 2 by 2 s was the indicator of the warming in the AIW observed in the 1994 TAP experiment. In addition to the long-term stability of the Arctic sound speed profile, the short-term stability results in exceptional amplitude (B1 dB rms) and phase stability (B0.01 cycles rms) of acoustic signals at low frequencies even over ranges approaching 3000 km for up to 1 h (the longest continuous observations made to date). This applies to essentially fixed terminals. This was first observed in 1980 at 15–30 Hz and 300 km and then again in 1994 and 1999 at 20 Hz and 2700 km. This implies that coherent integration of up to 1 h with optimal gain is achievable (and has been demonstrated). At very high frequencies (more than tens of kHz) reflection occurs at the ice–water interface. Attenuation is rapid, due not only to the severe scattering from the ice, but to volumetric absorption as well.
The ambient noise in the Arctic is highly variable, exhibiting some of the quietest as well as the noisiest ocean noise conditions of all the world’s oceans. A composite of various measurements of Arctic ambient noise is shown in Figure 6. In Figure 6, Knudsen Sea State Zero refers to the ambient noise level in the temperate oceans at sea state zero (the quietest conditions) for comparison (see Acoustic Noise). Icegenerated ambient noise is the dominant mechanism contributing to the general character of ambient noise in the Arctic Ocean from a few tenths of Hz up to 10 000 Hz. Episodic noise is also present in the form of seismic events, such as earthquakes along the Mid Arctic Ridge (Figure 1), biologics (mostly marine mammals in the marginal ice zones), and man-made noise from ice breakers, and seismic exploration. Ice noise is generated when the sea ice deforms, fractures, and breaks in response to environmental forcing such as wind, current, thermal, and internal and surface wave-induced stresses. In general, during stable or warming temperature conditions with low winds the quietest conditions obtain, particularly under shore-fast ice. During periods of rapid cooling ice-fracturing events resulting from thermal-induced tensile stresses can lead to higher noise levels. Higher noise levels also occur in the pack ice when the ice is in motion due to nonthermal forcing such as high winds causing bumping, grinding and rubbing of the ice flows. In the marginal ice zone where the ice concentrations are typically less than in the central Arctic pack, ice concentration and surface gravity wave-induced flexural floe failure are the primary correlates with the ambient noise. The actual noisegenerating mechanics are complicated but fall into at least two general categories. In the low frequency range (5–100 Hz) the important noise-generating mechanism is the unloading motion of the ice immediately following breaking. This has been shown to have a dipole radiation characteristic. The breaking process itself is important at intermediate frequencies (100–2000 Hz) and has an octopole radiation characteristic likely resulting from a slip– dislocation process.
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ACOUSTICS, ARCTIC
Noise from earthquakes and other seismic events occur with some regularity within the Arctic Ocean, concentrated along the Mid Arctic Ridge (Figure 1). These events have been recorded on acoustic arrays suspended from the ice within approximately 300 km of the Mid Arctic Ridge. The received frequencies are typically 20 Hz and below. Earthborne acoustic pressure and shear waves emanate from the source and couple through the ocean floor above the epicenter to compressional waterborne waves that propagate vertically and are reflected from the ice canopy into the Arctic sound channel. Most of the energy from these events arrives via this path, but there are weaker precursors associated with crustal propagating compressional and shear waves that couple to the water directly below the receive array. These arrivals are easily identified by their vertical arrival angle at the array as well as their arrival time. Biological noise in the Arctic is concentrated in the marginal ice zones and near the edge of the pack ice. The bowhead is the most numerous of the baleen whales in Arctic waters, migrating along the coast of Alaska in the Chukchi and Beaufort Seas in the spring and then exiting the Arctic waters in the fall. They can vocalize from 25 to 3500 Hz, but their dominant frequencies are between 100 and 400 Hz. Of the toothed whales the beluga and narwhal are the most common, with calls ranging from a few hundred Hz to as high as 20 kHz. Many species of pinnipeds, including hair seals such as the bearded, hooded, harp, ringed, and ribbon seal, and the walrus frequent the marginal ice zones. As a group their calls typically range from a few hundred Hz to 10 kHz. Unlike the temperate oceans where shipping typically dominates the ambient noise spectrum, manmade noise in the Arctic is a small contributor except for specific events that are isolated in time and space. Such events include icebreakers, seismic exploration, and some military and experimental activities. Icebreaker noise peaks in the 50–100 Hz range, but with broadband contributions up to 1000 Hz. Seismic exploration occurs during the summer and fall seasons when the ice extent is minimum (Figure 1), allowing easier access to the shelves. Most of the activity has been confined to the Beaufort Sea off the North Slope of Alaska near Prudhoe Bay.
Conclusions The presence of the year-round ice cap creates the upward refracting sound speed profile in the Arctic Ocean with the sound speed minimum at the ice–water interface. Reflection, scattering, mode
99
conversion, and absorption by the rough elastic sea ice cover causes high attenuation as frequency increases, limiting long-range propagation to very low frequencies. Bathymetric effects are important near the major ocean ridges, basin margins, and on the shelf areas where significant mode coupling can occur. The Arctic sound channel is very stable and predictable in the central Arctic basins and there is a close correspondence of propagating acoustic modes with the major water masses of the Arctic Ocean, especially the important AIW. This latter fact makes the use of acoustic thermometry for monitoring longterm Arctic Ocean temperature change particularly suitable. Ongoing research is exploring ways to relate changes in acoustic travel time and intensity to monitor other important variables in the Arctic Ocean including changes in the PW, DW, and the halocline, and average sea-ice thickness and roughness. The latter measurements, when combined with sea ice extent from satellite remote sensing, could provide an estimate of sea ice mass in the Arctic. The role of the Arctic Ocean in shaping and responding to global climate change is only beginning to be explored. Cost-effective, long-term, year-round synoptic observations in the Arctic Ocean require new measurement strategies. The year-round ice cover in the Arctic prevents the use of satellites for direct ocean observations common in the ice-free oceans. Shore-cabled mooring-based observations using advanced biogeochemical sensors and acoustic sources and hydrophone arrays, as well as instrumented autonomous underwater vehicles (AUVs) and under-ice drifters, represent new approaches for observing the Arctic Ocean. Interestingly the RAFOS concept (using nonexplosive sources) is being evaluated anew as a way to track AUVs and drifters in the Arctic as well as for acoustic communication of data. It is clear that Arctic acoustics will have as large a role to play in this important new endeavor in the future, as it has had in the submarine and military operations of the past.
See also Acoustics, Arctic. Acoustics in Marine Sediments. Acoustic Noise. Acoustics, Shallow Water. Platforms: Autonomous Underwater Vehicles. Bioacoustics. Ice–ocean interaction. Nepheloid Layers. North Atlantic Oscillation (NAO). 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.
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Further Reading Dyer I (1984) Song of sea ice and other Arctic Ocean melodies. In: Dyer I and Chryssostomidis C (eds.) Arctic Technology and Policy, pp. 11--37. Washington, DC: Hemisphere Publishing. Dyer I (1993) Source mechanisms of Arctic Ocean ambient noise. In: Kerman BR (ed.) Natural Physical Sources of Underwater Sound, pp. 537--551. Netherlands: Kluwer Academic. Leary WM (1999) Under Ice. College Station: Texas A&M University Press. LePage K and Schmidt H (1994) Modeling of lowfrequency transmission loss in the central Arctic. Journal of the Acoustic Society of America 96(3): 1783--1795. Mikhalevsky PN, Gavrilov AN, and Baggeroer AB (1999) The transarctic acoustic propagation experiment and
climate monitoring in the Arctic. IEEE Journal of Oceanic Engineering 24(2): 183--201. Newton JL (1989) Sound speed structure of the Arctic Ocean including some effects on acoustic propagation. US Navy Journal of Underwater Acoustics 39(4): 363--384. Richardson WJ, Greene CR Jr, Malme CI, and Thomson DH (1995) Marine Mammals and Noise. San Diego: Academic Press. Urick RJ (1975) Principles of Underwater Sound. New York: McGraw-Hill. Von Winkle WA (1984) Naval Underwater Systems Center (NUSC) Scientific and Engineering Studies: Underwater Acoustics in the Arctic. New London: Naval Underwater Systems Center Publisher.
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ACOUSTICS, DEEP OCEAN W. A. Kuperman, Scripps Institution of Oceanography, University of California, San Diego, CA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 1, pp 61–72, & 2001, Elsevier Ltd.
Introduction The acoustic properties of the ocean, such as the paths along which sound from a localized source will travel, are mainly dependent on its sound speed structure. The sound speed structure is dependent on the oceanographic environment described by variations in temperature, salinity, and density with depth or horizontal position. This article will review the ocean acoustic environment, sound propagation, ambient noise, scattering and reverberation, and the passive and active sonar equation.
Ocean Acoustic Environment Sound propagation in the ocean is governed by the spatial structure of the sound speed and the sound speed in the ocean is a function of temperature, salinity, and ambient pressure. Since the ambient pressure is a function of depth, it is customary to express the sound speed (c) in meters per second as an empirical function of temperature (T) in degrees celsius, salinity (S) in parts per thousand and depth (z) in meters, e.g. eqn [1]. c ¼ 14449:2 þ 4:6T 0:055T þ 0:00029T þ ð1:34 0:01T ÞðS 35Þ þ 0:016z
Units The decibel (dB) denotes a ratio of intensities (see Section 3.3) expressed in terms of a logarithmic (base _1
Sound speed (m s ) 1440 1460 1480 1500 1520 Warmer surface water profile
Surface duct profile
3
Mixed layer profile
½1
There exist more accurate formulas, if needed. Figure 1 shows a typical set of sound speed profiles, indicating greatest variability near the surface. In a warmer season (or warmer part of the day, sometimes referred to as the ‘afternoon effect’), the temperature increases near the surface and hence the sound speed increases toward the sea surface. In nonpolar regions where mixing near the surface due to wind and wave activity is important, a mixed layer of almost constant temperature is often created. In this isothermal layer, sound speed increases with depth because of the increasing ambient pressure, the last term in eqn [1]. This is the surface duct region. Below the mixed layer is the thermocline where the temperature and hence the sound speed decrease
1000
Water depth (m)
2
with depth. Below the thermocline, the temperature is constant and the sound speed increases because of increasing ambient pressure. Therefore, between the deep isothermal region and the mixed layer, there is a minimum sound speed; the depth at which this minimum takes place is referred to as the axis of the deep sound channel. However, in polar regions, the water is coldest near the surface so that the minimum sound speed is at the surface. Figure 2 is a contour display of the sound speed structure of the North and South Atlantic with the deep sound channel axis indicated by the heavy dashed line. Note that the deep sound channel becomes shallower toward the poles. Aside from sound speed effects, the ocean volume is absorbtive and will cause attenuation that increases with acoustic frequency. The ocean surface and bottom also have a strong influence on sound propagation. The ocean surface, though a perfect reflector when flat, causes scattering when its roughness becomes comparable in size with the acoustic wavelength. The ocean bottom, depending on its local structure will scatter and also attenuate the acoustic field.
Polar region profile
Thermocline Deep sound channel axis
2000
Deep isothermal layer
3000
4000
Figure 1 Generic sounds speed profiles.
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0
3000
1485 1495 1505
REYKJANES RIDGE
MID ATLANTIC RIDGE
4000
1515 1525
60˚N
50˚N
40˚N
30˚N
MID ATLANTIC RIDGE
CAPE VERDE BASIN
CANARY BASIN
5000 6000 70˚N
1475
1500
ANTARCTICA
2000
1510
GREENLAND
Depth (m)
1000
20˚N
10˚N
BRAZIL BASIN
0˚
10˚S
RIO GRANDE RISE
1535
20˚S
30˚S
ARGENTINE BASIN
40˚S
SCOTIA RIDGE EAST SCOTIA WEDDELL RIDGE SEA
50˚S
60˚S
70˚S
80˚S
Latitude Figure 2 Sound speed contours of the North and South Atlantic along 30.501W. Dashed line indicates axis of deep sound channel. (From Northrop and Colborn (1974).)
30
Depth (m)
Transmission loss (dB)
Ocean surface
0
40 50
Su
rfa
Source
Direct 100
ce
-re
fle
cte
dp
ath
path
Receiver range
60 ~r
70 ~r
_4
_2
80 90 100 0
1
2
3
4
5
6
7
8
9
10
Range (km) Figure 3 The insert shows the geometry of the Lloyd mirror effect. The plots show a comparison of Lloyd mirror to spherical spreading. Transmission losses are plotted in decibels corresponding to losses of 10 log r2 and 10 log r4, respectively, as explained in the test.
10) scale. The ratio of two intensities, I1/I2 is 10 log (I1/I2) in dB units. Absolute intensities are expressed using an accepted reference intensity of a plane wave having an rms pressure equal to 105 dyn cm2 or, equivalently, 1 mPa. Transmission loss is a decibel measure of relative intensity, the latter being proportional to the square of the acoustic amplitude.
Sound Propagation Very Short-range Propagation
The pressure amplitude from a point source in free space falls off with range r as r1; this geometric loss is called spherical spreading. Most sources of interest in the deep ocean are nearer the surface than the bottom. Hence, the two main short-range paths are the direct path and the surface-reflected path. When
these two paths interfere, they produce a spatial distribution of sound often referred to as a ‘Lloyd mirror pattern’ as shown in the insert of Figure 3. Basic Long-range Propagation Paths
Figure 4 is a schematic of propagation paths in the ocean resulting from the sound speed profiles (indicated by the dashed line) described above in Figure 1. These paths can be understood from Snell’s law eqn [2], which relates the ray angle y (z), with respect to the horizontal, to the local sound speed c(z) at depth z. cosyðzÞ ¼ constant cðzÞ
½2
The equation requires that the higher the sound speed, the smaller the angle with the horizontal,
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ACOUSTICS, DEEP OCEAN
Ocean basin
Arctic
Continental Continental shelf margin
Ice
F B C D
A E
A: Arctic B: Surface duct C: Deep sound channel
Ocean bottom
D: Convergence zone E: Bottom bounce F: Shallow water
Figure 4 Schematic representation of sound propagations paths in the ocean.
meaning that sound bends away from regions of high sound speed or, put another way, sound bends toward regions of low sound speed. Therefore, paths A, B, and C are the simplest to explain since they are paths that oscillate about the local sound speed minima. For example, path C depicted by a ray leaving a source near the deep sound channel axis at a small horizontal angle propagates in the deep sound channel. This path, in temperate latitudes where the sound speed minimum is far from the surface, permits propagation over distances of thousands of kilometers. Path D, which is at slightly steeper angles and is usually excited by a near surface source, is convergence zone propagation, a spatially periodic (35–65 km) refocusing phenomenon producing zones of high intensity near the surface due to the upward refracting nature of the deep sound speed profile. Regions between these zones are referred to as shadow regions. Referring back to Figure 1, there may be a depth in the deep isothermal layer at which the sound speed is the same as it is at the surface; this depth is called the critical depth and is the lower limit of the deep sound channel. A positive critical depth specifies that the environment supports longdistance propagation without bottom interaction, whereas a negative critical depth specifies that the ocean bottom is the lower boundary of the deep sound channel. The bottom bounce path E is also a periodic phenomenon but with a shorter cycle distance and shorter propagation distance because of losses when sound is reflected from the ocean bottom. An alternative way of describing paths is by denoting that they are composed of combinations of refraction (R), surface reflection (SR) and bottom reflection (BR) processes. Thus, Figure 4 also represents some of these paths. Note from Snell’s law that refractive paths involve a path turning around at the highest speed in its duct of confinement. Because such ray paths spend much of their propagation in
103
regions of high sound speed, larger launch angle paths with longer path lengths arrive earlier than shorter paths launched at shallower angles. This is just the opposite of boundary-limited propagation such a bottom bounce (or shallow water) in which a reflection occurs before refraction in a high-speed region can take place. Hence, for deep water refractive paths, those paths that penetrate to a deeper depth have a greater group speed (the horizontal speed of energy propagation) than those paths that do not go as deep, with the axial, most horizontal path along the deep sound channel axis having the slowest group speed. For pulse propagation, this axial arrival is the last arrival. Geometric Spreading Loss
The energy per unit time emitted by a sound source flows through a larger area with increasing range. Intensity is the power flux through a unit area, which translates to the energy flow per unit time through a unit area. Hence, the simplest example of geometric loss is spherical spreading for a point source in free space, for which the area increases as 4pr2, where r is the range from the point source. Thus spherical spreading results in an intensity decay proportional to r2. Since intensity is proportional to the square of the pressure amplitude, the fluctuations in pressure p induced by the sound decay as r1. For range-independent ducted propagation, that is, where rays are refracted or reflected back toward the horizontal direction (which is the case for most long-range propagation), there is no loss associated with the vertical dimension. In this case, the spreading surface is the area of a cylinder whose axis is in the vertical direction passing through the source: 2prH where H is the depth of the duct and is constant. Geometric loss in the near-field Lloyd mirror regime requires consideration of interfering beams from direct and surface reflected paths. To summarize, the geometric spreading laws for the pressure field (recall that intensity is proportional to the square of the pressure) are
• • •
Spherical spreading loss: ppr1 Cylindrical spreading loss: ppr1/2 Lloyd mirror loss: ppr2.
Volume Attenuation
Volume attenuation increases with frequency. In Figure 3, the losses associated with path C include only volume attenuation and scattering because this path does not involve boundary interactions. The volume scattering can be biological in origin or can arise from interaction with large internal wave activity in the vicinity of the upper part of the deep
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ACOUSTICS, DEEP OCEAN
sound channel where paths are refracted before they would interact with the surface. Both of these effects are small for low frequencies. This same internal wave region is also on the lower boundary of the surface duct, allowing scattering out of the surface duct and thereby also constituting a loss mechanism for the surface duct. This mechanism also leaks sound into the deep sound channel, a region that without scattering would be a shadow zone for a surface duct source. This type of scattering from internal waves is also thought to be a major source of fluctuation of the sound field. Attenuation is characterized by an exponential decay of the sound field. If A0 is the rms amplitude of the sound field at unit distance from the source, then the attenuation of the sound field causes the amplitude to decay with distance r along the path according to eqn [3], where the unit of a is nepers/ distance (nepers is a unitless quantity). A ¼ A0 expðarÞ
½3
This attenuation coefficient can be expressed in decibels per unit distance by the conversion a0 ¼ 0.686a. The frequency dependence of attenuation can be roughly divided into four regimes as displayed in Figure 5. In region I, leakage out of the sound channel is believed to be the main cause of attenuation. The main mechanisms associated with regions II and III are boric acid and magnesium sulfate chemical relaxation. Region IV is dominated by the shear and bulk viscosity associated with fresh
1000
Region I Leakage
Attenuation, (dB/km)
100
A' B'
Region II Chemical relaxation B (OH)3
30:1
0:11f 2 43f 2 þ a0 dBkm1 ¼ 3:3 103 þ 1 þ f 2 4100 þ f 2 þ 2:98 104 f 2
½4
Bottom Loss
The structure of the ocean bottom affects those acoustic paths that interact with the ocean bottom. This bottom interaction is summarized by bottom reflectivity, the amplitude ratio of reflected and incident plane waves at the ocean-bottom interface as a function of grazing angle, y (see Figure 6A). For a simple bottom that can be represented by a semiinfinite half-space with constant sound speed cb and density rb, the reflectivity is given by eqn [5] with the subscript w denoting water RðyÞ ¼
rb kwz rw kbz ; rb kwz þ rw kbz
½5
The wavenumbers are given by eqn [6]. kiz ¼ ðo=ci Þ sin yi ¼ k sin yi ;
i ¼ w; b
½6
The incident and transmitted grazing angles are related by Snell’s law according to eqn [7] and the incident grazing angle yw is also equal to the angle of the reflected plane wave. cb cos yw ¼ cw cos yb
½7
For this simple water-bottom interface for which we take cb >cw, there exists a critical grazing angle yc below which there is perfect reflection (eqn [8]).
3:1
10
water. A summary of the approximate frequency dependence (f in kHz) of attenuation (in units of dB km1) is given in eqn [4] with the terms sequentially associated with regions I–IV in Figure 5.
A B
cosyc ¼
cw cb
½8
1 _1
10
10:1
_2
10
_3
?
?
10
Region III MgSO4 relaxation
Region IV Shear & volume viscosity
_4
10
10
100 1000 10 K 100K
1M
10 M
Frequency (Hz) Figure 5 Regions of different dominant processes of attenuation of sound in sea water. (From Urick (1979).) The attenuation a is given in dB per 1000 yards.
For a lossy bottom, there is no perfect reflection, as also indicated in a typical reflection curve in Figure 6B. These results are approximately frequency independent. However, for a layered bottom, the reflectivity has a complicated frequency dependence as shown in the example in Figure 6C, where the contours are in decibels. This example shows the simple reflectivity result below 200 Hz and then a more complicated frequency dependence at higher frequencies. It should be pointed out that if the density of the second medium vanishes, the reflectivity reduces to the pressure release case of R(y) ¼ 1.
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ACOUSTICS, DEEP OCEAN
exp [i (k⊥ .r + k1z z)]
R exp [i (k⊥. r + k1z z)]
z 1c1 2c 2
r
i
R T
(A)
T exp [i (k⊥. r + k 2z z)]
Nonlossy bottom Lossy bottom Phase
0
0 c
0
90
Grazing angle (°)
5
_2
10
_4
_ 1
_3
_ 4_ 5
_ 3
15 20 0.05
_ 4
_43 _
_2
Grazing angle (°)
0
_ 4
(B)
_ 3
Phase (°)
Reflectivity R ()
1
_ _ 32
0.1
(C)
0.2
_3
0.4
_4
0.8
1.6
3.2
Frequency (kHz)
Figure 6 The reflection and transmission processes. Grazing angles are defined relative to the horizontal. (A) A plane wave is incident on an interface separating two media with densities and sound speeds rc. R(y) and T (y) are reflection and transmission coefficients. Snell’s law is a statement that k>, the horizontal component of the wavevector, is the same for all three waves. (B) Rayleigh reflection curve (eqn [5]) as a function of the grazing angle (y in (A)) indicating critical yc. The dashed curve shows that if the second medium is lossy, there is less than perfect reflection below the critical angle. (C) Examples of contour of reflection loss (20 log R) for a layered bottom, showing frequency and grazing angle dependence. The simpler reflectivity curve for each frequency is obtained from a vertical slice.
105
acoustically relevant parameters of sound speed, density, and attenuation. Sound propagation in the ocean is mathematically described by the wave equation whose coefficients and boundary conditions are derived from the ocean environment. There are essentially four types of models (computer solutions to the wave equation) to describe sound propagation in the sea: ray, spectral or fast field program (FFP), normal mode (NM), and the parabolic equation (PE). Ray theory is an asymptotic high-frequency approximation to the wave equation, whereas the latter three models are more or less direct solutions to the wave equations under an assortment of milder restrictions. The highfrequency limit does not include diffraction phenomena. All of these models can handle depth variation of the ocean acoustic environment. A model that also takes into account horizontal variations in the environment (i.e., sloping bottom or spatially variable oceanography) is termed rangedependent. For high frequencies (a few kHz or above), ray theory is the most practical. The other three model types are more applicable and usable at lower frequencies (below 1 kHz). The hierarchy of underwater acoustic models is shown in schematic form in Figure 7. The output of these models is typically propagation loss, which is the intensity relative to a unit source at unit distance, expressed in decibels. Transmission loss is the negative of propagation loss, and hence, a positive quantity. An example of the output of propagation models is shown in Figure 8, indicating agreement between the models. However, we also see a difference among the models in that ray theory predicts a sharper shadow zone than the wave theory model (i.e., the 10–30 km region in Figure 8B); this is an expected result from the infinite-frequency ray approximation.
Scattering and Reverberation Propagation Models
An ocean acoustic environment is often very complex, with range- and depth-dependent properties. Such an environment does not in general lend itself to simple analytic predictions of sound propagation. Even in range-independent environments there are many paths (multipaths) and these paths combine to form a complex interference pattern. For example, the convergence zones are an example of a more complex structure that cannot be described by a monotonic geometric spreading law. Acoustic models play an important role in predicting sound propagation; the inputs to these models are oceanographic quantities ultimately translated into the
Scattering caused by rough boundaries or volume inhomogeneities is a mechanism for loss (attenuation), reverberant interference, and fluctuation. Attenuation from volume scattering was addressed above. In most cases, it is the mean or coherent (or specular) part of the acoustic field that is of interest for a sonar or communications application, and scattering causes part of the acoustic field to be randomized. Rough surface scattering out of the ‘specular direction’ can be thought of as an attenuation of the mean acoustic field and typically increases with increasing frequency. A formula often used to describe reflectivity from a rough boundary is eqn [9], where R(y) is the reflection coefficient of the
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ACOUSTICS, DEEP OCEAN
Nonlinear wave equation
Linear wave equation
TDFFP FFP
NPE
TDPE
Frequency-domain wave equation
Normal modes
WKBJ approximation
Coupled modes
Adiabatic modes
Asymptotic
Rangeindependent
Ray theory
Rangedependent
PE
Asymptotic
Spectral Figure 7 Heirarchy of underwater acoustic models. TD refers to time domain. NPE is the nonlinear parabolic equation that describes high-amplitude (e.g., shockwave) propagation. The arrows are directed toward the flow of derivation of the model.
G2 2
½9
The scattered field is often referred to as reverberation. Surface, bottom, or volume scattering strength, SS,B,V is a simple parametrization of the production of reverberation and is defined as the ratio in decibels of the sound scattered by a unit surface area or volume referenced to a unit distance, (Insert equqtion), to the incident plane wave intensity, Iinc (eqn [10]). SS;B;V
Iscat ¼ 10 log Iinc
½10
The Chapman–Harris curves predicts the ocean surface scattering strength in the 400–6400 Hz region; eqn [11], where y is the grazing angle in degrees, w the wind speed in m s1 and f is the frequency in Hz.
Depth (m)
R0 ðyÞ ¼ RðyÞexp
1520 _ 1510 1530 m s 1
y 42:4 log b þ 2:6; 30 0:58 b ¼ 107 wf 1=3
½11
A more elaborate formula exists that is more accurate for lower wind speeds and lower frequencies. The simple characterization of bottom backscattering strength utilizes Lambert’s rule for diffuse scattering, given by eqn [12] where the first term is determined empirically. SB ¼ A þ 10 log sin2 y
½12
20
10
30
Range (km)
(A) 60 70 80
60 70 80 90 100
PE (250 Hz) Ray (2000 Hz) Experiment
110 120 0
(B)
SS ¼ 3:3b log
Source = 25 m, Receiver = 38 m
0 200 400 600 800 1000 0
Propagation loss (dB)
smooth interface and G is the Rayleigh roughness parameter defined as GR2ks sin y where k ¼ 2p/l, l is the acoustic wavelength, and s is the rms roughness (height).
10
20
30
Range (km)
Figure 8 Model and data comparison for a range-dependent deep water case. (A) Sound speed profiles as a function of range together with a ray trace showing the breakdown of surface duct propagation. (B) Parabolic equation comparison with data at 250 Hz and ray theory comparison with data at 2000 Hz.
Under the assumption that all incident energy is scattered into the water column with no transmission in to the bottom, A is 5 dB. Typical realistic values for A that have been measured are 17 dB for big Basalt Mid-Atlantic Ridge cliffs and 27 dB for sediment ponds.
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ACOUSTICS, DEEP OCEAN
Volume scattering strength is typically reduced to a surface scattering strength by taking (Insert equqtion) as an average volume scattering strength within some layer at a particular depth; then the corresponding surface scattering strength is given by eqn [13], where H is the layer thickness. SS ¼ Sv þ 10 log H
½13
The column or integrated scattering strength is defined as the case for which H is the total water depth. Volume scattering usually decreases with depth (about 5 dB per 300 m) with the exception of the deep scattering layer. For frequencies less than 10 kHz, fish with air-filled swimbladders are the main scatterers. Above 20 kHz, zooplankton or smaller animals that feed upon phytoplankton and the associated biological chain are the scatterers. The deep scattering layer (DSL) is deeper in the day than in the night, changing most rapidly during sunset and sunrise. This layer produces a strong scattering increase of 5–15 dB within 100 m of the surface at night and virtually no scattering in the daytime at the surface since it migrates down to hundreds of meters. Since higher pressure compresses the fish swimbladder, the backscattering acoustic resonance tend to be at a higher frequency during the day when the DSL migrates to greater depths. Examples of day and night scattering strengths are shown in Figure 9.
107
Finally, near-surface bubbles and bubble clouds can be thought of as either volume or surface scattering mechanisms acting in concert with the rough surface. Bubbles have resonances (typically greater than 10 kHz) and at these resonances, scattering is strongly enhanced. Bubble clouds have collective properties; among these properties is that a bubbly mixture, as specified by its void fraction (total bubble gas volume divided by water volume) has a considerable lower sound speed than water.
Ambient Noise There are essentially two types of ocean acoustic noise: man-made and natural. Generally, shipping is the most important source of man-made noise, though noise from offshore oil rigs is becoming more and more prevalent. Typically, natural noise dominates at low frequencies (below 10 Hz) and high frequencies (above a few hundred hertz). Shipping fills in the region between ten and a few hundred hertz. A summary of the spectrum of noise is shown in Figure 10. The higher-frequency noise is usually parametrized according to sea state (also Beaufort number) and/or wind. Table 1 summarizes the description of sea state. The sound speed profile affects the vertical and angular distribution of noise in the deep ocean. When there is a positive critical depth, sound from
≤100 m
≤ 850 m
_ 50
≤ 780 m
≤ 340 m
≤ 580 m
d
10 log ∫0 MZ d Z
_ 45
_ 55
≤ 140 m
_ 60
_ 65
Day
2
5
10
Night
15 2
5
10
15
Frequency (kHz) (A)
(B)
Figure 9 Day and night scattering strength measurements using an explosive source as a function of frequency. The spectra measured at various times after the explosion are labeled with the depth of the nearest scatterer that could have contributed to the reverberation. The ordinate corresponds to Sv in eqn [13]. (From Chapman and Marchall (1966))
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108
ACOUSTICS, DEEP OCEAN
Intermittent and local effects Earthquakes and explosions
146
Biologics Precipitation Ships, industrial activity Sea ice
KEY 126
Limits of prevailing noise Wind-dependent bubble and spray noise Low-frequency very-shallow-water wind dependence
Sound pressure spectrum level (dB re 1 Pa)
Heavy precipitation Heavy traffic noise Usual traffic noise _ deep water
106
Usual traffic noise _ shallow water Thermal noise General pattern of noise from earthquakes and explosions Extrapolations
86
Wind force (Beaufort)
66 8 5
46 3 2
Prevailing noises Turbulent-pressure fluctuations
26
1 Oceanic traffic Bubbles and spray (surface agitation)
Surface waves _ second-order pressure effects (seismic background)
Molecular Agitation
6 1
10
10
2
10
3
4
10
5
10
Frequency (Hz) Figure 10 Composite of ambient noise spectra. (From Wenz (1962).)
surface sources can travel long distances without interacting with the ocean bottom, but a receiver below this critical depth should sense less surface noise because propagation involves interaction with lossy boundaries, surface and/or bottom. This is illustrated in Figure 11, which shows a deep water environment with measured ambient noise. Figure 12 is an example of vertical directivity of noise that also follows the propagation physics discussed above. The shallower depth is at the axis of the deep sound channel, while the other is at the critical depth. The pattern is narrower at the critical depth where the sound paths tend to be horizontal since the rays are turning around at the lower boundary of the deep sound channel.
In a range-independent ocean, Snell’s law predicts a horizontal noise notch at depths where the speed of sound is less than the near-surface sound speed. Returning to eqn [2] and reading off the sound speeds from Figure 11 at the surface (c ¼ 1530 m s1) and say, 300 m (1500 m s1), a horizontal ray (y ¼ 0) launched from ocean surface would have an angle with respect to the horizontal of about 111 at 300 m depth. All other rays would arrive with greater vertical angles. Hence we expect this horizontal notch. However, the horizontal notch is often not seen at shipping noise frequencies. This is because shipping tends to be concentrated in continental shelf regions; range-dependent propagation couples such noise sources to the deep ocean. Thus, for example,
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ACOUSTICS, DEEP OCEAN
Table 1
109
Descriptions of the ocean sea surface. Approximate relation between scales of wind speed, wave height, and sea state
Sea criteria
Wind speed
12-h wind Fully risen sea
Beaufort Range scale (m s1)
Mirrorlike Ripples Small wavelets Large wavelets, scattered whitecaps Small waves, frequent whitecaps Moderate waves, many whitecaps Large waves, whitecaps everywhere, spray Heaped-up sea, blown spray, streaks Moderately high, longwaves, spindrift
0 1 2 3 4 5 6 7 8
o0.5 0.5–1.7 1.8–3.3 3.4–5.4 5.5–8.4 8.5–11.1 11.2–14.1 14.2–17.2 17.3–20.8
Mean Wave (m s1) height a, (m)
1.1 2.5 4.4 6.9 9.8 12.6 15.7 19.0
b
Wave height a, (m)
b
Durationb, (h)
c
Fetchb, (km)
c
Seastate scale
0 1/2 1 o19 2 19–74 3 74–185 4 185–370 5 370–740 6 740–1300 7
o0.30 o0.30 0.30–0.61 0.30–0.61 o2.5 0.61–1.5 0.61–1.8 2.5–6.5 1.5–2.4 1.8–3.0 6.5–11 2.4–3.7 3.0–5.2 11–18 3.7–5.2 5.2–7.9 18–29 5.2–7.3 7.9–11.9 29–42
a
The average height of the highest one-third of the waves (significant wave height). Estimated from data given in US Hydrographic Office (Washington, DC) publications HO 604 (1951) and HO 603 (1955). c The minimum fetch and duration of the wind needed to generate a fully risen sea. From Wenz (1962). b
50 Hz
25 Hz
100 Hz
150 Hz
1000
250 Hz
500 Hz
18 Hz
0
Depth (m)
2000
3000
4000
CRITICAL DEPTH = 4420 m
5000 BOTTOM 5322 m BOTTOM 5322 m
6000 1480
1500
1520
1540
50
_1
Sound velocity (m s ) (B)
(A)
60
70
80
90
Spectrum level (dB relative to 1 μPa)
Figure 11 (A) Sound speed profile and (B) noise level as a function of depth in the Pacific. (From Morris (1978))
propagation down a continental slope converts highangle rays to lower angles at each bounce. There are also deep sound channel shoaling effects that result in the same trend in angle conversion.
methodology of the sonar equation is analogous to an accounting procedure involving acoustic signal, interference, and system characteristics. It is instructive, beyond the specific application to conventional sonars, to understand this accounting methodology and below is a simplified summary.
Sonar Equation A major application of underwater acoustics is sonar technology. The performance of a sonar is often described simply in terms of the sonar equation. The
Passive Sonar Equation
A passive sonar system uses the radiated sound from a target to detect and locate the target. A radiating
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110
ACOUSTICS, DEEP OCEAN
object of source level SL (all units are in decibels) is received at a hydrophone of a sonar system at a lower signal level S because of the transmission loss ‘TL’ it suffers (e.g., cylindrical spreading plus attenuation or a TL computed from one of the propagation models) (eqn [14]). S ¼ SL TL
½14
The noise, N, at a single hydrophone is subtracted from eqn [14] to obtain the signal-to-noise ratio (SNR) at a single hydrophone (eqn [15]). SNR ¼ SL TL N
½15
Typically a sonar system consists of an array or antenna of hydrophones that provides signal-to-noise enhancement through a beamforming process; this process is quantified in decibels by array gain AG that is therefore added to the single hydrophone SNR to give the SNR at the output of the beamformer (eqn [16]). SNRBF ¼ SL TL N þ AG
½16
Because detection involves additional factors including sonar operator ability, it is necessary to specify a detection threshold (DT) level above the SNRBFat which there is a 50% (by convention) probability of detection. The difference between these two quantities is called signal excess (SE) (eqn [17]). SE ¼ SL TL N þ AG DT
½17
This decibel bookkeeping leads to an important sonar engineering descriptor called the figure of merit, FOM, which is the transmission loss that gives a zero signal excess (eqn [18]). FOM ¼ SL N þ AG DT
½18
Spectrum level (dB / µPa / steradian)
The FOM encompasses the various parameters a sonar engineer must deal with: expected source level, the noise environment, array gain and the detection
threshold. Conversely, since the FOM is a transmission loss, one can use the output of a propagation model (or, if appropriate, a simple geometric loss plus attenuation) to estimate the minimum range at which a 50% probability of detection can be expected. This range changes with oceanographic conditions and is often referred to as the ‘range of the day’ in navy sonar applications. Active Sonar Equation
A monostatic active sonar transmits a pulse to a target and its echo is detected at a receiver co-located with the transmitter. A bistatic active sonar has the receiver in a different location from the transmitter. The main differences between the passive and active cases is that the source level is replaced by a target strength, TS; reverberation and hence reverberation level, RL, is usually the dominant source of interference as opposed noise; and the transmission loss is over two paths: transmitter to target and target to receiver. In the monostatic case the transmission loss is 2TL, where TL is the one-way transmission loss; and in the bistatic case the transmission loss is the sum (in dB) over paths from the transmitter to the target and the target to the receiver, TL1 þ TL2. The concept of the detection threshold is useful for both passive and active sonars. Hence, for signal excess, we have eqn [19]. S ¼ SL TL1 þ TS TL2 ðRL þ N Þ þ AG DT
½19
The corresponding FOM for an active system is defined for the maximum allowable two-way transmission loss with TS ¼ 0 dB.
See also Acoustic Noise. Acoustic Scattering by Marine Organisms. Acoustics, Arctic. Acoustics, Shallow Water. Acoustics in Marine Sediments. Bioacoustics. Bubbles. Deep-Sea Drilling Methodology. Seismic Structure. Seismology Sensors. Sonar Systems. Tomography.
80 713 m 3781 m 70 _ 20 DOWN
_ 10
0
10
Elevation angle (°)
20 UP
Figure 12 The vertical directionality of noise at the axis of the deep sound channel and at the critical depth in the Pacific. (From Anderson (1979))
Further Reading Anderson VC (1979) Variations of the vertical directivity of noise with depth in the North Pacific. Journal of the Acoustical Society of America 66: 1446--1452. Brekhovskikh LM and Lysanov YP (1991) Fundamentals of Ocean Acoustics. Berlin: Springer-Verlag. Chapman RP and Harris HH (1962) Surface backscattering strengths measured with explosive sound sources.
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ACOUSTICS, DEEP OCEAN
Journal of the Acoustical Society of America 34: 1592-1597. Chapman RP and Marchall JR (1966) Reverberation from deep scattering layers in the Western North Atlantic. Journal of the Acoustical Society of America 40: 405--411. Collins MD and Siegmann WL (2001) Parabolic Wave Equations with Applications. New York: Springer-AIP. Dushaw BD, Worcester PF, Cornuelle BD, and Howe BM (1993) On equations for the speed of sound in seawater. Journal of the Acoustical Society of America 93: 255--275. Jensen FB, Kuperman WA, Porter MB, and Schmidt H (1994) Computational Ocean Acoustics. Woodbury: AIP Press. Keller and Papadakis JS (eds.) (1977) Wave Propagation in Underwater Acoustics. New York: Springer-Verlag. Makris NC, Chia SC, and Fialkowski LT (1999) The biazimuthal scattering distribution of an abyssal hill. Journal of the Acoustical Society of America 106: 2491--2512. Medwin H and Clay CS (1997) Fundamentals of Acoustical Oceanography. Boston: Academic Press. Morris GB (1978) Depth dependence of ambient noise in the Northeastern Pacific Ocean. Journal of the Acoustical Society of America 64: 581--590.
111
Munk W, Worcester P, and Wunsch C (1995) Acoustic Tomography. Cambridge: Cambridge University Press. Nicholas M, Ogden PM, and Erskine FT (1998) Improved empirical descriptions for acoustic surface backscatter in the ocean. IEEE-JOE 23: 81--95. Northrup J and Colborn JG (1974) Sofar channel axial sound speed and depth in the Atlantic Ocean. Journal of Geophysical Research 79: 5633--5641. Ross D (1976) Mechanics of Underwater Noise. New York: Pergamon. Ogilvy JA (1987) Wave scattering from rough surfaces. Reports on Progress in Physics 50: 1553--1608. Urick RJ (1979) Sound Propagation in the Sea. Washington, DC: US GPO. Urick RJ (1983) Principles of Underwater Sound. New York: McGraw Hill. Spiesberger JL and Metzger K (1991) A new algorithm for sound speed in seawater. Journal of the Acoustical Society of America 89: 2677--2688. Wenz GM (1962) Acoustics ambient noise in the ocean: spectra and sources. Journal of the Acoustical Society of America 34: 1936--1956.
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ACOUSTICS, SHALLOW WATER F. B. Jensen, SACLANT Undersea Research Centre, La Spezia, Italy Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 1, pp 89–96, & 2001, Elsevier Ltd.
away from the bottom and therefore can propagate to long ranges with little attenuation. Moreover, the environmental variability is much higher in coastal regions than in the deep ocean, with the result that there is much more acoustic variability in shallow water than in deep water.
Introduction
The Ocean Acoustic Environment
Using the commonly accepted definition of shallow water to mean coastal waters with depth up to 200 m, the shallow-water regions of the world constitute around 8% of all oceans and seas. These regions are particularly important since they are national economic zones and also more assessible. Sound waves in the sea play the role of light in the atmosphere, i.e. acoustics is the only means of ‘seeing’ objects at distances beyond a few hundred meters in seawater. All forms of electromagnetic waves (light, radar) are rapidly attenuated in seawater. Lowfrequency acoustic signals, on the other hand, propagate with little attenuation and can be heard over thousands of kilometers in the deep ocean. The use of sound in the sea is ubiquitous. It is employed by the military to detect mines and submarines, and ship-mounted sonars measure water depth, ship speed and the presence of fish shoals. Side-scan systems are used to map bottom topography, sub-bottom profilers for getting information about the deeper layering, and other sonar systems for locating pipelines and cables on and beneath the seafloor. Sound is also used for navigating submerged vehicles, for underwater communications and for tracking marine mammals. In an inverse sense sound is used for measuring physical parameters of the ocean environment and for monitoring oceanic processes through the techniques of acoustical oceanography and ocean acoustic tomography. Optimal sonar design for this great variety of applications demands using a wide range of acoustic frequencies. Practical shallow-water systems cover a frequency range from 50 Hz to 500 kHz, which, with a mean sound speed of 1500 m s1, correspond to acoustic wavelengths from 30 m down to 3 mm. The principal characteristic of shallow-water propagation is that the sound-speed profile is nearly constant over depth or downward refracting, meaning that long-range propagation takes place exclusively via lossy bottom-interacting paths. This is very different from deep-water scenarios, where the sound-speed structure is such that sound is refracted
The ocean is an acoustic waveguide limited above by the sea surface and below by the seafloor. The speed of sound in the waveguide plays the same role as the index of refraction does in optics. Sound speed is normally related to density and compressibility. In the ocean, density is related to static pressure, salinity and temperature. The sound speed in the ocean is an increasing function of temperature, salinity, and pressure, the latter being a function of depth. It is customary to express sound speed (c) as an empirical function of three independent variables: temperature (T) in degrees centigrade, salinity (S) in parts per thousand (%), and depth (D) in meters. A simplified expression for this dependence is
112
c ¼ 1449:2 þ 4:6T 0:055T 2 þ 0:00029T 3 ½1 þð1:34 0:010T ÞðS 35Þ þ 0:016D In shallow water, where the depth effect on sound speed is small, the primary contributor to sound speed variations is the temperature. Thus, for a salinity of 35%, the sound speed in seawater varies between 1450 m s1 at 01C and 1545 m s1 at 301C. Seasonal and diurnal changes affect the oceanographic parameters in the upper ocean. In addition, all of these parameters are a function of geography. In a warmer season (or warmer part of the day) in shallow seas where tidal mixing is weak, the temperature increases near the surface and hence the sound speed increases toward the sea surface. This near-surface heating (and subsequent cooling) has a profound effect on surface-ship sonars. Thus the diurnal heating causes poorer sonar performance in the afternoon – a phenomenon known as the afternoon effect. The seasonal variability, however, is much greater and therefore more important acoustically. A ray picture of propagation in a 100-m deep shallow water duct is shown in Figure 1. The soundspeed profile in the upper panel is typical of the Mediterranean in the summer. There is a warm surface layer causing downward refraction and hence repeated bottom interaction for all ray paths. Since
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ACOUSTICS, SHALLOW WATER
SV
SEA SURFACE
Depth
0 2.5˚ slope
Depth (m)
100
(A)
Range SV
113
200
Depth
300
400 0
Range
(B)
1
2
3
Range (km)
Figure 1 Ray paths in shallow water for typical Mediterranean summer and winter profiles. (A) In summer sound interacts repeatedly with the seabed but not with the sea surface. (B) In winter sound interacts with both the sea surface and the seabed, except for shallow rays emitted near the horizontal.
the seafloor is a lossy boundary, propagation in shallow water is dominated by bottom reflection loss at low and intermediate frequencies (o1 kHz) and scattering losses at high frequencies. The seasonal variation in sound-speed structure is significant with winter conditions being nearly iso-speed (Figure 1B). The result is that there is less bottom interaction in winter than in summer, which again means that propagation conditions are generally better in winter than in summer. Of course, the ocean sound-speed structure is neither frozen in time nor space. On the contrary, the ocean has its own weather system. There are currents, internal waves and thermal microstructure present in most shallow-water areas. Figure 2 illustrates the sound speed variability along a 15 km-long track in the Mediterranean Sea. The data were recorded on a towed thermistor chain covering depths between 5 and 90 m. In general, this type of timevarying oceanographic structure has an effect on sound propagation, both as a source of attenuation (acoustic energy being scattered into steeperangle SV
Depth
Summer
Figure 3 Seismic profile of bottom layering in coastal-water area of the Mediterranean.
propagation paths suffers increased bottom reflection loss) and of acoustic signal fluctuations with time. Turning to the upper and lower boundaries of the ocean waveguide, the sea surface is a simple horizontal boundary and a nearly perfect reflector. The seafloor, on the other hand, is a lossy boundary with varying topography. Both boundaries have smallscale roughness associated with them which causes scattering and hence attenuation of sound due to the increased bottom reflection loss associated with steep-angle propagation paths. In terms of propagation physics, the seafloor is definitely the most complex boundary, exhibiting vastly different reflectivity characteristics in different geographical locations. The structure of the ocean bottom in shallow water generally consists of a thin stratification of sediments overlying the continental crust. The nature of the stratification is dependent on many factors, including geological age and local geological activity. Thus, relatively recent sediments will be characterized by plane stratification parallel to the sea bed, whereas older sediments and sediments close to the crustal plate boundaries may have undergone significant deformation. An example of a complicated bottom layering is given in Figure 3, which displays a seismic section from the coastal Mediterranean. The upper stratification here is almost parallel to the seafloor, whereas deeper layers are strongly inclined.
Transmission Loss Range
Figure 2 Spatial variability of sound speed in shallow-water area of the Mediterranean. The depth covered is around 100 m and the range 15 km.
The decibel (dB) is the dominant unit in ocean acoustics and denotes a ratio of intensities (not pressures) expressed on a log10 scale. An acoustic signal traveling through the ocean becomes distorted due to multipath effects and
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ACOUSTICS, SHALLOW WATER
weakened due to various loss mechanisms. The standard measure in underwater acoustics of the change in signal strength with range is transmission loss defined as the ratio in decibels between the acoustic intensity I(r,z) at a field point and the intensity I0 at 1 m distance from the source, i.e. Iðr; zÞ I0 jpðr; zÞj ½dB re 1 m: ¼ 20 log j p0 j
TL ¼ 10 log
½2
Here use has been made of the fact that the intensity of a plane wave is proportional to the square of the pressure amplitude. The major contributors to transmission loss in shallow water are: geometrical spreading loss, water volume attenuation, bottom reflection loss, and various scattering losses.
The cylindrical spreading loss is therefore given by TL ¼ 10 log r
½dB re 1 m
½4
Note that for a point source in a waveguide, there is spherical spreading in the nearfield (rrD) followed by a transition region toward cylindrical spreading which applies only at longer ranges (rbD). As an example consider propagation in a shallowwater waveguide to a range of 20 km with spherical spreading applying on the first 100 m. The total propagation loss (neglecting attenuation) then becomes: 40 dB þ 23 dB ¼ 63 dB. This figure represents the minimum loss to be expected at 20 km. In practice, the total loss will be higher due both to the attenuation of sound in seawater, and to various reflection and scattering losses. Sound Attenuation in Seawater
Geometrical Spreading
The spreading loss is simply a measure of the signal weakening as it propagates outward from the source. Figure 4 shows the two geometries of importance in underwater acoustics. First consider a point source in an unbounded homogeneous medium (Figure 4A). For this simple case the power radiated by the source is equally distributed over the surface area of a sphere surrounding the source. If the medium is assumed to be lossless, the intensity is inversely to proportional the surface of the sphere, i.e. Ip1= 4pR2 . Then from eqn [2] the spherical spreading loss is given by TL ¼ 20 log r
½dB re 1 m
½3
where r is the horizontal range in meters. When the medium has plane upper and lower boundaries as in the waveguide case in Figure 4B, the farfield intensity change with horizontal range becomes inversely proportional to the surface of a cylinder of radius R and depth D, i.e. Ip1=ð2pRDÞ.
*
R
D
*
R
When sound propagates in the ocean, part of the acoustic energy is continuously absorbed, i.e. the energy is transformed into heat. Moreover, sound is scattered by different kinds of inhomogeneities, also resulting in a decay of sound intensity with range. As a rule, it is not possible in real ocean experiments to distinguish between absorption and scattering effects; they both contribute to sound attenuation in seawater. A simplified expression for the frequency dependence (f in kHz) of the attenuation is given by Thorp’s formula, a¼
0:11f 2 44f 2 þ ½dB km1 ; 1 þ f 2 4100 þ f 2
½5
where the two terms describe absorption due to chemical relaxations of boric acid, B(OH)3, and magnesium sulphate, MgSO4, respectively. According to eqn [5] the attenuation of low-frequency sound in seawater is indeed very small. For instance, at 100 Hz a tenfold reduction in sound intensity ( 10 dB) occurs over a distance of around 8300 km. Even though attenuation increases with frequency (r10 dB C150 km at 1 kHz and C9 km at 10 kHz), no other kind of radiation can compete with sound waves for long-range propagation in the ocean. Bottom Reflection Loss
I∝ (A)
1 I∝ 2πRD
1 4πR 2 (B)
Figure 4 Geometrical spreading laws. (A) Spherical spreading; (B) Cylindrical spreading.
Reflectivity, the ratio of the amplitudes of a reflected plane wave to a plane wave incident on an interface separating two media, is an important measure of the effect of the bottom on sound propagation. Ocean bottom sediments are often modeled as fluids which
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ACOUSTICS, SHALLOW WATER
rffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi ffi ðc1 =c2 Þ2 cos2 y1 rffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi Rðy1 Þ ¼ ffi 2 2 ðr2 =r1 Þsin y1 þ ðc1 =c2 Þ cos y1 ðr2 =r1 Þsin y1
½6
where y1 denotes the grazing angle of the incident plane wave of unit amplitude. The reflection coefficient has unit magnitude, meaning perfect reflection, when the numerator and denominator of eqn [6] are complex conjugates. This can only occur when the square root is purely imaginary, i.e. for cos y14c1/c2 (total internal reflection). The associated critical grazing angle below which there is perfect reflection is found to be c1 yc ¼ arccos c2
½7
Note that a critical angle only exists when the sound speed of the second medium is higher than that of the first. A closer look at eqn [6] shows that the reflection coefficient for lossless media is real for yc4yc, which means that there is loss ðjRj ¼ 1Þ but no phase shift associated with the reflection process. On the other hand, for ycoyc we have perfect reflection ðjRj ¼ 1Þ but with an angle-dependent phase shift. In the general case of lossy media (ci complex), the reflection coefficient is complex, and, consequently, there is both a loss and a phase shift associated with each reflection. The critical-angle concept is very important for understanding the waveguide nature of shallowwater propagation. Figure 5 shows bottom loss curves ðBL ¼ 10 log jRj2 Þ for a few simple fluid bottoms with different compressional wave speeds (Cp), densities and attenuations. Note that for a lossy bottom we never get perfect reflection. However, there is in all cases an apparent critical angle (yCC331 for cp ¼ 1800 m s1 in Figure 5), below which the reflection loss is much smaller than for supercritical incidence. With paths involving many bottom bounces such as in shallow-water propagation, bottom losses even as small as a few tenths of a decibel per bounce accumulate to significant total losses since the propagation path may involve many tens or even hundreds of bounces.
10
Bottom loss (dB)
means that they support only one type of sound wave – a compressional wave. The expression for reflectivity at an interface separating two homogeneous fluid media with density ri and sound speed ci, i ¼ 1, 2, was first worked out by Rayleigh as
115
Cp = 1550 m s−1 1600 1800
5
θc
0 0
60 30 Grazing angle (°)
90
Figure 5 Bottom reflection loss curves for different bottom types. Note that low-speed bottoms (clay, silt) are more lossy than high-speed bottoms (sand, gravel). Cp is the compressional wave speed.
Real ocean bottoms are complex layered structures of spatially varying material composition. A geoacoustic model is defined as a model of the real seafloor with emphasis on measured, extrapolated, and predicted values of those material properties important for the modeling of sound transmission. In general, a geoacoustic model details the true thicknesses and properties of sediment and rock layers within the seabed to a depth termed the effective acoustic penetration depth. Thus, at high frequencies (41 kHz), details of the bottom composition are required only in the upper few meters of sediment, whereas at low frequencies (o100 Hz) information must be provided on the whole sediment column and on properties of the underlying rocks. The information required for a complete geoacoustic model should include the following depthdependent material properties: the compressional wave speed, Cp; the shear wave speed, Cp; the compressional wave attenuation, Cp; the shear wave attenuation, Cp; and the density, Cp. Moreover, information on the variation of all of these parameters with geographical position is required. The amount of literature dealing with acoustic properties of seafloor materials is vast. Table 1 lists the geoacoustic properties of some typical seafloor materials, as an indication of the many different types of materials encountered just in continental shelf and slope environments. Boundary and Volume Scattering Losses
Scattering is a mechanism for loss, interference and fluctuation. A rough sea surface or seafloor causes
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ACOUSTICS, SHALLOW WATER
Table 1
Geoacoustic properties of continental shelf environments
Bottom type
p (%)
rb/rw –
Cp/Cw –
Cp (m s1)
Cs (m s1)
ap (aBl1 p )
as (aBl1 s )
Clay Silt Sand Gravel Moraine Chalk Limestone Basalt
70 55 45 35 25 – – –
1.5 1.7 1.9 2.0 2.1 2.2 2.4 2.7
1.00 1.05 1.1 1.2 1.3 1.6 2.0 3.5
1500 1575 1650 1800 1950 2400 3000 5250
o100 Csa Csb Csc 600 1000 1500 2500
0.2 1.0 0.8 0.6 0.4 0.2 0.1 0.1
1.0 1.5 2.5 1.5 1.0 0.5 0.2 0.2
a
Cs ¼ 80 z˜0.3 cs ¼ 110 z˜0.3 c cs ¼ 180 z˜0.3 Cw ¼ 1500 ms 1, rw ¼ 1000 kg m 3.
b
attenuation of the mean acoustic field propagating in the ocean waveguide. The attenuation increases with increasing frequency. The field scattered away from the specular direction, and, in particular, the backscattered field (called reverberation) acts as interference for active sonar systems. Because the ocean surface moves, it will also generate acoustic fluctuations. Bottom roughness can also generate fluctuations when the sound source or receiver is moving. The importance of boundary roughness depends on the sound-speed profile which determines the degree of interaction of sound with the rough boundaries. Often the effect of scattering from a rough surface is thought of as simply an additional loss to the specularly reflected (coherent) component resulting from the scattering of energy away from the specular direction. If the ocean bottom or surface can be modeled as a randomly rough surface, and if the roughness is small with respect to the acoustic wavelength, the reflection loss can be considered to be modified in a simple fashion by the scattering process. A formula often used to describe reflectivity from a rough boundary is: R0 ðyÞ ¼ RðyÞe0:5G
2
½8
where R0 (y) is the new reflection coefficient, reduced because of scattering at the randomly rough interface. G is the Rayleigh roughness parameter defined as G ¼ 2ks sin y
½9
where k ¼ 2p/l is the acoustic wavenumber and s is the rms roughness. Note that the reflection coefficient for the smooth ocean surface is simply 1 (the pressure-release condition is obtained from eqn [6] by setting r2 ¼ 0) so that the rough-sea-surface reflection coefficient for the coherent field is R0 ðyÞ ¼ exp 0:5G2 . For the ocean bottom, the
appropriate geoacoustic parameters (see Table 1) are used for evaluating R0 (y), and the rough-bottom reflection coefficient is then obtained from eqn [8]. Volume scattering is thought to arise primarily from biological organisms. For lower frequencies (less than 10 kHz), fish with air-filled swim bladders are the main scatterers whereas above 20 kHz, zooplankton or smaller animals that feed on the phytoplankton, and the associated biological food chain, are the scatterers. Many of the organisms undergo a diurnal migration rising towards the sea surface at sunset and descending to depth at sunrise. Since the composition and density of the populations vary with the environmental conditions, the scattering characteristics depend on geographical location, time of day, season and frequency. As an example, data from the Mediterranean Sea for volume scattering losses due to fish shoals show excess losses of 10–15 dB for a propagation range of 12 km and frequencies between 1 and 3 kHz. Finally, scattering off bubbles near the surface is sometimes referred to as either a volume- or surfacescattering mechanism. These bubbles arise not only from sea surface action, but also from biological origins and from ship wakes. Furthermore, bubbles are not the only scattering mechanism, but bubble clouds may have significantly different sound speed than plain seawater thereby altering local refraction conditions. At the sea surface, the relative importance of roughness versus bubble effects is not yet resolved.
Transmission-loss Data Figure 6 gives an example of transmission-loss variability in shallow water. The graph displays a collection of experimental data from different shallow-water areas (100–200 m deep) all over the
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ACOUSTICS, SHALLOW WATER
40
SD = 50.0 m RD = 50.0 m
3200 1600 Frequency (Hz)
50 60 70
800 400 200 100 50
80
25 (A)
90
SD = 50.0 m RD = 59.0 m 3200
100
1600 Frequency (Hz)
Transmission loss (dB)
117
110 120
800 400 200 100 50
130
25
140 1
2
5
10 20 Range (km)
50
100
Figure 6 Transmission loss variability in shallow water.
world. The data refer to downward-refracting summer conditions in the frequency band 0.5–1.5 kHz. Two features are of immediate interest. One is the spread of the data amounting to around 50 dB at 100 km and caused primarily by the varying bottomloss conditions in different areas of the world. The second feature is the fact that transmission is generally better than free-field propagation (20 log r) at short and intermediate ranges but worse at longer ranges. This peculiarity is due to the trapping of energy in the shallow-water duct, which improves transmission at shorter ranges (cylindrical versus spherical spreading), but, at the same time, causes increased boundary interaction, which degrades transmission at longer ranges. A second example of transmission-loss variability in shallow water is given in Figure 7, where broadband data from two different geographical areas are compared. The data set in Figure 7A, was collected in the Barents Sea in 60 m water depth. Note the high transmission losses recorded below 200 Hz, where energy levels fall off rapidly indicating that most of the acoustic energy emitted by the source is lost to the seabed. It is believed that this excess attenuation is caused by the coupling of acoustic energy into shear waves in the seabed. In contrast to the highloss environment in the Barents Sea, Figure 7B shows a data set from the English Channel in 90 m water depth. Here propagation conditions are excellent
10
0
(B)
20 Range (km)
30
40
Figure 7 Examples of frequency-dependent propagation losses measured in two shallow-water areas: (A) Barents Sea, (B) English Channel. Note the presence of an optimum frequency of propagation between 200 and 400 Hz.
over the entire frequency band. This second data set represents typical propagation conditions for thick sandy sediments with negligible shear-wave effects. Cutoff Frequency and Optimum Frequency
A common feature of all acoustic ducts is the existence of a low-frequency cutoff. Hence, there is a critical frequency below which the shallow-water channel ceases to act as a waveguide, causing energy radiated by the source to propagate directly into the bottom. The cutoff frequency is given by, f0 ¼
cw rffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi 4D 1 ðcw =cb Þ2
½10
This expression is exact only for a homogeneous water column of depth D and sound speed cw overlying a homogeneous bottom of sound speed cb. As an example, let us take D ¼ 100 m, cw ¼ 1500 m s1, and cb ¼ 1600 m s1 (sand–silt), which yields f0C11 Hz. Sound transmission in shallow water has the characteristic frequency-dependent behavior shown in Figure 7, i.e. there is an optimum frequency of propagation at longer ranges. Thus the 80 dB contour line extends farthest in range for frequencies around 400 Hz in Figure 7A and around 200 Hz in
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118
ACOUSTICS, SHALLOW WATER
Figure 7B, implying that transmission is best at these frequencies – the optimum frequencies of propagation for the two sites. Optimum frequency is a general feature of ducted propagation in the ocean. It occurs as a result of competing propagation and attenuation mechanisms at high and low frequencies. In the high-frequency regime we have increasing volume and scattering loss with increasing frequency. At lower frequencies the efficiency of the duct to confine sound decreases (the cutoff phenomenon). Hence propagation and attenuation mechanisms outside the duct (in the seabed) become important. In fact, the increased penetration of sound into a lossy seabed with decreasing frequency causes the overall attenuation of waterborne sound to increase with decreasing frequency. Thus we get high attenuation at both high and low frequencies, whereas intermediate frequencies have the lowest attenuation. It can be shown that the optimum frequency for shallow-water propagation is strongly dependent on water depth (foptp D1), has some dependence on the sound-speed profile, but is only weakly dependent on the bottom type. Typically, the optimum frequency is in the range 200–800 Hz for a water depth of 100 m.
differences, and individual pulse shapes being modified due to frequency-dependent amplitude and phase changes associated with each boundary reflection. From simple geometrical considerations, the time dispersion is found to be Dt C
R 1 1 c¯ cosy
½11
where R is the range between source and receiver, c¯ is the mean sound speed in the channel, and y is the maximum propagation angle with respect to the horizontal. This angle will be determined either by the source beamwidth or by the critical angle at the bottom (the smaller of the two). Since the dispersion considered here is solely due to the geometry of the waveguide, it is called geometrical dispersion. An example of measured pulse arrivals over a 10 h period in the Mediterranean is given in Figure 9. Note that the time-varying ocean (internal waves, currents, tides) causes strong signal fluctuations with time, particularly in the earlier part of the signal. The time dispersion is 15–20 ms and at least four main energy packets, each consisting of several ray arrivals can be identified.
Signal Transmission in the Time Domain
1
S
2 3 4
10 9 8
Geo-time (h)
Even though underwater acousticians have traditionally favored spectral analysis techniques for gaining information about the band-averaged energy distribution within a shallow-water waveguide, additional insight into the complication of multipath propagation can be obtained by looking at signal transmission in the time domain. Figure 8 indicates that the signal structure measured downrange will consist of a number of arrivals with time delays determined by the pathlength
7 6 5
4
D′ D
3
R 2 BOTTOM
1
Figure 8 Schematic of ray arrivals in shallow-water waveguide. A series of multipath arrivals are expected, represented by eigenrays connecting source (S) and receiver (R). The shortest path is the direct arrival D followed by the surface-reflected arrival D0 . Next comes a series of four arrivals all with a single bottom bounce; then four rays with two bottom bounces as illustrated in the figure, followed by four rays with three bottom bounces, etc.
0 0
10
20 Signal time (ms)
30
40
Figure 9 Measured pulse arrivals versus geo-time over a 10 km shallow-water propagation track in the Mediterranean Sea. The bandwidth is 200–800 Hz.
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ACOUSTICS, SHALLOW WATER
119
Wave eq.
WI
NM
Coupled WI
Coupled NM
Range independent
Adiabatic NM
Ray
PE
FD / FE
Range dependent
Figure 10 Hierarchy of numerical models in underwater acoustics. WI, wavenumber integration; NM, normal modes; PE, parabolic equation; FD, finite difference; FE, finite element.
Numerical Modeling The advent of computers has resulted in an explosive growth in the development and use of numerical models since the mid-1970s. Numerical models have become standard research tools in acoustic laboratories, and computational acoustics is becoming an evermore important branch of the ocean acoustic science. Only the numerical approach permits an analysis of the full complexity of the acoustic problem. An assortment of models has been developed over the past 25 years to compute the acoustic field in shallow-water environments in both the frequency and time domains. Entire textbooks are dedicated to the development of theoretical and numerical formalisms which can provide quantitative acoustic predictions for arbitrary ocean environments. Sound propagation is mathematically described by the wave equation, whose parameters and boundary conditions are descriptive of the ocean environment. As shown in Figure 10, there are essentially five types of models (computer solutions to the wave equation) to describe sound propagation in the sea: wavenumber integration (WI); normal mode (NM); ray; parabolic equation (PE) and direct finite-difference (FD) or finite-element (FE) solutions of the full wave equation. All of these models permit the ocean environment to vary with depth. A model that also permits horizontal variations in the environment, i.e. sloping bottom or spatially varying oceanography, is termed range dependent. As shown in Figure 10, an a priori assumption about the environment being range independent, leads to solutions based on spectral techniques (WI) or normal modes (NM); both of these techniques can, however, be extended to treat range dependence. Ray, PE and FD/FE solutions are applied directly to rangevarying environments. For high frequencies (a few kilohertz or above), ray theory, the infinite frequency approximation, is still the most practical, whereas the other five model types become more and more applicable below, say, a kilohertz in shallow water.
Models that handle ocean variability in three spatial dimensions have also been developed, but these models are used less frequently than two-dimensional versions because of the computational cost involved.
Conclusions The acoustics of shallow water has been thoroughly studied both experimentally and theoretically since World War II. Today the propagation physics is well understood and sophisticated numerical models permit accurate simulations of all processes (reflection, refraction, scattering) that contribute to the complexity of the shallow-water problem. Sonar performance predictability, however, is limited by knowledge of the controlling environmental inputs. The current challenge is therefore how best to collect relevant environmental data from the world’s enormously variable shallow-water areas.
See also Acoustic Noise. Acoustic Scattering by Marine Organisms. Acoustics, Arctic. Acoustics in Marine Sediments. Sonar Systems. Surface Gravity and Capillary Waves. Tomography.
Further Reading Brekhovskikh LM and Lysanov YP (1990) Fundamentals of Ocean Acoustics, 2nd edn. New York: SpringerVerlag. Etter PC (1996) Underwater Acoustic Modeling, 2nd edn. London: E & FN Spon. Jensen FB, Kuperman WA, Porter MB, and Schmidt H (2000) Computational Ocean Acoustics. New York: Springer-Verlag. Medwin H and Clay CS (1998) Fundamentals of Acoustical Oceanography. San Diego: Academic Press. Urick RJ (1996) Principles of Underwater Sound, 3rd edn. Los Altos: Peninsula Publishing.
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AEOLIAN INPUTS R. Chester, Liverpool University, Liverpool, Merseyside, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 1, pp 97–103, & 2001, Elsevier Ltd
Introduction The oceans are an important reservoir in the global biogeochemical cycles of many elements, but until the recent past it was thought that material fluxes to the reservoir were dominated by fluvial inputs. Over the last two or three decades, however, it has become apparent that the atmosphere is a major transport pathway in the land–sea exchange of material. This atmospherically transported material differs from that introduced by fluvial inputs in two important ways. (i) It is delivered, albeit at different flux magnitudes, to all areas of the sea surface, whereas river inputs are initially delivered to the land–sea margins. (ii) It does not pass through the biogeochemically dynamic estuarine filter; a region of intense dissolved/particulate reactivity, which, under present day conditions, retains B90% of fluvial particulate material. As a result, the atmosphere is the most important pathway for the long-range transport of much of the particulate material delivered directly to open-ocean regions. This material has an important influence on marine sedimentation; for example, in equatorial North Atlantic deep-sea sediments deposited to the east of the Mid-Atlantic Ridge and in central North Pacific deep-sea sediments, essentially all the land-derived components are aeolian in origin. Atmospheric aerosols can also exert an influence on climatic forcing by acting as cloud condensation nuclei and by processes such as the scattering of short-wave radiation by both anthropogenic and natural aerosols. Further, the presence of anthropogenic sulfate aerosols in the atmosphere can lead to an increase in albedo and so cool the planet; an effect, which, on a global scale, is comparable to that induced by the ‘greenhouse’ gases, but is opposite in sign. The aeolian material delivered to the sea surface by the atmosphere is dispersed from the source regions via the major wind systems, such as the Trades and the Westerlies, within which relatively smallscale winds (e.g. the Sirocco and the Mistral in the Mediterranean) can be important locally. Large-scale meteorological phenomena can also affect aeolian
120
transport; for example, long-term inter-annual variability in dust transport out of Africa to the Atlantic Ocean and the Mediterranean Sea has been linked to precipitation patterns induced by the North Atlantic Oscillation. Material in the marine atmosphere consists of gaseous and particulate components, both of which can originate from either natural or anthropogenic sources. Gas-to-particle conversions are important in the generation of particulate material, especially that derived from anthropogenic sources; however, air/sea gaseous exchange is covered in articles on air–sea interactions, and attention here is largely confined to the particulate aerosol. The sea surface itself is a major source of particulate material to the marine atmosphere in the form of sea salt. However, these sea salts are re-cycled components, and the globally important terrestrial sources of material to the marine atmosphere, i.e. those supplying material involved in land– sea exchange, are (i) the Earth’s crust (mineral dust), and (ii) anthropogenic processes (sulfates, nitrates, etc). Other terrestrial sources, which include volcanic activity and the biosphere (e.g. direct release from vegetation, biomass burning), can also supply components to the atmosphere. Material is removed from the atmosphere by a combination of two depositional modes; (i) the ‘dry’ mode, which does not involve an aqueous phase, and (ii) the ‘wet’ (precipitation scavenging) mode, either by cloud droplets (in-cloud processes) or by falling rain (below-cloud processes).
Land–sea Exchange of Individual Components Marine Aerosol: Major Components, Sources, and Distribution
Data are available on the concentrations of aerosols over many marine regions, and it is now apparent that there is an ‘aerosol veil’ over all oceans. Concentrations of material in the aerosol veil, however, vary from B103 ng m3 of air close to continental sources to B102 ng m3 of air over pristine oceanic regions. The aerosol veil is composed mainly of mineral dust and anthropogenic components. On a global scale, the anthropogenic material is dominated by sulfate aerosols, together with smaller amounts of nitrates. The mineral dust aerosol consists of a wide variety of minerals, with quartz, the clay minerals and feldspars usually being
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AEOLIAN INPUTS
predominant. The signatures of the major clay minerals (chlorite, kaolinite, illite, and montmorillonite) can be used as tracers to identify the sources of the dusts, the extent to which the material has been transported over the oceans, and its contribution to marine sedimentation. The principal continental sources of both the mineral dust aerosol and anthropogenic (mainly sulfate) aerosol are concentrated in specific latitudinal belts, predominantly in the northern hemisphere (see Figure 1A). Quantitatively, the crust-
121
derived mineral dust, which is derived mainly from the arid and semi-arid desert regions of the world (Figure 1A), imposes the strongest fingerprints on the marine aerosol. Mineral dust fluxes to the world ocean are listed in Table 1, and are illustrated in Figure 2. From this figure it can be seen that the highest dust fluxes to the sea surface are found off the major deserts, e.g. the Sahara in the North Atlantic and the Asian deserts in the North Pacific. Much of the material injected into the atmosphere from these arid sources is transported in the form of
40°N
0
40°S
(A)
120°E
60°E
60°E
60°N
180
120°E
_2
60°W
0
00
0
100
00
10
0 00
120°E
10
_1
y )
30°N 0
120°W
Dust flux (mg m
60°E
0
00
10
1000 10 000
100
1
1000 100
30°S 60°S
10 10
10
(B) Figure 1 Aerosols: terrestrial sources and fluxes to the world ocean. (A) Terrestrial sources of aerosol production; light gray areas indicate regions of anthropogenic emissions, and dark gray areas indicate regions of mineral aerosol production. (Reproduced with permission from Gilman C and Garrett C (1994) Heat flux parameterizations for the Mediterranean Sea: The role of atmospheric aerosols and constraints from the water budget. Journal of Geophysical Research 99: 5119–5134.) (B) Mineral aerosol fluxes to the world ocean (units, mg m2 y 1). (Reproduced with permission from Duce RA et al. (1991)).
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AEOLIAN INPUTS
Table 1
Dust deposition rates to the world ocean
Ocean region
Deposition rate
North Atlantic, north of trades North Atlantic trades South Atlantic Indian Ocean North Pacific: western North Pacific central and eastern South Pacific Entire Pacific
106g cm2y1
1012g y1
82 – 85 450
12 100–400 18–37 336
5000 11–62 5–64
300 30 18 350
All oceans (minimum–maximum)
816–1135
Adapted from Prospero JM (1981) and Prospero JM, Uematso M and Savoie DL (1989). In: Riley JP and Chester R (eds) Chemical Oceanography, vol. 10, pp. 137–218. London: Academic Press.
dust ‘pulses’; these are related to dust storms on the continental source regions, and are superimposed on background aerosol concentrations. Organic Matter and Organic Compounds
Various classes of particulate organic carbon (POC) and vapour-phase organic carbon (VOC) are present in the atmosphere. The non-methane atmospheric global VOC burden has been estimated to be B50 1012 g, and the total POC burden to be B1– 5 1012 g. The principal terrestrial sources of organic matter to the atmosphere are vegetation, soils, biomass burning, and the freshwater biomass, together with a variety of anthropogenic processes. The sea surface also contributes to the organic matter 60°E
120°E
180
burden in the marine atmosphere, with B14 1012 g y1 of organic carbon being produced by the ocean surface; 490% being on particles 41 mm in diameter. Particulate organic matter (POM) is removed from the air via ‘dry’ and ‘wet’ deposition, and in addition the removal of VOM includes conversion to POM and transformation to inorganic gaseous products. Estimates of the ‘wet’ atmospheric depositional flux of carbon to the ocean surface range between B2.2 1014 g y1 and B10 1014 g y1, and for the ‘dry’ flux a value of B6 1012 g y1 has been proposed. These are of the same order of magnitude as the estimates of fluvially transported POC entering the oceans (B1–2.5 1014 g y1). The estimates of atmospherically transported carbon must be regarded with extreme caution, but nonetheless, even when the marine source is taken into account, the ‘wet’ and ‘dry’ flux estimates indicate that the oceans act as a major sink for organic carbon in the atmosphere. Atmospherically transported carbon, however, makes up a maximum of only B2% of the carbon produced by primary productivity (B30–50 1015 g y1). Viable POC in the marine atmosphere includes material such as fungi, bacteria, pollen, algae, insects, yeasts, molds, mycoplasma, viruses, phages, protozoa, and nemotodes. Non-viable POC includes carbonaceous material (which has a refractory ‘soot’ component) and individual organic species. Concentrations of carbonaceous aerosols in the marine atmosphere vary over the range B0.05–1.20 mg C m3 of air and display a distinct latitudinal distribution. In the northern hemisphere the carbon, which ranges in concentration between B0.4 and B1.2 mg C m3 of air, has a predominantly anthropogenic continental source 120°W
60°W
0 100
60°N
Lead flux (μg m
_2
_1
y )
1000
1000
100
1000
30°N
100 100
0
10 100
100
10
10
30°S
1000
100
100
60°S
(A) Figure 2 Fluxes of lead to the world ocean (units, mg m2 y1). From Duce RA et al. (1991).
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10
AEOLIAN INPUTS
from combustion processes. In the southern hemisphere POC concentrations are lower (range B0.05– 0.30 mg C m3 of air), and natural continental and marine sources are about equal. A wide variety of individual species of both natural and synthetic organic compounds are transported from the continents to the sea surface via the atmosphere. The classes of organic compounds that have received particular attention include aliphatic hydrocarbons, wax esters, fatty alcohols, sterols, fatty acids, and long-chain unsaturated ketones. The most extensive investigation of atmospherically transported organic species on an ocean-wide basis was carried out in the Pacific as part of the SEXREX Program (Sea-Air Exchange Program). On the basis of data obtained from the Pacific it is apparent that in the remote marine atmosphere organic carbon accounts for B10% of the total aerosol, although only B1% of this has been characterized. The most abundant terrestrially derived components are the nalkanes and the C21–C36 fatty alcohols, which are common in the epicuticular waxes of vascular plants, and the most abundant marinederived species are the C13–C18 fatty acid salts. The atmosphere is a major pathway for the transport of a number of organic pollutants that enter sea water. These include the synthetic trace organics, such as high molecular-weight halogenated hydrocarbons of the following compounds, or compound classes; chlorobenzenes (e.g. hexachlorobenzene, HCB), chlorocyclohexanes (e.g. hexachlorocyclohexanes, HCHs), polychlorobiphenyls (PCBs), dichlorodiphenyltrichloroethanes (DDTs), and non-halogenated coumpounds (e.g. polynuclear aromatic hydrocarbons, PAHs). Global atmospheric fluxes of HCHs, HCBs, DDTs, and PCBs to the world ocean are listed in Table 2, and two important conclusions can be drawn from the data. (i) The dominant deposition of the organochlorines is to the North Atlantic and North Pacific, which is consistent with their source derivations; with HCH and DDT compounds having their highest deposition rates in the Table 2
123
North Pacific and PCBs in the North Atlantic. (ii) The atmospheric inputs of the organochlorines to the world ocean exceed those from fluvial inputs by 1 to 2 orders of magnitude. Nutrients
These include nitrate, phosphate, and micro-nutrients such as iron. The main features in the atmospheric input of nitrogen nutrients to the world ocean can be summarized as follows. (i) The atmospheric input of total nitrogen (N) to the global ocean is B30.2 1012 g N y1, made up of B13.4 1012 g N y1 oxidized nitrogen species and B16.8 1012 g N y1 reduced nitrogen species, which is similar in magnitude to the total (i.e. natural þ anthropogenic) fluvial nitrogen flux (B21–49 1012 g N y1). (ii) The overall flux of nitrogen to the sea surface is B87 mg N m2 y1; the largest fluxes being to the North Atlantic and North Pacific. (iii) The ‘wet’ removal of reduced nitrogen species may be an important source of nutrients to the oceans. The atmospheric fluxes of nitrogen species usually include only inorganic forms, such as nitrate and ammonia; however, if dissolved organic nitrogen (DON), most of which has an anthropogenic source, is included in ‘wet’ deposition it would increase the anthropogenic input of fixed nitrogen to the oceans by a factor of B1.5. The atmospheric input of phosphate has been studied in detail in the Mediterranean Sea. However, the role played by atmospherically transported phosphorus in nutrient cycles is not clearly understood; for example, it has been proposed that Saharan dust may act as a sink for the removal of dissolved phosphorus in the water column by adsorption onto iron-rich particles, or as a source with up to B8% of the phosphorus in the dusts being soluble in sea water. There is evidence that in summer months the atmosphere does provide a source of phosphorus to the Western Mediterranean and that this may account for new production, since the stratification of the surface waters prevents the input
The atmospheric input of some organochlorine compounds to the world oceans (units, 106 g y1)
Compounda
North Atlantic
South Atlantic
North Pacific
South Pacific
Indian Ocean
Global atmospheric
Global fluvial
HCHs HCB DDTs PCBs
850 17 16 100
97 10 14 14
2600 20 66 36
470 19 26 29
700 11 43 52
4800 77 170 240
40–80 4 4 40–80
a HCHs, hexachlorocyclohexanes; HCB, hexachlorobenzene; DDTs, dichlorodiphenyltrichloroethanes; PCBs, polychlorobiphenyls. Adapted from Duce RA et al. (1991).
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AEOLIAN INPUTS
of nutrients from deep waters. For example, in the Ligurian Sea a strong summer desert dust transport episode was followed 10 days later by a significant increase in phytoplankton concentration, which could have resulted from the solubilization of phosphorus from the dusts. In recent years there has been a renewed interest in the role played by iron as a limiting nutrient in primary production, especially in high-nitrogen, lowproductivity (HNLP) regions in which there is sufficient light and nutrient concentrations but low productivity. The concentrations of iron in openocean waters are generally low, and it is thought that the element would run out before nitrate is exhausted. To provide sufficient iron it therefore has been suggested that it must be added to sea water from other sources, one of which is the long-range transport of atmospheric dust. However, iron oxyhydroxide particles and iron colloids are not directly available to phytoplankton, and the bioavailable forms of the metal are thought to be dissolved Fe(II), which is rapidly converted into Fe(III), and Fe(III) itself. As a result, iron must be dissolved from atmospheric dust before it becomes available to phytoplankton, and although adsorption onto particulate matter is the dominant control on the concentrations of dissolved iron in open-ocean waters, the deposition of mineral dust from high concentration episodic atmospheric events can result in a net addition of dissolved iron to surface waters. Further, during ‘wet’ deposition iron may undergo reductive dissolution to Fe(II), a form of iron that would be immediately available to phytoplankton. It also has been suggested that siderophores (compounds with a high affinity for ferric iron which are secreted by organisms) may play a role in the bioavailability of iron. Various small-scale laboratory simulations have shown that phytoplankton growth rates increase in response to the addition of iron to the system, but there is considerable disagreement over the interpretation of the results. Laboratory experiments also have been criticized on the grounds that they do not represent planktonic community response on an ocean-wide scale. To overcome this, the Fe-limitation hypothesis has been tested by largescale intervention experiments, such as IronEx I and IronEx II, which were carried out in the equatorial Pacific. In these experiments patches of sea (defined by SF6 tracer) were seeded with iron in the concentrations expected from natural events. Although there was a doubling of the plant biomass following iron addition, interpretation of the data from IronEx I was hampered as a result of the subduction of the seeded seawater patch below a layer of less dense water. During IronEx II, however, a massive
phytoplankton bloom was triggered, providing direct evidence that in these HNLP waters phytoplankton growth is iron limited.
Trace Metals
Particulate trace metals in the marine atmosphere that are involved in land–sea exchange are derived from two principal terrestrial sources. The first is the Earth’s crust. Crustal weathering involves low-temperature generation processes, and crust-derived elements are found on particles with mass median diameters (MMDs) in the range B1–3 mm. Relative to other sources, crust-derived elements are referred to as the non-enriched elements (NEEs). The second source comprises a variety of anthropogenic processes, which often involve high temperatures (e.g. fuel combustion, ore smelting). Anthropogenically derived elements (e.g. Pb, Cu, Zn, Cd, As, Hg) are largely found on smaller particles with MMDso0.5 mm, and are termed the anomalously enriched elements (AEEs). Less important terrestrial sources of trace metals to the marine atmosphere include volcanic activity and the biosphere. In addition, the sea surface supplies recycled elements to the marine atmosphere; this is a low-temperature source, and sea salt-associated elements are located on particles with MMDs in the range B3–7 mm. Global elemental atmospheric emission rates from natural and anthropogenic sources are listed in Table 3. The overall trace metal composition of the marine aerosol is dependent on the extent to which material from the various sources are mixed together in the atmosphere, and trace metal concentrations are a function of factors such as the distance the air mass Table 3 Global elemental emission rates to the atmosphere (units, 109 g y1) Element
Natural
Anthropogenic
Al As Cd Co Cr Cu Fe Hg Mn Pb V Zn
20940 12 1.4 6.1–7.3 44 22–28 10370 2.5 221–317 11.5–12 41–45 45–280
4000 19 7.7 2.9 30.5 35–52 6000 3.6 39–408 332–404 22–86 132–280
Adapted from Chester R (2000) (full references given in original table).
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AEOLIAN INPUTS
Table 4
125
Concentrations of particulate trace metals in the marine atmosphere (units, ng m3 of air) Coastal regions
Open-ocean regions
Close to anthropogenic sources
Close to crustal sources
Trace metal
North Sea
W. Black Sea
N. Atlantic N. Arabian north-east Sea trades
Tropical Tropical Indian Tropical N. Atlantic Ocean N. Pacific
Tropical S. Pacific
Al Fe Mn Ni Cr V Cu Zn Pb
294 353 14.5 3.8 4.7 – 6.3 41 34.5
540 420 17 4.9 9.0 3.2 – 46 60
5925 3685 65 6.6 10 15 4.5 16 6.9
160 100 2.2 0.64 0.43 0.54 0.79 4.4 9.9
– – – 0.013 0.07 0.016
1227 790 17 2.0 3.0 6.3 2.6 10 4.3
11 8.8 0.16 0.043 0.066 0.023 0.077 0.10 0.17
21 17 0.29 – 0.09 0.08 0.045 0.17 0.12
Adapted from Chester R (2000) (full references given in original table).
transporting them has travelled from the source, the ‘aging’ of the aerosol in the air, and the relative effectiveness of the processes that remove material from the atmosphere. As a result, trace metals in the marine atmosphere have concentrations ranging over several orders of magnitude – see Table 4. It is apparent from the data in this table that, in general, the trace metal concentrations decrease with increasing remoteness from continental sources in the general rank order: coastal seas4North Atlantic4North Pacific and tropical Indian Ocean4South Pacific. This is reflected in trace metal fluxes, which decrease in the same sequence; lead atmospheric fluxes to the world ocean are illustrated in Figure 2. Despite the fact that trace metal concentrations decrease towards more pristine oceanic environments, atmospheric inputs can be the dominant source of some trace metals to the mixed layer in open-ocean regions where inputs from other sources are minimal. This is especially the case for the ‘scavenged-type’ metals (e.g. Al, Mn, Pb), which have a surface source and a relatively short residence time in sea water. Aerosol-associated trace metals are removed from the air by ‘dry’ or ‘wet’ depositional processes, and once deposited at the sea surface the initial constraint on the manner in which the metals enter the major marine biogeochemical cycles is a function of the extent to which they undergo solubilization in sea water. There is considerable size-dependent fractionation between the parent aerosol and the deposited material in both the ‘dry’ and the ‘wet’ depositional modes. However, with respect to the seawater solubility of trace metals, the major difference between the two depositional modes is that they
follow separate geochemical routes. In ‘dry’ deposition material is delivered directly to the sea surface and trace metal solubility is largely constrained by particle 2 seawater reactivity; estimates of the seawater solubility of trace metals from aerosols are listed in Table 5. In contrast, in ‘wet’ deposition there is an initial particle 2 rainwater reactivity; much of this is pH-dependent, and can involve the dissolution of some trace metals prior to the deposition of the scavenged aerosol at the sea surface. Data on the trace metal composition of marine rainwaters from a number of regions are now available, and a selection are listed in Table 6, from which it can be seen that trace metals in rainwaters, like those in aerosols, have their highest concentrations in coastal regions and decrease towards more pristine oceanic regions.
Table 5 Seawater solubility of trace metals from particulate aerosols; % total element soluble Trace metal
Solubility (%)
Al Fe Mn Ni Cr V Cu Zn Pb
B1–10 B1–50 B20–50 t20–50 t10–20 t20–85 t10–85 t10–75 t10–90
Data, which are from various sources, include both crustal and anthropogenic aerosols.
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AEOLIAN INPUTS
Table 6
Trace metal concentrations in marine-influenced rainwaters (units, ng l1) Mediterranean Sea
North Atlantic
Trace metal
North Sea coast (VWM)
Irish Sea coast
South coast, France (VWM)
Sardinia (VWM)
Bermuda (VWM)
Bantry Bay, Ireland (VWM)
North Pacific
South Pacific
Al Fe Mn Cu Zn Pb
– 88 3.8 2.3 13 4.0
43 48 2.0 8.7 9.3 5.2
144 – – 2.8 – 3.7
883 519 8.0 2.9 16 1.6
– 4.8 0.27 0.66 1.15 0.77
3.62 8.06 0.13 0.86 8.05 0.51
2.1 1.0 0.012 0.013 0.052 0.035
16 0.42 0.020 0.021 1.6 0.014
VWM, volume-weighted mean concentrations; this normalizes the trace metal concentration in a rain to the total amount of rainfall over the sampling period. Adapted from Chester R (2000) (full references given in original table).
Table 7 Atmospheric and fluvial trace metal fluxes to the world ocean (units, 109 g y1) Atmospheric input
Fluvial input
Element
Dissolved
Particulate
Dissolved
Particulate
Fe P Ni Cu Pb Zn Cd As
3.2 103 310 8–11 14–45 80 33–170 1.9–3.3 2.3–5.0
29 103 640 14–17 2–7 10 11–60 0.4–0.7 1.3–2.9
1.1 103 Total 300a 11 10 2 6 0.3 10
110 103 Total 300 1400 1500 1600 3900 15 80
Total phosphorus input to marine sediments. Adapted from Duce RA et al. (1991).
Estimates of particulate and dissolved atmospheric and fluvial trace metal fluxes to the world ocean are listed in Table 7, from which a number of overall conclusions can be drawn. (i) Rivers are the principal source of particulate trace metals to the oceans; phosphorus being an exception. (ii) For iron, nickel, copper, and phosphorus the dissolved atmospheric and the dissolved fluvial inputs are the same order of magnitude. (iii) For lead, zinc, and cadmium the dissolved atmosphere fluxes are dominant; this will still be the case for lead even when allowance is made for the phasing out of leaded gasoline, which has accounted for a large fraction of the anthropogenic lead previously released into the atmosphere.
Conclusions There is an ‘aerosol veil’ over all marine regions, and the atmosphere is a major transport route for the supply of mineral dust, organic matter, nutrients, and
trace metals to the world ocean. Atmospheric fluxes decrease in strength away from the continental source regions, but in remote open-ocean areas they can be the dominant supply route for the deposition of land-derived particulate material and some trace metals to the ocean surface. Unlike fluvial inputs, which are delivered to the land–sea margins, the atmosphere supplies material to the ‘mixed layer’ over the whole ocean surface, and the material plays an important role in oceanic biogeochemical cycles and in the formation of marine sediments.
See also Air–Sea Gas Exchange. Air–Sea Transfer: Dimethyl Sulfide, COS, CS2, NH4, Non-Methane Hydrocarbons, Organo-Halogens. Air–Sea Transfer: N2O, NO, CH4, CO; Anthropogenic Trace Elements in the Ocean. Atmospheric Input of Pollutants. Metal Pollution. Nitrogen Cycle. Phosphorus Cycle. Photochemical Processes. Refractory Metals. Transition Metals and Heavy Metal Speciation.
Further Reading Buat-Menard P (ed.) (1986) The Role of Air-Sea Exchange in Geochemical Cycling. Dordrecht: Kluwer Academic Publishers. Charlson RJ and Heintzenberg (eds.) (1995) Aerosol Forcing of Climate. Berlin: Dahlem Workshop. Chester R (2000) Marine Geochemistry 2nd edn. Oxford: Blackwell Science. Duce RA, Mohnen VA, Zimmerman PR, et al. (1983) Organic material in the global troposphere. Reviews of Geophysics and Space Physics 21: 921--952. Duce RA, Liss PS, Merrill JT, et al. (1991) The atmospheric input of trace species to the World Ocean. Global Biogeochemical Cycles 5: 193--529.
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AEOLIAN INPUTS
Guerzoni S and Chester R (eds.) (1996) The Impact of Desert Dust Across the Mediterranean. Dordrecht: Kluwer Academic Publishers. Knap AH (ed.) (1990) The Long Range Atmospheric Transport of Natural and Contaminant Substances. Dordrecht: Kluwer Academic Publishers. Peltzer ET and Gagosian RB (1989) Organic chemistry of aerosols over the Pacific Ocean. In: Riley JP and Chester R (eds.) Chemical Oceanography, vol. 10, pp. 282--338. London: Academic Press.
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Prospero JM (1981) Eolian transport to the World Ocean. In: Emiliani C (ed.) The Sea, vol. 7, pp. 801--874. New York: John Wiley, Interscience. Prospero JM (1996) The atmospheric transport of particles to the ocean. In: Ittekkot V, Schafer P, Honjo S, and Depetris PJ (eds.) Particle Flux to the Oceans, pp. 19--52. New York: John Wiley.
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AGULHAS CURRENT J. R. E. Lutjeharms, University of Cape Town, Rondebosch, South Africa Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 1, pp 104–113, & 2001, Elsevier Ltd.
Introduction The greater Agulhas Current forms the western boundary system of the circulation in the South Indian Ocean. Contrary to the flow of comparable subtropical gyres in other ocean basins, the sources of the Agulhas Current are interrupted by a substantial barrier, the island of Madagascar. This leads to the formation of two minor western boundary flows, the East Madagascar Current and the Mozambique drift. Once fully constituted off the coast of south-eastern Africa, the Agulhas Current proper can be considered to consist of two distinct parts: the northern and the southern current. The northern part flows along a steep continental shelf and its trajectory is extremely stable. The southern part flows along the wide shelf expanse of the Agulhas Bank and by contrast meanders widely. South of the African continent the Agulhas Current retroflects in a tight loop, with most of its waters subsequently flowing eastward as the Agulhas Return Current. This loop configuration is unstable and at irregular intervals it is pinched off to form a detached Agulhas ring. These rings, carrying warm and salty Indian Ocean water, drift into the South Atlantic Ocean. Some cross the full width of this ocean in the next 2– 3 years, whereas many are dissipated within 5 months of being spawned. The Agulhas Return Current flows back into the South Indian Ocean along the Subtropical Convergence. This juxtaposition generates considerable mesoscale turbulence in the form of meanders and an assortment of eddies. Water from the Agulhas Return Current leaks northward, back into the subtropical gyre, along its full length. By about 701E all Agulhas water has been lost to the eastward flow that subsequently continues as the South Indian Ocean Current.
Importance Historically the Agulhas Current was one of the first ocean currents to receive a great deal of scientific attention. It was described in some detail as early as 1766 by Major James Rennell, preeminent British
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geographer at the time. This was followed by wideranging investigations by Dutch mariners such as Van Gogh and Andrau in the 1850s. This early interest was motivated purely by nautical concerns, the Agulhas Current constituting a formidable impediment to vessels sailing to India and to the East. Fundamental studies by German investigators dominated research on the Agulhas Current region in the 1930s, but this endeavour was terminated by the Second World War. During the past few decades a renaissance in interest in this current has occurred for totally different reasons. It has been demonstrated that the greater Agulhas Current system (Figure 1) has a marked influence on the climate variability over the southern African subcontinent. It has also been shown that this current is a key link in the exchanges of water between ocean basins and thus probably has a special role in the oceans’ influence on global climate. This renewed interest has stimulated a number of research cruises, the placement of current meter moorings, investigations by satellite remote sensing, as well as theoretical and modeling studies, all leading to an enormous increase in knowledge of the Agulhas Current.
Large-scale Circulation The Agulhas Current forms part of the overall circulation of surface waters in the South Indian Ocean that is anticyclonic, i.e., anticlockwise in the Southern Hemisphere (Figure 2). On its eastern side, the equatorward flow is weak and dispersed, whereas the recirculation in the South West Indian Ocean is particularly strongly developed, penetrating to 1000 m depth at its centre. The southern border to the circulation is the Subtropical Convergence. This strong thermohaline front at roughly 411S separates the characteristic flows and water masses of the subtropical gyre and those of the Antarctic Circumpolar Current that lies to the south. Along the Subtropical Convergence, the Agulhas Return Current and the South Indian Ocean Current carry their respective water masses eastward. To the north the gyral circulation is closed by the South Equatorial Current that is found between about 101 and 251S and carries water from east to west. At the eastern shores of Madagascar, the South Equatorial Current splits into a northern and a southern limb of the East Madagascar Current; about 70% going north along this shoreline, 30% heading south. Most of that heading north eventually
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AGULHAS CURRENT
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Figure 1 A conceptual portrayal of the greater Agulhas Current. Ocean regions shallower than 3000 m have been shaded. Intense currents are black, whereas the general background circulation is shown by open arrows. Cyclonic eddies are open; anticyclonic rings and eddies are black. Note the stability of the northern Agulhas Current, the meanders of the southern Agulhas Current, the tight retroflection loop, and the continuously weakening eastward flow of the Agulhas Return Current along the Subtropical Convergence. Agulhas rings (black) are advected by the Benguela Current past the extensive coastal upwelling off south-western Africa.
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E Figure 2 The baroclinic volume transport for the upper 1000 m of the South Indian Ocean. Values are in units of 106 m3 s1. Note the very small contribution coming from the Mozambique Channel and the concentration of the recirculation in the South West Indian Ocean west of 701E. This transport pattern, averaged over a long period, is not to be confused with the depictions of instantaneous currents in Figures 1 and 4. (After Stramma and Lutjeharms (1997) Journal of Geophysical Research 102(C3): 5513–5530. ^ American Geophysical Union.)
reaches the east coast of the African continent, where it forms the East African Coastal Current. There is some leakage from the Mozambique Channel and from the southern limb of the East Madagascar Current into the Agulhas Current, but most of the Agulhas Current’s waters come from the subgyre of the South West Indian Ocean (see Figure 2). The
Agulhas Current itself is narrow, deep, and fast; a typical western boundary current. All these surface currents are driven largely by the reigning wind systems. The wind systems over the South West Indian Ocean fall largely outside the influence of the seasonally varying monsoonal winds of the North
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Indian Ocean. The average air motion over the South Indian Ocean is dominated by a large-scale, anticyclonic circulation around a high-pressure system centered south-east of the island of Madagascar. This flow generally is stronger in austral summer than in winter. Strongest winds (24 m s1) are found at 501S latitude in summer, and weakest at 351S (2 m s1). A band of minimum wind stress extends across this ocean at 351S. Next to the continental land masses the winds are usually aligned with the coasts. Apart from driving the surface currents, the atmosphere also has a considerable effect on the formation of certain water masses (Figure 3) that are typical for the region. This temperature–salinity portrayal indicates the characteristic for each specific water mass in this ocean region. The fresher Tropical Surface Water is formed north of 201S where there is an excess of precipitation over evaporation; Subtropical Surface Water is formed between 281 and 381S, where this ratio is reversed. Subtropical Surface Water is found as a shallow subsurface salinity maximum in regions to the north and south of its region of formation. Intermediate waters lie at depths between 1000 and 2000 m and consist of Antarctic Intermediate Water and North Indian Intermediate Water (also called Red Sea Water). The former subducts between the Antarctic Polar Front and the Subtropical Convergence;
30 Tropical Surface Water
26
Potential temperature (˚C)
22 Tropical Thermocline Water
18
Subtropical Surface Water
the latter is formed owing to very high rates of evaporation in the Red Sea, the Arabian Sea and the Persian Gulf. Central Water, lying between the surface and the intermediate waters, is a mixture of these two. Below the intermediate waters are Indian Deep Water and North Atlantic Deep Water, each formed by subduction in the respective ocean regions after which they are named. All these respective water masses are involved in some way or other in the source currents of the Agulhas Current.
Sources of the Agulhas Current According to the transport portrayed in Figure 2, 30% of the volume flux of the Agulhas Current derives from east of Madagascar, only 13% comes through the Mozambique Channel, and 67% is recirculated in a South West Indian Ocean subgyre. Note, however, that the inflow from east of Madagascar does not necessarily come from the East Madagascar Current (see Figure 1). The southern limb of the East Madagascar Current starts at the bifurcation point of the South Equatorial Current at about 171S along the east coast of Madagascar. Its surface speed here is roughly 1 m s1, increasing downstream to about 1.5 m s1. The current is very stable in both flux and trajectory and exhibits no clear seasonality in any of its characteristics. Using the 0.5 m s1 isotach as the outer limits of the current, it is 75 km wide, 200 m deep, and its core lies 50 km offshore. It carries Tropical as well as Subtropical Surface Water with a total volume transport of 21 106 m3 s1 (i.e., 21 Sv). Where it overshoots the end of the shelf of Madagascar it retroflects, with most of its waters subsequently heading eastwards (Figure 4). There may be some E
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Figure 3 The relationship of potential temperature and salinity for waters in the western Indian Ocean. Some of the characteristic water masses to be found here are SICW, South Indian Central Water; SAMW, Subantarctic Mode Water; NADW, North Atlantic Deep Water; AABW, Antarctic Bottom Water; AASW, Antarctic Surface Water; AAIW, Antarctic Intermediate Water; and SAASW, Subantarctic Surface Water. (After Gordon et al. (1987) Deep-Sea Research 34(4): 565–599. ^ Elsevier Science.)
S
Retroflection Agulhas Current
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Figure 4 A conceptual portrayal of the flow regime in the source regions of the Agulhas Current. Narrow, intense currents are shown by black arrows. The East Madagascar Current is a miniature western boundary current that retroflects south of Madagascar. (After Lutjeharms et al. (1981) Deep-Sea Research 28(9): 879–899. ^ Elsevier Science.)
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AGULHAS CURRENT
Northern Agulhas Current Water from the South West Indian Ocean subgyre feeds into the Agulhas Current along its full length, but at a latitude of 271S this current is nonetheless thought to be fully constituted (Figure 5). This northern part of the Agulhas Current is characterized in particular by an extremely stable
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leakage of East Madagascar Water into the Agulhas Current by way of rings and filaments, but this constitutes an insignificant contribution. The wider flow east of Madagascar, up to 240 km offshore, is 41 106 m3 s1. A substantial part of this more extensive flow may eventually make its way into the Agulhas Current (Figures 2 and 4). The flow through the Mozambique Channel was once thought to be the major contributor to the flux of the Agulhas Current. Now it is considered to be minor (Figure 2). The entire existence of a consistent, continuous Mozambique Current, flowing along the African coastline, has in fact been called into question. The northern mouth of the Mozambique Channel is largely closed to subsurface flow by bottom ridges shallower than 2000 m, except at its western side. The surface circulation in the northern part of the channel is anticyclonic to an estimated depth of 1000 m. The eastern side of this flow, which might be the start of a Mozambique Current, is 250 km wide, has a surface speed of 0.3 m s1 and a volume flux of 6 106 m3 s1. The water masses here are characteristic of the monsoonal regime to the north with no Antarctic Intermediate Water. There is some evidence for the presence of this water mass in the central part of the channel, but there still is no North Atlantic Deep Water. The rest is essentially Subtropical Surface Water of the South Indian variety. Occasional strong flows of 2 m s1 have been observed off the African coastline in this central part of the channel, but this is extremely variable. The southern third of the channel has all the thermohaline characteristics of the South West Indian Ocean, including the presence of North Atlantic Deep Water. Net volume flux through the southern mouth of the channel seems to be very changeble. Calculations have varied from 26 106 m3 s1 southward to 5 106 m3 s1 northward, above 1000 m. The best-substantiated characteristic of the circulation in the Mozambique Channel is therefore its very high mesoscale variability. This argues against a persistent western boundary current and instead suggests a series of eddies moving southward down the channel. This scenario is consistent with most available observations and also some numerical models.
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2000
0 10 20 Di sta 30 nc 40 e of fs ho 50 re (k 60 m ) 70
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3000
80
Figure 5 The spatial velocity structure of the northern Agulhas Current at Durban based on direct current measurements during a research cruise. Speeds below 1000 m have been estimated using a geostrophic calculation and are less certain. (After Duncan (1970) PhD dissertation, University of Hawaii.)
trajectory; its core meanders less than 15 km to either side. This stability is thought to be due to the strong slope of the continental shelf along which it flows. The current has a well-developed, inshore thermal front that meanders somewhat more extensively. The surface characteristics of the current are given in Table 1. Surface temperatures decrease by about 21C downstream. In the north they are at a maximum of 281C in February and a minimum of 231C in July. At Port Elizabeth, where the southern Agulhas Current starts, the maximum temperature is 251C in January, with a minimum of 211C in August. Surface salinities decrease from 35.5 PSU in the north to 35.3 PSU in the south. At Durban the core of the current usually lies 20 km offshore and penetrates to a depth of 2500 m (see Figure 5). Between the 0.5 m s1 isotachs it is 90 km wide; its offshore termination being more disperse than its strong inshore edge. Its core slopes so that at 900 mdepth it lies 65 km offshore. Its total volume flux is 73 106 m3 s1 and this increases by an estimated 6 106 m3 s1 for every 100 km distance downstream in the northern Agulhas Current. Surface speeds at its core usually lie between 1.4 and 1.6 m s1, with occasional peaks of up to 2.6 m s1.
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Table 1
Kinematic characteristics of the upper layers of the northern Agulhas Current
Peak speed in current core (m s1) Current core offshore distance (km) Distance offshore 0.5 m s1 (km) Distance offshore 1.0 m s1 (km) Core width, between 1.0 m s1 isotachs (km) Distance offshore temperature max. (km) Distance offshore 151C/200 m intersection (km) Distance offshore 35.35 PSU salinity/200 m (km)
Mean
SD
Minimum
1.36 52 35 42 34 58 50 47
0.30 14 14 14 15 20 15 13
30 10 25 10 35 25 23
Maximum 2.45 >100 70 95 >60 >100 90 >100
After Pearce (1997), Journal of Marine Research 35(4): 731–753.
Neither these velocities nor the volume fluxes show any discernible seasonality. The northern Agulhas Current is underlain by an opposing undercurrent at 1200 m that carries 6 106 m3 s1 water equatorward at a rate of about 0.3 m s1. It consists partially of modified North Indian Intermediate Water. The invariant path of the northern Agulhas Current is interrupted during about 20% of the time by an intermittent, solitary meander – the Natal Pulse – that originates at the Natal Bight, an offset in the coast north of 301S (see Figure 6). It translates downstream at a very steady 20 km per day, continuously growing in its lateral dimensions. On its
landward side it encloses a cyclonic eddy that creates a strong coastal countercurrent as the Natal Pulse passes. This meander is triggered at the Natal Bight whenever the current intensity there exceeds a certain threshold, allowing baroclinic instability to develop away from the constraining shelf slope. The Natal Pulse is an important component of the current system since it may cause upstream retroflection at the Agulhas Plateau (see Figure 1) and may precipitate ring shedding at the Agulhas retroflection, far downstream. Flow over the shelf adjacent to the northern Agulhas Current is dependent on the shelf morphology.
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Figure 6 Current-induced upwelling in the Natal Bight off south-eastern Africa. The left panel gives the distribution of nitrate (in mmol l1) at 10 m depth and the right hand panel the simultaneous distribution of chlorophyll a. The active upwelling cell off Cape St Lucia with high nitrates and chlorophyll a is well circumscribed. (After Lutjeharms et al. (2000) Continental Shelf Research 20(14): 1907–1939. ^ Elsevier Science.)
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AGULHAS CURRENT
Where the shelf is narrow, the flow is mostly parallel to the current. Over the Natal Bight, where the shelf is wider, the flow consist of cyclonic eddies. At the northern end of the Natal Bight, the current forces inshore upwelling (Figure 6). The water in this upwelling cell may be 51C colder than the adjacent current, have a high nutrient content, and exhibit enhanced biological primary productivity. The cold water thus upwelled flows over the bottom of the whole Natal Bight, strengthening the vertical layering over this shelf region. Off Durban a recurrent lee eddy is often observed. Seaward of the Agulhas Current, off the Mozambique Ridge (see Figure 1), a large number of very intense deep-sea eddies have been observed. They may be at least 2000 m deep, 100 km in diameter, have surface speeds of 1 m s1 and circular transports of between 6 106 and 18 106 m3 s1. Their lifetimes are estimated to be 1–3 years. Most of those observed are cyclonic, although a few anticyclonic ones have been seen. They seem to come from both the Mozambique Channel and from east of Madagascar, but their true origins remain unknown.
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Meanders usually have a trailing plume and an embedded, cyclonic lee eddy (Figure 7). There is evidence that these eddies are preferentially clustered in the eastern bight of the Agulhas Bank (see Figure 1). The dimensions of all these shear edge features change markedly with distance downstream (see Table 2). Sea surface temperatures also change more readily with distance downstream here than they do in the northern Agulhas Current. Sea surface temperatures in the southern Agulhas Current reach a maximum of 261C at Algoa Bay in February; 231C off the southern tip of the Agulhas Bank. In August these temperatures are 211C and 171C, respectively. The volume flux of the Agulhas Current off the southern tip of the Agulhas Bank has been estimated at 70 106 m3 s1 down to 1500 m, i.e., about the same as that of the current to its full depth at Durban. Even though there is this increase, it seems that the Table 2 Dimensions of shear edge features along the landward edge of the southern Agulhas Current
Southern Agulhas Current In contrast to the northern Agulhas Current, the southern Agulhas Current is characterized by wide meanders as it flows past the Agulhas Bank south of Africa (see Figure 1). Meanders are present along this shelf edge at least 65% of the time. They have an average wavelength of 300 km and a phase speed that varies from 5 to 23 km per day.
Plume lengths at surface (km) Plume widths at surface (km) Diameter of enclosed eddy (km) Plume dispersion from current (km)
Port Elizabeth
Tip of Agulhas Bank
100 27 27 50
162 37 51 150
After Lutjeharms et al. (1989) Continental Shelf Research 9(7): 597–616.
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18
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Eddy
200 Figure 7 Three vertical temperature sections across the Agulhas Bank picture the thermal composition of the southern Agulhas Current, a shear-edge eddy, and its associated plume. The Agulhas Current lies outside the 181C envelope. The plume has water warmer than 181C while the core of the eddy has water colder than 101C. (After Lutjeharms et al. (1989) Continental Shelf Research 9(7): 1570–1583. ^ Elsevier Science.)
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increase per unit distance downstream found in the northern Agulhas Current is not maintained in its southern part. Water masses in the current are generally the same in the northern and the southern part. The presence of Tropical Surface Water is maintained in the southern part, as are remnants of North Indian Intermediate Water (or Red Sea Water). Tropical Surface Water is mostly found at the inshore side of the current and derives from the Mozambique Current. Its presence and volume seem to be intermittent. Some of the surface plumes generated by meanders in the far southern reaches of the Agulhas Current are advected past the western edge of the Agulhas Bank as Agulhas filaments (see Figure 1). They are present about 60% of the time and carry substantial amounts of heat that are rapidly lost to the much colder atmosphere. They also carry about 3–9 1012 kg of salt per year into the South Atlantic; salt in excess to that of the waters already present there. On average they are 50 km wide and 50 m deep. The southern Agulhas Current influences the water masses over the adjacent Agulhas Bank in three ways. First, plumes of warm Agulhas surface water may extend over the bank, heating the top layers (see Figure 7). Second, current-driven upwelling takes place off the far eastern side of the Agulhas Bank (see Figure 1) and this water flows westward and covers the greater part of the bottom of the shelf. This process cools the water column from below, leading to intense seasonal thermoclines over the Agulhas Bank. Third, most of the mean flow over the eastern part of this shelf is parallel to the current. At the southern tip of the Agulhas Bank, the current detaches from the shelf edge.
The Agulhas Retroflection The region where the Agulhas Current then terminates south of Africa is characterized by its extremely high levels of mesoscale variability (Figure 8). The measured eddy kinetic energy is higher here than in any comparable western boundary current such as the Kuroshio or the Gulf Stream. This is due to a number of dynamical traits of the current retroflection. First, the continuous progradation, or westward penetration, of the Agulhas retroflection loop into the South Atlantic Ocean causes substantial levels of variability. The loop has an average diameter of 340 km (770 km) and progrades westward at about 10 km per day. The outer limits to its movement are 101 and 211E. At its furthest extent a ring is shed by loop occlusion. This happens between 4 and 9 times per year. This ring spawning activity adds considerably to the general variability. Ring shedding events are usually preceded by the arrival of a Natal
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25˚ E
30˚ 35˚ 40˚
Figure 8 The high levels of mesoscale variability that are characteristic of the Agulhas Current retroflection and the Agulhas Return Current are here portrayed by the superimposed thermal borders at the sea surface as observed for a period of one year. The inshore border of the northern Agulhas Current is particularly stable, that of the southern Agulhas Current less so, whereas the Agulhas Return Current exhibits a tendency to prefer certain meanders. The Agulhas retroflection has a range of locations and is attended by a host of rings (to the north) and eddies (to the south). (After Lutjeharms and van Ballegooyen (1988) Journal of Physical Oceanography 18(11): 1570–1580.)
Pulse on the Agulhas Current. The average lag time between initiation of a Natal Pulse off the Natal Bight and the shedding of a ring at the retroflection is 165 days. Between a newly formed ring and the reconstituted retroflection loop, a wedge of Subtropical Surface Water usually penetrates northward. Its water has a temperature of 171C and salinity lower than 34.9 over the top 100 m, also adding to the variability of the region. Newly formed rings retain the hydrographic and kinematic characteristics of the southern Agulhas Current (Table 3). Since the heat loss from such a ring may be between 80 and 160 W m2 and since there is substantial evaporation in the region, the temperature and salinity of the upper layers of the features are considerably modified near the retroflection region. Rings are about 320 km (7100 km) in diameter at the sea surface. Estimated by the location of their maximum azimuthal speeds, the diameters are a reduced 240 km (740 km). These radial speeds lie between 0.3 and 0.9 m s1. The mean depth of the 101C isotherm in Agulhas rings, a proxy for the geostrophic speed of their water masses, is 650 m (7130 m). Further properties, as estimated by a number of investigators, are given in Table 4. Rings move off into the South Atlantic Ocean at speeds of 4–8 km per day. Maximum translation rates of up to 16 km per day have been observed. They lose 50% of their sea surface height – and therefore by inference of their energy – during the
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AGULHAS CURRENT
first 4 months of their lifetime. A full 40% of rings never seem to leave the Cape Basin, off the southeastern coast of Africa, at all but totally disintegrate here. This decay may well be enhanced by the splitting of rings. This process seems to be largely induced by rings passing over prominent features of the bottom topography such as seamounts. This rapid dissipation of these features means that a considerable part of all the excess salt, heat, energy, and vorticity carried by the rings is deposited exclusively in this corner of the South Atlantic. The remaining rings seem to have lifetimes between 2 and 3 years. They move westward across the full width of the South Atlantic Ocean, slightly to the left of the general background flow. A few of them interact with the upwelling front off the south-eastern coast of Africa, with upwelling filaments occasionally Table 3 Thermohaline characteristics of the principal water masses found at the Agulhas Current retroflection and vicinity Temperature range (1C) Surface Water Central Water South East Atlantic Ocean South West Indian Ocean Antarctic Intermediate Water South East Atlantic Ocean South West Indian Ocean Deep Water North Atlantic Deep Water (SE Atlantic) Circumpolar Deep Water (SW Indian) Antarctic Bottom Water
being wrapped around passing rings. These Agulhas rings play a crucial role in the interbasin exchange of waters between the South Indian and South Atlantic Oceans. This is partially quantified in Table 5.
Agulhas Return Current That part of the Agulhas Current not involved in ring production flows back in an easterly direction on having successfully negotiated the retroflection. Here also there are very high levels of mesoscale variability with substantial meandering and eddies being shed to both sides of the Agulhas Return Current/ Subtropical Convergence. Much of this meandering
Table 5 Estimates of interbasin volume transport between the South Indian and the South Atlantic Oceans caused by ring shedding. The values were calculated by the investigators named here. For full references see De Ruijter et al. (1999) Investigators
Volume transport per ring (106m3s1)
Referenced to
Olson and Evans (1986) Duncombe Rae et al. (1989) Gordon and Haxby (1990)
0.5–0.6 1.2 1.0–1.5 2.0–3.0 0.4–1.1
T > 101C Total T > 101C Total Total
1.1 0.8–1.7 0.45–0.90 0.65 1.0
T > 101C 1000 db 1500 db Total T > 101C
Salinity range (PSU)
16.0 to 26.0
>35.50
6.0 to 16.0 8.0 to 15.0
34.50 to 35.50 34.60 to 35.50
2.0 to 6.0 2.0 to 10.0
33.80 to 34.80 33.80 to 34.80
1.5 to 4.0
34.80 to 35.00
0.1 to 2.0
34.63 to 34.73
0.9 to 1.7
135
McCartney and Woodgate-Jones (1991) Van Ballegooyen et al. (1994) Byrne et al. (1995) Clement and Gordon (1995) Duncombe Rae et al. (1996) Goni et al. (1997)
34.63 to 34.72
After Valentine et al. (1993) Deep-Sea Research 40(6): 1285–1305. ^ Elsevier Science.
After De Ruijter WPM et al. (1999) Indian–Atlantic inter-ocean exchange: dynamics, estimation and impact. Journal of Geophysical Research 104(C9): 20885–20911. ^ American Geophysical Union.
Table 4 Physical properties of Agulhas rings as furnished by a number of independent investigators, calculated with respect to the characteristics of water in the South East Atlantic Ocean Investigators
Olson and Evans (1986) Duncombe Rae et al. (1989) Duncombe Rae et al. (1992) Van Ballegooyen et al. (1994) Byrne et al. (1995) Clement and Gordon (1995) Duncombe Rae et al. (1996) Goni et al. (1997) Garzoli et al. (1996)
Heat flux (103 PW)
Salt flux (105 kg/s)
25
6.3
7.5
Available potential energy (1015 J)
Kinetic energy (1015 J)
30.5–51.4
6.2–8.7
38.8
2.3
18 7.0 11.3 24 2.8–3.8
4.5 7.0 2.0
4.2
1.7
1.1
1.0–1.6
0.7–1.0
After De Ruijter WPM et al. (1999) Indian–Altantic inter-ocean exchange: dynamics, estimation and impact. Journal of Geophysical Research 104(C9): 20885–29911. ^ American Geophysical Union, where full references can be found.
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AGULHAS CURRENT
velocity has reduced to 0.4 m s1 and the volume flux to about 19 106 m3 s1. By a longitude of 701E, at the Kerguelen Plateau, little of the Agulhas characteristics remain along the Subtropical Convergence, all of the Agulhas water having leaked off into the South West Indian Ocean subgyre. The speed, volume transport, as well as mesoscale variability associated with the Agulhas Return Current will all have declined here to values observed in the southeastern Atlantic Ocean, upstream of any influence from the Agulhas Current. The Agulhas Return Current can therefore be considered to have terminated here. The continuing flow along the Subtropical Convergence east of here is known as the South Indian Ocean Current.
80 _1
is brought about by the variable bathymetry over which the current has to pass. The first obstacle to a purely zonal flow for the Agulhas Return Current is the Agulhas Plateau (see Figure 1). Here it carries out a northward meander of 290765 km. Cold eddies are frequently formed here with diameters of 280750 km, but rapidly warm to become indistinguishable from ambient surface waters. Warm eddies may in turn be shed to the south. One such eddy that has been observed closely remained in roughly the same position for 2 months, rotated every 3 days, and had a volume flux of 32 106 m3 s1 to a depth of 1500 m. Downstream of the Agulhas Plateau the next meander lies at a distance of 4507110 km. These meanders move upstream at about half a wavelength per season. Upstream of the Plateau there are also westwardpropagating Rossby waves on the Agulhas Return Current/Subtropical Convergence. However, the direct correlation between the Agulhas Return Current and the Subtropical Convergence is not always straightforward. Sometimes they are in close juxtaposition, sometimes not. When they are not, two separate fronts, or even multiple fronts, may be formed (Table 6). The Agulhas Current has a considerable effect on the Subtropical Convergence, forcing it to lie 51 of latitude farther south than in the South Atlantic Ocean and increasing its surface gradients so that a meridional gradient of 51C in 35 km is not unknown. The Agulhas Return Current starts off with characteristics nearly identical to those of the Agulhas Current. South of Africa it may exhibit a surface speed of 1.3 m s1 and a volume transport of 40 106 m3 s1 to 1000 m. The parallel flow along the Subtropical Convergence would be 16–20 106 m3 s1 at the same time. All these flow characteristics decrease rapidly as the current progresses eastward (Figure 9). By about 551E, the surface
Volume transport (106 m3 s )
136
AJAX SCARC ARC Marathon Discovery SUZIL
70 60 50 40 30 20 10 0
0
10
20
30
40
50
60
70
80
E
Figure 9 The nature of the volume flux along the Subtropical Convergence south of Africa, as established by a number of individual research cruises (shown in the box). Transport is in units of 106 m3 s1 and has been calculated to a depth of 1500 m. The influence of the Agulhas Current is felt from 201E eastwards, but by the 701E meridian it has been dissipated completely. This may therefore be considered the termination of the Agulhas Return Current and the start of the South Indian Ocean Current. (After Lutjeharms and Ansorge Journal of Marine Systems, in press.)
Table 6 The geographic location and the thermal characteristics of the surface expressions of the Agulhas Front as well as the Subtropical Convergence south of Africa. Values for the Agulhas Front were based on 24 crossings, that of the Subtropical Convergence on 70. Values in parentheses denote standard deviations for the calculated averages Latitudinal position From
Agulhas Front Subtropical Convergence
391090 (011160 ) 401350 (011230 )
To
401010 (011060 ) 421360 (011320 )
Temperature Middle
Width (km)
From
To
Middle
Range
Gradient (C km1)
391370 (011140 ) 411400 (011190 )
96.3 69.1 225.1 140.6
21.0 (1.6) 17.9 (2.1)
15.7 (1.5) 10.6 (1.8)
18.4 (1.2) 14.2 (1.7)
5.4 (1.6) 7.3 (1.9)
0.102 0.106 0.047 (0.043)
After Lutjeharms and Valentine (1984) Deep-Sea Research, 31(12): 1461–1476. ^ Elsevier Science.
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AGULHAS CURRENT
Conclusion
137
Further Reading
The Agulhas Current is unusual as a western boundary current for a number of reasons. First, because the African continent terminates at relatively low latitudes, the current penetrates freely into the adjacent ocean basin and a substantial leakage between basins is feasible. Second, through the process of ring and filament shedding, an interaction between a western boundary current and an extensive coastal upwelling regime is brought about that is geographically not possible elsewhere. Third, the very stable nature of the northern Agulhas Current and its characteristic Natal Pulse creates a dynamic environment in which mesoscale disturbances can have profound circulatory effects downstream. The contemporary ignorance about the East Madagascar Current, about the circulation of the Mozambique Channel, and about the origin of midocean eddies in the South West Indian Ocean needs to be eliminated. Only then will a more realistic concept of the interactions between elements of the greater Agulhas Current system become possible.
De Ruijter WPM, Biastoch A, and Drijfhout SS et al. (1999) Indian–Atlantic inter-ocean exchange: dynamics, estimation and impact. Journal of Geophysical Research 104: 20885--20911. Lutjeharms JRE (1996) The exchange of water between the South Indian and the South Atlantic. In: Wefer G, Berger WH, Siedler G, and Webb D (eds.) The South Atlantic: Present and Past Circulation, pp. 125--162. Berlin: Springer-Verlag. Lutjeharms JRE (2001) The Agulhas Current. Berlin: Springer-Verlag. Shannon LV (1985) The Benguela Ecosystem. 1. Evolution of the Benguela, physical features and processes. Oceanography and Marine Biology, An Annual Review 23: 105--182. Shannon LV and Nelson G (1996) The Benguela: large scale features and processes and system variability. In: Wefer G, Berger WH, Siedler G, and Webb D (eds.) The South Atlantic: Present and Past Circulation, pp. 163--210. Berlin: Springer-Verlag.
See also Mesoscale Eddies. Arctic Ocean Circulation. Water Types and Water Masses
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AIRCRAFT REMOTE SENSING L. W. Harding, Jr and W. D. Miller, University of Maryland, College Park, MD, USA R. N. Swift, and C. W. Wright, NASA Goddard Space Flight Center, Wallops Island, VA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 1, pp 113–122, & 2001, Elsevier Ltd.
Introduction The use of aircraft for remote sensing has steadily grown since the beginnings of aviation in the early twentieth century and today there are many applications in the Earth sciences. A diverse set of remote sensing uses in oceanography developed in parallel with advances in aviation, following increased aircraft capabilities and the development of instrumentation for studying ocean properties. Aircraft improvements include a greatly expanded range of operational altitudes, development of the Global Positioning System (GPS) enabling precision navigation, increased availability of power for instruments, and longer range and duration of missions. Instrumentation developments include new sensor technologies made possible by microelectronics, small, high-speed computers, improved optics, and increased accuracy of digital conversion of electronic signals. Advances in these areas have contributed significantly to the maturation of aircraft remote sensing as an oceanographic tool. Many different types of aircraft are currently used for remote sensing of the oceans, ranging from balloons to helicopters, and from light, single engine piston-powered airplanes to jets. The data and information collected on these platforms are commonly used to enhance sampling by traditional oceanographic methods, giving increased spatial and temporal resolution for a number of important properties. Contemporary applications of aircraft remote sensing to oceanography can be grouped into several areas, among them ocean color, sea surface temperature (SST), sea surface salinity (SSS), wave properties, near-shore topography, and bathymetry. Prominent examples include thermal mapping using infrared (IR) sensors in both coastal and open ocean studies, lidar and visible radiometers for ocean color measurements of phytoplankton distributions and ‘algal blooms’, and passive microwave sensors to make observations of surface salinity structure of estuarine plumes in the coastal ocean. These topics
138
will be discussed in more detail in subsequent sections. Other important uses of aircraft remote sensing are to test instruments slated for deployment on satellites, to calibrate and validate space-based sensors using aircraft-borne counterparts, and to make ‘under-flights’ of satellite instruments and assess the efficacy of atmospheric corrections applied to data from space-based observations. Aircraft have some advantages over satellites for oceanography, including the ability to gather data under cloud cover, high spatial resolution, flexibility of operations that enables rapid responses to ‘events’, and less influence of atmospheric effects that complicate the processing of satellite data. Aircraft remote sensing provides nearly synoptic data and information on important oceanographic properties at higher spatial resolution than can be achieved by most satellite-borne instruments. Perhaps the greatest advantage of aircraft remote sensing is the ability to provide consistent, high-resolution coverage at larger spatial scales and more frequent intervals than are practical with ships, making it feasible to use aircraft for monitoring change. Disadvantages of aircraft remote sensing include the relatively limited spatial coverage that can be obtained compared with the global coverage available from satellite instruments, the repeated expense of deploying multiple flights, weather restrictions on operations, and lack of synopticity over large scales. Combination of the large-scale, synoptic data that are accessible from space with higher resolution aircraft surveys of specific locations is increasingly recognized as an important and useful marriage that takes advantages of the strengths of both approaches. This article begins with a discussion of sensors that use lasers (also called active sensors), including airborne laser fluorosensors that have been used to measure chlorophyll (chl-a) and other properties; continues with discussions of lidar sensors used for topographic and bathymetric mapping; describes passive (sensors that do not transmit or illuminate, but view naturally occurring reflections and emissions) ocean color remote sensing directed at quantifying phytoplankton biomass and productivity; moves to available or planned hyper-spectral aircraft instruments; briefly describes synthetic aperture radar applications for waves and wind, and closes with a discussion of passive microwave measurements of salinity. Readers are directed to the Further Reading section if they desire additional information on individual topics.
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AIRCRAFT REMOTE SENSING
Active Systems Airborne Laser Fluorosensing
The concentrations of certain waterborne constituents, such as chl-a, can be measured from their fluorescence, a relationship that is exploited in shipboard sensors such as standard fluorometers and flow cytometers that are discussed elsewhere in this encyclopedia. NASA first demonstrated the measurement of laser-induced chl-a fluorescence from a low-flying aircraft in the mid-1970s. Airborne laser fluorosensors were developed shortly thereafter in the USA, Canada, Germany, Italy, and Russia, and used for measuring laser-induced fluorescence of a number of marine constituents in addition to chl-a. Oceanic constituents amenable to laser fluorosensing include phycoerythrin (photosynthetic pigment in some phytoplankton taxa), chromophoric dissolved organic matter (CDOM), and oil films. Airborne laser fluorosensors have also been used to follow dyes such as fluorescein and rhodamine that are introduced into water masses to trace their movement. The NASA Airborne Oceanographic Lidar (AOL) is the most advanced airborne laser fluorosensor. The transmitter portion features a dichroic optical device to spatially separate the temporally concurrent 355 and 532 nm pulsed-laser radiation, followed by individual steering mirrors to direct the separated beams to respective oceanic targets separated by B1 m when flown at the AOL’s nominal 150 m operational altitude. The receiver focal plane containing both laser-illuminated targets is focused onto the input slits of the monochromator. The monochromator output focal planes are viewed by custommade optical fibers that transport signal photons from the focal planes to the photo-cathode of each photo-multiplier module (PMM) where the conversion from photons to electrons takes place with a substantial gain. Time-resolved waveforms are collected in channels centered at 404 nm (water Raman) and 450 nm (CDOM) from 355 nm laser excitation, and at 560 nm and 590 nm (phycoerythrin), 650 (water Raman), and 685 nm (chl-a) from the 532 nm laser excitation. The water Raman from the respective lasers is the red-shifted emission from the OH bonds of water molecules resulting from radiation with the laser pulse. The strength of the water Raman signal is directly proportional to the number of OH molecules accessed by the laser pulse. Thus, the water Raman signal is used to normalize the fluorescence signals to correct for variations in water attenuation properties in the surface layer of the ocean. The AOL is described in more detail on http:// lidar.wff.nasa.gov.
139
The AOL has supported major oceanographic studies throughout the 1980s and 1990s, extending the usefulness of shipboard measurements over wide areas to permit improved interpretation of the shipderived results. Examples of data from the AOL show horizontal structure of laser-induced fluorescence converted to chl-a concentration (Figure 1). Prominent oceanographic expeditions that have benefited from aircraft coverage with the AOL include the North Atlantic Bloom Experiment (NABE) of the Joint Global Ocean Flux Study (JGOFS), and the Iron Enrichment Experiment (IRONEX) of the Equatorial Pacific near the Galapagos Islands. This system has been flown on NASA P-3B aircraft in open-ocean missions that often exceeded 6 h in duration and collected hundreds of thousands of spectra. Each ‘experiment’ is able to generate both active and passive data in ‘pairs’ that are used for determining ocean color and recovering chl-a and other constituents, and that are also useful in the development of algorithms for measuring these constituents from oceanic radiance spectra. Airborne Lidar Coastal Mapping
The use of airborne lidar (light detection and ranging) sensors for meeting coastal mapping requirements is a relatively new and promising application of laser-ranging technology. These applications include high-density surveying of coastline and beach morphology and shallow water bathymetry. The capability to measure distance accurately with lidar sensors has been available since the early 1970s, but their application to airborne surveying of terrestrial features was seriously hampered by the lack of knowledge of the position of the aircraft from which the measurement was made. The implementation of the Department of Defense GPS constellation of satellites in the late 1980s, coupled with the development of GPS receiver technology, has resulted in the capability to provide the position of a GPS antenna located on an aircraft fuselage in flight to an accuracy approaching 5 cm using kinematic differential methodology. These methods involve the use of a fixed receiver (generally located at the staging airport) and a mobile receiver that is fixed to the aircraft fuselage. The distance between mobile and fixed receivers, referred to as the baseline, is typically on the order of tens of kilometers, and can be extended to hundreds of kilometers by using dual frequency survey grade GPS receivers aided by tracking the phase code of the carrier from each frequency. Modern airborne lidars are capable of acquiring 5000 or more discrete range measurements per second. Aircraft attitude and heading information are
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AIRCRAFT REMOTE SENSING
Figure 1 Cross-section profiles flown across a large oceanic front west of the Galapagos Islands on 25 October and 3 November 1993. (A) A dramatic horizontal displacement of the front as measured with an infrared radiometer; (B) and (C) show corresponding changes in laser-induced fluorescence of chl-a and phycoerythrin. This flight was made as part of the original IRONEX investigation in late 1993. NFU, normalized fluorescence units.
used along with the GPS-determined platform position to locate the position of the laser pulse on the Earth’s surface to a vertical accuracy approaching 10 cm with some highly accurate systems and o30 cm for most of these sensors. The horizontal accuracy is generally 50–100 cm. Depending on the pulse repetition rate of the laser transmitter, the off-nadir pointing angle, and the speed of the aircraft platform, the density of survey points can exceed one sample per square meter. NASA’s Airborne Topographic Mapper (ATM) is an example of a topographic mapping lidar used for coastal surveying applications. An example of data from ATM shows shoreline features off the east coast of the USA (Figure 2). ATM was originally developed to measure changes in the elevation of Arctic ice sheets in response to global warming. The sensor was applied to measurements of changes in coastal morphology beginning in 1995. Presently, baseline
topographic surveys exist for most of the Atlantic and Gulf coasts between central Maine and Texas and for large sections of the Pacific coast. Affected sections of coastline are re-occupied following major coastal storms, such as hurricanes and ‘Nor’easters’, to determine the extent of erosion and depositional patterns resulting from the storms. Additional details on the ATM and some results of investigations on coastal morphology can be found on websites (http:// lidar.wff.nasa.gov and http://aol.wff.nasa.gov/aoltm/ projects/beachmap/98results/). Other airborne lidar systems have been used to survey coastal morphology, including Optec lidar systems by Florida State University and the University of Texas at Austin. Beyond these airborne lidar sensors, there are considerably more instruments with this capability that are currently in use in the commercial sector for surveying metropolitan areas, flood plains, and for other terrestrial
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AIRCRAFT REMOTE SENSING
141
Figure 2 (Left) Map of coastal topography around Pacifica, CA, USA, derived from lidar data obtained in April 1998, after a winter of severe storms associated with El Nin˜o. Map insert shows the Esplanade Drive area of Pacifica rendered from lidar data gridded at 2 m resolution and colored according to elevation. (Right) Cross-sections derived from lidar data of October 1997 and April 1998 at locations marked in the map inset. The profiles in (A) show a stable cliff and accreting beach, whereas about 200 m to the south the profiles in (B) show erosion of the sea cliff and adjacent beach resulting in undermining of houses. Each profile shows individual laser spot elevations that fall within a 2 m wide strip oriented approximately normal to shoreline. (Reproduced with permission from Sallenger et al., 1999.)
applications. At the last count (early 2000) there were approximately 60 airborne lidars in operation worldwide, with most engaged in a variety of survey applications generally outside the field of coastal mapping. Pump and Probe Fluorometry
Several sections in this chapter describe recoveries of phytoplankton biomass as chl-a by active and passive measurements. Another recent accomplishment is an active, airborne laser measurement intended to aid in remote detection of photosynthetic performance, an important ingredient of primary productivity computations. Fluorometric techniques, such as fast repetition rate (FRR) fluorometry (explained in another article of this encyclopedia), have provided an alternative approach to 14C assimilation and O2 evolution in the measurement of primary productivity. This
technology has matured with the commercial availability of FRR instruments that can give vertical profiles or operate in a continuous mode while underway. There have been several attempts to develop airborne lidar instruments to determine phytoplankton photosynthetic characteristics from aircraft in the past decade. NASA scientists have deployed a pump and probe fluorometer from aircraft, wherein the AOL laser (described above) acts as the pump and a second laser with variable power options and rapid pulsing capabilities (10 ns) functions as the probe.
Passive Systems Multichannel Ocean Color Sensor (MOCS)
Passive ocean color measurements using visible radiometers to measure reflected natural sunlight from the ocean have been made with a number of
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AIRCRAFT REMOTE SENSING
instruments in the past two decades. These instruments include the Multichannel Ocean Color Sensor (MOCS) that was flown in studies of Nantucket Shoals in the early 1980s, the passive sensors of the AOL suite that have been used in many locations around the world, and more recently, simple radiometers that have been deployed on light aircraft in regional studies of Chesapeake Bay (see below). MOCS was one of the earliest ocean color sensors used on aircraft. It provided mesoscale data on shelf and slope chl-a in conjunction with shipboard studies of physical structure, nutrient inputs, and phytoplankton primary productivity. Ocean Data Acquisition System (ODAS) and SeaWiFS Aircraft Simulator (SAS)
Few aircraft studies have obtained long time-series sufficient to quantify variability and detect secular trends. An example is ocean color measurements made from light aircraft in the Chesapeake Bay region for over a decade, providing data on chl-a and SST from 4250 flights. Aircraft over-flights of the Bay using the Ocean Data Acquisition System (ODAS) developed at NASA’s Goddard Space Flight Center commenced in 1989. ODAS was a nadirviewing, line-of-flight, three-band radiometer with spectral coverage in the blue-green region of the visible spectrum (460–520 nm), a narrow 1.51 fieldof-view, and a 10 Hz sampling rate. The ODAS instrument package included an IR temperature sensor (PRT-5, Pyrometrics, Inc.) for measuring SST. The system was flown for B7 years over Chesapeake Bay on a regular set of tracks to determine chl-a and SST. Over 150 flights were made with ODAS between 1989 and 1996, coordinated with in situ observations from a multi-jurisdictional monitoring program and other cruises of opportunity. ODAS was flown together with the SeaWiFS Aircraft Simulator (SAS II, III, Satlantic, Inc., Halifax, Canada) beginning in 1995 and was retired soon thereafter and replaced with the SAS units. SAS III is a multi-spectral (13-band, 380–865 nm), line-offlight, nadir viewing, 10 Hz, passive radiometer with a 3.51 field-of-view that has the same wavebands as the SeaWiFS satellite instrument, and several additional bands in the visible, near IR, and UV. The SAS systems include an IR temperature sensor (Heimann Instruments, Inc.). Chl-a estimates are obtained using a curvature algorithm applied to water-leaving radiances at wavebands in the blue-green portion of the visible spectrum with validation from concurrent shipboard measurements. Flights are conducted at B50–60 m s1 (100–120 knots), giving an alongtrack profile with a resolution of 5–6 m averaged to
50 m in processing, and interpolated to 1 km2 for visualization. Imagery derived from ODAS and SAS flights is available on a web site of the NOAA Chesapeake Bay Office for the main stem of the Bay (http://noaa.chesapeakebay.net), and for two contrasting tributaries, the Choptank and Patuxent Rivers on a web site of the Coastal Intensive Sites Network (CISNet) (http://www.cisnet-choptank.org). Data from ODAS and SAS have provided detailed information on the timing, position, and magnitude of blooms in Chesapeake Bay, particularly the spring diatom bloom that dominates the annual phytoplankton cycle. This April–May peak of chl-a represents the largest accumulation of phytoplankton biomass in the Bay and is a proximal indicator of over-enrichment by nutrients. Data from SeaWiFS for spring 2000 show the coast-wide chl-a distribution for context, while SAS III data illustrate the high-resolution chl-a maps that are obtained regionally (Figure 3). A well-developed spring bloom corresponding to a year of relatively high freshwater flow from the Susquehanna River, the main tributary feeding the estuary, is apparent in the main stem Bay chl-a distribution. Estimates of primary productivity are now being derived from shipboard observations of key variables combined with high-resolution aircraft measurements of chl-a and SST for the Bay.
Hyper-spectral Systems Airborne Visible/Infrared Imaging Spectrometer (AVIRIS)
The Airborne Visible/Infrared Imaging Spectrometer (AVIRIS) was originally designed in the late 1980s by NASA at the Jet Propulsion Laboratory (JPL) to collect data of high spectral and spatial resolution, anticipating a space-based high-resolution imaging spectrometer (HIRIS) that was planned for launch in the mid-1990s. Because the sensor was designed to provide data similar to satellite data, flight specifications called for both high altitude and high speed. AVIRIS flies almost exclusively on a NASA ER-2 research aircraft at an altitude of 20 km and an airspeed of 732 km h1. At this altitude and a 301 field of view, the swath width is almost 11 km. The instantaneous field of view is 1 mrad, which creates individual pixels at a resolution of 20 m2. The sensor samples in a whiskbroom fashion, so a mirror scans back and forth, perpendicular to the line-of-flight, at a rate of 12 times per second to provide continuous spatial coverage. Each pixel is then sent to four separate spectrometers by a fiber optic cable. The spectrometers are arranged so that they each cover a part of the
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AIRCRAFT REMOTE SENSING
Chesapeake Bay chlorophyll 14 May 2000
Mid-Atlantic chlorophyll 8 May 2000 64.0
_ Chl [mg m 3]
39˚30'
40 39
10.0
80 60
38 39˚00'
1.0
40
37
30
0.1
36
25
38˚30'
35 0.01
34
143
76
74
72
20
70 16
(A)
12
38˚00'
10 8 37˚30'
6 4 2
37˚00' 77˚00'
(B)
76˚30'
76˚00'
75˚30'
Patuxent chlorophyll 1 May 2000 38˚35' Choptank chlorophyll 1 May 2000 38˚45'
38˚30'
38˚40'
38˚25'
38˚35'
38˚20'
(C)
76˚40'
76˚30'
(D)
76˚20'
76˚10'
76˚00'
Figure 3 Spring chl-a (mg m3) in: (A) the mid-Atlantic region from SeaWiFS; (B) Chesapeake Bay; (C) Patuxent R; (D) Choptank R from SAS III.
spectrum from 0.40 to 2.4 mm, providing continuous spectral coverage at 10 nm intervals over the entire spectrum from visible to near IR. Data are recorded to tape cassettes for storage until rectification, atmospheric correction, and processing at JPL. A typical AVIRIS ‘scene’ is a 40 min flight line. At ER-2
flight parameters, this creates an image roughly 500 km long and 11 km wide. Data are encoded at 12-bits for a high degree of discrimination. The physical dimensions of AVIRIS are quite large, 84 cm wide 160 cm long 117 cm tall at a weight of 720 pounds.
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AIRCRAFT REMOTE SENSING
Data collected with AVIRIS have been used for terrestrial, marine, and atmospheric applications. Accomplishments of AVIRIS include separation of the chl-a signature from bottom reflectance for clear lake waters of Lake Tahoe and turbid waters near Tampa Bay, interpretation of spectral signals from resuspended sediment and dissolved organic materials in W. Florida, and of suspended sediment and kelp beds in S. California. Recent efforts have focused on improving atmospheric correction procedures for both AVIRIS and satellite data, providing inputs for bio-optical models which determine inherent optical properties (IOPs) from reflectance, algorithm development, and sporadic attempts at water quality monitoring (e.g., chl-a, suspended sediment, diffuse attenuation coefficient, kd). AVIRIS data have recently been used as an input variable to a neural network model developed to estimate water depth. The model was able to separate the contributions of different components to the total waterleaving radiance and to provide relatively accurate estimates of depth (rms error ¼ 0.48 m). Compact Airborne Spectrographic Imager (CASI)
The Compact Airborne Spectrographic Imager (CASI) is a relatively small, lightweight hyper-spectral sensor that has been used on a variety of light aircraft. CASI was developed by Itres Research Ltd (Alberta, Canada) in 1988 and was designed for a variety of remote sensing applications in forestry, agriculture, land-use planning, and aquatic monitoring. By allowing user-defined configurations, the 12-bit, push-broom-type sensor (333 scan lines per second) using a charge-coupled detector (CCD) can be adapted to maximize either spatial (37.81 across track field of view, 0.0771 along-track, 512 pixels – pixel size varies with altitude) or spectral resolution (288 bands at 1.9 nm intervals between 400 and 1000 nm). Experiments have been conducted using CASI to determine bottom type, benthic cover, submerged aquatic vegetation, marsh type, and in-water constituents such as suspended sediments, chl-a, and other algal pigments. Portable Hyper-spectral Imager for Low-Light Spectroscopy (PHILLS)
The Portable Hyper-spectral Imager for Low-Light Spectroscopy (PHILLS) has been constructed by the US Navy (Naval Research Laboratory) for imaging the coastal ocean. PHILLS uses a backside-illuminated CCD for high sensitivity, and an all-reflective spectrograph with a convex grating in an Offner configuration to produce a distortion-free image. The instrument benefits from improvements in large-
format detector arrays that have enabled increased spectral resolution and higher signal-to-noise ratios for imaging spectrographs, extending the use of this technology in low-albedo coastal waters. The ocean PHILLS operates in a push-broom scanned mode whereby cross-track ground pixels are imaged with a camera lens onto the entrance slit of the spectrometer, and new lines of the along-track ground pixels are attained by aircraft motion. The Navy’s interest in hyper-spectral imagers for coastal applications centers on the development of methods for determining shallow water bathymetry, topography, bottom type composition, underwater hazards, and visibility. PHILLS precedes a planned hyper-spectral satellite instrument, the Coastal Ocean Imaging Spectrometer (COIS) that is planned to launch on the Naval Earth Map Observer (NEMO) spacecraft.
Radar Altimetry Ocean applications of airborne radar altimetry systems include several sensors that retrieve information on wave properties. Two examples are the Radar Ocean Wave Spectrometer (ROWS), and the Scanning Radar Altimeter (SRA), systems designed to measure long-wave directional spectra and near-surface wind speed. ROWS is a Ku-band system developed at NASA’s Goddard Space Flight Center in support of present and future satellite radar missions. Data obtained from ROWS in a spectrometer mode are used to derive two-dimensional ocean spectral wave estimates and directional radar backscatter. Data from the pulse-limited altimeter mode radar yield estimates of significant wave height and surface wind speed.
Synthetic Aperture Radar (SAR) Synthetic Aperture Radar (SAR) systems emit microwave radiation in several bands and collect the reflected radiation to gain information about sea surface conditions. Synthetic aperture is a technique that is used to synthesize a long antenna by combining signals, or echoes, received by the radar as it moves along a flight track. Aperture refers to the opening that is used to collect reflected energy and form an image. The analogous feature of a camera to the aperture would be the shutter opening. A synthetic aperture is constructed by moving a real aperture or antenna through a series of positions along a flight track. NASA’s Jet Propulsion Laboratory and Ames Research Center have operated the Airborne SAR (AIRSAR) on a DC-8 since the late 1980s. The radar
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AIRCRAFT REMOTE SENSING
of AIRSAR illuminates the ocean at three microwave wavelengths: C-band (6 cm), L-band (24 cm), and Pband (68 cm). Brightness of the ocean (the amount of energy reflected back to the antenna) depends on the roughness of the surface at the length scale of the microwave (Bragg scattering). The primary source of roughness, and hence brightness, at the wavelengths used is capillary waves associated with wind. Oceanographic applications derive from the responsiveness of capillary wave amplitude to factors that affect surface tension, such as swell, atmospheric stability, and the presence of biological films. For example, the backscatter characteristics of the ocean are affected by surface oil and slicks can be observed in SAR imagery as a decrease of radar backscatter; SAR imagery appears dark in an area affected by an oil spill, surface slick, or biofilm, as compared with areas without these constituents.
Microwave Salinometers Passive microwave radiometry (L-band) has been tested for the recovery of SSS from aircraft, and it may be possible to make these measurements from space. Salinity affects the natural emission of EM radiation from the ocean, and the microwave signature can be used to quantify SSS. Two examples of aircraft instruments that have been used to measure SSS in the coastal ocean are the Scanning Low-Frequency
145
Microwave Radiometer (SLFMR), and the Electronically Thinned Array Radiometer (ESTAR). SLFMR was used recently in the estuarine plume of Chesapeake Bay on the east coast of the USA to follow the buoyant outflow that dominates the nearshore density structure and constitutes an important tracer of water mass movement. SLFMR is able to recover SSS at an accuracy of about 1 PSU (Figure 4). This resolution is too coarse for the open ocean, but is quite suitable for coastal applications where significant gradients occur in regions influenced by freshwater inputs. SLFMR has a bandwidth of 25 MHz, a frequency of 1.413 GHz, and a single antenna with a beam width of approximately 161 and six across-track positions at 761, 7221 and 7391. Tests of SLFMR off the Chesapeake Bay demonstrated its effectiveness as a ‘salinity mapper’ by characterizing the trajectory of the Bay plume from surveys using light aircraft in joint operations with ships. Flights were conducted at an altitude of 2.6 km, giving a resolution of about 1 km. The accuracy of SSS in this example is B0.5 PSU. ESTAR is an aircraft instrument that is the prototype of a proposed space instrument for measuring SSS. This instrument relies on an interferometric technique termed ‘aperture synthesis’ in the across-track dimension that can reduce the size of the antenna aperture needed to monitor SSS from space. It has been described as a ‘hybrid of a real and a synthetic aperture radiometer.’ Aircraft surveys of
(PSU)
Figure 4 Sea surface salinity from an airborne microwave salinity instrument for (A) 14 September 1996; (B) 20 September 1996. Images reveal strong onshore-offshore gradients in salinity from the mouth of Chesapeake Bay to the plume and shelf, and the effect of high rainfall and freshwater input on the salinity distribution over a 1-week interval. (Adapted with permission from Miller et al., 1998.)
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AIRCRAFT REMOTE SENSING
SSS using ESTAR in the coastal current off Maryland and Delaware showed good agreement with thermosalinograph measurements from ships in the range of 29–31 PSU.
See also Beaches, Physical Processes Affecting. Bio-Optical Models. Fluorometry for Biological Sensing. Inherent Optical Properties and Irradiance. IR Radiometers. Iron Fertilization. Optical Particle Characterization. Phytoplankton Blooms. Primary Production Distribution. River Inputs. Satellite Oceanography, History and Introductory Concepts. Satellite Remote Sensing: Ocean Color. Satellite Remote Sensing of Sea Surface Temperatures. Satellite Remote Sensing: Salinity Measurements. Upper Ocean Time and Space Variability.
Further Reading Blume H-JC, Kendall BM, and Fedors JC (1978) Measurement of ocean temperature and salinity via
microwave radiometry. Boundary Layer Meteorology 13: 295--308. Campbell JW and Esaias WE (1985) Spatial patterns in temperature and chlorophyll on Nantucket Shoals from airborne remote sensing data, May 7–9, 1981. Journal of Marine Research 43: 139--161. Harding LW Jr, Itsweire EC, and Esaias WE (1994) Estimates of phytoplankton biomass in the Chesapeake Bay from aircraft remote sensing of chlorophyll concentrations, 1989–92. Remote Sensing Environment 49: 41--56. Le Vine DM, Zaitzeff JB, D’Sa EJ, et al. (2000) Sea surface salinity: toward an operational remote-sensing system. In: Halpern D (ed.) Satellites, Oceanography and Society, pp. 321--335. Elsevier Science. Miller JL, Goodberlet MA, and Zaitzeff JB (1998) Airborne salinity mapper makes debut in coastal zone. EOS Transactions of the American Geophysical Union 79: 173--177. Sallenger AH Jr, Krabill W, Brock J, et al. (1999) EOS Transactions of the American Geophysical Union 80: 89--93. Sandidge JC and Holyer RJ (1998) Coastal bathymetry from hyper-spectral observations of water radiance. Remote Sensing Environment 65: 341--352.
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AIR–SEA GAS EXCHANGE a wavy water surface are still not known. A number of new imaging techniques are described which give direct insight into the transfer processes and promise to trigger substantial theoretical progress in the near future.
B. Ja¨hne, University of Heidelberg, Heidelberg, Germany & 2009 Elsevier Ltd. All rights reserved.
Introduction
Theory
The exchange of inert and sparingly soluble gases, including carbon dioxide, methane, and oxygen, between the atmosphere and oceans is controlled by a 20–200-mm-thick boundary layer at the top of the ocean. The hydrodynamics in this layer is significantly different from boundary layers at rigid walls since the orbital motion of the waves is of the same order as the velocities in the viscous boundary layer. Laboratory and field measurements show that wind waves and surfactants significantly influence the gastransfer process. Because of limited experimental techniques, the details of the mechanisms and the structure of the turbulence in the boundary layer at
Mass Boundary Layers
Table 1
The transfer of gases and volatile chemical species between the atmosphere and oceans is driven by a concentration difference and the transport by molecular and turbulent motion. Both types of transport processes can be characterized by ‘diffusion coefficients’, denoted by D and Kc, respectively (Table 1). The resulting flux density jc is proportional to the diffusion coefficient and the concentration gradient. Thus, jc ¼ ðD þ Kc ðzÞÞrc
½1
Diffusion coefficients for various gases and volatile chemical species in deionized water and in some cases in seawater
Species
Heat 3 Heb,c 4 He 4 Hea Ne Kr Xe 222 Rnb H2 H2a CH4 CO2 DMSb CH3Brb F12b (CCl2F2) F11b (CCl3F) SF6b
Molecular mass
3.02 4.00 20.18 83.80 131.30 222.00 2.02 16.04 44.01 62.13 94.94 120.91 137.37 146.05
A (10 5 cm2 s 1)
379.2 941 818 886 1608 6393 9007 15 877 3338 1981 3047 5019 2000 3800 4100 3400 2900
Ea (kJ mol 1)
2.375 11.70 11.70 12.02 14.84 20.20 21.61 23.26 16.06 14.93 18.36 19.51 18.10 19.10 20.50 20.00 19.30
s(Fit) (%)
2.1 2.1 1.8 3.5 1.6 3.5 11 1.6 4.3 2.7 1.3
Diffusion coefficient (10 5 cm2 s 1) 5 1C
15 1C
25 1C
35 1C
135.80 5.97 5.10 4.86 2.61 1.02 0.77 0.68 3.17 3.05 1.12 1.07 0.80 0.98 0.58 0.60 0.69
140.72 7.12 6.30 5.88 3.28 1.41 1.12 0.96 4.10 3.97 1.48 1.45 1.05 1.31 0.79 0.81 0.92
145.48 8.39 7.22 7.02 4.16 1.84 1.47 1.34 5.13 4.91 1.84 1.91 1.35 1.71 1.05 1.07 1.20
150.08 9.77 8.48 8.03 4.82 2.40 1.94 1.81 6.23 5.70 2.43 2.43 1.71 2.20 1.37 1.38 1.55
a
In seawater. Values of diffusion coefficients from fit, not measured values. c Set 15% higher than 4He. Columns 3 and 4 contain the parameters for the fit of the diffusion coefficient: D ¼ A exp[ Ea/(RT )], the last four columns the diffusion coefficients for 5, 15, 25, and 35 1C. Data collected from Ja¨hne B, Heinz G, and Dietrich W (1987) Measurement of the diffusion coefficients of sparingly soluble gases in water. Journal of Geophysical Research 92: 10767–10776; and King DB, De Bryun WJ, Zheng M, and Saltzman ES (1995) Uncertainties in the molecular diffusion coefficient of gases in water for use in the estimation of air–sea exchange. In: Ja¨hne B and Monahan E (eds.) Air–Water Gas Transfer, pp. 13–22. Hanau: Aeon. b
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147
148
AIR–SEA GAS EXCHANGE
In a stationary homogeneous case and without sinks and sources by chemical reactions, the flux density j is in vertical direction and constant. Then integration of [1] yields vertical concentration profiles Z 0
Zr
1 dz D þ Kc ðzÞ
The molecular diffusion coefficient is proportional to the velocity of the molecules and the free length between collisions. The same concept can be applied to turbulent diffusion coefficients. Far away from the interface, the free length (called ‘mixing length’) is set proportional to the distance from the interface and the turbulent diffusion coefficient Kc for mass transfer is Kc ¼
k u z Sct
Air-side mass boundary layer (100–1000 μm)
½2 Kc < D
Ca
C Water C w s = Ca s surface Aqueous mass boundary layer (20–200 μm)
s
Viscous boundary layer (600–2000 μm)
Water phase
zr
Cw
Reference level b
z
½3
where k ¼ 0.41 is the von Ka´rma´n constant, u , the friction velocity, a measure for the velocity fluctuations in a turbulent flow, and Sct ¼ Km/Kc the turbulent Schmidt number. Closer to the interface, the turbulent diffusion coefficients are decreasing even faster. Once a critical length scale l is reached, the Reynolds number Re ¼ u l=v (n is the kinematic viscosity, the molecular diffusion coefficient for momentum) becomes small enough so that turbulent motion is attenuated by viscosity. The degree of attenuation depends on the properties of the interface. At a smooth solid wall, Kcpz3, at a free water interface it could be in the range between Kcpz3 and Kcpz2 depending on surface conditions. Boundary layers are formed on both sides of the interface (Figure 1). When the turbulent diffusivity becomes equal to the kinematic viscosity, the edge of the ‘viscous boundary layer’ is reached. As the name implies, this layer is dominated by viscous dissipation and the velocity profile becomes linear because of a constant diffusivity. The edge of the ‘mass boundary layer’ is reached when the turbulent diffusivity becomes equal to the molecular diffusivity. The relative thickness of both boundary layers depends on the dimensionless ratio Sc ¼ v/D (Schmidt number). The viscous and mass boundary layers are of about the same thickness in the air, because values of D for various gaseous species and momentum are about the same (Scair is 0.56 for H2O, 0.63 for heat, and 0.83 for CO2). In the liquid phase the situation is completely different. With Schmidt numbers in the range from 100 to 3000 (Figure 2 and Table 2), molecular diffusion for a dissolved volatile chemical species is two to three orders of magnitude slower than diffusion of momentum. Thus the mass boundary layer
z∼
Kc < υ
Figure 1 Schematic graph of the mass boundary layers at a gas–liquid interface for a tracer with a solubility a ¼ 3. 1
102
10
103
104
0
105
0
Air-side control
H2O
Atrazine
30
0 Pentachlorophenol
Momentum
DDT
103
3)
Sm
Transition zone
n= ce (
2/ (n =
SO2 (pH<3)
1/2)
urfa
gh s
Rou
ce
urfa
hs
oot
102
10
104
Aniline
Heat
103
105
SO2 (pH=6) Phenol DEHP
40 40 Chloroaniline
104
Solubility
Cðzr Þ Cð0Þ ¼ jc
Air phase
Ethylacetate 0
102
Naphthalene
40
Trichlorobenzene Benzene Hexachlorobenzene H2S Trichloroethylene 0
1
1
CO2
40
10
0 Rn
40
Waterside control
O2 40
1
35 40
He
CH4
5 0
CO 0
25
SF6
10 102 103 Schmidt number Sc = /D
5
104
Figure 2 Schmidt number/solubility diagram including various volatile tracers, momentum, and heat for a temperature range (1C) as indicated. Filled circles refer to only a temperature of 20 1C. The regions for air-side, mixed, and waterside control of the transfer process between gas and liquid phase are marked. At the solid lines the transfer resistance is equal in both phases. The following dimensional transfer resistances were used: ra ¼ 31, rw ¼ 12Sc2/3 (smooth), rw ¼ 6.5Sc1/2 (wavy surface) with ra ¼ Rau * a and rw ¼ Rwu * w. Adapted from Ja¨hne B and HauXecker H (1998) Air–water gas exchange. Annual Review of Fluid Mechanics 30: 443–468.
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AIR–SEA GAS EXCHANGE
Table 2
Schmidt numbers of various gases and volatile species in the temperature range from 0 to 40 1C
Species
Heat 3 He 4 He 4 Hea Ne Kr Xe 222 Rn H2 H2a CH4 CO2 DMS CH3Br F12 (CCl2F2) F11 (CCl3F) SF6 a
149
Schmidt number 0 1C
5 1C
10 1C
15 1C
20 1C
25 1C
30 1C
35 1C
40 1C
13.45 329 379 402 768 2045 2701 3168 633 648 1908 1922 2593 2120 3639 3521 3033
11.19 254 293 310 579 1478 1930 2235 473 488 1400 1397 1905 1545 2624 2549 2208
9.46 200 230 244 445 1090 1409 1611 360 375 1047 1036 1428 1150 1931 1883 1640
8.10 160 184 194 347 819 1047 1182 278 293 797 782 1089 870 1447 1416 1239
7.02 130 149 157 276 625 791 883 219 232 616 600 844 669 1102 1082 952
6.14 107 122 129 221 483 606 669 174 186 483 466 662 522 851 839 741
5.42 88 102 107 180 379 471 514 140 151 383 367 527 412 666 658 585
4.82 74 85 89 148 301 370 400 114 124 308 293 423 329 527 523 467
4.32 63 72 75 122 241 294 314 94 103 250 236 344 266 422 420 376
In seawater.
is significantly thinner than the viscous boundary layer in the liquid phase. This means that the transfer of gases is much slower and almost entirely controlled by the tiny residual turbulence in the small top fraction of the viscous boundary layer. This basic fact makes it difficult to investigate the mechanism of air– water gas transfer both theoretically and experimentally. In addition, the transfer process depends strongly on the water temperature because the Schmidt number decreases by about a factor of 6 from 0 to 35 1C (Table 2). Transfer velocity and transfer resistance The amount of species exchanged between the air and water across the interface can be described by a quantity with the units of a velocity. It represents the velocity with which a tracer is pushed by an imaginary piston across the surface. This quantity is known as the ‘transfer velocity’ k (also known as the ‘piston velocity’, ‘gas exchange rate’, or ‘transfer coefficient’). It is defined as the flux density divided by the concentration difference between the surface and the bulk at some reference level zr: k¼
jc Cs Cb
Cs Cb jc
1 R¼ ¼ k
½5
Z
Zr
0
1 dz D þ Kc ðzÞ
½6
Thus the transfer resistances over several height intervals can be added in the same way as electrical resistances that are connected in series. Typical values of the transfer velocity across the waterside mass boundary layer are 10 6–10 5 m s 1 (1–10 m d 1). With respect to typical mixed layer depths in the ocean of about 100 m, gas transfer is a very slow process. It takes a time constant t ¼ h/k ¼ 10–100 days for the concentration of dissolved gases in the mixed layer to come into equilibrium with the atmosphere. Boundary layer thickness The ‘boundary layer thickness’ z˜ is defined as the thickness of a fictional layer in which the flux is maintained only by molecular transport: j ¼ D(Cs Cb)/z˜. Then with [4] the boundary layer thickness is given by
½4
The inverse of the transfer velocity is known as the ‘transfer resistance’ R: R¼
The indices s and b denote the surface and bulk. Both quantities can directly be related to vertical concentration profiles by introducing [4] and [5] into [2]:
z˜ ¼
D k
½7
Geometrically, z˜ is given as the intercept of the tangent to the concentration profile at the surface and the bulk (Figure 1). With thicknesses between 20 and 200 mm, the mass boundary layer is extremely thin.
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AIR–SEA GAS EXCHANGE
Boundary layer time constant The ‘time constant’ t˜ for the transport across the mass boundary layer is given by t˜ ¼
z˜ D ¼ k k2
½8
Typical values for t˜ are 0.04–4 s. Any chemical reaction with a time constant larger than t˜ does not significantly affect the transfer process. Therefore, CO2 can be regarded as an inert gas and not fast hydrating acid gases such as SO2. The definitions of the three parameters k, z˜, and t˜, are generally valid and do not depend on any models of the boundary layer turbulence. According to [7] and [8] they are coupled via the molecular diffusion coefficient. Therefore only one of them needs to be measured to get knowledge of all three parameters, provided the molecular diffusion coefficient of the species is known. Partitioning of Transfer between Air and Water
Because a mass boundary layer exists on both sides, it is important to determine which one controls the transfer, that is, exhibits the largest transfer resistance (or lowest transfer velocity). At the surface itself the thermodynamic solubility equilibrium is assumed to be established between the tracer concentrations ca in the gas phase and cw in the liquid phase cws ¼ acas
½9
where a is the dimensionless solubility (partition coefficient). A solubility aa1 causes a concentration jump at the surface (Figure 1). Thus, the resulting total transfer velocity kt or transfer resistance Rt can be viewed from either the gas phase or the liquid phase. Adding them up, the factor a must be considered to conserve the continuity of the concentration profile: 1 1 1 ¼ þ ; Rat ¼ Ra þ Rw =a kat ka akw 1 a 1 waterside: ¼ þ ; Rwt ¼ aRa þ Rw kwt ka kw air side:
½10 The total transfer velocities in air and water differ by the factor a: kat ¼ akwt. The ratio akw/ka determines which boundary layer controls the transfer process. A high solubility shifts control of the transfer process to the gas-phase boundary layer, and a low solubility to the aqueous layer. The solubility value for a transition from air-sided to watersided control depends on the ratio of the transfer velocities. Typically, kw is about
100–1000 times smaller than ka. Thus, the transfer of even moderately soluble volatile chemical species with solubilities up to 30 is controlled by the waterside. Some environmentally important species lie in a transition zone where it is required to consider both transport processes (Figure 2). The transfer of highly soluble volatile and/or chemically reactive gas is controlled by the air-side transfer process and thus analogous to the transfer of water vapor. The following considerations concentrate on the waterside transfer process. Gas Exchange at Smooth Water Surfaces
At smooth water surfaces, the theory of mass transfer is well established because it is equivalent to mass transfer to a smooth solid wall. The turbulent diffusivity can be described by the classic approach of Reichhardt with an initial z3 increase that smoothly changes to a linear increase in the turbulent layer as in [3]. Then integration of [6] yields the following approximation for Schmidt numbers higher than 60: kw ¼ uw
1 Sc2=3 ; 12:2
Sc > 60
½11
This equation establishes the basic analogy between momentum transfer and gas exchange. The transfer coefficient is proportional to the friction velocity in water, which describes the shear stress (tangential force per unit area) t ¼ rw u2w applied by the wind field at the water surface. Assuming stress continuity at the water surface, the friction velocity in water is related to the friction velocity in air by uw ¼ ua
ra rw
1=2 ½12
The friction velocity in air, ua , can further be linked via the drag coefficient to the wind speed UR at a reference height: cD ¼ ðua =UR Þ2 . Depending on the roughness of the sea surface, the drag coefficient has values between 0.8 and 2.4 10 3. In this way the gas exchange rate is directly linked to the wind speed. The gas exchange further depends on the chemical species and the water temperature via the Schmidt number. Gas Exchange at Rough and Wavy Water Surfaces
A free water surface is neither solid nor is it smooth as soon as short wind waves are generated. On a free water surface velocity fluctuations are possible. Thus, there can be convergence or divergence zone at the surface; surface elements may be dilated or contracted. At a clean water surface dilation or
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AIR–SEA GAS EXCHANGE
contraction of a surface element does not cause restoring forces, because surface tension only tries to minimize the total free surface area, which is not changed by this process. As a consequence of this hydrodynamic boundary condition, the turbulent diffusivity normal to the interface can now increase with the distance squared from the interface, Kcpz2. Then 1 kw ¼ uw Sc1=2 b
½13
where b is a dimensionless constant. In comparison to the smooth case in [11], the exponent n of the Schmidt number drops from 2/3 to 1/2. This increases the transfer velocity for a Schmidt number of 600 by about a factor of 3. The total enhancement depends on the value of the constant b. Influence of Surface Films
A film on the water surface creates pressure that works against the contraction of surface elements. This is the point at which the physicochemical structure of the surface influences the structure of the near-surface turbulence as well as the generation of waves. As at a rigid wall, a strong film pressure at the surface maintains a two-dimensional continuity at the interface just as at a rigid wall. Therefore, [11] should be valid for a smooth film-covered water surface and has indeed been verified in wind/ wave tunnel studies as the lower limit for the transfer velocity. As a consequence, both [11] and [13] can only be regarded as limiting cases. A more general approach is required that has not yet been established. One possibility is a generalization of [11] and [13] to kw ¼ uw
1 ScnðsÞ bðsÞ
½14
where both b and n depend on dimensionless parameters describing the surface conditions s. Even films with low film pressure may easily decrease the gas transfer rate to half of its value at clean water surface conditions. But still too few measurements at sea are available to establish the influence of surfactants on gas transfer for oceanic conditions more quantitatively. Influence of Waves
Wind waves cannot be regarded as static roughness elements for the liquid flow because their characteristic particle velocity is of the same order of magnitude as the velocity in the shear layer at the surface.
151
This fact causes a basic asymmetry between the turbulent processes on the air and on the water sides of the interface. Therefore, the wave effect on the turbulent transfer in the water is much stronger and of quite different character than in the air. This basic asymmetry can be seen if the transfer velocity for CO2 is plotted against the transfer velocity for water vapor (Figure 3(a)). At a smooth water surface the points fall well on the theoretical curve predicted by the theory for a smooth rigid wall. However, as soon as waves occur at the water surface, the transfer velocity of CO2 increases significantly beyond the predictions. Even at high wind speeds, the observed surface increase is well below 20%. When waves are generated by wind, energy is not only transferred via shear stress into the water but a second energy cycle is established. The energy put by the turbulent wind into the wave field is transferred to other wave numbers by nonlinear wave–wave interaction and finally dissipated by wave breaking, viscous dissipation, and turbulence. The turbulent wave dissipation term is the least-known term and of most importance for enhanced near-surface turbulence. Evidence for enhanced turbulence levels below wind waves has been reported from field and laboratory measurements. Experimental results also suggest that the gas transfer rate is better correlated with the ‘mean square slope’ of the waves as an integral measure for the nonlinearity of the wind wave field than with the wind speed. It is not yet clear, however, to what extent ‘microscale wave breaking’ can account for the observed enhanced gas transfer rates. A gravity wave becomes unstable and generates a steep train of capillary waves at its leeward face and has a turbulent wake. This phenomenon can be observed even at low wind speeds, as soon as wind waves are generated. At higher wind speeds, the frequency of microscale wave breaking increases. Influence of Breaking Waves and Bubbles
At high wind speeds, wave breaking with the entrainment of bubbles may enhance gas transfer further. This phenomenon complicates the gas exchange between atmosphere and the oceans considerably. First, bubbles constitute an additional exchange surface. This surface is, however, only effective for gases with low solubility. For gases with high solubility, the gas bubbles quickly come into equilibrium so that a bubble takes place in the exchange only for a fraction of its lifetime. Thus, bubble-mediated gas exchange depends – in contrast to the exchange at the free surface – on the solubility of the gas tracer.
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AIR–SEA GAS EXCHANGE
(a)
(b)
+ with waves (unlimited fetch) • no waves
0.001 0.80
10−2 Schmidt number exponent n
10−4
0.75 0.70
0.65
0.65
0.60
0.60
0.55
0.55
0.50
0.50
0.45
0.45
0.40 0.001
1
0.3 0.80
0.70
kH2O (cm s−1) 0.1
0.1 Small circular facility Large circular facility
0.75
kCO2 (Sc = 600) (cm s−1)
10−3
0.01
0.01
0.1
0.40 0.3
Mean square surface slope s 2
10
Figure 3 (a) Transfer velocity of CO2 plotted against the transfer velocity of water vapor as measured in a small circular wind wave facility. (b) Schmidt number exponent n as a function of the mean square slope. (a) From Ja¨hne B (1980) Zur Parameterisierung des Gasaustausches mit Hilfe von Laborexperimenten. Dissertation, University of Heidelberg. (b) From Ja¨hne B and HauXecker H (1998) Air–water gas exchange. Annual Review of Fluid Mechanics 30: 443–468.
Second, bubble-mediated gas transfer shifts the equilibrium value to slight supersaturation due to the enhanced pressure in the bubbles by surface tension and hydrostatic pressure. Third, breaking waves also enhance near-surface turbulence during the breaking event and the resurfacing of submerged bubbles. Experimental data are still too sparse for the size and depth distribution of bubbles and the flux of the bubbles through the interface under various sea states for a sufficiently accurate modeling of bubblemediated air–sea gas transfer and thus a reliable estimate of the contribution of bubbles to the total gas transfer rate. Some experiments from wind/wave tunnels and the field suggest that significant enhancements can occur, other experiments could not observe a significant influence of bubbles. Empiric Parametrization
Given the lack of knowledge all theories about the enhancement of gas transfer by waves are rather speculative and are not yet useful for practical application. Thus, it is still state of the art to use semiempiric or empiric parametrizations of the gas exchange rate with the wind speed. Most widely used is the parametrization of Liss and Merlivat. It identifies three physically well-defined regimes (smooth,
wave-influenced, and bubble-influenced) and proposes a piecewise linear relation between the wind speed U and the transfer velocity k: k ¼ 106 8 > 0:472UðSc=600Þ2=3 ; Ur3:6 m s1 > < 7:917ðU 3:39ÞðSc=600Þ1=2 ; U > 3:6 m s1 and Ur13 m s1 > > : 16:39ðU 8:36ÞðSc=600Þ1=2 ; U > 13 m s1
½15 At the transition between the smooth and wavy regime, a sudden artificial jump in the Schmidt number exponent n from 2/3 to 1/2 occurs. This actually causes a discontinuity in the transfer rate for Schmidt number unequal to 600. The empiric parametrization of Wanninkhof simply assumes a quadratic increase of the gas transfer rate with the wind speed: k ¼ 0:861 106 ðs m1 ÞU2 ðSc=600Þ1=2 ½16 Thus, this model has a constant Schmidt number exponent n ¼ 1/2. The two parametrizations differ significantly (see Figure 4). The Wanninkhof parametrization predicts significantly higher values. The discrepancy between the two parametrizations
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AIR–SEA GAS EXCHANGE
80
14
60
C
SF6 − 3He 222
k 600 (cm h−1)
where Fw and hw are the surface area and the effective height Vw/Fw of a well-mixed water body, respectively. The time constant tw ¼ hw/k is in the order of days to weeks. It is evident that the transfer velocities obtained in this way provide only values integrated over a large horizontal length scales and timescales in the order of tw. Thus, a parametrization of the transfer velocity is only possible under steady-state conditions over extended periods. Moreover, the mass balance contains many other sources and sinks besides air–sea gas exchange and thus may cause severe systematic errors in the estimation of the transfer velocity. Consequently, mass balance methods are only poorly suited for the study of the mechanisms of air–water gas transfer.
Wanninkhof relationship Liss−Merlivat relationship
70
50
Rn
Heat (CFT)
40 30 20 10 0
0
5
153
10 15 Wind speed u10 (m s−1)
20
Tracer Injection
Figure 4 Summary of gas exchange field data normalized to a Schmidt number of 600 and plotted vs. wind speed together with the empirical relationships of Liss and Merlivat and Wanninkhof. Adapted from Ja¨hne B and HauXecker H (1998) Air–water gas exchange. Annual Review of Fluid Mechanics 30: 443–468.
(up to a factor of 2) mirrors the current uncertainty in estimating the air–sea gas transfer rate.
Experimental Techniques and Results Laboratory Facilities
Laboratory facilities play an important role in the investigation of air–sea gas transfer. Only laboratory studies allow a systematic study of the mechanisms and are thus an indispensible complement to field experiments. Almost all basic knowledge about gas transfer has been gained by laboratory experiments in the past. Among other things this includes the discovery of the influence of waves on air–water gas exchange (Figure 3(a)) and the change in the Schmidt number exponent (Figure 3(b)). Many excellent facilities are available worldwide (Table 3). Some of the early facilities are no longer operational or were demolished. However, some new facilities have also been built recently which offer new experimental opportunities for air–water gas transfer studies. Geochemical Tracer Techniques
The first oceanic gas exchange measurements were performed using geochemical tracer methods such as the 14C, 3He/T, or 222Rn/226Ra methods. The volume and time-average flux density is given by mass balance of the tracer concentration in a volume of water Vw: Vw c_w ¼ Fw j or j ¼ hw c_w
½17
The pioneering lake studies for tracer injection used sulfur hexafluoride (SF6). However, the tracer concentration decreases not only by gas exchange across the interface but also by horizontal dispersion of the tracer. This problem can be overcome by the ‘dual tracer technique’ (Watson and co-workers) simultaneously releasing two tracers with different diffusivities (e.g., SF6 and 3He). When the ratio of the gas transfer velocities of the two tracers is known, the dilution effect by tracer dispersion can be corrected, making it possible to derive gas transfer velocities. But the basic problem of mass balance techniques, that is, their low temporal resolution, remains also with artificial tracer approaches. Eddy Correlation Flux Measurements
Eddy correlation techniques are used on a routine basis in micrometeorology, that is, for tracers controlled by the boundary layer in air (momentum, heat, and water vapor fluxes). Direct measurements of the air–sea fluxes of gas tracers are very attractive because the flux densities are measured directly and have a much better temporal resolution than the mass balance-based techniques. Unfortunately, large experimental difficulties arise when this technique is applied to gas tracers controlled by the aqueous boundary layer. The concentration difference in the air is only a small fraction of the concentration difference across the aqueous mass boundary layer. But after more than 20 years of research has this technique delivered useful results. Some successful measurements under favorable conditions have been reported and it appears that remaining problems can be overcome in the near future. The Controlled Flux Technique
The basic idea of this technique is to determine the concentration difference across the mass boundary
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(c) 2011 Elsevier Inc. All Rights Reserved. 2.0 0.8 104 83 15 N Y Y Y Y Y N
1.5 0.3 27 8 25 N N Y N N N N
3.0 0.8 800 768 15 N Y Y N N N N
100 8.0
D
2.4 1.5 96 144 12 Y Y N N N N N
40 2.4
SIO
0.85 0.25 24.8 8 25 N Y 70.6 Y N N Y
33 0.76
C
1.0 0.5 15 10 30 Y Y 70.5 Y N N Y
15 1.0
UM
1.22 0.76 16.7 ? 25 N Y 70.5 Y Y N Y
18.3 0.91
W 119 2.0 40 36 5.6 3.0 239 716 19 Y N N N N N N
SU 1.57 0.10 0.60 0.40 0.50 0.08 0.16 0.01 11 Y N N 5–35 5–35 Y Y
HD1
11.6 0.20 4.0 3.4 0.70 0.25 3.5 0.87 12 Y N N N N N Y
HD2
29.2 0.62 9.92 8.68 2.40 1.20 18.0 20.7 15 Y N o 0.6 5–35 5–35 Y Y
HD3
0.33 0.10 1.44 0.14 8 Y N o 0.1 Y N N Y
3.90 0.37
HD4
HH, Bundesanstalt fu¨r Wasserbau, Hamburg; M, IMST, Univ. Marseille, France; D, Delft Hydraulics, Delft, The Netherlands (no longer operational); SIO, Hydraulic Facility, Scripps Institution of Oceanography, La Jolla, USA; C, Canada Center for Inland Waters (CCIW); UM, University of Miami; W, NASA Air–Sea Interaction Research Facility, Wallops; SU, Storm basin, Marine Hydrophysical Institute, Sevastopol, Ukraine (no longer operational), HD1, Small annular wind/wave flume, Univ. Heidelberg (no longer operational), HD2, Large annular wind/wave flume, Univ. Heidelberg (dismantled); HD3, Aeolotron, Univ. Heidelberg (HD3, in operation since June 2000); HD4, Teflon-coated small Heidelberg linear wind/wave flume (N ¼ No, Y ¼ Yes).
40 2.6
M
15 1.8
HH
Comparison of the features of some major facilities for small-scale air–sea interaction studies (operational facilities are typeset in boldface)
Length (mean perimeter) (m) Width of water channel (m) Outer diameter (m) Inner diameter (m) Total height (m) Max. water depth (m) Water surface area (m2) Water volume (m3) Maximum wind speed (m s 1) Suitable for sea water Wave maker Water current generator (m s 1) Water temperature control (1C) Air temperature control (1C) Air humidity control Gastight air space
Table 3
14
28
13 12
Wind (20 Hz)
26
Wind (60 s average)
24
11 Gas transfer rates 10
(k600, CFT data)
22 20
9
18
8
16
7
14
6
12
5
10
4
8
3
6
2
4
1
2
0 03:00
155
03:30
04:00
k600 (cm h−1)
Wind (u10) (m s−1)
AIR–SEA GAS EXCHANGE
0 04:30
Time (UTC) (hh:mm) Figure 5 Wind speeds and gas transfer velocities computed with the controlled flux technique (CFT) during the 1995 MBL/CoOP West Coast experiment (JD133) for a period of 90 min. The transfer velocities are normalized to Schmidt number 600 and averaged over 4 min each. From Ja¨hne B and HauXecker H (1998) Air–water gas exchange. Annual Review of Fluid Mechanics 30: 443–468.
layer when the flux density j of the tracer across the interface is known. The local transfer velocity can be determined by simply measuring the concentration difference Dc across the aqueous boundary layer (cold surface skin temperature) according to [4] with a time constant t˜ for the transport across the boundary layer [8]. This technique is known as the ‘controlled flux technique’ (CFT). Heat proves to be an ideal tracer for the CFT. The temperature at the water surface can then be measured with high spatial and temporal resolution using IR thermography. A known and controllable flux density can be applied by using infrared radiation. Infrared radiation is absorbed in the first few 10 mm at the water surface. Thus, a heat source is put right at top of the aqueous viscous boundary layer. Then the CFT directly measures the waterside heat transfer velocity. A disadvantage of the CFT is that the transfer velocity of gases must be extrapolated from the transfer velocity of heat. The large difference in the Schmidt number (7 for heat, 600 for CO2) casts some doubt whether the extrapolation to so much higher Schmidt numbers is valid. Two variants of the technique proved to be successful. Active thermography uses a CO2 laser to heat a spot of several centimeters in diameter on the water surface. The heat transfer rates are estimated
from the temporal decay of the heated spot. Passive thermography uses the naturally occurring heat fluxes caused by latent heat flux jl, sensible heat flux js, and long wave emission of radiation jr . The net heat flux jn ¼ jl þ js þ jr results according to [4] in a temperature difference across the interface of DT ¼ jh/(rcpkh). Because of the turbulent nature of the exchange process any mean temperature difference is associated with surface temperature fluctuations which can be observed in thermal images. With this technique the horizontal structure of the boundary layer turbulence can be observed. Surface renewal is directly observable in the IR image sequences, which show patches of fluid being drawn away from the surface. With some knowledge about the statistics of the temperature fluctuations, the ‘temperature difference’ DT across the interface as well as the time constant t˜ of heat transfer can be computed from the temperature distribution at the surface. Results obtained with this technique are shown in Figure 5 and also in the overview graph (Figure 4). Summary of Field Data
A collection of field data is shown in Figure 4. Although the data show a clear increase of the transfer
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velocity with wind speed, there is substantial scatter in the data that can only partly be attributed to uncertainties and systematic errors in the measurements. Thus, in addition, the field measurements reflect the fact that the gas transfer velocity is not simply a function of the wind speed but depends significantly on other parameters influencing nearsurface turbulence, such as the wind-wave field and the viscoelastic properties of the surface film.
Outlook In the past, progress toward a better understanding of the mechanisms of air–water gas exchange was hindered by inadequate measuring technology. However, new techniques have become available and will continue to become available that will give a direct insight into the mechanisms under both laboratory and field conditions. This progress will be achieved by interdisciplinary research integrating different research areas such as oceanography, micrometeorology, hydrodynamics, physical chemistry, applied optics, and image processing. Optical- and image-processing techniques will play a key role because only imaging techniques give direct insight to the processes in the viscous, heat, and mass boundary layers on both sides of the air– water interface. Eventually all key parameters including flow fields, concentration fields, and waves will be captured by imaging techniques with sufficient spatial and temporal resolution. The experimental data gained with such techniques will stimulate new theoretical and modeling approaches.
Nomenclature 2
D (cm s 1) jc (Mol cm 2 s 1) k (cm s 1) Kc (cm2 s 1) R (cm 1 s) Sc ¼ n/D t˜ ¼ z˜/k (s) u (cm s 1) z˜ (cm) a n (cm2 s 1)
molecular diffusion coefficient concentration flux density transfer velocity turbulent diffusion coefficient transfer resistance Schmidt number boundary layer time constant friction velocity boundary layer thickness dimensionless solubility kinematic viscosity
Sea Transfer: N2O, NO, CH4, CO. Breaking Waves and Near-Surface Turbulence. Bubbles. Carbon Dioxide (CO2) Cycle. Pollution: Effects on Marine Communities. Surface Gravity and Capillary Waves.
Further Reading Borger AV and Wanninkhof R (eds.) (2007) Special Issue: 5th International Symposium on Gas Transfer at Water Surfaces. Journal of Marine Systems 66: 1--308. Businger JA and Kraus EB (1994) Atmosphere–Ocean Interaction. New York: Oxford University Press. Donelan M, Drennan WM, Saltzman ES, and Wanninkhof R (eds.) (2001) Gas Transfer at Water Surfaces. Washington, DC: American Geophysical Union. Duce RA and Liss PS (eds.) (1997) The Sea Surface and Global Change. Cambridge, UK: Cambridge University Press. Garbe C, Handler R, and Ja¨hne B (eds.) (2007) Transport at the Air–Sea Interface, Measurements. Models, and Parameterization. Berlin: Springer. Ja¨hne B (1980) Zur Parameterisierung des Gasaustausches mit Hilfe von Laborexperimenten. Dissertation, University of Heidelberg. Ja¨hne B and HauXecker H (1998) Air–water gas exchange. Annual Review of Fluid Mechanics 30: 443--468. Ja¨hne B, Heinz G, and Dietrich W (1987) Measurement of the diffusion coefficients of sparingly soluble gases in water. Journal of Geophysical Research 92: 10767--10776. Ja¨hne B and Monahan E (eds.) (1995) Air–Water Gas Transfer. Hanau: Aeon. King DB, De Bryun WJ, Zheng M, and Saltzman ES (1995) Uncertainties in the molecular diffusion coefficient of gases in water for use in the estimation of air–sea exchange. In: Ja¨hne B and Monahan E (eds.) Air–Water Gas Transfer, pp. 13--22. Hanau: Aeon. Liss PS and Merlivat L (1986) Air–sea gas exchange rates: Introduction and synthesis. In: Buat-Menard P (ed.) The Role of Air–Sea Exchange in Geochemical Cycles, pp. 113--127. Dordrecht: Reidel. McGilles WR, Asher WE, Wanninkhof R, and Jessup AT (eds.) (2004). Special Issue: Air Sea Exchange. Journal of Geophysical Research 109. Wanninkhof R (1992) Relationship between wind speed and gas exchange over the ocean. Journal of Geophysical Research 97: 7373--7382. Wilhelms SC and Gulliver JS (eds.) (1991) Air–Water Mass Transfer. New York: ASCE.
Relevant Websites See also Air–Sea Transfer: Dimethyl Sulfide, COS, CS2, NH4, Non-Methane Hydrocarbons, Organo-Halogens. Air–
http://www.ifm.zmaw.de – Institute of Oceanography, Universita¨t Hamburg. http://www.solas-int.org – SOLAS.
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AIR–SEA TRANSFER: DIMETHYL SULFIDE, COS, CS2, NH4, NON-METHANE HYDROCARBONS, ORGANO-HALOGENS J. W. Dacey, Woods Hole Oceanographic Institution, Woods Hole, MA, USA H. J. Zemmelink, University of Groningen, Haren, The Netherlands Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 1, pp 131–137, & 2001, Elsevier Ltd.
The oceans, which cover 70% of Earth’s surface to an average depth of 4000 m, have an immense impact on the atmosphere’s dynamics. Exchanges of heat and momentum, water and gases across the sea surface play major roles in global climate and biogeochemical cycling. The ocean can be thought of as a vast biological soup with myriad processes influencing the concentrations of gases dissolved in the surface waters. The quantities of mass flux across the surface interface, though perhaps small on a unit area basis, can be very important because of the extent of the ocean surface and the properties of the gases or their decomposition products in the atmosphere. Gas exchange across the sea–air surface depends, in part, on differences in partial pressures of the gases between the ocean surface and the atmosphere. The partial pressure of a gas in the gas phase can be understood in terms of its contribution to the pressure in the gas mixture. So the partial pressure of O2, for example, at 0.21 atm means that at 1 atmosphere total pressure, O2 is present as 21% of the gas, or mixing, volume. Trace gases are present in the atmosphere at much lower levels, usually expressed as parts per million (106 atm), parts per billion (109 atm) or parts per trillion (1012 atm, pptv). Dimethylsulfide (DMS), when present at 100 pptv, accounts for about 100 molecules per 1012 molecules of mixed gas phase, or about 1010 of the gas volume. In solution, a dissolved trace gas in equilibrium with the atmosphere would have the same partial pressure as the gas in the air. Its absolute concentration in terms of molecules or mass per unit volume of water depends on its solubility. Gas solubility varies over many orders of magnitude depending on the affinity of water for the gas molecules and the volatility of the gas. Gases range widely in their solubility in sea water, from the permanent gases like nitrogen (N2), oxygen (O2), nitrous oxide (N2O) and methane (CH4) that have a low solubility in sea
water to the moderately soluble carbon dioxide (CO2) and dimethylsulfide (CH3)2S, to highly soluble ammonia (NH3 and its ionized form NHþ 4 ) and sulfur dioxide (SO2). Sulfur dioxide is more than 106 times more soluble than O2 or CH4. Using the example above of an atmospheric DMS concentration of 100 pptv, the equilibrium concentration of DMS in surface water would be about 0.07 nmol l1. Generally the solubility of any individual gas increases at cooler water temperatures, and solubility of gases in sea water is somewhat less than for fresh water because of the so-called ‘salting out’ effect of dissolved species in sea water. At any moment the partial pressure difference between surface water and the atmosphere depends on an array of variables. The gases in this article are biogenic, meaning that their mode of formation is the result of one or more immediate or proximate biological processes. These dissolved gases may also be consumed biologically, or removed by chemical processes in sea water, or they may flux across the sea surface to the atmosphere. The rates at which the source and sink processes occur determines the concentration of the dissolved gas in solution as well as the turnover, or residence time, of each compound. Similarly, there can be several source and sink processes for the gases in the atmosphere. Long-lived compounds in the atmosphere will tend to integrate more global processes, whereas short-lived compounds are concentrated near their source and reflect relatively short-term influences of source and sink. In this sense, carbonyl sulfide is a global gas. At it has a residence time of several years in the atmosphere, its concentration does not vary in the troposphere to any appreciable degree. On the other hand, the concentration of DMS varies on a diel basis and with elevation, with higher concentrations at night when atmospheric oxidants (most notable hydroxyl) are relatively depleted. The extent of disequilibrium between the partial pressures of a gas in the surface water and in the atmosphere determines the thermodynamic gradient which drives gas flux. The kinetics of flux ultimately depend on molecular diffusion and larger-scale mixing processes. Molecular diffusivity is generally captured in a dimensionless parameter, the Schmidt number (ratio of viscosity of water to molecular
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diffusivity of gas in water), and varies widely between gases depending primarily on the molecular cross-section. From moment to moment, the flux of any particular gas is dependent on interfacial turbulence which is generated by shear between the wind and the sea surface whereby higher wind speed causes increasing turbulence and thus stimulating the onset of waves and eventually the production of bubbles and sea spray. There are considerable uncertainties relating gas exchange to wind speed. These arise due to the various sea-state factors (wave height, swell, breaking waves, bubble entrainment, surfactants, and others) whose individual dependencies on actual wind speed and wind history are not well quantified. The fluxes of gases across the air–sea interface are usually calculated using a wind-speed parameterization. These estimates are considered to be accurate to within a factor of 2 or so. This article summarizes the characteristics of several important trace gases – dimethylsulfide, carbonyl sulfide, carbon disulfide, nonmethane hydrocarbons, ammonia and methylhalides – focusing on their production and fate as it is determined by biological and chemical processes.
Dimethylsulfide Natural and anthropogenic sulfur aerosols play a major role in atmospheric chemistry and potentially in modulating global climate. One theory holds that a negative feedback links the emission of volatile organic sulfur (mostly as DMS) from the ocean with the formation of cloud condensation nuclei, thereby
regulating, in a sense, the albedo and radiation balance of the earth. The direct (backscattering and reflection of solar radiation by sulfate aerosols) and indirect (cloud albedo) effects of sulfate aerosols may reduce the climatic forcing of trace greenhouse gases like CO2, N2O and CH4. The oxidation products of DMS which also contribute to the acidity of rain, particularly in marine areas, result from industrialized and/or well-populated land. Dimethylsulfide (DMS) is the most abundant volatile sulfur compound in sea water and constitutes about half of the global biogenic sulfur flux to the atmosphere. Studies of the concentration of DMS in the ocean have shown that average surface water concentrations may vary by up to a factor of 50 between summer and winter in mid and high latitudes. Furthermore, there are large-scale variations in DMS concentration associated with phytoplankton biomass, although there are generally poor correlations between local oceanic DMS concentrations and the biomass and productivity of phytoplankton (due to differences between plankton species in ability to produce DMS). The nature and rates of the processes involved in the production and consumption of DMS in sea water are important in determining the surface concentrations and the concomitant flux to the atmosphere. The biogeochemical cycle of DMS (Figure 1) begins with its precursor, b-dimethylsulfoniopropionate (DMSP). DMSP is a cellular component in certain species of phytoplankton, notably some prymnesiophytes and dinoflagellates. The function of DMSP is unclear, although there is evidence for an
Photochemistry SO2 Stratosphere COS DMS CS2
SO2 MSA
Troposphere Photochemistry DMS
Ocean
DMSO Photochemistry 2_ 4
DMSP
SO Organic matter
Figure 1 Fate and production of dimethylsulfide (DMS), carbonylsulfide (COS) and carbon disulfide. DMSO, dimethylsulfoxide; DMSP, dimethylsulfoniopropionate; MSA, methane sulfonic acid.
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AIR–SEA TRANSFER
osmoregulatory role as its cellular concentrations have been found to vary with salinity. It is generally thought that healthy algal cells do not leak either DMSP or DMS, although mechanical release into the surrounding sea water can lead to DMS production during cell senescence and grazing by zooplankton or as a consequence of viral attack on phytoplankton cells. Oceanic regions dominated by prolific DMSPproducing phytoplankton tend to have high DMS and DMSP concentrations. Breakdown of DMSP, presumably after transfer from the particulate algal (pDMSP) form to a dissolved (dDMSP) form in sea water, can proceed in different ways, mostly depending on microbiological conditions. One major pathway involves cleavage of DMSP to DMS and acrylic acid. Bacterial metabolism of dDMSP may be a major mechanism for DMS production in sea water, with acrylic acid residue acting as a carbon source for heterotrophic growth. Sulfonium compounds are vulnerable to attack by hydroxide ion; the resulting chemical elimination reaction occurs rapidly and quantitatively in strong base but only slowly at the pH of sea water. DMS in sea water has many potential fates. The volatility of DMS and the concentration gradient across the sea–air interface lead to the ocean being the major source of DMS to the atmosphere. Estimates of the annual sulfur release (as DMS) vary from 13–37 Tg S y1 (Kettle and Andreae, 1999). However, whereas the absolute flux of DMS from sea to air may be large on a global scale, sea–air exchange may represent only a minor sink for seawater DMS. It has been estimated that DMS loss to the atmosphere is only a very small percentage of the DMS sink, but this undoubtedly depends on the biogeochemical conditions in the water column at the time. Photochemical oxidation of DMS, either to dimethylsulfoxide (DMSO) or to other products, occurs via photosensitized reactions. The amount of photochemical decomposition depends on the amount of light of appropriate wavelengths and the concentration of colored organic compounds in solution to convert light energy into reactive radicals. Light declines exponentially with depth; the distribution of colored dissolved organic materials exhibits depth and seasonal variability. Microbial consumption of DMS, although extremely variable in both time and space in the ocean, appears to be a significant sink for oceanic DMS. The residence time of DMS is probably of the order of a day or two in most seawater systems. Since the atmospheric residence time of DMS is about a day or two, the atmospheric consequences of DMS flux are mostly confined to the troposphere. In the troposphere, DMS is oxidized primarily by hydroxyl radical. The main atmospheric oxidation
159
products are methane sulphonic acid, SO2 and DMSO.
Carbonyl Sulfide Carbonyl sulfide (COS, OCS) is the major sulfur gas in the atmosphere, present throughout the troposphere at 500 pptv. COS has a long atmospheric residence time (B4 years). Because of its relative inertness COS diffuses into the stratosphere where it oxidizes to sulfate particles and contributes in reactions involving stratospheric ozone chemistry. Unlike DMS which is photochemically oxidized in the troposphere, the major sink for COS is terrestrial vegetation and soils. COS is taken up by plants by passing through the stomata and subsequently hydrolyzing to CO2 and H2S through the action of carbonic anhydrase inside plant cells. There is no apparent physiological significance to the process; it appears to just occur accidentally to the normal physiology of plants. COS is produced in the ocean by photochemical oxidation of organic sulfur compounds whereby dissolved organic matter acts as a photosensitizer. The aqueous concentration of COS manifests a strong diel cycle, with the highest concentrations in daytime (concentration range on the order of 0.03– 0.1 nmol l1). COS hydrolyzes in water to H2S at rates dependent on water temperature and pH. The flux of oceanic COS to the atmosphere may represent about one-third of the global COS flux.
Carbon Disulfide Concentrations in surface water are around 1011 mol 11. Although a number of studies have indicated that the ocean forms an important source for atmospheric CS2, the underlying biochemical cycles still remain poorly understood. CS2 is formed by photochemical reactions (possibly involving precursors such as DMS, DMSP and isothiocyanates). CS2 formation has been observed to occur in bacteria in anoxic aquatic environments and in cultures of some marine algae species. The residence time of CS2 in the atmosphere is relatively short (about one week). Although CS2 might contribute directly to SO2 in the troposphere, its main significance is in the formation of COS via photochemical oxidation which results in the production of one molecule each of SO2 and COS per molecule of CS2 oxidized. The resulting COS may contribute to the stratospheric aerosol formation. Concentrations around 14 pmol l1 of carbon disulfide in the mid-Atlantic Ocean were first observed
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(1974); higher concentrations have been found in coastal waters. More than a decade later CS2 concentrations in the North Atlantic were found to be comparable to the earlier observations. However, in coastal waters CS2 concentrations were found to be a factor 10 lower, respectively 33 and 300 pmol l1. The global CS2 flux has been estimated on 6.7 Gmol S y1, and it has been concluded that the marine emission of CS2 provides a significant indirect source of COS, but it forms an insignificant source of tropospheric SO2.
Nonmethane Hydrocarbons Nonmethane hydrocarbons (NMHCs) are important reactive gases in the atmosphere since they provide a sink for hydroxyl radicals and play key roles in the production and destruction of ozone in the troposphere. NMHCs generally refer to the C2–C4 series, notably ethane, ethene, acetylene, propane, propene, and n-butane, but also the five-carbon compound isoprene. Of these, ethene is generally the most abundant contributing 40% to the total NMHC pool in sea water. Published data of concentrations of NMHCs in sea water vary widely sometimes exceeding a factor 100. For example, in one extensive study, ethene and propane were found to be the most abundant species in the intertropical South Pacific, with mixing ratios of 2.7 to 58 and 6 to 75 pptv, respectively; whereas in the equatorial Atlantic these species showed mixing ratios of 20 pptv and 10 pptv, respectively. The water-column dynamics of NMHCs are poorly understood. NMHCs have been detected in the surface sea and with maxima in the euphotic zone and tend to be present at concentrations in sea water at around 1010 mol l1. Evidence suggests that photochemical oxidation of dissolved organic matter results in the formation of NMHCs. There can be very little doubt that the physiology of planktonic organisms is also involved in NMHC formation. Ethene and isoprene are freely produced by terrestrial plants where the former is a powerful plant hormone but the function of the latter less well understood. It is likely that similar processes occur in planktonic algae. NMHC production tends to correlate with light intensity, dissolved organic carbon and biological production. A simplified scheme of marine NMHC production is shown in Figure 2. The flux of NMHCs to the atmosphere (with estimates ranging from o10 Mt y1 to 50 Mt y1) is minor on a global scale, but has a potential significance in local atmospheric chemistry. Although oceans are known to act as sources of NMHCs, the
NMHC (R-H) + OH
Products (R + H2O) Troposphere
NMHC (R-H) Photosynthetic organisms
Ocean Photolysis and chemical conversion Dissolved organic matter
Zooplankton Bacteria
Figure 2 Simplified scheme of marine nonmethane hydrocarbon (NMHC) production. In the marine troposphere NMHC acts as a sink for hydroxyl (OH) radicals and thereby plays a key role in ozone chemistry.
sources of individual NMHCs in the marine boundary layer are not always clear. Those NMHCs with a life time of more than a week (e.g., ethane, ethyne, propane, cyclopropane) show latitudinal gradients consistent with a continental source, whereas variations of NMHCs with life times shorter than a week (all alkenes and pentane) are more consistent with a marine source.
Ammonia Ammonia is an extremely soluble gas, reacting with water and dissociating into an ammonium ion at ambient pH. At pH 8.2, about one-tenth of dissolved ammonia is present as NH3. Ammonium is also a rapidly cycling biological nutrient; it is taken up by bacteria and phytoplankton as a source of fixed nitrogen, and released by sundry physiological and decompositional processes in the food web. Anthropogenic loading of ammonium (and other nutrients) into the coastal marine environment results in increased phytoplankton growth in a phenomenon called eutrophication. Ammonium is oxidized to nitrate by bacteria in a process known as nitrification (Figure 3). Conversely, in anoxic environments, ammonium can be formed by nitrate-reducing bacteria. Ammonia plays an important role in the acid–base chemistry in the troposphere where the unionized ammonia (NH3) is converted into ionized ammonia (NHþ 4 ) via a reaction that neutralizes atmospheric acids as HNO3 and H2SO4. This leads to the formation of ammonium aerosols such as the stable ammonium sulfate. Eventually the ammonia returns to the surface by dry or wet deposition. Few data exist on the fluxes of NH3 over marine environments. Evidence suggests that most of the ocean surface serves as a source of NH3 to the
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AIR–SEA TRANSFER
NOx
Oxidation
NH3(g)
Reduction
161
NH4-aerosol
Dry and wet deposition
Troposphere NH3(g)
Ammonification
NH4+
Water
Nitrification
NO3
Denitrification
Organic nitrogen Nitrogen fixation
N2
Figure 3 Simplified scheme of marine NHx chemistry. In the marine boundary layer NO3 acts as an initiator for the degradation of many organic compounds, in particular dimethylsulfide (DMS).
atmosphere, even in regions of very low nutrients. In the North Sea, an area situated in the middle of densely populated and industrialized countries of Western Europe, air from nearby terrestrial sources may act as a source of NH3 into surface waters. It has been estimated that the annual biogenic emission of ammonia from European seas is around 30 kt N y1, which is comparable to the emissions of smaller North European countries, leading to the conclusion (amongst others) that seas are among the largest sources of imported ammonium for maritime countries. The net emission of ammonia from coastal waters of the north-east Pacific Ocean to the atmosphere has been shown to be in the order of 10 mmol m2 d1.
Organohalogens Halogenated compounds, such as methyl chloride (CH3Cl), methyl bromide (CH3Br) and methyl iodide (CH3I) are a major source of halogens in the atmosphere, and subsequently form sources of reactive species capable of catalytically destroying ozone. Among these CH3I is likely to play an important role in the budget of tropospheric ozone, through production of iodine atoms by photolysis. Due to their higher photochemical stability methyl chloride and
methyl bromide are more important in stratospheric chemistry; it has been suggested that BrO species are responsible for losses of tropospheric ozone in the Arctic (Figure 4). Atmospheric methyl halides, measured over the ocean by several cruise surveys, have been shown to have average atmospheric mixing ratios of: CH3Cl, 550–600 pptv; CH3Br, 10–12 pptv; CH3I, 0.5–1 pptv. Their temporal and spatial variations are not well understood, neither is their production mechanism in the ocean known. Measurements of atmospheric and seawater concentrations of CH3Cl and CH3I have indicated that the oceans form natural sources of these methyl halides. In contrast, CH3Br appears to be undersaturated in the open ocean and exhibits moderate to 100% supersaturation in coastal and upwelling regions, leading to a global atmosphere to ocean flux of 13 Gg y1. Coastal salt marshes, although they constitute a minor area of the global marine environment, may produce roughly 10% of the total fluxes of atmospheric CH3Br and CH3Cl and thus contribute significantly to the global budgets. Macrophytic and phytoplanktonic algae produce a wide range of volatile organohalogens including di- and tri-halomethanes and mixed organohalogens. There is evidence for the involvement of enzymatic synthesis of methyl halides, but the metabolic
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Temporary reservoirs HBr, BrONO2, etc. O3 Hv
Hv
CH3Br
O2
Br Hv
Coupling with CIOx
BrO HO2
O3 HOBr
Stratosphere
O3
Troposphere
NO2 HO2
Ozone loss CH3Br
OH
Photolabile fraction
Br O2 BrO
Hv
_
Br
BrO RCHO
and different dynamics spatially and with depth in the ocean. As conditions change in an apparently warming world, changes in the dynamics of surface ocean gases can be expected. The behavior of these trace gases or even the dynamics of the planktonic community are not understood sufficiently to allow good quantitative predictions about changes in trace gas flux to be made. Changes in flux of some gases could lead to an acceleration of warming, while changes in others could lead to cooling. It is, thus, important to understand the factors controlling trace gas dynamics in the surface ocean.
See also
Soluble fraction (HBr) Dry deposition
Wet deposition
Air–Sea Gas Exchange. Air–Sea Transfer: N2O, NO, CH4, CO. Atmospheric Input of Pollutants. Carbon Cycle. Chlorinated Hydrocarbons. Nitrogen Cycle.
Ocean CH3Br
Photochemistry
Dissolved organic matter
Further Reading
Phytoplankton/Bacteria
Figure 4 Schematic illustration of the circulation of methyl bromide.
production pathways are not well known. In free sea water, photochemical processes, ion substitution and, possibly the alkylation of halide ions (during the oxidation of organic matter by an electron acceptor such as Fe(III)) are also potential formation mechanisms. Sunlight or microbial mediation are not required for these reactions. In the ocean, chemical degradation of CH3Br occurs by nucleophilic substitution by chloride and hydrolysis. Microbial consumption is also a likely sink for halogenated compounds.
Conclusions The biogenic trace gases are influenced by the complete range of biological processes – from the biochemical and physiological to the ecological level of food web dynamics. The gases that are influenced directly by plant physiology (probably the light NMHCs and isoprene, for example) tend to be most closely related to phytoplankton biomass or primary productivity. Other gases produced during grazing and decomposition (e.g., DMS, NH3), or gases formed by photochemical reactions in dissolved organic material show differing temporal dynamics
Andreae MO (1990) Ocean–atmosphere interactions in the global biogeochemical sulfur cycle. Marine Chemistry 10: 1--29. Andreae MO and Crutzen PJ (1997) Atmospheric aerosols: biogeochemical sources and role in atmospheric chemistry. Science 276: 1052--1058. Barrett K (1998) Oceanic ammonia emissions in Europe and their transboundary fluxes. Atmospheric Environment 32(3): 381--391. Chin M and Davis DD (1993) Global sources and sinks of OCS and CS2 and their distributions. Global Biogeochemical Cycles 7: 321--337. Cox RA, Rattigana OV, and Jones RL (1995) Laboratory studies of BrO reactions of interest for the atmospheric ozone balance. In: Bandy RA (ed.) The Chemistry of the Atmosphere; Oxidants and Oxidation in the Earth’s Atmosphere. Cambridge: The Royal Society of Chemistry. Crutzen PJ (1976) The possible importance of COS for the sulfate layer of the stratosphere. Geosphysical Research Letters 3: 73--76. Graedel TE (1995) Tropospheric budget of reactive chlorine. Global Biogeochemical Cycles 9: 47--77. Kettle AJ and Andreae MO (1999) Flux of dimethylsulfide from the oceans: a comparison of updated datasets and flux models. Journal of Geophysical Research 105: 26793--26808. Lovelock JE (1974) CS2 and the natural sulfur cycle. Nature 248: 625--626. Turner SM and Liss PS (1985) Measurements of various sulfur gases in a coastal marine environment. Journal of Atmospheric Chemistry 2(3): 223--232.
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AIR–SEA TRANSFER: N2O, NO, CH4, CO C. S. Law, Plymouth Marine Laboratory, The Hoe, Plymouth, UK Copyright & 2001 Elsevier Ltd.
identifies the marine contribution to total atmosphere budgets. There is also a brief examination of the approaches used for determination of marine trace gas fluxes and the variability in current estimates.
This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 1, pp 137–144, & 2001, Elsevier Ltd.
Nitrous Oxide (N2O) Introduction The atmospheric composition is maintained by abiotic and biotic processes in the terrestrial and marine ecosystems. The biogenic trace gases nitrous oxide (N2O), nitric oxide (NO), methane (CH4) and carbon monoxide (CO) are present in the surface mixed layer over most of the ocean, at concentrations which exceed those expected from equilibration with the atmosphere. As the oceans occupy 70% of the global surface area, exchange of these trace gases across the air–sea interface represents a source/sink for global atmospheric budgets and oceanic biogeochemical budgets, although marine emissions of NO are poorly characterized. These trace gases contribute to global change directly and indirectly, by influencing the atmospheric oxidation and radiative capacity (the ‘greenhouse effect’) and, together with their reaction products, impact stratospheric ozone chemistry (Table 1). The resultant changes in atmospheric forcing subsequently influence ocean circulation and biogeochemistry via feedback processes on a range of timescales. This article describes the marine sources, sinks, and spatial distribution of each trace gas and
The N2O molecule is effective at retaining long-wave radiation with a relative radiative forcing 280 times that of a CO2 molecule. Despite this the relatively low atmospheric N2O concentration results in a contribution of only 5–6% of the present day ‘greenhouse effect’ with a direct radiative forcing of about 0.1 Wm2. In the stratosphere N2O reacts with oxygen to produce NO radicals, which contribute to ozone depletion. N2O is a reduced gas which is produced in the ocean primarily by microbial nitrification and denitrification. N2O is released during ammonium (NHþ 4 ) oxidation to nitrite (NO2 ) (Figure 1), although the exact mechanism has yet to be confirmed. N2O may be an intermediate of nitrification, or a byproduct of the decomposition of other intermediates, such as nitrite or hydroxylamine. Nitrification is an aerobic process, and the N2O yield under oxic conditions is low. However, as the nitrification rate decreases under low oxygen, the relative yield of N2O to nitrate production increases and reaches a maximum at 10–20 mmol dm3 oxygen (mmol ¼ 1 106 mol). Conversely, denitrification is an anaerobic process in which soluble oxidized nitrogen
Table 1 The oceanic contribution and atmospheric increase and impact for methane, nitrous oxide, nitric oxide, and carbon monoxidea Trace gas
Atmospheric concentration (ppbv)
Atmospheric lifetime (years)
Major impact in atmosphere
Increase in atmosphere (1980– 90)
Oceanic emission as % of total global emissions
Nitrous oxide (N2O) Nitric oxide (NO)
315
110–180
0.25% (0.8 ppbv y1)
7–34%
0.01
o0.2
Not known
Not known
Methane (CH4)
1760
10
0.8% (0.6 ppbv y1)
1–10%
Carbon monoxide (CO)
120
0.2–0.8
Infrared active Ozone sink/source Ozone sink/source OH sink/oxidation capacity Infrared active OH sink/oxidation capacity Ozone sink/source OH sink/oxidation capacity Ozone sink/source Infrared active
13 to 0.6%
0.9–9%
a
ppbv, parts per billion by volume. (Adapted from Houghton et al., 1995.)
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Nitrification +
NH 4 H2O O2
_
NO2
_
N2O
NO
NO3
N2 Denitrification Figure 1 ‘Leaky Pipe’ flow diagram of nitrification and denitrification indicating the potential exchange and intermediate role of NO and N2O (Reprinted by permission from Nature copyright (1990), Macmillan Magazines Ltd.)
compounds, such as nitrate and nitrite, are converted to volatile reduced compounds (N2O and N2) in the absence of oxygen. Oxygen availability inhibits denitrification at ambient levels, and also determines the products of denitrification. An enzymatic gradient of sensitivity to oxygen results in the accumulation of N2O under sub-oxia (3–10 mmol dm3) due to the inhibition of the enzyme nitrous oxide reductase. At lower oxygen (o3 mmol dm3) the reaction continues through to N2 and so anoxic environments are sinks for N2O. N2O yields from nitrification are 0.2–0.5%, whereas denitrification yields may be as high as 5% at optimal levels of sub-oxia. An inverse correlation between N2O and oxygen, and associated linear relationship between nitrate and N2O, suggest that N2O in the ocean originates primarily from nitrification. This may not be the case for sediments, in which denitrification is the dominant source of N2O under variable oxygen tension, with nitrification only contributing in a narrow suboxic band. Attribution of source is difficult as nitrification and denitrification may occur simultaneously and interact, with exchange of products and intermediates (Figure 1). This is further complicated, as denitrification will be limited to some extent by nitrate supply from nitrification. Isotopic data from the surface ocean in oligotrophic regions imply that N2O originates primarily from nitrification. However, recent evidence from waters overlying oxygen-deficient intermediate layers suggests that the elevated surface mixed-layer N2O arises from coupling between the two processes, as the observed isotope signatures cannot be explained by nitrification or denitrification alone. An additional N2O source from the dissimilatory reduction of nitrate to ammonium is restricted to highly anoxic environments such as sediments. The oceanic N2O distribution is determined primarily by the oxygen and nutrient status of the water
column. Estuaries and coastal waters show elevated supersaturation in response to high carbon and nitrogen loading, and the proximity of sub-oxic zones in sediment and the water column. As a result the total marine N2O source tends to be dominated by the coastal region. The N2O flux from shelf sea sediments is generally an order of magnitude lower than estuarine sediments, although the former have a greater spatial extent. A N2O maximum at the base of the euphotic zone is apparent in shelf seas and the open ocean, and is attributed to production in suboxic microzones within detrital material. Oceanic surface waters generally exhibit low supersaturations (o105%), although N2O supersaturations may exceed 300% in surface waters overlying low oxygen intermediate waters and upwelling regions, such as the Arabian Sea and eastern tropical North Pacific. These ‘natural chimney’ regions dominate the open ocean N2O source, despite their limited surface area (Table 2). The surface N2O in upwelling regions such as the Arabian Sea originates in part from the underlying low-oxygen water column at 100– 1000 m, where favorable conditions result in the accumulation of N2O to supersaturations exceeding 1200%. N2O transfer into the surface mixed layer will be limited by vertical transport processes and a significant proportion of N2O produced at these depths will be further reduced to N2. The oceans account for 1–5 Tg N-N2O per annum (Tg ¼ 1 1012 g) or 6–30% of total global N2O emissions, although there is considerable uncertainty attached to this estimate (Figure 2). A recent estimate with greater representation of coastal sources has resulted in upward revision of the marine N2O source to 7–10.8 Tg N-N2O per annum; although this may represent an upper limit due to some bias from inclusion of estuaries with high N2O supersaturation. However, this estimate is in agreement with a total
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165
Table 2 N2O and CH4 regional surface water supersaturations (from Bange et al., 1996; 1998) (supersaturation is >100%, undersaturation is o100% with equilibrium between atmosphere and water at 100%)
Estuaries Coastal/shelf Oligotrophic/transitional ocean Upwelling ocean
Surface % N2O saturation mean (range)
Surface % CH4 saturation mean (range)
607 (101–2500) 109 (102–118) 102.5 (102–104) 176 (108–442)
1230 (146–29 000) 395 (85–42 000) 120 (80–200) 200 (86–440)
N2O sources _1 (Tg N yr ) Oceans 3 (1_ 5) Soil natural 6 (3.3 _ 9.7)
Soil anthropogenic 3.3 (0.6 _ 14.8)
Biomass burning 0.5 (0.2 _1) Industrial 1.3 (0.7 _1.8)
½2
NO2 þ hn-NO þ O 3 P
½3
O þ O2 þ M-O3 þ M
½4
At high concentrations (>50 ppbv; ppbv ¼ parts per billion by volume), O3 in the atmospheric boundary layer becomes a toxic pollutant that also has important radiative transfer properties. The production of nitric acid from NO influences atmospheric pH, and contributes to acid rain formation. In addition, the oxidation of NO to the nitrate (NO3) radical at night influences the oxidizing capacity of the lower troposphere. Determination of the magnitude and location of NO sources is critical to modeling boundary layer and free tropospheric chemistry. NO cycling in the ocean has received limited attention, as a result of its thermodynamic instability and high reactivity. Photolysis of nitrite in surface waters occurs via the formation of a nitrite radical with the production of NO:
Livestock feed 2.1 (0.6 _ 3.1)
N2O sinks _1 (Tg N yr ) Atmospheric increase 3.9 (3.1_ 4.7)
Stratospheric photolysis 12.3 (9 _ 16) Figure 2 Atmospheric nitrous oxide sources and sinks (adapted from Houghton et al., 1995). Units: Tg ¼ 1 1012 g.
oceanic production rate of 11 Tg N-N2O per annum calculated from new production and nitrification.
Nitric Oxide (NO) Nitric oxide (NO) plays a central role in atmospheric chemistry, influencing both ozone cycling and the tropospheric oxidation capacity through reactions with hydroperoxy- and organic peroxy-radicals. When the NO concentration exceeds B40 pptv (pptv ¼ parts per trillion by volume) it catalyzes the production of ozone (O3): CO þ OH * þ O2 -HO2* þ CO2
HO2* þ NO-OH * þ NO2
½1
* þ HOH-NO þ OH þ OH NO 2 þ hn-NO2
This reaction may account for 10% of nitrite loss in surface waters of the Central Equatorial Pacific, resulting in a 1000-fold increase in dissolved NO at a steady-state surface concentration of 5 pmol dm3 during light periods (pmol ¼ 1 1012 mol). This photolytic production is balanced by a sink reaction with the superoxide radical (O 2 ) to produce peroxynitrite: O 2 þ NO- OONO
This reaction will be dependent upon steady-state concentration of the superoxide radical; however, as the reaction has a high rate constant, NO is rapidly turned over with a half-life on the order of 10–100 seconds.
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Soils 15 (10 _ 20) Aircraft emmisions 0.6
Biomass burning 5.5 (2.5 _ 8.5)
Lightning 6 (2 _ 8) Transport from stratosphere 0.5
Fossil fuel combustion 21
Figure 3 Atmospheric nitric oxide sources (from Graedel and Crutzen, 1992). Units: Tg ¼ 1 1012 g.
As with N2O, NO may also be produced as a byproduct or intermediate of denitrification and nitrification (Figure 1). NO production by soils is better characterized than in marine systems, and is significant both in terms of nitrogen loss and the global NO budget (Figure 3). The greater oxygen availability in soils limits reduction of NO via denitrification and so enhances NO efflux. Sediment pore water NO maxima have been attributed to denitrification, although, as this process also represents a sink for NO (Figure 1), this may reflect poising at an optimal redox potential for NO production. Conversely, the NO maximum in low oxygen intermediate waters in the east tropical North Pacific derives from nitrification. Current understanding of the oceanic NO distribution is that it is limited to the surface ocean and intermediate low oxygen water column. There is potential for higher NO concentrations in coastal and estuarine waters from sediment and photolytic sources, and nitrite photolysis to NO may also be significant in upwelling regions. Despite the short half-life of NO in surface waters, the maintenance of steady-state NO concentration suggests that photolytic production may support an, as yet unquantified, source of atmospheric NO. Surface concentrations in the Central Equatorial Pacific suggest that the oceanic NO source would not exceed 0.5 Tg N per annum, which is relatively insignificant when compared with other sources (Figure 3).
Methane (CH4) CH4 is the most abundant organic volatile in the atmosphere and, next to CO2, is responsible for 15% of the current greenhouse radiative forcing, with a direct radiative forcing of 0. 5 Wm2. CH4 reacts with OH and so limits the tropospheric oxidation capacity and influences ozone and other greenhouse
gases. The reaction with OH generates a feedback that leads to a reduction in the rate of CH4 removal. CH4 is a reduced gas which, paradoxically, is supersaturated in the oxidized surface waters of the ocean (see Table 2). CH4 is produced biotically and abiotically, although its oceanic distribution is controlled primarily by biological processes. Methanogenesis is classically defined as the formation of CH4 from the fermentation and remineralization of organic carbon under anoxic conditions. Methanogens require a very low reducing potential and are generally obligate anaerobes, although there is evidence that they can tolerate some exposure to oxygen. However, methanogens cannot utilize complex organic molecules and often coexist with aerobic consortia to ensure a supply of simple C1 substrates. Methanogens utilize formate, acetic acid, CO2, and hydrogen in sulfate-rich anoxic environments, although they are generally out-competed by sulfatereducing bacteria which have a greater substrate affinity. However, the methanogens can also utilize other noncompetitive substrates such as methanol, methylamines, and reduced methylated compounds, when out-competed for the C1 compounds. A significant fraction of CH4 is oxidized before exiting the marine system and so the oxidation rate is critical in determining the air–sea flux. This is accomplished by methanotrophs that obtain their carbon and energy requirements from CH4 oxidation under aerobic conditions via the following reactions: CH4 -CH3 OH -HCHO-HCOOC -CO2 methane - methanol - formaldehyde - formate - carbon dioxide
Methanotrophs are found in greater numbers in sediments than in oxic sea water, and consequently the oceanic water column CH4 oxidation is an order of magnitude lower than in sediments. Methanotrophs have a high inorganic nitrogen requirement and so methanotrophy is highest at the oxic–anoxic interface where ammonium is available. Anaerobic CH4 oxidation also occurs but is less well characterized. It is generally restricted to anaerobic marine sediments, utilizing sulfate as the only oxidant available, and is absent from anaerobic freshwater sediments which lack sulfate. A significant proportion of CH4 produced in anaerobic subsurface layers in sediments is oxidized during diffusive transport through the sulfate-CH4 transition zone by anaerobic oxidation and subsequently by aerobic oxidation in the overlying oxic layers. Anaerobic oxidation represents the main sink for CH4 in marine sediments, where it may account for 97% of CH4 production.
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AIR–SEA TRANSFER: N2O, NO, CH4, CO
CH4 production is characteristic of regions with high input of labile organic carbon such as wetlands and sediments, but is usually restricted to below the zone of sulfate depletion. The oceanic CH4 source is dominated by coastal regions, which exhibit high CH4 fluxes as a result of bubble ebullition from anoxic carbon-rich sediments, and also riverine and estuarine input. Some seasonality may result in temperate regions due to increased methanogenesis at higher temperatures. The predominant water column source in shelf seas and the open ocean is CH4 production at the base of the euphotic zone. This may arise from lateral advection from sedimentary sources, and in situ CH4 production. The latter is accomplished by oxygen-tolerant methanogens that utilize methylamines or methylated sulfur compounds in anoxic microsites within detrital particles and the guts of zooplankton and fish. Lateral advection and in situ production may be greater in upwelling regions, as suggested by the increased CH4 supersaturation in surface waters in these regions. Oceanic CH4 concentration profiles generally exhibit a decrease below 250 m due to oxidation. Methanogenesis is elevated in anoxic water columns, although these are not significant sources of atmospheric CH4 due to limited ventilation and high oxidation rates. Other sources include CH4 seeps in shelf regions from which CH4 is transferred directly to the atmosphere by bubble ebullition, although their contribution is difficult to quantify. Abiotic CH4 originating from high-temperature fluids at hydrothermal vents also elevates CH4 in the deep and intermediate waters in the locality of oceanic ridges. A significant proportion is oxidized and although the contribution to the atmospheric CH4 pool may be significant in localized regions this has yet to be constrained. Hydrates are crystalline solids in which methane gas is trapped within a cage of water molecules. These form at high pressures and low temperatures in seafloor sediments generally at depths below 500 m. Although CH4 release from hydrates is only considered from anthropogenic activities in current budgets, there is evidence of catastrophic releases in the geological past due to temperature-induced hydrate dissociation. Although oceanic hydrate reservoirs contain 14 000 Gt CH4, there is currently no evidence of significant warming of deep waters which would preempt release. Other aquatic systems such as rivers and wetlands are more important sources than the marine environment. Shelf regions are the dominant source of CH4 from the ocean (14(11–18) Tg CH4 per annum), accounting for 75% of the ocean flux (Table 2). The ocean is not a major contributor to the atmospheric
167
CH4 sources _1 (Tg CH4 yr )
Other natural 35 (20 _ 90) Wetland 115 (55 _ 150) Anthropogenic (other) 275 (200 _ 350)
Oceans 14 (11_ 18) Anthropogenic (fossil fuel) 100 (70 _ 120)
CH4 sinks _1 (Tg CH4 yr )
Tropospheric OH oxidation 445 (360 _ 530)
Loss to stratosphere 40 (32 _ 48) Soils 30 (15 _ 45) Atmospheric increase 37 (35 _ 40)
Figure 4 Atmospheric methane sources and sinks (adapted from Houghton et al. (1995)). Units: Tg ¼ 1 1012 g.
CH4 budget, as confirmed by estimates of the oceanic CH4 source (Figure 4).
Carbon Monoxide (CO) The oxidation of CO provides the major control of hydroxyl radical content in the troposphere and limits the atmospheric oxidation capacity. This results in an increase in the atmospheric lifetime of species such as CH4, N2O, and halocarbons, and enhances their transfer to the stratosphere and the potential for subsequent ozone destruction. It has been suggested that decreasing stratospheric ozone and the resultant increase in incident ultraviolet (UV) radiation may increase marine production and efflux of CO, thereby generating a positive feedback loop. However, this may be compensated by a negative feedback in which increased UV reduces biological production and dissolved organic matter, so reducing the CO source. CO also influences tropospheric ozone by its interaction with NOx, and is a minor greenhouse gas with a radiative forcing of 0.06 Wm2 at current atmospheric concentrations.
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The principal source of dissolved CO is the abiotic photodegradation of dissolved organic matter (DOM) by UV-R, and CO represents one of the major photoproducts of DOM in the ocean. Quantum yields for CO are highest in the UV-B range (280–315 nm) and decrease with increasing wavelengths. However, the UV-A (315–390 nm) and blue portion of the visible spectrum contribute to marine CO production as a greater proportion of radiation at these wavelengths reaches the Earth’s surface. Humics represent approximately half of the DOM and account for the majority of the chromophoric dissolved organic matter (CDOM), the colored portion of dissolved organic matter that absorbs light energy. The CO photoproduction potential of humics is dependent upon the degree of aromaticity. Terrestrial humics are characterized by an increased prevalence of phenolic groups, and addition of precursor compounds containing phenolic moieties to natural samples stimulates CO production. Direct photo-oxidation of humics and compounds containing carbonyl groups, such as aldehydes, ketones, and quinones, occurs via the production of a carbonyl radical during a-cleavage of an adjacent bond:R þ ðCOR0 Þ * -R0 þ CO RCOR0 þ hv-hskip200pt ðRCOÞ * þR0 -R0 þ CO CO production may also occur indirectly by a photosensitized reaction in which light energy is transferred via an excited oxygen atom to a carbonyl compound. This may occur with ketonic groups via the photosensitized production of an acetyl radical. Whereas light and CDOM are the primary factors controlling CO production there may also be additional influence from secondary factors. For example, organo-metal complexes have increased light absorption coefficients and their photo-decomposition will enhance radical formation and CO production at higher levels of dissolved metals such as iron. There is also minor biotic production of CO by methanogens, but this does not appear to be significant. CO can be oxidized to carbon dioxide by selected microbial groups including ammonia oxidizers and methylotrophs that have a broad substrate specificity and high affinity for CO. However, only the carboxidotrophs obtain energy from this reaction, and these may be unable to assimilate CO efficiently at in situ concentrations. CO turnover times of 4 hours are typical for coastal waters, whereas this varies between 1 and 17 days in the open ocean. The lower oxidation rate in the open ocean may be due to light inhibition of CO oxidation. Extrapolation from laboratory
measurements suggests that only 10% of photochemically produced CO is microbially oxidized. Dissolved CO exhibits diurnal variability in the surface ocean in response to its photolytic source, although this is also indicative of a strong sink term. The decline in the surface mixed-layer CO concentration in the dark results from a combination of CO oxidation, vertical mixing, and air–sea exchange. As the equilibration time between atmosphere and oceanic surface mixed layer is on the order of a month, this suggests that the former two processes dominate. Superimposed upon the diurnal cycle of CO in the surface ocean are spatial and seasonal gradients that result from the interaction of photoproduction and the sink processes. Below the euphotic zone CO is uniformly low throughout the intermediate water column. CO production potential is highest in wetland regions, which are characterized by high CDOM and enhanced light attenuation. Photochemical production of CO represents a potential sink for terrestrial dissolved organic carbon (DOC) in estuaries and coastal waters. This pathway may account for some of the discrepancy between the total terrestrial DOC exported and the low proportion of terrestrial DOC observed in the marine pool. Although a strong lateral gradient in CDOM exists between rivers and the open ocean, estuarine CO production may be limited by reduced UV light penetration. CO photoproduction may occur down to 80 m in the open ocean, and 20 m in the coastal zone, but is restricted to the upper 1 m in wetlands and estuaries. In addition, estuarine and coastal CO flux may also be restricted by the higher CO oxidation rates. There is evidence that upwelling regions may support enhanced CO production, in response to upwelled CDOM that is biologically refractory but photolabile. The presence of a CO gradient in the 10 m overlying the surface ocean suggests that the photolytic source of CO may influence the marine boundary layer. The marine source of CO is poorly constrained, with estimates varying from 10 to 220 Tg CO per annum. A flux of 1200 Tg CO per annum was estimated on the assumption that low rates of oceanic CO oxidation would only remove a small proportion of photoproduced CO, and that the residual would be ventilated to the atmosphere. The discrepancy between this and other flux estimates implies that a significant CO sink has been overlooked, although this may reflect shortcomings of different techniques. The oceanic contribution to the global source is between 1 and 20%, although Extrapolation of photochemical production rates from wetlands, estuaries, and coasts suggests that these
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Table 3
169
Atmospheric CO sources and sinksa (adapted from Zuo et al., 1998)
CO sources (Tg CO y1) Industrial/fuel combustion Biomass burning Vegetation and soils Methane oxidation NMHC oxidation Ocean (Coast/Shelf Total sources
CO sinks (Tg CO y1) 400–1000 300–2200 50–200 300–1300 200–1800 10–220 300–400) 1260–6720
Tropospheric hydroxyl oxidation Soils Flux to stratosphere
1400–2600 250–530 80–140
Total sinks
1730–3270
a
Note that a separate estimate of the coastal/shelf CO source is shown for comparison, but does not contribute to the total source. Tg ¼ 1 1012 g.
alone may account for 20% of the total global CO flux. Although the marine source is responsible for o10% of the total global flux (Table 3), it may still dominate atmospheric oxidation conditions in remote regions at distance from land.
Air–Sea Exchange of Trace Gases The flux of these trace gases across the air–sea interface is driven by physical transfer processes and the surface concentration anomaly, which represents the difference between the partial pressure observed in surface water and that expected from equilibrium with the atmosphere. Direct determination of the oceanic emission of a trace gas is difficult under field conditions. Atmospheric gradient measurements above the ocean surface require enhanced analytical resolution, whereas more advanced micrometeorological techniques have yet to be applied to these trace gases. Determination of the accumulation rate in a floating surface flux chamber is a simpler approach, but may generate artefactual results from the damping of wave- and wind-driven exchange, and enhanced transfer on the inner chamber surfaces. Consequently the majority of flux estimates are calculated indirectly rather than measured. The surface anomaly is derived from the difference between the measured surface concentration (Cw), and an equilibrium concentration calculated from the measured atmospheric concentration (Cg) and solubility coefficient (p) at ambient temperature and salinity. This is then converted to a flux by the application of a dynamic term, the gas transfer velocity, k: F ¼ kðCw aCgÞ The transfer velocity k is the net result of a variety of molecular and turbulent processes that operate at different time and space scales. Wind is the primary driving force for most of these turbulent processes,
and it is also relatively straightforward to obtain accurate measurements of wind speed. Consequently, k is generally parameterized in terms of wind speed, with the favored approaches assuming tri-linear and quadratic relationships between the two. These relationships are defined for CO2 at 201C in fresh water and sea water and referenced to other gases by a Schmidt number (Sc) relationship: k gas ¼ k ref ðSc gas=Sc ref Þn where n is considered to be 1/2 at most wind speeds. This dependency of k is a function of the molecular diffusivity (D) of the gas and the kinematic viscosity of the water (m), and is expressed in terms of the Schmidt number (Sc ¼ m/D). Determination of marine trace gas fluxes using different wind speed–transfer velocity relationships introduces uncertainty, which increases at mediumhigh wind speeds to a factor of two. Furthermore, additional uncertainty is introduced by the extrapolation of surface concentration gradient measurements to long-term climatological wind speeds. Current estimates of oceanic fluxes are also subject to significant spatial and temporal bias resulting from the fact that most studies focus on more productive regions and seasons. This uncertainty is compounded by the extrapolation of observational data sets to unchartered regions. With the exception of N2O, the ocean does not represent a major source for these atmospheric trace gases, although spatial variability in oceanic source strength may result in localized impact, particularly in remote regions. In the near future, advances in micrometeorological techniques, improved transfer velocity parameterizations and the development of algorithms for prediction of surface ocean concentrations by remote sensing should provide further constraint in determination of the oceanic source of N2O, NO, CH4, and CO.
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AIR–SEA TRANSFER: N2O, NO, CH4, CO
See also Air–Sea Gas Exchange. Air–Sea Transfer: Dimethyl Sulfide, COS, CS2, NH4, Non-Methane Hydrocarbons, Organo-Halogens. Carbon Dioxide (CO2) Cycle. Gas Exchange in Estuaries. Photochemical Processes. Plankton and Climate. Surface Films. Upwelling Ecosystems.
Further Reading Bange HW, Bartel UH, et al. (1994) Methane in the Baltic and North Seas and a reassessment of the marine emissions of methane. Global Biogeochemical Cycles 8: 465--480. Bange HW, Rapsomanikis S, and Andreae MO (1996) Nitrous oxide in coastal waters. Global Biogeochemical Cycles 10: 197--207.
Carpenter EJ and Capone DG (eds.) (1983) Nitrogen in the Marine Environment. London: Academic Press. Graedel TE and Crutzen PJ (eds.) (1992) Atmospheric Change: An Earth System Perspective. London: W. H. Freeman and Co. Houghton JT, Meira Filho M, Bruce J, et al. (1995) Climate Change 1994. Radiative Forcing of Climate Change and an Evaluation of the IPCC IS92 Emission Scenarios, Intergovernmental Panel on Climate Change. Cambridge: Cambridge University Press. Liss PS and Duce RA (eds.) (1997) The Sea Surface and Global Change. Cambridge: Cambridge University Press. Zuo Y, Guerrero MA, and Jones RD (1998) Reassessment of the ocean to atmosphere flux of carbon monoxide. Chemistry and Ecology 14: 241--257.
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ALCIDAE T. Gaston, National Wildlife Research Centre, Quebec, Canada Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 1, pp 145–151, & 2001, Elsevier Ltd.
Introduction The auks are seabirds of the northern hemisphere. They form a family (Alcidae) or subfamily (Alcinae) of the order Charadriiformes, which includes the gulls, terns, jaegers, sandpipers, and plovers. They are often considered the northern hemisphere equivalents of the penguins, being well-adapted to underwater swimming. However, all extant species retain the power of flight. They are almost entirely confined to Arctic, Subarctic and Boreal waters and are marine throughout the year, being most common in waters of the continental shelf and slope. Twentythree species are currently recognized and one other became extinct in historic times (great auk Pinguinus impennis). Six species of five genera occur in the North Atlantic and 20, of 10 genera in the North Pacific, south of the Chukchi Sea. Two murres (Uria spp.), the dovekie (Alle alle) and the black guillemot (Cepphus grylle) are circumpolar in distribution. Many species of auks are extrememly abundant, with common and thick-billed murres (Uria aalge, U. lomvia), least and crested auklets (Aethia pusilla, A. cristatella), dovekies and Atlantic puffins (Fratercula arctica) all having world populations of 410 million individuals. They form the dominant avian biomass over large areas of Arctic and Subarctic waters throughout the year and may be significant predators on large zooplankton and small fishes in areas around their breeding colonies.
Miocene, about 15 million years ago (Ma), although putative proto-auks have been described from as far back as the Eocene. A subfamily of flightless auks, the Mancallinae, was present in the Pacific during the late Miocene (7–4 Ma). Deposits from California have yielded at least five species of the flightless Mancalla, ranging in size from about 1 to 4 kg. Auks probably originated in the Pacific, where all but two (Alca, Alle) of the extant genera are widely distributed, colonizing the Atlantic soon after. Since then, Pacific and Atlantic auks evolved largely in isolation. The Pliocene auk fauna from California was similar in diversity to the recent fauna, but the diversity of auks represented in the Pliocene sediments of eastern North America suggests that the current, relatively small, community of auks in the Atlantic (six species, including the great auk) is not a historical consequence of the lack of connections with the Pacific. Instead, it is a consequence of extinctions during the Pleistocene, presumably as a result, directly or indirectly, of the ice ages.
Characteristics and Adaptations Auks are well-adapted to underwater swimming. They have compact, streamlined bodies, short wings and very short tails. The feet are placed far back on the body, webbed, with no hind toe; claws are narrow and the tarsus is laterally compressed. The bill is variable in shape and may be highly ornamented in the breeding season. There are 11 primary wing feathers and 16–21 secondaries, the outermost primary being very small and the longest being, usually, the tenth. The feather tracts of the back and belly are Table 1 Tribes and genera of the Alcidae (after Gaston and Jones 1998)
Evolution and Systematics
Tribe
Genus
Recent work on mitochondrial DNA and allozymes divides the family Alcidae into 5 tribes and 12 genera (Table 1), with the puffins (Fraterculini) and auklets (Aethiini) being sister tribes, more closely related to one another than to the other auks. Otherwise, the molecular data suggest that the tribes originated from an initial rapid divergence among early members of the family, possibly about 10–12 million years ago. There is extensive fossil material on the family. The first definite records of auks are from the middle
Fraterculini
Fratercula Cerorhinca Aethia Cyclorhynchus Ptychoramphus Brachyramphus Cepphus Synthliboramphus Alle Pinguinus Alca Uria
Aethiini
Brachyramphini Cepphini Alcini
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continuous and, beneath the contour feathers, down feathers are present and dense all over the body. There are 6–8, occasionally 9, pairs of tail feathers. The fusiform body, with the feet set far back, is common to all underwater swimming birds. Other characters that are common to underwater swimmers, but absent in Charadriiformes that do not dive, are the presence of strongly developed vertebral hypophyses on the last cervical vertebra, and the enlarged number of thoracic vertebrae (8–10 compared to 5–7 in other Charadriiformes). An increase in the number of vertebrae allows for a longer body, while retaining flexibility of movement. Auks fly with rapid wing-beats and without gliding or soaring, using the spread feet for steering and braking. They generally take off from land with difficulty and use the feet to taxi when taking off from water. Underwater, they swim using the wings as paddles, like penguins. Maneuvering underwater is achieved mainly by asymmetrical strokes of the wings. Air is released from plumage before diving by forcing it from the breast feathers using subcutaneous muscles. On land, the Alcini and Brachyramphini rest on the belly, or on the tarsi, rising to their feet only in walking, whereas Aethiini and Fraterculini are more agile and normally stand with the tarsus erect, rather than horizontal. The anatomy of the auks is similar to that of other Charadriiformes, except for their specializations towards underwater swimming. They have shorter wings and legs than gulls and shore birds, and the relative length of the humerus is reduced. The covert feathers are stiffer and more extensive, both above and below the wing, an arrangement that reinforces the trailing edge of the wing and closes gaps between adjacent flight feathers. Primaries 6–10 form a closely knit unit with little independent movement, making the wing more effective as an underwater paddle. The rigidity is caused by a greater development of connective tissue, compared to other Charadriiformes. Despite the fact that the primary feathers are very stiff, the leading edge of the wing is very curved on the downstroke, presumably because of water resistance. The articulating surface of the humerus is much larger in auks than in other Charadriiformes, allowing greater force to be transmitted in swimming. However, like other flighted birds, but unlike penguins, the auks retain considerable flexibility of movement at the articulation of the humerus. The supracoracoideus muscle, which raises the wing during the recovery stroke (2–3% of body mass) is larger than in shore birds and gulls (1.5%) and the pectoralis muscles, the main source of power on the downstroke, although similar in size to those of gulls
and shorebirds, are more elongated, with the supracoracoideus lying directly below them. This arrangement improves streamlining, and also improves insulation by distributing stored fat in a thin subcutaneous layer over most of the body, rather than depositing it in discrete pockets, as occurs in many birds. Underwater, auks swim very jerkily, with a rapid acceleration at each downstroke. Hence, forward propulsion on the upstroke, if any, is relatively small. The wings are held sharply bent at the wrist, reducing the functional surface area to slightly less than half that when spread in flight. The fact that the whole wing area is unnecessary for effective underwater propulsion is emphasized by the fact that all genera except Aethia and Cyclorhynchus lose their flight feathers simultaneously during the annual postbreeding moult, becoming flightless for a month or more. The auks have developed many physiological adaptations for prolonged diving, including high blood volume and high levels of myoglobin (41% body mass), both of which enhance the bird’s ability to store oxygen. Despite these enhancements, many long dives undertaken by thick-billed murres, and probably other auks, exceed the estimated limit for aerobic respiration, requiring the birds to respire anaerobically for a portion of the time underwater. Heart rate and peripheral blood flow also may be restricted. The digestive system includes a well-developed proventriculus, a muscular stomach and a relatively short intestine. A functional crop is present only in auklets, although there is some croplike development of the lower proventriculus in the puffins. Auklets and dovekies develop diverticulae in the throat while breeding, in which they transport food for their young. Nasal glands, which are used to excrete salt and maintain ion balance, are very well developed. The dominant plumage color of auks is black above and white below, but Cepphus species are mainly black in summer, mainly white in winter, and Brachyramphus species are mottled brown in summer (Figure 1). Temporary nuptial ornaments, including ornamental plumes on the head and deciduous, brightly colored, horny plates on the bill, are present in the Aethiini and Fraterculini. Sexual dimorphism is very small: males of most species are slightly larger in some dimensions, especially bill depth. There is no plumage dimorphism and no distinctive immature plumage. First-winter birds generally appear similar to winter-plumage adults, although young Cepphus are readily distinguished by their greater mottling.
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Figure 1 Examples of the Alcidae. (1) Razorbill (Alca torda). Adult breeding plumage. Length: 42 cm; approximate body mass: 710 g. Range: North Atlantic Ocean. (2) Brunnich’s guillemot (Uria lomvia). Other name: thick-billed murre. (a) Adult breeding plumage; (b) adult non-breeding plumage; (c) transitional plumage. Length: 40 cm; approximate body mass: 950 g. Range: North Pacific, North Atlantic and Arctic Oceans. (3) Least auklet (Aethia pusilla). Length: 15 cm; approximate body mass: 85 g. Range: North Pacific Ocean. (4) Marbled murrelet (Brachyramphus marmoratus). Adult breeding plumage. Length: 25 cm; approximate body mass: 235 g. Range: eastern North Pacific Ocean. (5) Crested auklet (Aethia cristatella). (a) adult breeding plumage; (b) adult, non-breeding plumage. Length: 25 cm; approximate body mass: 260 g. Range: North Pacific Ocean. (6) Atlantic puffin (Fratercula arctica). Other name: puffin. Adult, breeding plumage. Length: 35 cm; approximate body mass: 470 g. Range: North Atlantic Ocean. (7) Little auk (Alle alle). Other name: dovekie. Adult breeding plumage. Length: 20 cm; approximate body mass: 165 g. Range: North Pacific and North Atlantic Oceans, Arctic Ocean. Illustrations from Harrison P (1985) Seabirds, an identification guide. Revised edition. Boston, Massachusetts: Houghton Mifflin.
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Diet and Foraging Ecology The diet of the smaller auks consists largely of zooplankton, especially copepods and euphausiids and larval fishes, although the parakeet auklet Cyclorhynchus psittacula takes substantial quantities of jellyfish medusae. The larger auks take a mixture of zooplankton and small fishes and squid, with common murres and Cepphus spp., and possibly Brachyramphus spp., being mainly piscivorous. Nestling Uria, Alca, Brachyramphus, Cepphus, Cerorhinca, and Fratercula species are all fed predominantly on fish. Auks that feed mainly on fish tend to have narrower, more pointed bills (low width/gape ratio), and narrower, more cornified, tongues than those that feed on plankton. Planktivorous auks, apart from having broader, shorter bills, have large numbers of fleshy denticles on the palate and the upper surface of the tongue. The convergence of these characters for auks feeding exclusively on plankton is dramatically demonstrated by the parallels between the auklets and the dovekie. These adaptations are also shown within the genus Uria, with the predominantly fish-eating common murre having a narrower tongue and fewer palatal denticles than the thick-billed murre, which feeds on a greater variety of prey. Bill depth seems to be unrelated to diet, suggesting that the very deep bills of puffins and auklets may have evolved for secondary sexual purposes, rather than as feeding adaptations. Auks feed predominantly in continental shelf waters, with species of Brachyramphus and Cepphus feeding entirely in inshore waters and most other genera occurring mainly within 50–100 km of land. However, Fratercula species are found far offshore in winter, including waters beyond the continental shelf. All species feed entirely underwater, with the smaller species diving to maximum depths of 30–40 m and the larger species to over 100 m. The deepest dives are achieved by Uria species, which may reach 200 m on occasions. Normal foraging depths for smaller auks are 10–30 m and for larger auks 30–60 m. Many species take advantage of prey aggregations caused by oceanographic fronts and tidally induced upwellings. Striking examples occur in the passes among Aleutian Islands and other complex archipelagos, where feeding auklets may reach great densities. The prey involved are usually slow-swimming zooplankton such as pteropod mollusks, copepods, euphausiids, and amphipods.
Reproduction Breeding sites are coastal, except in the marbled and Kittlitz’s murrelets, which are also the only solitary
nesters, and in the dovekie. Most auks breed on islands, and some species breed exclusively on remote islands well offshore, because islands offer refuge from terrestrial predators and at the same time may be close to foraging areas. Hence, the distribution of breeding auks is much influenced by the distribution of suitable breeding islands. Most species are highly social while breeding, with a prolonged prebreeding period and extensive courtship activity in which both vocal and visual displays are prominent. All species are socially monogamous and the sexes play equal roles in incubation and in rearing chicks to the age when they leave the colony. Most show a strong tendency to return to the colony where they were reared. Auks are highly variable in the type of nesting site that they use. None builds a nest: the Aethia auklets and some Synthliboramphus murrelets use crevices under boulders or among scree, and murres lay their eggs on open cliff ledges or on flat rocky islets (common murre), while puffins, ancient murrelet, and Cassin’s auklet often dig extensive burrows in soil. Most of the smaller auks make use of sites that protect them from surface predators, such as gulls and crows: only the largest, the murres, nest in the open in colonies, where their densely packed ranks form a defense against nest predators. Cliffs are frequently used, especially in the Arctic. The other open nesters, the Brachyramphus murrelets, are solitary nesters, either on the horizontal limbs of mature trees (marbled murrelet Brachyramphus marmoratus, commonly) or on the ground on remote mountain tops (Kittlitz’s murrelet B. brevirostris). Like many other small seabirds, some of the auks are nocturnal in their coming and goings to their breeding sites. The Synthliboramphus murrelets and whiskered and Cassin’s auklets (Aethia pygmaea, Ptychoramphus aleuticus) are invariably nocturnal, and the rhinoceros auklet (Cerorhinca monocerata) is largely nocturnal, but diurnal or crepuscular in some parts of its range. Brachyramphus species are normally crepuscular. Everything about the breeding strategies of the auks suggest the dominant influences of predation and kleptoparasitism. Egg formation is a lengthy process and females spend most of their time away from the colony during the last 10 days before they lay. Clutches consist of one or two eggs and only one brood is reared annually (California populations of Cassin’s auklet sometimes rear two). Eggs are thick-shelled and either white, buff, tan, or bluish in ground colour, sometimes with prominent black or brown markings. After clutch completion, the sexes alternate incubation duty, taking equal shares. Laying, in practically all populations, is confined to a span of
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ALCIDAE
about 6 weeks, with most eggs being laid within 2–3 weeks. Incubation periods range from 29 to 45 days. The time from the start of incubation to chick departure varies from a minimum of 35 days in the precocial Synthliboramphus species to a maximum of more than 80 days in the Cerorhinca. Most species, including the puffins, auklets and razorbill (Alca torda) that lay only a single egg, have two lateral brood patches. The murres and the dovekie have a single, central patch, set far back on the belly. Brood patches are a potential site of heat loss. They defeather rapidly at the start of incubation, refeather as soon as hatching occurs, and are small relative to the size of the eggs, so that during incubation only part of the egg’s surface is in contact with the patch. Nonbreeders either do not form brood patches or develop only partial patches. Notwithstanding these adaptations, the insulation of the auks appears to be relatively poor, considering the climates in which they live. To maintain normal avian body temperature, they rely principally on a very high rate of metabolism: basal metabolic rates (the metabolic rate of resting, nondigesting birds) among auks are exceptionally high, compared to those of other birds of similar size. Hatchlings are covered with a dense, woolly, down plumage. They are active within 1–2 days after hatching and capable of thermoregulation either immediately (Synthliboramphus) or within a few days (the rest). All genera except Synthliboramphus are semiprecocial and chicks are reared at the nest site for a minimum of 2 weeks. There is no postfledging parental care except in the dovekie, although in Synthliboramphus, Uria, and Alca the young are cared for until some time after departing the colony, which they leave when only partly grown and incapable of sustained flight. Most breeders leave the colony either with their chicks (Alcini and Synthliboramphus spp., where there is parental care after departure), or within 1–2 weeks following chick departure. Juveniles, family parties, and postbreeding adults disperse rapidly from the colony area (except young whiskered auklets).
Postbreeding and Wintering A complete prebasic molt (except in the puffins, where it involves the body plumage only) follows rapidly on the termination of breeding and involves the shedding of nuptial ornaments and the adoption of a generally distinctive nonbreeding (winter) plumage (no change in Craveri’s and Xantus’ murrelets). In the auklets, primary molt is sequential and begins during chick-rearing, but in other species the primaries are dropped simultaneously, making flight
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impossible for a period. Most species remain scattered offshore while molting. Following the annual prebasic molt, most species disperse or shift toward distant wintering areas. However, some populations of common murres and black guillemots return to the area of the breeding colony and commence periodic attendance at the breeding site. This also applies to some populations of marbled murrelets. These birds feed in the same areas practially year-round. However, among most species breeding in Arctic and Subarctic waters, the bulk of the population moves substantially farther south. Young birds tend to disperse farthest and there is usually a disproportionately high representation of first-year birds in ‘wrecks’; – the periodic casting ashore of large numbers of weakened birds, often during prolonged storms. Wintering areas and the behavior of auks on the wintering grounds have been less studied than their activities during the breeding season. Detailed studies of feeding ecology in winter have mostly been carried out on species and populations that occur in inshore waters and the winter ecology of many is essentially unknown. Movement toward the breeding colonies begins in February– April, depending on latitude.
Population Dynamics Annual survival of adult auks is generally greater than 85%, and greater than 95% in some populations of common murres and Atlantic puffins (Fratercula arctica), making them among the longestlived birds. In the longer-lived species, average age at first breeding is 5 years or more, whereas Cassin’s auklets, ancient murrelets, and probably some Aethia spp. may begin to breed at 2, many at 3 years. Populations contain substantial numbers of nonbreeders. Those in their second summer or older often attend the breeding colony to select mates and breeding sites. Reproductive success increases with age for the first few years of breeding. Comparisons of reproductive success are complicated by the fact that different species depart from the colony at different stages of development. The maximum productivity is about 1.5 young/pair per year and most average o1.
Auks and People Auks and their eggs have been harvested by people from the earliest times and their bones are frequent constituents of middens throughout the coastal areas of the northern hemisphere. The remains of the great auk have been discovered in 40 middens in Norway alone, as well as some in Britain, Iceland, Greenland, and the United States. Excavations in Newfoundland
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dating to 4000 BP contain many Great Auk bones and they are also found in middens in Florida dating to 3000 BP. All major colonies of thick-billed murres in the eastern Canadian Arctic show traces of Eskimo occupation nearby, usually with associated remains of thick-billed murres. In areas surrounding the Straits of Georgia, British Columbia, common murre remains are widespread in Indian middens dating from the pre-European period, while middens in the Aleutians, some dating to 4000 BP, contain the remains of auklets and puffins, as well as murres. A variety of techniques were developed for catching auks. One of the most widespread, used by Bering Sea Inuit, Icelanders, and Faeroese to catch puffins, and by the Inuit of north-west Greenland to catch dovekies, was a net at the end of a long pole. The hunter sheltered behind a stone wall, or depression in the ground, and suddenly raised the net in the path of low-flying birds, which were unable to turn in time to avoid it. In Iceland, snares placed on rafts floating offshore were also used to good effect, trapping murres, puffins, and razorbills. The technique of netting birds flying over the colony is a very efficient way to harvest auks, as the prebreeding component of the population often circles constantly, making them much more vulnerable than the breeders. Similarly, snares placed on boulders or floating rafts used by displaying birds, a method used on St. Lawrence Island to capture least and crested auklets, and in Thule, Greenland, to capture dovekies, select mainly prebreeders. The removal of these birds has much less effect on the population than the killing of breeders and may partly explain how early societies managed to coexist successfully with their prey. It has been estimated that 150 000–200 000 Atlantic puffins were taken annually in Iceland, without any apparent effect on population levels. In contrast, the shooting of breeding birds at colonies that became widespread in Greenland in this century has been the main cause of the drastic decline suffered by thick-billed murre populations there over the past 50 years. The harvesting of auk eggs has been very common throughout their range and may well have been a factor controlling their distribution on inshore islands accessible to permanent human settlements. On St. Kilda, harvesting of puffin and murre eggs was a regular activity, while in the Queen Charlotte Islands, and throughout the Alaskan islands, the excavation of ancient murrelet and auklet burrows for eggs was a routine spring harvest. Thick-billed murre eggs are harvested in Greenland, Canada, the Pribilof Islands, and Russia. In addition to their use as food, the skins and ornaments of certain auks were valued for clothing and
decoration. Inuit on St. Lawrence Island and Aleuts in the Aleutian chain sewed parkas out of auk skins, especially crested auklets and horned puffins. Elsewhere, puffin and dovekie skins were sewn into inner garments, to be worn under furs. In north-west Greenland, dovekie skins were made into undershirts, while farther south, in Upernavik District, murre skins were made into capes. Dovekie skins had to be softened by chewing; only elderly women did this, as their teeth were worn smooth enough not to damage the delicate skins. The spectacular beaks of puffins and auklets were also used as ornaments by the Aleuts and Inuit of the Bering Sea region, hundreds sometimes being sewn on the outside of a garment, along with the golden crests of tufted puffins. Commercial exploitation of auk colonies by postindustrial societies resulted in substantial declines. In the Gulf of St. Lawrence, the huge auk colonies visited by Audubon in 1827 were reduced to a mere remnant by the late nineteenth century, while at Funk Island, Newfoundland, the great auk was exterminated largely for its feathers. The common murre population at the Farallon Islands, California, and several large thick-billed murre populations in the Russian Far East and Novaya Zemlya were decimated by egg harvesting. Substantial harvesting of auks still occurs in several areas. Traditional harvests of murres and puffins continue in Iceland and the Faroes, although reduced from former levels. Relatively small numbers of thick-billed murres are taken by Inuit in the Canadian Arctic, although they form important components of the summer diet at a few settlements. Much larger numbers are taken in West Greenland, where regulations prohibiting the shooting of birds at their colonies were introduced only in 1978 and were still more or less unenforced in 1987. Shooting away from the colonies is still permitted throughout the year in some districts, although subject to seasonal limits in the more populated areas. The same populations are hunted more heavily off Newfoundland and Labrador in winter, with the annual kill estimated at about 200 000 since 1993. Although it is only legal to kill murres, some razorbills and dovekies, and a few puffins, are also shot. Although direct harvests have affected several auk species and were responsible for the extermination of the great auk, the effects of mammalian predators, introduced either deliberately, or accidentally, have probably had a much greater impact on auk populations worldwide. The main agents of destruction were foxes, introduced throughout the Alaskan islands for fur farming. Rats have also caused many declines and extirpations. Raccoons and mink have an important impact in some areas, and rabbits,
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ALCIDAE
through their effects on vegetation and soil, may also have caused problems for some burrowing species. Japanese, Craveri’s, Xantus’ and marbled murrelets are all considered endangered or threatened in one way or another. It is certain that the majority of auk populations are smaller, in many cases much smaller, than they would have been a few centuries ago. Probably, we will see little change in that situation, although programs to eliminate introduced predators from certain important Pacific islands may improve the situation for some species. All auks are very susceptible to contamination by oil and they have formed the majority of seabirds killed in oil spills off Europe and North America. Unlike gulls, they have not profited at all from fisheries wastes. Protection from egging has led to increases of some species in the twentieth century. However, overall, the auks remain precariously dependent on human goodwill for their future survival.
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Dynamics. Seabirds and Fisheries Interactions. Seabirds as Indicators of Ocean Pollution.
Further Reading Gaston AJ and Jones IL (1998) The Auks:: Family Alcidae. Oxford: Oxford University Press. Birkhead TR and Nettleship DN (1985) The Atlantic Acidae. London: Academic Press. Harris MP (1989) The Puffin. London: T & AD Poyser. Gaston AJ and Elliot RD (eds.) (1991) Conservation Biology of the Thick-billed Murre in the Northwest Atlantic. Ottawa: Canadian Wildlife Service. Sealy SG (ed.) (1990) Auks at Sea. Studies in Avian Biology 14, Cooper Ornithological Society. Friesen VL, Baker AJ, and Piatt JF (1996) Phylogenetic relationships within the Alcidae (Aves: Charadriiformes) inferred from total molecular evidence. Molecular Biology and Evolution 13: 359--367.
See also Fish Predation and Mortality. Laridae, Sternidae and Rynchopidae. Network Analysis of Food Webs. Plankton. Seabird Conservation. Seabird Foraging Ecology. Seabird Migration. Seabird Population
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ANTARCTIC CIRCUMPOLAR CURRENT S. R. Rintoul, CSIRO Antarctic Climate and Ecosystems Cooperative Research Centre, Hobart, TAS, Australia & 2009 Elsevier Ltd. All rights reserved.
Introduction The Drake Passage between the South American and Antarctic continents is the only band of latitudes where the ocean circles the Earth, unblocked by land. The existence of this oceanic channel has profound implications for the global ocean circulation and climate. The Drake Passage gap permits the Antarctic Circumpolar Current (ACC) to exist, a system of ocean currents which flows from west to east along a roughly 25 000-km-long path circling Antarctica. In terms of transport, the ACC is the largest current in the world ocean, carrying about 13778 106 m3 s 1 through the Drake Passage. The wind-driven ocean circulation theories that explain much of the ocean current patterns observed at other latitudes do not apply in an unbounded channel and the unique dynamics of the ACC have long been a puzzle for oceanographers. The threedimensional circulation in the ACC belt is now understood to reflect the interplay of wind and buoyancy exchange with the atmosphere, water mass modification, eddy fluxes of heat and momentum, and strong interactions between the flow and bathymetry. The strong eastward flow of the ACC has several important implications for the global ocean circulation and its influence on regional and global climate. By transporting water between the major ocean basins, the ACC tends to smooth out differences in water properties between the basins. The interbasin connection allows a global-scale pattern of ocean currents to be established, known as the thermohaline circulation (see Ocean Circulation: Meridional Overturning Circulation), which transports heat, moisture, and carbon dioxide around the globe and strongly influences the Earth’s climate. The strong flow of the ACC is associated with steeply sloping density surfaces, which shoal to the south across the current and bring dense waters to the surface in the high-latitude Southern Ocean. Where the dense waters are exposed at the sea surface, they exchange heat, moisture, and gases like oxygen and carbon dioxide with the atmosphere. In this sense, the Southern Ocean provides a window to the deep sea. The ecology and biogeography of the Southern Ocean
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are influenced strongly by the ACC and the overturning circulation plays an important role in the marine carbon and nutrient cycles. The west-to-east flow of the ACC inhibits north–south exchanges across the current and isolates Antarctica from the warm waters to the north; the present glacial climate of Antarctica was not established until the South American and Antarctic continents began to drift apart about 30 Ma, opening a circumpolar channel.
Structure of the Antarctic Circumpolar Current A schematic view of the major currents of the Southern Hemisphere oceans south of 201 S is shown in Figure 1. The flow of the ACC is focused in several jets, associated with sharp cross-stream gradients (or fronts) in temperature, salinity, and other properties. The three main fronts of the ACC – the Subantarctic Front, Polar Front, and southern ACC front – are indicated by the arrows circling Antarctica. To the south of the circumpolar flow of the ACC, clockwise gyres are found in the Weddell Sea, Ross Sea, and the Australian–Antarctic Basin (see Current Systems in the Southern Ocean). A westward flow associated with the Antarctic Slope Front and Antarctic Coastal Current is found near the continental shelf break around much of Antarctica. To the north of the ACC, water flows to the east in the southern limb of the large anticlockwise subtropical gyres in each basin. Exchanges of water masses between the ACC and the gyre circulations to the north and south are important components of the global circulation. The distribution of water properties on a transect crossing the ACC is illustrated in Figure 2. The temperature and salinity of water masses are largely set at the sea surface, where there is active exchange with the atmosphere; nutrient and oxygen concentrations are also influenced by biological processes. Water masses carry these characteristics with them as they sink from the surface into the ocean interior. The major water masses of the Southern Ocean are associated with various property extrema that reflect their circulation and formation history. For example, the Subantarctic Mode Water is formed by deep convection in winter on the northern flank of the ACC, producing deep well-mixed layers that are rich in oxygen. The Antarctic Intermediate Water is the name given to the prominent salinity minimum layer north of the ACC. The Circumpolar Deep Water
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90° W Figure 1 A schematic view of the major ocean currents of the Southern Hemisphere oceans south of 201 S. Depths shallower than 3500 m are shaded. C, current; G, gyre; F, front; ACC, Antarctic Circumpolar Current.
(CDW) is often divided into two layers: Upper CDW corresponds to the oxygen minimum layer, and Lower CDW corresponds to the deep salinity maximum layer. The relatively fresh layer near the seafloor is Antarctic Bottom Water, which sinks near Antarctica and carries water rich in oxygen and chlorofluorocarbons into the deep ocean (see Bottom Water Formation). Water properties at a given depth change dramatically as the Southern Ocean is crossed from north to south (Figure 2). Surfaces of constant temperature, salinity, density, and other properties slope upward to the south. As a result, density layers found at 3000-m depth in subtropical latitudes approach the sea surface near Antarctica. The shoaling of density surfaces to the south is associated with the strong eastward flow of the ACC. Tilted density surfaces create pressure forces which drive ocean currents and an accompanying Coriolis force to balance the pressure force. (This balance of forces,
known as geostrophy, describes the dynamics of all large-scale ocean currents (see Ocean Circulation).) In the Southern Hemisphere, an increase in density to the south supports an eastward flow (relative to the seafloor), as observed in the ACC. The rise of temperature, salinity, and density surfaces to the south occurs in a series of steps, or rapid transitions, rather than as a uniform slope across the Southern Ocean (Figure 2). These rapid transitions are known as fronts. Because the strength of an ocean current is proportional to the magnitude of the horizontal density gradient, each of the fronts is associated with a maximum in velocity. Most of the flow of the ACC is concentrated in the fronts, with smaller transports observed between the fronts (Figure 2(f)). The zones between the fronts also coincide with regions of relatively uniform water properties at each depth. Unlike many fronts at lower latitudes, the ACC fronts extend from the sea surface to the seafloor. The current jets associated
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Figure 2 Property distributions along a roughly north–south section across the Southern Ocean at 1401 E south of Australia (World Ocean Circulation Experiment section SR3): (a) potential temperature (1C), (b) salinity (PSS78), (c) neutral density (kg m 3), (d) oxygen (mmol kg 1), (e) chlorofluorocarbon 11 (CFC-11) (pM kg 1), and (f) transport at each station pair (solid line, left axis) and cumulative transport from south to north (dashed line, right axis). Transport is in units of sverdrups (1 Sv ¼ 106 m3 s 1). Contours slope upward from north to south; regions where the slope of the contours is steep correspond to the ACC fronts and to transport maxima. SAMW, Subantarctic Mode Water; LCDW, Lower Circumpolar Deep Water; AAIW, Antarctic Intermediate Water; UCDW, Upper Circumpolar Deep Water; CDW, Circumpolar Deep Water CDW; AABW, Antarctic Bottom Water. SAF, Subantarctic Front; PF, Polar Front; sACCf, southern ACC front; SB, southern boundary; ASF, Antarctic Slope Front; N and S indicate northern and southern branches of the primary fronts. The positive and negative peaks in transport labeled SAZ indicate a strong recirculation in the Subantarctic Zone north of the ACC at the time this section was occupied. Sections are adapted from Orsi AH and Whitworth T, III (2005) In: Sparrow M, Chapman P, and Gould J (eds.) Hydrographic Atlas of the World Ocean Circulation Experiment (WOCE), Vol. 1: Southern Ocean. Southampton: International WOCE Project Office (ISBN 0-904175-49-90), with permission (http://www.soc.soton.ac.uk).
with the fronts therefore also extend throughout the water column. The deep-reaching nature of the ACC fronts reflects the weak stratification of the Southern Ocean, where the change in density from surface
waters to deep waters is small compared to lower latitudes. The mean current speeds of the ACC jets are relatively modest, typically less than 0.5 m s 1 (about 1 knot), with much weaker flow between the
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Figure 2 Continued.
jets. However, because the current is broad and deep, it carries a large transport. Most studies of the ACC now recognize three main fronts. From north to south, these fronts are the Subantarctic Front, the Polar Front, and the southern ACC front. The southern limit of the ACC domain is defined by the ‘southern boundary of the ACC’, the southernmost streamline to pass through Drake Passage, which is also often associated with a weak front and current core. These fronts of the ACC are circumpolar in extent and can be found on any north–south transect across the Southern Ocean. Because the fronts extend to the seafloor and the
stratification in the Southern Ocean is relatively weak, the position of the fronts is strongly influenced by bathymetry. Between the fronts lie zones with more or less uniform physical and chemical characteristics: the Subantarctic Zone between the Subantarctic and Subtropical Fronts, the Polar Frontal Zone between the Subantarctic and the Polar Fronts, and the Antarctic Zone south of the Polar Front. Within each zone, the water properties tend to be similar at each depth and follow a similar seasonal cycle. For example, the Subantarctic Zone is characterized by very deep surface mixed layers in winter and low
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Figure 2 Continued.
nutrient concentrations (close to zero for silicic acid and somewhat higher concentrations of nitrate and phosphate). The Antarctic Zone is characterized by fresh surface waters, shallow summer mixed layers, and high concentrations of major nutrients like nitrate and silicic acid, but low concentrations of micronutrients like iron. The zones delimited by the fronts of the ACC also define biogeographic zones populated by distinct species assemblages. For example, waters south of the Polar Front tend to be dominated by large phytoplankton such as diatoms (who need silicic acid) and large zooplankton, while coccolithophores
and small zooplankton dominate north of the Subantarctic Front. The distribution and foraging patterns of larger animals (e.g., fish, seabirds, and marine mammals) are also influenced by the frontal structure of the ACC. In some cases, the fronts themselves tend to be associated with higher primary productivity. The higher productivity near fronts can be caused by advection of micronutrients by the current or by upwelling caused by eddies or by topographic interactions. The currents of the ACC can also play a direct role in ecosystem dynamics, for example, by carrying krill and larvae from the Antarctic Peninsula to South Georgia.
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Circulation of the Antarctic Circumpolar Current A map of the elevation of the sea surface (dynamic height) shows how the position and intensity of the current varies along its circumpolar path (Figure 3). Contours of dynamic height are approximate streamlines for the flow, so the current flow is rapid in regions where the contours are closely spaced and weak in regions where the contours are widely separated. The steering of the fronts by large bathymetric features is clearly illustrated in Figure 3. Recent studies using high-resolution sampling from ships and satellites have revealed a more complex structure to the ACC than previously appreciated.
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The ACC fronts consist of multiple branches, which merge and diverge in different regions and at different times along the circumpolar path of the current system (Figures 3 and 4). The multiple jets in the ACC reflect the tendency for geophysical flows on a sphere to self-organize into narrow, elongated, persistent zonal flows; similar circulation patterns form in the Earth’s atmosphere and on other planets. The position of the fronts varies with time, but generally over a relatively small latitude range at any given longitude (typically 711 of latitude). The variability is larger downstream of major bathymetric features and in regions where the ACC fronts interact with the strong boundary currents of the subtropical gyres to the north (e.g., south of Africa).
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Figure 2 Continued.
Transport of the Antarctic Circumpolar Current During the World Ocean Circulation Experiment (WOCE) in the 1990s, the transport through Drake Passage was estimated to be 13778 106 m3 s 1, similar to the estimates made in the late 1970s. Because the Atlantic basin is nearly closed to the north of the Southern Ocean, the net transport between Africa and Antarctica must be very close to the Drake Passage transport (to within c. 1 106 m3 s 1). The transport between Australia and Antarctica must be somewhat greater, to compensate for the flow from the Pacific to the Indian Oceans through the
Indonesian archipelago. Repeat transects during WOCE showed that the baroclinic transport south of Australia is 147710 106 m3 s 1, consistent with estimates that about 10–15 106 m3 s 1 flows through the Indonesian passages. The fronts, in particular the Subantarctic and Polar Fronts, carry most of the ACC transport. The relative contribution of these two fronts to the total transport varies around the circumpolar path. For example, south of Australia the Subantarctic Front carries 4 times more water to the east than the Polar Front, while in Drake Passage the transport carried by the two fronts is roughly equal in magnitude.
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ANTARCTIC CIRCUMPOLAR CURRENT
(f)
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Latitude (°S) Figure 2 Continued.
The transport of water masses by the ACC also changes with longitude. For example, the ACC carries an excess of intermediate density water into the Atlantic through Drake Passage, which is compensated by an excess of deep water leaving the basin south of Africa. These changes in water mass transports by the ACC reflect water mass transformations in the Atlantic basin, where relatively light Antarctic Intermediate Water is converted to denser North Atlantic Deep Water. The ACC also transports vast amounts of heat, fresh water, nutrients, carbon, and other properties between the ocean basins. The transport of the ACC varies over a range of timescales. Multiyear deployments of bottom pressure recorders in Drake Passage during the late 1970s and 1990s suggest a standard deviation in net transport of about 8–10 106 m3 s 1. For periods shorter than about 6 months, most of the variability is due to changes in sea level (i.e., changes in the barotropic, or depth-independent, flow). Models and sea level measurements suggest these barotropic motions are highly correlated with changes in wind stress and tend to follow bathymetric contours (more precisely, the flow is along lines of constant planetary vorticity, where planetary vorticity is given by the Coriolis parameter (equal to twice the rotation rate of the Earth multiplied by the sine of the latitude) divided by the ocean depth). For longer periods, variations in the density field (and hence the
baroclinic, or depth-varying, flow) also become important.
Dynamics of the Antarctic Circumpolar Current The absence of continental barriers in the latitude band of Drake Passage makes the dynamics of the ACC distinctly different in character from those of currents at other latitudes. Simple wind-driven ocean circulation theory (the Sverdrup balance), which generally does a good job of describing the circulation of the upper ocean in basins bounded by continents, cannot be applied in the usual way in a continuous ocean channel. The dynamical balance of the ACC has therefore been a topic of great interest for many years. The strong westerly winds over the Southern Ocean have long been recognized to help drive the ACC. The winds drive surface waters to the left of the wind (see Ekman Transport and Pumping), causing upwelling to the south of the wind stress maximum and downwelling to the north. This pattern of upwelling and downwelling helps to establish the tilt of density surfaces across the ACC and therefore its geostrophic flow. Despite the absence of continental barriers, the Sverdrup theory of wind-driven currents has been applied in the Southern Ocean by assuming
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ANTARCTIC CIRCUMPOLAR CURRENT ∇ SSH , 129−130° E 1994
1994.5
1995
1995.5
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1997 40
50
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Latitude (° S)
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Figure 4 The north–south gradient of sea surface height (in m per 100 km) at 1301 E, south of Australia. Large height gradients indicate strong currents. The multiple bands of large gradient correspond to the jets of the ACC. Note that the jets merge and diverge, and change in intensity with time, and also persist for many months at the same latitude. Reproduced with permission from Sokolov S and Rintoul SR (2007) Multiple jets of the Antarctic Circumpolar Current. Journal of Physical Oceanography 37: 1394–1412. & Copyright [2007] AMS.
that relatively shallow bathymetric features act as ‘effective continents’. However, such calculations assume no interaction between the current and the bathymetry, which we know to be a poor
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assumption for the deep-reaching ACC. In addition, wind is not the only factor driving the ACC. The atmosphere also drives ocean currents by exchanging heat and fresh water with the ocean, causing the density of seawater to change. Exchange of fresh water can result from precipitation, evaporation, and the freezing and melting of ice, both sea ice and glacial ice in the form of icebergs and ice shelves. Because the speed of ocean currents is proportional to the horizontal gradient of density, any process that produces horizontal density gradients will drive ocean currents. In the case of the ACC, both the strong westerly winds and the air– sea exchange of buoyancy play a part in driving the current. The momentum supplied to the Southern Ocean by the wind needs to be compensated in some way. The question of what balances the wind forcing has been a topic of debate for many decades. Recent studies have confirmed the early hypothesis by W. Munk and E. Palme´n that interaction of the ACC with seafloor topography provides a force to balance the wind. This force, known as the bottom form stress, results when the ocean currents are organized such that there is higher pressure on one side of a ridge on the seafloor than is found on the other side. In the case of the ACC, higher pressure is generally found on the west side of topographic ridges or hills, providing a force from the solid Earth to the ocean that balances the wind stress at the sea surface. While these pressure differences are too small to observe directly, realistic numerical simulations clearly show this force balance in action. Eddies produced by dynamical instabilities of the ACC fronts play a crucial role in establishing the momentum and heat balance of the Southern Ocean. The ACC has some of the most vigorous eddy activity observed in the ocean (Figure 5). Eddies are produced when dynamical processes release some of the energy stored in the sloping of density surfaces across the ACC, converting some of the energy in the mean flow into motions that vary with time, or eddies. This process is called baroclinic instability. The eddies transfer momentum vertically from the sea surface to the deep ocean, helping to set up the system of deep currents that interact with bathymetry to provide the bottom form stress. Both transient eddies (motions that vary with time) and standing eddies (deviations of the flow from the east–west average) contribute to the momentum transport. At the same time, the eddies carry heat poleward across the ACC, to compensate the heat lost to the cold atmosphere near Antarctica. Attempts to relate ACC transport variability to variations in wind have been inconclusive. It is now
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1/6° HIM
Day 365 Africa
S Am outh eri ca
The Antarctic Circumpolar Current and the Overturning Circulation
Au st ra
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Antarctica
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ACC starts to speed up, but this causes more vigorous eddy fluxes that dissipate the extra energy in the mean flow.
−2 −1.8 −1.6 −1.4 −1.2 −1 −0.9 −0.8 −0.7 −0.6 −0.5 −0.4 −0.3 −0.2 −0.1 0
0.2 0.4 0.6
log10 of magnitude of velocity averaged over top 100m in ms−1
Figure 5 A snapshot of surface speed from a high-resolution numerical simulation of the ACC. The filamented, eddy-rich structure of the ACC stands out clearly. Adapted with permission from Hallberg R and Gnanadesikan A (2006) The role of eddies in determining the structure and response of the Southern Ocean overturning: Results from the Modelling Eddies in the Southern Ocean project. Journal of Physical Oceanography 36: 2232–2252. & Copyright [2007] AMS.
understood that the dynamical balance of the ACC depends on a number of factors, including wind and buoyancy forcing, eddy–mean flow interaction, topographic form stress and the ocean stratification, so a simple relationship between transport and wind should not be expected. Recent improvements in ocean observing systems have allowed changes in the ACC to be assessed for the first time. Comparison of temperature profiles collected since the 1950s suggests that much of the Southern Ocean has warmed. The warming is largest in the ACC belt and is consistent with a southward shift of the current, allowing warm water to move south into areas previously occupied by cooler water. The southward movement of the ACC has been linked to a southward shift and strengthening of the westerly winds. The change in the winds, in turn, has been linked both to loss of ozone in the polar stratosphere and to enhanced greenhouse warming. The transport of the ACC has apparently not changed much over this time period, despite the change in wind forcing. Recent studies suggest the ACC may be in a regime in which the transport is insensitive to changes in wind: as the wind forcing increases, the
The eastward flow of the ACC is dynamically linked to a weaker circulation in the north–south plane. The distribution of water properties on transects across the Southern Ocean clearly reveals water masses spreading across the ACC. For example, the salinity maximum of the Lower CDW and oxygen minimum of the Upper CDW can be traced as they shoal from depths of 2000–3500 m north of the ACC to approach the sea surface south of the Polar Front (Figure 2). The high-oxygen, low-salinity waters formed in the Southern Ocean (Antarctic Intermediate Water and Antarctic Bottom Water) can be followed as they cross the ACC and enter the basins to the north. These distributions reflect an ocean circulation pattern known as the overturning circulation (Figure 6). Deep water spreads to the south across the ACC and upwells at the sea surface. Some of the upwelled deep water is driven north beneath the westerly winds, gains heat and fresh water from the atmosphere, and therefore becomes less dense, and ultimately sinks to form Antarctic Intermediate Water and Subantarctic Mode Water. Deep water that upwells further south and closer to Antarctica is converted to denser Antarctic Bottom Water and returns to the north. The result of these water mass transformations is a circulation in the north–south plane that consists of two counter-rotating cells. According to the residual mean theory, the strength of the net overturning circulation (mean flow plus the eddy contribution) is determined by the surface buoyancy forcing. The Southern Ocean overturning cells play an important part in the global-scale overturning circulation. The Southern Ocean imports deep water from the basins to the north, and exports bottom water and intermediate water. Recent studies suggest the conversion of deep water to intermediate water in the Southern Ocean is a key link in the global overturning circulation. For decades it has been assumed that the sinking of dense water in the polar regions was balanced by widespread upwelling at lower latitudes. However, measurements of mixing rates and large-scale tracer budgets suggest that mixing in the interior of the ocean is an order of magnitude too weak to support the upwelling required. The transformation of deep water to
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ANTARCTIC CIRCUMPOLAR CURRENT
Buoyancy loss
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Latitude Figure 6 A schematic view of the Southern Ocean overturning circulation. Upper Circumpolar Deep Water (UCDW) upwells at high latitude and gains buoyancy (from heating, precipitation, and ice melt) as it is driven north, to ultimately sink as Antarctic Intermediate Water (AAIW) or Subantarctic Mode Water (SAMW). Lower Circumpolar Deep Water (LCDW) and North Atlantic Deep Water (NADW) upwell closer to the Antarctic continent, are made more dense (by cooling and salt rejected during sea ice formation), and sink to form Antarctic Bottom Water (AABW). The major southern ocean fronts are indicated. STF, Subtropical Front; SAF, Subantarctic Front; PF, Polar Front. Reproduced with permission from Speer K, Rintoul SR, and Sloyan B (2000) The diabatic Deacon cell. Journal of Physical Oceanography 30: 3212–3222. & Copyright [2007] AMS.
intermediate water by air–sea buoyancy exchange in the Southern Ocean provides an alternative means of connecting the upper and lower limbs of the global overturning circulation. Mixing likely also makes a contribution in regions of rough bathymetry, including the Southern Ocean, where elevated mixing rates have been measured. Eddies spawned by the ACC make an important contribution to the overturning circulation. The eddies transfer mass across the Drake Passage gap, where the absence of land barriers means that there can be no net east–west pressure gradient and therefore no net north–south flow. Furthermore, the same forces that drive the overturning circulation (wind and buoyancy exchange) also drive the ACC. The west-to-east flow of the ACC and the overturning circulations cannot be understood in isolation. The two are intimately linked, and eddy fluxes, topographic interactions, and wind and buoyancy forcing are all important ingredients of the dynamical coupling between them. The Southern Ocean overturning also has a large influence on global biogeochemical cycles. Upwelling of nutrient-rich deep water south of the ACC returns nutrients to the surface layer. The nutrients are exported from the Southern Ocean to lower latitudes by the overturning circulation, ultimately supporting a large fraction of global primary production. Water
masses at the surface of the Southern Ocean exchange oxygen and carbon dioxide with the atmosphere and carry ‘ventilated’ water to the interior of the ocean when the water masses sink. Where deep water upwells south of the ACC, carbon dioxide is released to the atmosphere; where water sinks from the sea surface, carbon dioxide is carried from the atmosphere into the ocean. As a result of the overturning circulation, more of the carbon released by human activities is accumulating just north of the ACC than in any other latitude band of the ocean.
Summary The ACC is the largest current in the world ocean, carrying about 13778 106 m3 s 1 from west to east around Antarctica. By connecting the ocean basins, the ACC allows water masses and climate anomalies to propagate between the basins. The current flow is concentrated in a number of circumpolar fronts, which extend from the sea surface to the seafloor. The fronts also mark the boundaries between zones with distinct physical, chemical, and ecological characteristics. Eddies produced by dynamical instabilities of the fronts play an important part in the dynamics of the ACC by transporting momentum vertically and heat and mass poleward.
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Both wind and buoyancy forcing contribute to driving the ACC. Interaction between the deep-reaching flow and the bottom topography establish bottom form stresses to balance the wind forcing. The strong eastward flow of the ACC is intimately connected to an overturning circulation made up of two counterrotating cells. Water mass transformations driven by exchange of heat and moisture with the atmosphere connect the upper and lower limbs of the thermohaline circulation. The transport and storage of heat, fresh water, and carbon dioxide by the ACC have a significant influence on global and regional climate.
See also Bottom Water Formation. Carbon Cycle. Current Systems in the Southern Ocean. Ekman Transport and Pumping. Energetics of Ocean Mixing. Heat Transport and Climate. Mesoscale Eddies. Ocean Circulation. Ocean Subduction. Satellite Altimetry. Shelf Sea and Slope Sea Fronts. Ocean Circulation: Meridional Overturning Circulation. Water Types and Water Masses. Weddell Sea Circulation. Wind- and Buoyancy-Forced Upper Ocean. Wind Driven Circulation.
Further Reading Cunningham S, Alderson SG, King BA, and Brandon MA (2003) Transport and variability of the Antarctic Circumpolar Current in Drake Passage. Journal of Geophysical Research 108: 8084 (doi:10.1029/2001 JC001147). Deacon G (1984) The Antarctic Circumpolar Ocean. London: Cambridge University Press. Gille ST (2002) Warming of the Southern Ocean since the 1950s. Science 295: 1275--1277. Gordon AL and Molinelli E (1986) Southern Ocean Atlas. Washington, DC and New Delhi: National Science Foundation and Amerind Publishing. Hallberg R and Gnanadesikan A (2006) The role of eddies in determining the structure and response of the Southern Ocean overturning: Results from the
Modelling Eddies in the Southern Ocean project. Journal of Physical Oceanography 36: 2232--2252. Munk WH and Palme´n E (1951) Note on the dynamics of the Antarctic Circumpolar Current. Tellus 3: 53--55. Niiler PP, Maximenko NA, and McWilliams JC (2003) Dynamically balanced absolute sea level of the global ocean derived from near-surface velocity observations. Geophysical Research Letters 30: 2164. Nowlin WD Jr. and Klinck JM (1986) The physics of the Antarctic Circumpolar Current. Review of Geophysics and Space Physics 24: 469--491. Olbers D, Borowski D, Volker C, and Wolff J-O (2004) The dynamical balance, transpor‘t and circulation of the Antarctic Circumpolar Current. Antarctic Science 16: 439--470. Orsi AH and Whitworth T, III (2005) In: Sparrow M, Chapman P, and Gould J (eds.) Hydrographic Atlas of the World Ocean Circulation Experiment (WOCE), Vol. 1: Southern Ocean. Southampton: International WOCE Project Office (ISBN 0-904175-49-9). Orsi AH, Whitworth T, III, and Nowlin WD (1995) On the meridional extent and fronts of the Antarctic Circumpolar Current. Deep-Sea Research I 42: 641--673. Rintoul SR, Hughes C, and Olbers D (2001) The Antarctic Circumpolar Current system. In: Siedler G, Church J, and Gould J (eds.) Ocean Circulation and Climate, pp. 271--302. London: Academic Press. Sokolov S and Rintoul SR (2007) Multiple jets of the Antarctic Circumpolar Current. Journal of Physical Oceanography 37: 1394--1412. Speer K, Rintoul SR, and Sloyan B (2000) The diabatic Deacon cell. Journal of Physical Oceanography 30: 3212--3222. Whitworth T, III (1980) Zonation and geostrophic flow of the Antarctic Circumpolar Current at Drake Passage. Deep-Sea Research 27: 497--507. Whitworth T, III and Nowlin WD (1987) Water masses and currents of the Southern Ocean at the Greenwich Meridian. Journal of Geophysical Research 92: 6462--6476.
Relevant Website http://www.soc.soton.ac.uk – Electronic Atlas of WOCE Data.
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ANTARCTIC FISHES I. Everson, Anglia Ruskin University, Cambridge, UK & 2009 Elsevier Ltd. All rights reserved.
Introduction Antarctica is a continental landmass much of which is covered by an ice cap, consequently the ichthyofauna is totally marine. Surrounding the continent is the Southern Ocean, approximately 36 million km2, continuous with the Atlantic, Indian, and Pacific Ocean basins to the north and whose northern limit is generally taken as the Antarctic Polar Frontal Zone (APFZ). There is a clear separation between the Antarctic and the Southern Hemisphere continents, the nearest connection being with South America via the Scotia Arc, a series of islands separated from each other by deep water. The Antarctic Circumpolar Current (ACC) and the general oceanographic regime mean that marine isotherms are more or less concentric around the continent. Close to the continent the seasonal variation in temperature is rarely more than 1 1C while even at the northern limit, as for example at South Georgia, the range is little more than 4 1C. These two factors, geographical isolation and constant low temperature, have a major effect on Antarctic fish.
marine environment. Antarctic notothenioids with their high species diversity and endemism form a species flock comparable to that of Lake Baikal. Early taxonomic studies were based on the traditional methods of morphometric and meristic analyses. Recent studies have used molecular biological analyses not only of nuclear material but also of antifreeze compounds to indicate phylogeny.
Adaptations Cold Adaptation
Some of the earliest studies on the physiology of Antarctic fish concerned the measurement of oxygen uptake rates. Initially it had been assumed that, since many biochemical processes are temperaturedependent, the metabolic rates of Antarctic fish might be very low. The initial experiments indicated that rates were substantially higher than those of temperate fish when studied at low temperature and the degree of elevation of the metabolic rate in Antarctic fish was attributed to a phenomenon termed ‘cold adaptation’. Subsequent studies demonstrated that the greater part of this elevation was caused by handling stress and the extended recovery Table 1
Composition of Southern Ocean ichthyofauna
Taxon
Fish Fauna The Southern Ocean ichthyofauna is relatively sparse and unusual in composition, consisting of 213 species belonging to only 18 families (Table 1 and Figure 1). Nearly half the species belong to one group, the perciform notothenioids, which make up 45% of the fish fauna. Restricting consideration to the shelf, and particularly in the highest latitudes, notothenioids make up 77% of the species and 90–95% of the biomass of fish. Notothenioids are morphologically and ecologically diverse, and have variegated into a wide variety of niches, mainly demersal, and also in the water column and even within sea ice. As a group, this makes them more diverse than, for example, the finches of the Gala´pagos archipelago. The concept of species flocks has been developed for freshwater fish to identify groups that have a close affinity; typically such flocks are to be found in ancient lake systems and it is extremely unusual for such a flock to be identified from a large
Agnatha Chondrichthyes Osteichthyes Notacanthiformes Anguilliformes Salmoniformes Stomiiformes Aulopiformes Myctophiforms Gadiformes Ophidiiformes Lophiiformes Lampriformes Beryciformes Zeiformes Scorpaeniformes Perciformes Zoarcidei Notothenioidei Blennioidei Scombroidei Stromateoidei Pleuronectiformes
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Benthic 2 8
Benthopelagic
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Pelagic
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2 2 4
9 1
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5 12 9 35 1 1 3 2 6
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Anchor ice (to 30 m)
Annual ice Platelet ice Trematomus nicolai
Pagothenia borchgrevinki Pleuragramma antarcticum
Dissostichus mawsoni Trematomus bernacchii Trematomus loennbergii
Figure 1 These six species from McMurdo Sound demonstrate some of the life-history types included in the Nototheniidae. Pelagic, cryopelagic, epibenthic, and benthic species are illustrated. Dots indicate typical habitat, although most species have considerable depth ranges. Modified from Eastman and DeVries (1986).
time, of the order of 24 h or more, following introduction into respirometers. In spite of this, it is now accepted that some slight elevation of metabolic rate remains that cannot be explained wholly by experimental technique. Consideration of the phenomenon has raised some controversy between different workers. The existence of the phenomenon has been demonstrated experimentally, although it does not appear to confer any evolutionary advantage because it implies a higher energy requirement on the part of the fish. All these studies have been undertaken on whole fish; the overall oxygen uptake rate being the balance between all the component metabolic pathways that are present. As such it has been argued that the term ‘cold adaptation’ has little meaning and that it is more sensible to consider each metabolic component separately to provide an overall balance. Antifreeze
Pure water freezes at 0 1C, but the presence of salts causes the freezing point to be depressed such that normal seawater freezes at around 1.85 1C. At McMurdo Sound the annual mean water temperature is 1.87 1C and varies within the range 1.40 to 2.15 1C. Body fluids, such as the blood plasma, of most teleost fish have a freezing point of c. 0.7 1C. Even though this difference is small it is important, because living in waters close to the freezing point of seawater, Antarctic fish require some mechanism to prevent their body fluids from freezing. In the absence of ice, fish could live in a supercooled state. Unfortunately this is not a stable state because very few ice crystals are required to cause a supercooled liquid to freeze. An alternative adaptation is required. The ionic concentration of the blood of most marine teleosts is 320–380 mOsm kg 1, only about one-third of that of Antarctic seawater (1050 mOsm kg 1). The freezing
point depression of some notothenioids at McMurdo Sound is 2.2 1C, although their blood osmolality is 550–625 mOsm kg 1, equivalent to a freezing point depression of 1.02 to 1.16 1C. Thus although there appears to be some compensation as measured by the osmolality, it is insufficient to explain all of the depression in freezing point. Compensation for this difference comes in the form of antifreeze glycopeptides (AFGPs) which exert their effect by a mechanism known as adsorption-inhibition (Figures 2 and 3). Even though ice crystals can form, their further growth is prevented when AFGPs are adsorbed onto them because the AFGP molecule prevents growth of the ice crystal along its main axis. Thus the AFGPs have an antifreeze function, lowering the freezing point beyond that which would be expected from the osmolality. The AFGPs however do not lower the melting point. The AFGP molecules are of such a size that they would be lost through the glomeruli of normal teleost kidneys. In glomerular nephrons of normal teleosts, molecules with a molecular weight of o68 000 Da pass through the filtration barrier. As the urine passes through the different parts of the nephron, it is modified by reabsorption of nonwaste products and secretion of waste products. The AFGP molecules are of such a size that they would pass through the glomeruli but would need to be reabsorbed later on in the nephron. The kidneys of all Antarctic fish which possess AFGPs are aglomerular, obviating this requirement. Thus the evolution of the aglomerular trait in Antarctic fish complements that of the presence of antifreeze. Cardiovascular
A continuous low water temperature means that the oxygen-carrying potential of seawater is high. Thus, as long as the partial pressure of oxygen in the
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ANTARCTIC FISHES
seawater remains high, so will the available oxygen. It is against this background that further cardiovascular adaptations have evolved. Early taxonomic studies relied on specimens preserved in alcohol or formalin, both of which affect the color of the fish. Because fish typically possess red H
O H
N
C C
CH3
H
C
N
N
O H C
C
C
CH 3 H CH2OH
O H
C
O
O
H
OH
CH3
H H
CH 2OH O
O
HO H OH H
H H H
H
NH C
O
CH 3
OH
Figure 2 Basic repeating structural unit of the AFGPs of notothenoids. The peptide consists of amino acids in the sequence [alanyl-alanyl-threonine]n. Each threonine is joined to a disaccharide through a glycosidic linkage. In low-molecularweight AFGPs 6–8, proline is periodically substituted for alanine at position 1 of the tripeptide. Reproduced from Eastman JT (1993) Antarctic Fish Biology: Evolution in a Unique Environment. London: Academic Press.
th grow Ice sence b in a GPs h owt e of AF r g Ice esenc r in p GPs F of A
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blood, until the 1950s no mention was made of the anemic appearance of the gills of some species of Antarctic fish. At that time, it was noticed that members of the Channichthyidae (at that time called Chaenichthyidae) were white, as a result of which they were called ‘white-blooded fish’ or ‘icefish’. The blood of channichthyids is devoid of hemoglobin, although small numbers of nonfunctional erythrocytes have been described in a few species. Initial consideration was given to determine whether, because channichthyids do not possess scales, cutaneous respiration might be a major factor in oxygen uptake. However, the absorptive area and vascularization relative to the gills mitigated against that mechanism. Alternatively it was thought possible that channichthyids possessed either a more efficient oxygen utilization mechanism or else lowered oxygen requirement. This second consideration was being examined at a time when the concept of metabolic cold adaptation was under discussion. The viscosities of the plasma of red-blooded notothenioids and channichthyid fish are very close, although the blood of the former is approximately 25% higher than the latter. Studies on oxygen uptake rates indicated that channichthyids utilized oxygen at a slightly lower rate as compared with equivalent red-blooded notothenioids. In the absence of hemoglobin, the oxygen-carrying capacity of channichthyid blood is only about one-tenth that of
c-axis
a-axis a-axis Step
Long straight growth fronts
Basal plane
Adsorbed AFGP molecules
Prism face
Small, highly curved growth fronts
Figure 3 Model of an ice crystal depicting adsorption-inhibition as a mechanism for the freezing point depression of water by antifreezes. In the absence of AFGPs, ice crystal growth occurs as water molecules are added to the crystal in a regular fashion at steps on the basal planes. When the AFGPs are adsorbed, ice cannot propagate over them and long straight fronts become divided into many small curved fronts. Reproduced from Eastman JT (1993) Antarctic Fish Biology: Evolution in a Unique Environment. London: Academic Press.
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red-blooded fish. Two mechanisms are possible to compensate for this effect: either channichthyid blood is circulated at a much faster rate or there is much more of it in the system. The latter has proven to be the case and channichthyid blood takes up 8–9% of the total volume of the fish (2–4 times that of other teleosts); the heart rate and blood pressure are low but the stroke volume and resultant cardiac output are large. To reduce the resistance to flow, the capillaries are larger than in other teleosts and the blood is less viscous. Even though the hemoglobin-less condition is clearly effective, it is a feature that confines the fish to areas of high oxygen tension such as those present in Antarctic waters. Only one channichthyid species, Champsocephalus esox, is found outside of the Antarctic zone. Experimental studies have demonstrated that channichthyids are particularly sensitive to hypoxia, indicating that in their natural habitat the oxygen saturation is always consistently high.
Weddell Sea Circulation. Wind- and BuoyancyForced Upper Ocean.
Further Reading Cheng C-HC, Cziko PA, and Evans CW (2006) Nonhepatic origin of notothenioid antifreeze reveals pancreatic synthesis as common mechanism in polar fish freezing avoidance. Proceedings of the National Academy of Sciences of the United States of America 103: 10491--10496. Clarke A (1991) What is cold adaptation and how should we measure it? American Zoologist 31: 81--92. Di Prisco G, Pisano E, and Clarke A (eds.) (1998) Fishes of Antarctica; A Biological Overview. Milan: Springer. Eastman JT (1993) Antarctic Fish Biology: Evolution in a Unique Environment. London: Academic Press. Gon O and Heemstra PC (eds.) (1990) Fishes of the Southern Ocean. Grahamstown: JLB Smith Institute of Ichthyology. Kock K-H (1992) Antarctic Fish and Fisheries. Cambridge, MA: Cambridge University Press.
See also Antarctic Circumpolar Current. Current Systems in the Southern Ocean. Fish Ecophysiology. Fish Feeding and Foraging. Fish Reproduction.
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ANTHROPOGENIC TRACE ELEMENTS IN THE OCEAN E. A. Boyle, Massachusetts Institute of Technology, Cambridge, MA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 1, pp 162–169, & 2001, Elsevier Ltd.
Introduction Human activities have increased the fluxes of several chemical elements into the ocean above natural levels. Despite convincing evidence for this enhancement of elemental fluxes (see Further Reading section for references relevant to the discussion in this article), there is only one element – lead (Pb) – where abundant evidence proves that open-ocean seawater concentrations are substantially higher than they were in preindustrial times. For a few other elements – e.g. cadmium (Cd) and mercury (Hg) – there is some evidence suggesting a detectable anthropogenic impact (or models indicating that an anthropogenic enhancement must exist even if it has not been observed). For most other elements, the size of the oceanic reservoir of these elements overwhelms relatively large anthropogenic fluxes, and it may require centuries of further inputs before the human impact can be discerned.
Anthropogenic Lead in the Ocean Sampling and analysis for Pb have been difficult because of low concentrations and abundant contamination sources: lead paint, lead weights, and gasoline exhausts. The latter source is now substantially reduced, so perhaps the current Pb contamination problem is less serious than it was in the 1980s. Patterson and co-workers were the first to call attention to the overwhelming anthropogenic augmentation of Pb fluxes into the environment, particularly that resulting from alkyl leaded gas utilization. Patterson’s evidence began with the demonstration that Pb deposition in remote Greenland snows had increased by two orders of magnitude. In the late 1970s/early 1980s, his laboratory obtained the first valid data (uncontaminated and properly analyzed) for the vertical distribution of Pb in water sample profiles from the North Atlantic, North Pacific, and South Tropical Pacific (Figure 1). These data demonstrated that the highest concentrations of
Pb occurred in the surface ocean and that concentrations decreased with increasing depth in the water column. At that time, the highest Pb concentrations were found in the North Atlantic Ocean (160 10 12 moles kg 1 at the surface decreasing to 26 10 12 moles kg 1 at 3000 m water depth). The high Pb concentrations in this basin are emitted from the major industrial nations surrounding the basin. Lower Pb concentrations were seen in North Pacific surface waters (60 10 12 moles kg 1 at the surface decreasing to 5 10 12 moles kg 1 at 3500 m depth), and the lowest concentrations in the south tropical Pacific (20 10 12 moles kg 1 at the surface decreasing to 4 10 12 moles kg 1 at 4000 m). In Patterson’s view, this evidence proved the anthropogenic origin of Pb in the modern ocean. The lead industry attempted to discredit Patterson’s evidence on environmental Pb pollution by many spurious arguments. Although they were not clever enough, they might have attempted to discredit Patterson’s interpretation of his oceanic Pb data by pointing out that similar comparative concentration variations occur for aluminum (Al) in the ocean, even though oceanic Al is entirely of natural origin. Al is released from terrestrial dusts blown into the ocean and ‘scavenged’ onto sinking biologically produced particles that remove it from the deep ocean. Therefore Al is high in the surface ocean, highest downwind of major dust sources such as north-west Africa, and lowest in the South Pacific because of low dust inputs to the surface and cumulative scavenging in the deep waters. Lead was phased out of gasoline in the USA (followed soon by Canada and Japan, and somewhat later by western Europe and a few other countries) beginning in 1970 when the US Environmental Protection Agency mandated emissions controls on gasoline exhausts. Originally, controls on lead emissions were not the goal of the regulations; the regulations were formulated to minimize emissions of hydrocarbons, nitrogen oxides, and carbon monoxide. As it turned out, the technological fix for those problems was to use catalytic converters on the exhaust stream, and the activity of the catalysts was destroyed by lead exhausts. Hence leaded gasoline could not be used with catalytic converters, and regulations mandated the elimination of leaded gasoline. Later, regulations specifically directed at minimizing lead emissions into the environment were also enacted.
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Pb (10 50
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moles kg 1)
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5000 Figure 1 Oceanic Pb profiles obtained by Patterson and co-workers on samples collected between 1976 and 1981.
Changing leaded gasoline utilization patterns offer an opportunity to establish that Pb decreases in the ocean in response to the decreasing Pb emissions from leaded gasoline. Pb gas utilization increased from its introduction in the 1920s until the 1970s, then decreased rapidly – resulting in nearly complete elimination from the USA by 1990 and from western Europe by the turn of the century (Figure 2). Surface waters of the ocean should respond within a few years to changes in fluxes from the atmosphere, and the upper layers of the ocean should respond on decadal timescales. These expectations are based upon the penetration of the nuclear bomb tritium into the ocean and natural radioisotope 210Pb (halflife 22.3 years). Studies of tritium penetration show that surface waters ventilate the upper thermocline of the ocean on a timescale of several years to several decades. 210Pb is supplied to the ocean from windblown aerosols which have acquired 210Pb from the decay of the radioactive noble gas 222Rn. 222Rn is released into the atmosphere as a decay product of
naturally occurring crystal 238U which decays in several stages to the immediate parent of 222Rn, 226 Ra. The flux of 210Pb out of the atmosphere has been measured at numerous sites, therefore the flux of 210Pb into the ocean can be estimated reasonably well. In the surface ocean, 210Pb is rapidly scavenged from sea water by newly formed biological material whose residues eventually sink out of the surface waters carrying 210Pb. The steady-state concentration of 210Pb in the mixed layer of the ocean is thus determined by a balance between 210Pb supply from the atmosphere and 210Pb removal by sinking particles. The time constant for this process is estimated by dividing the measured surface water 210Pb reservoir by the flux of 210Pb from the atmosphere. The rate of removal varies with the rate of biological activity, but in typical open-ocean ‘deserts’ such as the Sargasso Sea, 210Pb is removed from surface waters every 2 years. Thus as the flux of anthropogenic lead from the atmosphere decreases, the concentration of lead in surface waters should follow this decrease with a 2 year lag.
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US and European gasoline Pb consumption, 1930 _93 300
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Gasoline Pb consumption (10 metric tonnes per year)
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Year Figure 2 Alkyl Pb gas consumption, 1930–93. Four European countries (accounting for 70% of EC gasoline consumption) stacked to compare with the much higher US consumption.
Decreasing concentrations of Pb in surface waters of the Atlantic Ocean have been documented by observations from 1979 until the present (Figure 3). The decreasing concentration of Pb in the upper layers of the ocean is also observed (Figure 4), with a slower response in the deeper waters that are replaced on decadal timescales. These decreases are seen in other parts of the Atlantic and North Pacific as well. This evidence amply demonstrates that the phasing out of leaded gasoline in the USA has been closely followed by decreases of the Pb concentration of the ocean on the appropriate timescale. In the period before direct observations of Pb in the ocean, the response of oceanic Pb to earlier changes in emissions are documented by the Pb content of reef-building corals. Corals precipitate calcium carbonate skeletons with annual variations in their density that can be counted in the same way
as tree rings, hence providing a chronology of the time of deposition. In corals from near Bermuda, concentrations of Pb increased from very low levels in the 1880s to higher levels in the 1920s, in pace with the emissions of Pb from high temperature industrial activities (smelting, coal combustion, etc.) (Figure 5). With the introduction of leaded gasoline, Pb increased far more until the mid-1970s, when coralline Pb began to decrease with the phasing out of leaded gasoline, confirming the picture provided by direct observations in this period. Other information on the sources and dispersion of anthropogenic lead can be derived from the stable isotope ratios of lead (e.g. 206Pb/207Pb) which vary from one mining source to another because of differing Pb/U and Pb/Th for geological sources of lead. In particular, US leaded gasoline for many years had a 206Pb/207Pb ratio of 41.20, whereas European
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_
Pb (pmol kg 1)
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Figure 3 Pb in Bermuda surface waters, 1979–99. Squares (near surface waters) and triangles (40–50 m depth) are samples collected and analyzed by the MIT group (see Wu and Boyle 1997 for data through 1996; 1997–2000 from Boyle et al., unpublished); circular symbols are samples collected and analyzed by Schaule and Patterson 1981 and Veron et al 1993.
leaded gasoline was o1.15. Eventually anthropogenic Pb will be removed from the water column and reside in marine sediments, and there is already evidence from surficial sediments of the North Atlantic Ocean for significant quantities of anthropogenic Pb.
Anthropogenic Mercury in the Ocean Volatile mercury (Hg) and organomercury compounds are emitted from the land into the atmosphere naturally from wildfires, volcanoes, and microbial activity. Mercury is emitted into the atmosphere by humans as a result of high temperature processes (e.g. smelting, coal combustion, incineration) combined with commercial uses of elemental mercury (e.g. thermometers, batteries), as well as disposal of mercury-laden wastes (e.g. from gold mining operations) that are then converted into volatile forms in the environment. Estimates indicate that the anthropogenic mercury emissions from the land into the atmosphere exceeded natural sources by approximately a factor of three during the past century, not nearly so large as the Pb emission enhancement, but nonetheless substantial. A large fraction of this Hg is carried over great
distances by the atmosphere and deposited into the ocean. As for Pb, contamination is a major problem in Hg sampling and analysis (perhaps even more so), because Hg0 is volatile and broken mercury thermometer residues exist in most chemical laboratories. Evidence for anthropogenic perturbation of the oceanic Hg reservoir is less straightforward than it is for Pb. Observations in the modern Atlantic marine atmosphere show that the Northern Hemisphere has two- to three-fold higher Hg concentrations than the Southern Hemisphere, as might be expected for an anthropogenic source. However, land area is also higher in the Northern Hemisphere, and it is likely that even natural emissions are higher in the Northern Hemisphere, so observation of an interhemisphere difference does not by itself prove anthropogenic perturbation. Measurements of Hg in sea water have attempted to determine elemental (Hg0), organometallic (e.g. CH3Hg, (CH3)2Hg), and ‘reactive’ forms (ionic and inorganically complexed Hg, organically complexed Hg, and labile organic and particulate Hg). Reported Hg concentrations range from 10 11 to 10 9 moles kg 1. There are no direct seawater measurements or indirect seawater
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1000 Figure 4 Pb in ocean profiles in the Sargasso Sea near Bermuda: 1979 data from Schaule and Patterson 1983; 1984 data from Boyle et al. (1986); data from Veron et al 1993; 1993 data from Wu and Boyle 1997; 1997 data (Wu and Boyle, unpublished).
80 Bermuda D. strigosa
Pb/Ca, nmol mol
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Figure 5 Pb in corals near Bermuda in annually laminated layers deposited between 1880 and 1997 (Shen and Boyle, 1987).
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indicators such as corals which have yet documented an anthropogenic enhancement, although it has been estimated that Hg concentrations in the surface ocean have increased from 0.5 to 1.5 pmol kg 1 during the past century.
Other Anthropogenic Elements in the Ocean Several other elements have had significant anthropogenic enhancements of global fluxes, but in most cases, the size of the oceanic reservoir of the elements is too large for an observable shift in open-ocean concentrations. A two- to three-fold increase in the concentration of cadmium (Cd) has been observed in corals near Bermuda during the past century. However, the concentrations of Cd in surface waters near Bermuda are low even today (certainlyo10 11 moles kg 1 and sometimes as low as 10 12 moles kg 1; because concentrations in deep waters are two or three orders of magnitude higher, the oceanic reservoir is large compared with the anthropogenic enhancement of Cd in surface waters. Although concentrations of Zn in corals were not measured, a similar anthropogenic enhancement of emissions to the atmosphere may well have caused slight enhancements to the Zn concentrations of the surface ocean, but evidence of such an increase has not been reported. Because of the use of tributyl tin as a stabilizer in PVC plastic (hence an incineration source for inorganic Sn) and as an antifouling agent in marine paints, there has been some interest in the fate of anthropogenic Sn in the oceanic environment. Inorganic Sn is very low in the surface waters of the Sargasso Sea (B3 10 12 moles kg 1). Tributyl tin has been observed in enclosed harbors, but not in the open ocean. On the whole, however, the size of the ocean precludes major enhancements of oceanic trace metal concentrations. Even where the enhancement is clear, such as for Pb, the concentrations are quite low and there is no reason to suppose that these enhancements pose a threat to marine biota.
Anthropogenic Trace Metal Enhancements in Coastal Waters and Embayments As opposed to minimal evidence for anthropogenic trace metal contamination of open-ocean waters, it is clear that major impacts on the trace metal chemistry of coastal sea water have occurred and that these have sometimes had serious consequences. The discharge of untreated industrial and human wastes into coastal waters, and the release of toxic metals and
metal compounds from antifouling paints have clearly increased metal concentrations in shallow ocean waters in many coastal regions. One of the most notorious examples of coastal pollution was in Minimata Bay, Japan where from the 1930s until the 1960s a chemical company dumped tonnes of mercury, which contaminated the bay sediments. This Hg pollution entered the marine food chain. Thousands of people developed severe neurological symptoms as a result of eating Hgcontaminated fish. This is perhaps the most extreme example where pollution in the marine environment has demonstrably harmed humans directly. More recently, it has been demonstrated that high Cd, Zn, Cu, and Pb concentrations in Atlantic coastal waters of southern Spain have risen by many orders of magnitude due to acid mine drainage from the Rio Tinto massive sulfides. As yet, little is known about any human health effects from this trace element pollution, although the region has an active coastal fishing industry. There is too much detail in the many local examples to discuss here, but there is little doubt that many coastal embayments proximate to big cities have elevated metal levels, and that an understanding of the consequences of this pollution is still being developed. Unfortunately, proper measurement capabilities did not exist until the past two decades, so there is usually no direct evidence for the degree of enhancement of dissolved metals relative to preindustrial times. One exception to this situation is in San Francisco Bay, where the anthropogenic enhancement of dissolved cadmium has been documented by measuring the Cd content of benthic foraminifera. An anthropogenic impact on metals in coastal marine sediments has been demonstrated repeatedly. Lead has been shown to be enriched in coastal sediments. Even more exotic elements such as silver (photography) and osmium (electron microscopy) are enriched in the sewage of some cities and can be traced from their source.
Conclusions The anthropogenic dominance of lead in the openocean environment has been conclusively demonstrated. Although there is evidence that the fluxes of other elements to the ocean have been enhanced to a significant degree, there is less direct evidence documenting their enhancements in the open-ocean environment. Coastal waters and sediments near major cities show large enhancements in several trace metals.
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ANTHROPOGENIC TRACE ELEMENTS IN THE OCEAN
See also Metal Pollution. Radioactive Wastes. Refractory Metals. Transition Metals and Heavy Metal Speciation.
Further Reading Bruland KW and Franks RP (1983) Mn, Ni, Cu, Zn, and Cd in the western North Atlantic. In: Wong EBCS, Bruland KW, Burton JD, and Goldberg ED (eds.) Trace Metals in Seawater, pp. 395--414. New York: Plenum. Byrd JT and Andreae MO (1982) Tin and methyltin species in seawater; concentrations and fluxes. Science 218: 565--569. Chow TJ, Bruland KW, Bertine K et al. (1973) Lead pollution: records in Southern California coastal sediments. Science 181: 551–552. Fitzgerald WF, Engstrom DR, Mason RP, and Nater EA (1997) The case for atmospheric mercury contamination in remote areas. Environmental Science and Technology 32: 1--7. Flegal AR and Patterson CC (1983) Vertical concentration profiles of lead in the Central Pacific at 15N and 20S. Earth and Planetary Science Letters 64: 19--32. Hamelin B, Grousset F, and Sholkovitz ER (1990) Pb isotopes in surficial pelagic sediments from the North Atlantic. Geochimica et Cosmochimica Acta 54: 37--47. Hamelin B, Ferrand JL, Alleman L, and Nicolas E (1997) Isotopic evidence of pollutant lead transport from North America to the subtropical North Atlantic gyre. Geochimica et Cosmochimica Acta 61: 4423. Harrison RM and Laxen DPH (1981) Lead Pollution Causes and Control. London: Chapman and Hall. Helmers E and van bulleted Loeff MMR (1993) Lead and aluminum in Atlantic surface waters (501N to 501S) reflecting anthropogenic and natural sources in the eolian transport. Journal of Geophysical Research 98: 20261--20273. Mason RP, Fitzgerald WF, and Morel FMM (1994) The biogeochemical cycling of elemental mercury: anthropogenic influences. Geochimica et Cosmochimica Acta 58: 3191--3198. Jenkins WJ (1980) Tritium and He-3 in the Sargasso Sea. Journal of Marine Research 38: 533--569. Measures CI and Edmond JM (1988) Aluminum as a tracer of the deep outflow from the Mediterranean. Journal of Geophysical Research 93: 591--595. Measures CI and Edmond JM (1990) Aluminium in the south Atlantic: steady state distribution of a short residence time element. Journal of Geophysical Research 95: 5331--5340. Measures CI, Edmond JM, and Jickells T (1986) Aluminum in the northwest Atlantic. Geochimica et Cosmochimica Acta 50: 1423--1429.
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Orians KJ and Bruland KW (1985) Dissolved aluminum in the central North Pacific. Nature 316: 427--429. Orians KJ and Bruland KW (1986) The biogeochemistry of aluminum in the Pacific Ocean. Earth and Planetary Science Letters 78: 397--410. Murozumi M, Chow TJ, and Patterson C (1969) Chemical concentrations of pollutant lead aerosols, terrestrial dusts and sea salts in Greenland and Antarctic snow strata. Geochimica et Cosmochimica Acta 33: 1247--1294. Nriagu JO (1989) A global assessment of natural sources of atmospheric trace metals. Nature 338: 47--49. Ravizza GE and Bothner MH (1996) Osmium isotopes and silver as tracers of anthropogenic metals in sediments from Massachusetts and Cape Cod bays. Geochimica et Cosmochimica Acta 60: 2753--2763. Schaule B and Patterson CC (1983) Perturbations of the natural lead depth profile in the Sargasso Sea by industrial lead. In: Wong CS, Boyle EA, Bruland KW, Burton JD, and Goldberg ED (eds.) Trace Metals in Seawater, pp. 487--504. New York: Plenum. Schaule BK and Patterson CC (1981) Lead concentrations in the northeast Pacific: evidence for global anthropogenic perturbations. Earth and Planetary Science Letters 54: 97--116. Shen GT and Boyle EA (1987) Lead in corals: reconstruction of historical industrial fluxes to the surface ocean. Earth and Planetary Science Letters 82: 289--304. Shen GT and Boyle EA (1988) Thermocline ventilation of anthropogenic lead in the western North Atlantic. Journal of Geophysical Research 93: 15715--15732. Shen GT, Boyle EA, and Lea DW (1987) Cadmium in corals as a tracer of historical upwelling and industrial fallout. Nature 328: 794--796. Sherrell RM, Boyle EA, and Hamelin B (1992) Isotopic equilibration between dissolved and suspended particulate lead in the Atlantic Ocean: evidence from Pb-210 and stable Pb isotopes. Journal of Geophysical Research 97: 11257--11268. Slemr F and Langer E (1992) Increase in global atmospheric concentrations of mercury inferred from measurements over the Atlantic Ocean. Nature 355: 434--437. Turekian KK, Benninger LK, and Dion EP (1983) Be-7 and Pb-210 total deposition fluxes at New Haven, Connecticut and at Bermuda. Journal of Geophysical Research 88: 5411--5415. van Geen A and Luoma SN (1999) A record of estuarine water contamination from the Cd content of foraminiferal tests in San Francisco Bay, California. Marine Chemistry 64: 57. van Geen A, Adkins JF, Boyle EA, Nelson CH, and Palanques A (1997) A 120 year record of metal contamination on an unprecedented scale from mining of the Iberian Pyrite Belt. Geology 25: 291--294. Veron AJ, Church TM, Flegal AR, Patterson CC, and Erel Y (1993) Response of lead cycling in the surface Sargasso Sea to changes in tropospheric input. Journal of Geophysical Research 98: 18269--18276.
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Veron AJ, Church TM, Rivera-Duarte I, and Flegal AR (1999) Stable lead isotope ratios trace thermohaline circulation in the subarctic North Atlantic. Deep-Sea Research II 46: 919--935. Wolff EW and Peel DA (1985) The record of global pollution in polar snow and ice. Nature 313: 535--540.
Wu JF and Boyle EA (1997) Lead in the western North Atlantic Ocean: completed response to leaded gasoline phaseout. Geochimica et Cosmochimica Acta 61: 3279--3283.
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ANTIFOULING MATERIALS D. J. Howell and S. M. Evans, Newcastle University, Newcastle, UK & 2009 Elsevier Ltd. All rights reserved.
Introduction Any unprotected surface which is introduced into the sea, such as a buoy, a ship’s hull, an oil rig support, or a fish cage, will become fouled by the growth of marine organisms. Colonization of a newly immersed surface involves a typical ecological succession of different flora and fauna. Initially, the surface undergoes biochemical changes as macromolecules, such as glycoproteins, proteoglycans, and polysaccharides, become adsorbed on to it. Bacteria then colonize it within c. 1 h of submersion, followed by diatoms, yeasts, and protozoa. Subsequently, invertebrate larvae and algal spores settle on the surface and metamorphose and/or grow, often forming a dense covering over it. In the case of wooden structures, there are even some organisms, such as the shipworm (actually a mollusk, Teredo), which bore beneath the surface. The climax community can be highly diverse. Overall, there are 44000 known species of fouling organisms, including barnacles, hydroids, tube worms, bryozoans, seaweeds, and others. This article considers the problem of fouling on ships’ hulls, and measures taken to prevent it. Coatings containing the biocide tributyltin, more commonly known as TBT, were particularly effective antifoulants but TBT leached from them into the marine environment, harming nontarget species. Their use was initially regulated, then eventually banned altogether. The extent to which the paint industry has been able to meet the challenge of finding equally effective, but less environmentally damaging, alternative antifoulants is discussed.
Historical Development Fouling has caused problems to shipping throughout history. Teredo have been responsible for the destruction of many wooden sailing ships, even causing them to break up at sea. Modern steel hulls are not penetrated by these mollusks; nevertheless, their surfaces are highly vulnerable to fouling. Coverings by the growth of organisms can be remarkable if they are left unchecked. The total weight of fouled organisms can be as high as 150 kg m 2, representing
6000 t of fouling attached to a large commercial vessel which has an underwater surface area of 40 000 m2. However, it is not solely the increased weight which causes the problem to shipping. Fouling leads to an increased friction between the hull and seawater, causing the so-called ‘hull roughness’. As a result of these combined effects, badly fouled ships suffer from loss of speed and maneuverability. They are also expensive to operate, incurring fuel penalties of Z50%. Not surprisingly, there have been extensive efforts to reduce fouling, and its impacts, down the ages. The ancient Carthaginians and Phoenicians used pitch, and possibly copper sheathing, on the bottoms of ships to prevent fouling, and later coated hulls with sulfur and arsenical compounds. The Greeks and Romans both introduced lead sheathing. Copper cladding was then ‘reinvented’ in the seventeenth and eighteenth centuries, but it became redundant after the introduction of steel vessels at the end of the eighteenth century, because of problems caused by galvanic corrosion. A new technology was needed to protect ships’ hulls from fouling, but it took some 60 or so years until the mid-nineteenth century for the development of antifouling paints to provide it. There have been substantial developments in paint technology since the first antifouling paints were introduced, and the search for more effective coatings still continues today. However, the large majority of antifouling paints incorporate biocides into the matrix and, in this respect, the technology has not changed. The biocides leach slowly from the paints, killing organisms which attempt to settle on the hull. Cuprous oxide was used in early formulations, but numerous other toxins – including organomercury, lead and arsenical compounds, and DDT – were added as boosters, in order to enhance the effectiveness and life expectancy of the paints. However, such compounds pose severe environmental and human health risks, and they were withdrawn voluntarily by the paint industry during the 1960s. They were replaced largely by the highly effective biocide tributyltin (TBT). The properties of the paint matrix are another important consideration because they determine qualities such as the rate of biocide release and therefore the long-term effectiveness of a paint as an antifoulant. Originally, active biocidal ingredients were dispersed in a soluble resinous matrix, which released the biocide as it dissolved slowly in seawater. A problem with these free association paints,
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as they became known, was that the release rates of biocides from them were uncontrolled. The initial rate was rapid so that the paint was highly effective when it was first applied (Figure 1). However, subsequent release of the biocide from the matrix declined steadily and antifouling performance diminished with time; it might be virtually exhausted within c. 12 months. Consequently, free association paints typically offered a maximum period of only 1–2 years of service before dry-docking and repainting was necessary. There were limited improvements when matrices which were insoluble, using chlorinated rubber and vinyl resins, were developed in the 1940s. They had greater mechanical strength and allowed thicker coatings of paint to be applied, thereby providing effective antifouling performance over a full 2-year period. However, a major breakthrough in antifouling paint technology occurred with the introduction of so-called self-polishing copolymer paints in the late 1960s. The term self-polishing was used because the biocide was released slowly as the paint surface was gradually worn away. The essential difference from free association paints was that the biocide was chemically bonded in a copolymer resin system, via an organotin–ester linkage. This ester group hydrolyzed at the surface of the paint where it was in contact with seawater, and this resulted in a slow and controlled release of the biocide. The remaining surface of the paint was mechanically weakened by breakage of these bonds, and was eroded by moving seawater, resulting in the exposure of a fresh surface layer. The hydrolysis/erosion process was then repeated until there was no paint left.
The technology has been particularly successful with TBT copolymer (poly)tributyltinmethacrylatemethylmethacrylate, although cuprous oxide or other boosters were included in the formulations. They were needed even though TBT is active against most fouling organisms, since some slime-forming diatoms are resistant to it. TBT-based self-polishing copolymer antifouling systems (TBT SPC systems) brought enormous benefits to the shipping industry. They could provide effective antifouling cover for Z5 years, more than doubling the performance of free association paints. Reduced fuel costs and less frequent need to drydock and repaint vessels were estimated to be worth some US$5.7 billion per annum to the industry during the mid-1990s. There were also huge environmental benefits because lower fuel consumption by the world’s shipping fleet reduced the release of ‘greenhouse’ gases and emissions which were responsible for acid rain. Annual fuel savings, which could be attributed to the use of TBT-based antifoulants, were believed to be 7.2 million tons, reducing carbon dioxide emissions by 22 million tons, and sulfur dioxide by 0.6 million tons, per year. An additional benefit was that effective antifouling prevented the transport of invasive (nonnative) organisms across the world on ships’ hulls. Such species can cause enormous economic and ecological damage once they become established. It was originally thought that these organisms hitchhiked their way across the oceans in ballast water, but it is becoming increasingly clear that fouled hulls are an additional means of transport. The effect on local economies can be devastating. It has been estimated that the
Free association paints
Copolymer paints
Leaching
Polishing
Leaching rate
Leaching rate
Time
Time
Figure 1 A diagrammatic comparison of the mode of action of free association paints and self-polishing copolymer paints, and the rates at which they release tributyltin into the water. Reproduced from Stebbing ARD (1985) Organotins and water quality: Some lessons to be learned. Marine Pollution Bulletin 16: 383–390.
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cost of control measures that were required to limit the destruction caused by the introduced European zebra mussel Dreissena polymorpha to the USA, which had infested over 40% of internal waterways, was US$1 billion between 1989 and 2000. TBT SPC systems came into widespread use. They were, for example, first used by major shipping lines in Europe and the Far East in the mid-1970s, and were registered by the Environment Protection Agency in the USA in 1978. Subsequently, it was estimated that they were used on 70% of the world’s commercial shipping fleet, and on high proportions of fishing vessels and pleasure craft.
The Downside of TBT-Based Antifoulants: Their Regulation and Eventual Ban Unfortunately, there are environmental costs of using TBT-based paints. The ideal antifouling biocide would be one which degraded into harmless residues immediately after release into the water column so that the toxic effects occurred at the ship’s hull but nowhere else in the marine environment. In fact, TBT does degrade reasonably rapidly in seawater, where it has a residence time of only a few days. Nevertheless, this is sufficient time for it to become adsorbed on to particles, and to aggregate in sediments in which there are high levels of organic matter. Here degradation processes are considerably slower, and the half-life of TBT may then be a matter of months or years. Regrettably, TBT is so toxic that even low concentrations can harm marine life. For example, the lethal dose concentration (15-day LC50) for larvae of the mussel Mytilus edulis is 10 ng l 1, and that (96-h LC50) for larvae of the sole Solea solea is 21 ng l 1. Sublethal effects occur at even lower concentrations. A dose of 40.4 ng l 1 can affect the growth and reproduction of phytoplankton and zooplankton, and one of 42 ng l 1 can affect the process of shell formation in the oyster Crassostrea gigas. This latter concentration is also sufficient to cause the development of the condition known as imposex in the dogwhelk Nucella lapillus. In this case, TBT acts as a hormone disruptor affecting gender differentiation. Female dogwhelks develop male genital organs, including a pseudopenis and vas deferens, which become superimposed on their reproductive systems (Figure 2). Not surprisingly, the widespread and uncontrolled use of TBT-based paints worldwide during the 1970s and 1980s (especially free association paints) resulted in unacceptably high levels of TBT in some coastal areas of intense boat use. The main problems were in
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enclosed bodies of water with poor flushing characteristics, such as harbors, dry-docks, marinas, estuaries, and bays. The first documented cases of serious biological impact on nontarget organisms came from Arcachon Bay in west France. The bay is a center of both oyster culture and of high yachting activity, and TBT originating from paints used on yacht hulls was held responsible for abnormal growth and reproductive failure of cultured oysters. The oyster farming industry was in a state of near collapse. Production of oysters fell in the early 1980s to c. 33–50% of the normal harvest (Table 1). At about the same time, TBT pollution was also linked to serious decline in populations of dogwhelks (N. lapillus). They were suffering from imposex and it was so severe that populations of these organisms in areas of high boating activity had become partially or totally sterile. Subsequently, imposex was described in more than 100 species of gastropods worldwide, and the condition in many of them has been used as a biological indicator of TBT contamination. It became evident from surveys using this indicator that TBT contamination had become a global problem. It was clearly necessary to control the use of TBT-based paints. Pleasure craft, especially yachts, were believed to be the main source of pollution in coastal waters and several countries banned the use of these antifoulants on vessels o25 m in length. The French Government was the first to react, due to the oyster crisis, introducing controls in 1982. The UK followed in 1987, the USA in 1988, and subsequently Canada, New Zealand, Australia, South Africa, Hong Kong, and most European countries reacted in similar ways. The regulations were effective in reducing TBT contamination. Declining concentrations of TBT were reported during the 1990s in water samples, sediments (although less rapidly due to their persistence in them), and tissues of mollusks in monitoring programs worldwide. More dramatically, there have been enormous improvements in the health of marine biota which had been affected by TBT pollution in the previous decade. Oyster production in Arcachon Bay recovered spectacularly to former levels almost immediately after the introduction of regulations in France in 1982 (Table 1), and there was widespread recovery of populations of dogwhelks suffering from imposex. Nevertheless, commercial harbors, especially those with dry-docking and repair facilities, were still hot spots of TBT contamination (see Table 2), due to the accumulation of butyltins themselves and paint chips containing them in sediments (Figure 3). This has remained an unfortunate legacy of the TBT era and there are still management implications with respect to the disposal of contaminated sediments. TBT may
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p
gp vd
a
v Stage
cg
1
2
3
4
n b A
B
A
6
5
B
Figure 2 Six stages in the development of imposex in the dogwhelk N. lapillus. Abbreviations: a, anus; b, ‘blister’; cg, capsule gland; gp, genital papilla; n, ‘nodule’; p, penis; v, vulva; vd, vas deferens. Reproduced from Gibbs PE, Bryan GW, Pascoe PL, and Burt GR (1987) The use of the dogwhelk (Nucella lapillus) as an indicator of TBT contamination. Journal of the Marine Biological Association of the United Kingdom 67: 507–524.
be remobilized from them either as a result of dredging or natural events, such as storms or tidal flow. A ban on the use of TBT as the active biocide in antifoulants became inevitable. The International Maritime Organization’s Convention on the Control of Harmful Antifouling Systems (AFS 2001) resolved that the application of TBT-based paints be prohibited from 1 January 2003. Furthermore, it stated that by 1 January 2008 (effective date), ships either (1) shall not bear such compounds on their hulls or
external parts or surfaces; or (2) shall bear a coating that forms a barrier to such compounds leaching from the underlying noncompliant antifouling system.
Meeting the Challenge: Alternative Antifoulants The ban on the use of TBT created a need for new coating systems that have the same antifouling
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efficiency as TBT-based coatings, but are more environmentally acceptable. The success of TBT-based coatings inhibited research on alternatives and this did not commence in earnest until the 1990s when the ban became probable. Since then, there has been substantial research effort on the biology and chemistry of marine fouling and natural antifouling systems. It has greatly increased our understanding of how to reduce and perhaps prevent fouling in the future. Research on alternative coatings has followed two major lines: (1) the development of antifouling Table 1 Production of oysters in Arcachon Bay, France, between 1978 and 1985 Period
Production (t)
1978–79 1979–80 1980–81 1981–82 1982–83 1983–84 1984–85
10 000 6000 3000 5000 8000 12 000 12 000
Reproduced from Alzieu C (1996) Biological effects of tributyltin on marine organisms. In: de Mora SJ (ed.) Tributyltin: Case Study of an Environmental Contaminant, pp. 167–211. Cambridge: Cambridge University Press.
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coatings that use the same binder delivery mechanism as the banned TBT coatings but contain different biocidal products; and (2) the development of nonbiocidal alternatives. Biocidal Antifouling Coatings
There are, broadly, two types of biocidal coatings on the market today. First, those that have hydrating binders, also known as ablative or solid matrix coatings, and, second, those that have hydrolyzing binders and mimic TBT SPC coatings. Copper compounds, such as cuprous oxide (Cu2O), copper thiocyanate (CuSCN), or metallic copper, are now used as the principal biocide. There is nevertheless concern about levels of this metal in the marine environment, and regulatory bodies are already turning their attention to copper contamination in ports, harbors, and coastal zones. Some actions have already been taken. For instance, the use of copper has been banned as an antifouling biocide on pleasure craft in Baltic Sea and west coast areas of Sweden, and release rate limits have been imposed in Canada, Denmark, and are about to be enforced in the USA. However, the situation is complex due to the issue relating to the bioavailablity of copper. Copper, when released into the marine environment, is only bioavailable in the ionic form. Due to the high
Table 2 Two ‘hot spots’ of TBT contamination in Puget Sound (USA): gradients of diminishing concentrations of organotins (total TBT and its metabolites DBT and MBT) in tissue samples, and of measures of imposex (VDSI and RPSI), in the dogwhelks Nucella emarginata and N. lamellosa from a shipyard at Anacortes and the complex of commercial harbors and shipyards at Seattle Source
Species
Distance of site from source (km)
Total organotins % with imposex (ng g 1)
Anacortes shipyard
N. lamellosa
0
668
a
1.3 2.4 6.5 8.0
212 61 43 24
0.9 1.9 2.3 8.0 9.0 0.9 1.9 2.3 8.0 9.0
Seattle
Seattle
N. emarginata
N. lamellosa
VDSI
RPSI
a 100 100 100
4.0 2.8 3.3
1.2 o0.1 o0.1
146 65 42 72
100 100 100 94
4.0 3.9 1.9 1.9
14.1 9.1 o0.1 o0.1
131 141 68 66 42
100 100 100 100 100
4.0 4.0 3.1 3.5 3.2
19.3 4.3 6.7 0.5 o0.1
a
Population consisted of males only. DBT, dibutyltin; MBT, monobutyltin; VDSI, vas deferens sequence index; RPSI, relative penis size index. Reproduced from Evans SM, Barnes N, Birchenough AC, Brancato MS, and Hardman E (2001) Tributyltin contamination in two estuaries and adjacent coasts: Puget sound, Washington and Narragansett Bay, Rhode Island (USA). Invertebrate Reproduction and Development 39: 221–229.
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Figure 3 Residual paint flakes in sediments may be the cause of long-term pollution problems. The illustration shows (from left), tin and copper elemental maps, and scanning electron microscope microphotograph, of a paint flake recovered from a dredge in a working shipyard. Twelve layers of paint are visible in the flake, reflecting the coatings cycle of a commercial vessel with the newest layers at the top. Antifouling layers are typified by high copper content. Two layers of coating containing high tin and copper levels are present. Published with permission from Dickon Howell.
proportion of humic substances in seawater, this ionic copper is quickly bound to forms that are not bioavailable. The result of this is that environmental copper may exceed Environmental Quality Standards (EQS) levels if it is measured using the total dissolved form, even though the bioavailable copper is actually below the approved limits. In any case, copper as the sole biocide may not give complete fouling protection. There are several algal species (e.g., Enteromorpha spp., Ectocarpus spp., Achnanthes spp.) that show marked physiological tolerance to copper, and effective coatings must include organic booster biocides to combat them. These boosters include both copper and zinc pyrithione (antifungal agents used in anti-dandruff shampoos and outdoor paints), diuron (a substituted urea herbicide), ChlorothalonilTM (tetrachloroisophthalonitrile), dichlofluanid (a phenylsulfamide acaricide), Irgarol 1051TM (a triazine), Sea Nine 211TM (dichloro-isothiazolone), zineb and maneb (both dithiocarbomate fungicides containing zinc and manganese, respectively). Many of them, including zineb, maneb, ChlorothalonilTM, and dichlofluanid have a history of use as agricultural pesticides. When the new biocidal coating formulations were brought to market, there was initial concern that little research had been done on the environmental persistence of these booster biocides in the marine environment in terms of their degradation, partitioning, and bioavailability. This concern has been justified as both diuron and Irgarol 1051TM have been shown to have relatively high environmental persistence. Following monitoring studies, Irgarol 1051TM, a photosystem II inhibitor, has been shown to have very slow degradation (half-life of 100 days in marine sediments) and a high bioaccumulation
factor. The latter was shown, for example, in monitoring studies of the seagrass Zostera marina, a species of high international conservation importance. Concentrations of Irgarol 1051TM were found to be 25 000 times higher in Zostera tissues than in ambient (o0.003 mg l 1) seawater. It is for these reasons that Irgarol 1051TM has also been withdrawn from use by the UK Health and Safety Executive and also in Denmark. Similarly, diuron has since been withdrawn from use by the UK Health and Safety Executive and is being considered for withdrawal by the European Commission. The ban may spread to some of the other formulations, even those that seem to be more environmentally friendly. For example, Sea Nine 211TM and both copper and zinc pyrithione have relatively low environmental persistence, with half-lives of under 48 h, but some people believe that their introduction into the marine biosphere is still undesirable. Even a short pulse of a rapidly degradable pollutant at low concentrations can affect plankton populations significantly, thus causing effects that can cascade through a community. It can also be argued, with the benefit of hindsight, that some of the biocides that have been used as antifoulants should never have been allowed to come on to the market in the first place. Some of them had been used successfully in terrestrial applications but their behavior, once released in the marine environment, was unforeseen. This is well illustrated by experiences with TBT, whose sublethal impacts at extremely low, almost undetectable concentrations in seawater were not appreciated until they were in widespread use in marine coatings. A major problem when TBT coatings were introduced was that there was no standardized method for testing antifoulants
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that gave results that were comparable to a ‘realworld situation’. Regrettably this is still the case. Particular biocides can still ‘pass’ ecotoxicological assays carried out in isolation of other assessments. However, such tests provide no information on, for example, how a biocidal coating system performs when it is applied to a ship’s hull, what synergistic effects there may be between active biocides, and precisely what the biocide’s environmental fate will be. Fortunately, progress is being made and considerable attention is now given to the environmental persistence and toxicity of new biocides and evidence is accumulating that the non-TBT SPC coatings are performing well. There are reports that there are intervals of 4–5 years between recoatings of new formulations, rivaling the achievements of the old TBT SPC coatings. Nonbiocidal Alternatives
In an ideal world, it would be unacceptable to release any toxic compound into the environment, and the ultimate aim of research workers and the regulators must be to eliminate the use of biocides in antifouling coatings altogether. There are many promising areas of research and development of nonbiocidal alternatives. These nontoxic coating systems fall into three categories: (1) foul release coatings; (2) socalled ‘smart’ coatings; and (3) hard marine coatings in conjunction with regular cleaning. Foul release coatings prevent the settlement of marine biota by having very low surface energy and lower surface roughness than biocidal coatings, providing a surface onto which organisms have great difficulty attaching themselves. Settlement can occur if vessels are stationary for extended periods but the adhesion between the fouling and the coating surface is very weak, so hydrodynamic forces will remove the organism when the vessel is traveling at relatively high speed of 15 knots or more. The majority of foul release systems are based on silicone as it has a low surface energy, a low elastic modulus, a microphase separation structure of hydrophilic and hydrophobic layers, and a nonfixed surface layer containing oil materials, which plasticize the coating and reduce adhesion strength. Silicones also provide a smoother surface than SPCs, making them more hydrodynamically efficient. Nevertheless, they are not yet seen as a viable option for slow-moving ocean-going container ships and are primarily used on high-speed coastal vessels. ‘Smart’ coatings are at the cutting edge of antifouling research. They include super-hydrophobic/ hydrophilic surfaces, use of magnetic fields, piezoelectric stimulation, bubble curtains to prevent fouling while the vessel is stationary, and biomimetic
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surfaces. A promising approach is in the field of biomimetics. Coatings have been developed mimicking natural surfaces such as the shells of mollusks and crustaceans or the skin of fish and marine mammals. These surfaces have distinct nanoprofiles which inhibit fouling due to their bioenergetic properties. The underlying principle is that of attachment theory which states that attachment of a sphere (spore) to a flat substratum leads to a net increase in total surface area and an increase in work or energy exerted on the system. Attachment therefore will be preferential in valleys that are of similar dimension to the body size of the spore, resulting in a net decrease in surface area and therefore a reduction in hydrodynamic stress. If a surface can be created that is energetically unsuitable to a settling spore (i.e., has few attachment points), settlement will be inhibited (e.g., Figure 4). Although these technologies have not been tested on a working vessel as yet, initial results are encouraging and commercial-scale testing will become a reality in the near future, particularly in the case of the SharkletTM coating. The use of hard marine coatings that require regular in-service cleaning (every 2–3 weeks) is a trend that is rising, especially in the luxury cruise liner market. However, this process brings with it an increased threat of the spreading of invasive species, because organisms picked up in one part of the world may be deposited at another when the vessel hull is cleaned. This threat has not been well publicized because legislative drivers for regulating against the translocation of invasive species have mainly focused on ballast water. It is, nevertheless, a real one, and is causing particular concern in countries, such as Australia and New Zealand, where threats from
Figure 4 A scanning electron microscope microphotograph of the surface profile of the newly developed SharkletTM coating. Courtesy of Dr. Tony Brennan, University of Florida.
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invasive species present particularly acute problems. Further development of this particular method of fouling control will probably depend on significant improvement of underwater hull cleaning facilities that can ensure containment of all organic matter cleaned from the hull.
See also Eutrophication. Metal Pollution. Pollution: Approaches to Pollution Control. Pollution: Effects on Marine Communities. Pollution, Solids. Shipping and Ports. Ships. Transition Metals and Heavy Metal Speciation.
Further Reading Alzieu C (1996) Biological effects of tributyltin on marine organisms. In: de Mora SJ (ed.) Tributyltins: Case Study of an Environmental Contaminant, pp. 167--211. Cambridge, UK: Cambridge University Press. Alzieu C (1998) Tributyltin: Case study of a chronic contaminant in the coastal environment. Ocean and Coastal Management 40: 23--36. de Mora SJ (ed.) (1996) Tributyltin: Case Study of an Environmental Contaminant. Cambridge, UK: Cambridge University Press.
Evans SM (1999) TBT contamination: The catastrophe that never happened. Marine Pollution Bulletin 38: 629--636. Evans SM (1999) TBT or not TBT? That is the question. Biofouling 14: 117--129. Evans SM, Barnes N, Birchenough AC, Brancato MS, and Hardman E (2001) Tributyltin contamination in two estuaries and adjacent coasts: Puget sound, Washington and Narragansett Bay, Rhode Island (USA). Invertebrate Reproduction and Development 39: 221--229. Gibbs PE, Bryan GW, Pascoe PL, and Burt GR (1987) The use of the dogwhelk (Nucella lapillus) as an indicator of TBT contamination. Journal of the Marine Biological Association of the United Kingdom 67: 507--524. Hunter JE and Cain P (1996) Shipping and the environment: Is compromise inevitable? Anti-Fouling Coatings in the 1990s – Environmental, Economic and Legislative Aspects, Paper 16, IMAS 96, 22–24 October 1996. London: The Institute of Marine Engineers. Stebbing ARD (1985) Organotins and water quality: Some lessons to be learned. Marine Pollution Bulletin 16: 383--390. Yebra DM, Kiil S, and Dam-Johansen K (2004) Antifouling technology – past, present and future steps towards efficient and environmentally friendly antifouling coatings. Progress in Organic Coatings 50: 75--104.
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ARCTIC OCEAN CIRCULATION B. Rudels, Finnish Institute of Marine Research, Helsinki, Finland & 2009 Elsevier Ltd. All rights reserved.
Introduction The Arctic Ocean is the northernmost part of the Arctic Mediterranean Sea, which also comprises the Greenland Sea, the Iceland Sea, and the Norwegian Sea (the Nordic Seas) and is separated from the North Atlantic by the 500–850-m-deep Greenland– Scotland Ridge. The Arctic Ocean is a small, 9.4 106 km2, enclosed ocean. Its boundaries to the south are the Eurasian continent, Bering Strait, North America, Greenland, Fram Strait, and Svalbard. The shelf break from Svalbard southward to Norway closes the boundary. The Arctic Ocean lies almost entirely north of and occupies most of the region north of the polar circle. More than half of its area, 53%, consists of large, shallow shelves, the broad Eurasian marginal seas: the Barents Sea (200–300 m), the Kara Sea (50–100 m), the Laptev Sea (o50 m), the East Siberian Sea (o50 m) and the Chukchi Sea (50–100 m), and the narrower shelves north of North America and Greenland. The deep Arctic Ocean comprises two major basins, the Eurasian Basin and the Canadian (also called Amerasian) Basin, separated by the approximately 1600-m-deep Lomonosov Ridge. The Eurasian Basin is further divided into the Nansen and Amundsen basins by a mid-ocean ridge (the Gakkel Ridge), while the Canadian Basin is separated by the Alpha Ridge and the Mendeleyev Ridge into the Makarov and the Canada Basins. The Amundsen Basin is the deepest (B4500 m), while the maximum depths of the Makarov and the Nansen Basins are B4000 m. The Canada Basin is slightly shallower (B3800 m) but by far the largest (Figure 1). The Arctic Ocean water masses are primarily of Atlantic origin. Atlantic waters (AWs) enter the Arctic Ocean from the Nordic Seas through the 2600-m-deep Fram Strait and over the B200-m-deep sills in the Barents Sea. The Arctic Ocean also receives low-salinity Pacific water through the shallow (45 m) and narrow (50 km) Bering Strait. The outflows occur through Fram Strait and through the shallow (150–230 m) and narrow channels in the Canadian Arctic Archipelago.
The physical oceanography of the Arctic Ocean is shaped by the severe high-latitude climate and the large freshwater input from runoff, B0.1 Sv (1 Sv ¼ 1 106m3 s 1), and net precipitation, B0.07 Sv. The Arctic Ocean is a strongly salinity stratified ocean that allows the surface water to cool to freezing temperature and ice to form in winter and to remain throughout the year in the central, deep part of the ocean. The water masses in the Arctic Ocean can, because of the stratification, be identified by different vertical layers. Here five separate layers will be distinguished: 1. The B50-m-thick upper, low-salinity polar mixed layer (PML) is homogenized in winter by freezing and haline convection, while in summer the upper 10–20 m become diluted by sea ice meltwater. The salinity, S, ranges from 30 to 32.5 in the Canadian Basin and between 32 and 34 in the Eurasian Basin. 2. The 100–250-m-thick halocline, with salinity increasing with depth, while the temperature remains close to freezing, 32.5oSo34.5 (Figure 2). 3. The 400–700-m-thick Atlantic layer historically defined as subsurface water with potential temperature, y, above 0 1C, 34.5oSo35. 4. The intermediate water below the Atlantic layer that communicates freely across the Lomonosov Ridge, 0.5 1Coyo0 1C, 34.87oSo34.92. 5. The deep and bottom waters in the different basins, 0.55 1Coyo 0.5 1C, 34.92oSo34.96 (Canadian Basin), 0.97 1Coyo 0.5 1C, 34.92o So34.945 (Eurasian Basin). There are large lateral variations of the characteristics in these layers that depend upon the circulation and upon the mixing processes in the Arctic Ocean (Figures 2 and 3). A more detailed classification, especially for the deeper water masses, is given in Table 1. It should be kept in mind that this classification is not unique and that several others exist in the literature.
Circulation The circulation of the uppermost layers of the Arctic Ocean has mainly been inferred from the ice drift, determined from satellites and from drifting buoys, while at deeper levels primarily the distributions of temperature and salinity and more recently of other tracers have been used to deduce the movements of the different water masses, often with
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135° E
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180°
45° E
35° W
0°
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45° W
Figure 1 Map of the Arctic Mediterranean Sea showing geographical and bathymetric features. The bathymetry is from IBCAO (the International Bathymetric Chart of the Arctic Ocean; Jakobsson et al., 2000; 2008) and the projection is Lambert equal area. The 500 and 2000 m isobaths are shown. All maps used here are made by Martin Jakobsson (personal communication). BIT, Bear Island Trough; CB, Canadian Basin; EB, Eurasian Basin; GFZ, Greenland Fracture Zone; MJP, Morris Jessup Plateau; JMFZ, Jan Mayen Fracture Zone; SAT, St. Anna Trough; YM, Yermak Plateau; VC, Victoria Channel; VS, Vilkiltskij Strait; FJL, Franz Josef Land; BS, Barrow Strait; HG & CS, Hell Gate and Cardigan Sound.
the assumption that the circulation is largely controlled by the bathymetry. Direct, moored current measurements have been scarce and mostly confined to the continental slope and to the Lomonosov Ridge. These measurements have confirmed the importance of the bathymetry for the circulation. In the deep basins current measurements have been made from drifting ice camps and more recently also from autonomous ice-tethered platforms, relaying the observations via satellite to shore. Subsurface drifters are just beginning to be used and observational efforts during the International Polar Year (IPY) 2007– 09 are likely to significantly increase the knowledge of the circulation in the Arctic Ocean. The motions of the ice cover and the surface water are predominantly forced by the wind, and the atmospheric high-pressure cell over the Arctic creates the anticyclonic Beaufort gyre in the Canada Basin. Ice leaks from the offshore side of the gyre and joins the Transpolar Drift (TPD) that brings ice from the Canada Basin toward Fram Strait. A second branch originating from the Siberian shelves, mainly from
the Laptev Sea, carries ice across the Eurasian Basin. About 90% of the ice export (0.09 Sv) passes through Fram Strait (Figure 4). The PML and the halocline are maintained by river runoff, ice melt, and the inflow of low-salinity water through Bering Strait. The lowest surface salinities and the thickest halocline are therefore observed in the Canada Basin. The inflow through Bering Strait, although affected by local winds, is in the last instance driven by a higher sea level in the North Pacific as compared to the Arctic Ocean. This creates a pressure gradient that forces the Pacific water northward into the Arctic Ocean. The Bering Strait inflow continues across the Chukchi Sea in four branches. One branch enters the East Siberian Sea, while one of the central inflow branches enters the Arctic Ocean along the Herald Canyon west of the Chukchi Plateau, and the other passes via the Central Gap east of the Chukchi Plateau into the Canada Basin. The easternmost branch reaches the Canada Basin along the Barrow Canyon close to Alaska. River runoff, mainly from the Mackenzie
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0 PML
CB 400
600 −2
1 −1 0 Potential temperature (°C)
32.0
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-27.
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Potential temperature (°C)
Halocline
AB MB
26
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NB 200
−2 32.0
32.5
33.0
33.5 Salinity
34.0
34.5
35.0
Figure 2 Potential temperature and salinity profiles and yS curves from the upper layers of the Nansen Basin (NB, dark yellow), Amundsen Basin (AB, green), Makarov Basin (MB, magenta), and Canada Basin (CB, blue). The PML and the halocline are indicated in the salinity profiles. Above the PML, the low-salinity layer due to seasonal ice melt is seen. The temperature maximum in the Canada Basin is due to the presence of Bering Strait Summer Water (BSSW) and the temperature minimum below indicates the upper halocline with S B 33.1, deriving from the colder, more saline Bering Strait Winter Water (BSWW) and from brine release and haline convection in the Chukchi Sea. No halocline is present in the Nansen Basin, only a deep winter mixed layer between the thermocline and the seasonal ice melt layer with temperature close to freezing. The curved shape of the Nansen Basin thermocline as seen in the yS diagram suggests wintertime haline convection with dense, saline parcels penetrating into the thermocline. Adapted from Rudels B, Jones EP, Schauer U, and Eriksson P (2004) Atlantic sources of the Arctic Ocean surface and halocline waters. Polar Research 23: 181–208.
River, adds freshwater to the Canada basin. The runoff peaks in early summer (June). The river runoff as well as most of the Pacific water becomes trapped in the anticyclonic Beaufort gyre, forming an oceanic high-pressure cell in the southern Canada Basin. The Pacific water leaves the Beaufort gyre and the Arctic Ocean mainly through the Canadian Arctic Archipelago, but a smaller fraction also exits, at least intermittently, through Fram Strait. Warm AW crosses the Greenland–Scotland Ridge and continues toward the Arctic Ocean in the Norwegian Atlantic Current. The Norwegian Atlantic Current splits as it reaches the Bear Island Trough. One part enters the Barents Sea together with the Norwegian Coastal Current, which carries low-salinity water from the Baltic Sea and runoff from the Norwegian coast. The remaining part continues as the West Spitsbergen Current to Fram Strait. There the current again splits. Some AW
recirculates westward in the strait to join the southward-flowing East Greenland Current, and the rest enters the Arctic Ocean in two streams. One stream flows over the Svalbard shelf and slope, the other passes west and north around the Yermak Plateau and then continues eastward, eventually joining the inner stream at the continental slope east of Svalbard. The deeper outer stream also transports intermediate and deep waters from the Nordic Seas into the Arctic Ocean. As the AW enters the Arctic Ocean, it encounters, and melts, sea ice, and the upper part becomes colder and less saline. In winter this upper layer is homogenized by convection and mechanical stirring and cooled to freezing temperature. The salinity of the upper layer is 34.2–34.4, and it is advected with the warm AW core eastward along the Eurasian continental slope. In summer, seasonal ice melt creates a low-salinity upper layer, which is removed in
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0
0
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4000 34.7
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.70
27
AW
.97
0.0
27 2
dAW(\) 1
AAW
dAAW(/) CBDW
0
uPDW(\) 0.5 30.444 1.5 35.142 2.5 39.738
uPDW(\) −1
CBDW −0.5
AIW(/)
−1.0
EBDW
EBDW AIW(/)
NDW
NDW −2 34.7
35.1
Salinity
4 3
35
−1.5 34.8
34.9 Salinity
35
35.1
34.86 34.88
34.9
34.92 34.94 34.96
Salinity
Figure 3 Characteristics of the water columns in different parts of the Arctic Mediterranean. Upper row: potential temperature and salinity profiles; lower row: yS curves. Green: The Greenland Sea, the ultimate source of the Arctic Intermediate Water (AIW) and the Nordic Seas Deep Water (NDW) entering the Arctic Ocean. Red: The West Spitsbergen Current in Fram Strait carrying warm Atlantic Water (AW), AIW, and NDW into the Arctic Ocean. Dark yellow: The Fram Strait branch at the continental slope of the Nansen Basin. Magenta: The interior Nansen Basin. Cyan: The Amundsen Basin. Black: The Makarov Basin. Blue: The Canada Basin. The shift in depth of the temperature maximum between the Makarov Basin and the Canada Basin is due to the stronger presence of Pacific water (PW) in the Canada Basin, displacing the deeper part of the water column. Note that the Canadian Basin deep waters becomes warmer than the Eurasian Basin deep waters below 1000 m and more saline below 1500 m, above the sill depth of the Lomonosov Ridge. In the yS diagrams, (/) indicates a stratification unstable in temperature or salinity, (\) indicates a stratification stable in both components. AAW, Arctic Atlantic Water; dAW, dense Atlantic Water; dAAW, dense Arctic Atlantic Water; uPDW, upper Polar Deep Water; CBDW, Canadian Basin Deep Water; EBDW, Eurasian Basin Deep Water. The s1.5 isopycnal shows the density at the sill depth of the Lomonosov Ridge and the s2.5 isopycnal the density at sill depth in Fram Strait.
winter by freezing and the upper layer is homogenized down to the thermocline. No cold halocline is present between the mixed layer and the AW in the Nansen Basin (see Figure 2). The Arctic Intermediate Water (AIW) and the Nordic Seas Deep Water (NDW) that enter the Arctic Ocean in the
West Spitsbergen Current can be identified west and north of the Yermak Plateau as less saline and colder anomalies. Farther to the east these signals in temperature and salinity gradually disappear, but signs of the Nordic Seas waters can still be detected by other tracers, for example, CFCs.
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ARCTIC OCEAN CIRCULATION
Table 1
215
Simplified water mass classification for the Arctic Ocean
Water masses Upper waters (syo27.70) Polar mixed layer (Upper) halocline (Lower) halocline Intermediate waters I (27.70osyo27.97) Atlantic Water Arctic Atlantic Water
Abbreviation
Definition
Main source or origin
PML
32oSo34 32.5oSo33.5 33.5oSo34.5
Arctic Ocean Chukchi Sea Nansen Basin, Barents Sea
AW AAW
2oy 0oyo2
West Spitsbergen Current Arctic Ocean (transformed)
Intermediate waters II (27.97osy, s0.5o30.444) Dense Atlantic Water DAW Dense Arctic Atlantic Water DAAW Arctic Intermediate Water AIW Upper Polar Deep Water UPDW
0oy, 0oy, yo0, yo0,
Deep waters (30.444os0.5) Nordic Seas Deep Water Canadian Basin Deep Water Eurasian Basin Deep Water
So34.915 0.6oy, 34.915oS yo 0.6, 34.915oS
NDW CBDW EBDW
In the Barents Sea, the AW remains in contact with the atmosphere during most of its transit. It is cooled significantly and becomes freshened by net precipitation and by the melting of sea ice. Some of the water entering the Barents Sea returns as colder, denser water to the Norwegian Sea in the Bear Island Channel, but the major part reaches the eastern Barents Sea. In the Barents Sea, the AW becomes separated into three different water masses. (1) The bulk of the inflow is cooled and freshened by air–sea interaction and becomes denser. (2) Some of the AW reaches the shallow areas west of Novaya Zemlya and becomes transformed into saline, dense bottom water by the ice formation and brine rejection. (3) The upper part of the AW interacts with sea ice and a less saline, upper layer is formed by ice melt, which in winter becomes homogenized down to the thermocline. These waters all enter the Arctic Ocean, mainly by passing between Novaya Zemlya and Franz Josef Land into the Kara Sea and then continuing in the St. Anna Trough to the Arctic Ocean. However, some water reaches the Arctic Ocean directly from the Barents Sea along the Victoria Channel. The Barents Sea branch follows the eastern side of the St. Anna Trough and then continues along the continental slope as an almost 1000-m-thick, cold, and weakly stratified water column, displacing the warmer, more saline Fram Strait inflow branch from the slope and deflecting the denser basin waters toward larger depths. Strong isopycnal mixing between the two branches takes place and inversions and irregular intrusive layers are formed (Figure 5, right column).
unstable stable in unstable stable in
in S (/) y and S (\) in S or y (/) S and y (\)
West Spitsbergen Current Arctic Ocean (transformed) Greenland Sea Arctic Ocean Greenland Sea Canadian Basin Eurasian Basin
North of the Laptev Sea the two inflow branches have largely merged and bottom temperatures above 0 1C indicate that Fram Strait branch water again is present at the slope. The boundary current splits at the Lomonosov Ridge with one part continuing into the Canadian Basin, the rest returning along the ridge and the Amundsen Basin toward Fram Strait. The AW temperature in the boundary current is significantly reduced already north of the Laptev Sea. This could partly be due to mixing between the two branches and partly be the result of heat loss to the upper layers and the ice. However, it may also be an indication that some of the Fram Strait branch recirculates already in the Nansen Basin. The returning Fram Strait branch then becomes more prominent in the northern Nansen Basin, while the Barents Sea branch dominates in the Amundsen Basin and along the Lomonosov Ridge (Figure 5, left column). The bulk of the Barents Sea branch in the St. Anna Trough is denser, colder, and less saline than the Atlantic layer and forms a distinct salinity minimum beneath the AW in the Amundsen Basin and above the Gakkel Ridge. This minimum identifies the upper Polar Deep Water (uPDW) in the Eurasian Basin and is located higher in the water column and is more distinct than the AIW minimum in the Nansen Basin deriving from the Nordic Seas. This is about the only area of the Arctic Ocean where a part of the water column is unstably stratified in salinity (Figures 3 and 5). One part of the Barents Sea inflow, mostly comprising water from the Norwegian Coastal Current, enters the Kara Sea through the Kara Strait south of Novaya Zemlya. It mixes with the runoff from Ob
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90° E
135° E 180° SCC AC
45° E SB MC TPD ACC
ESC BI C
BG
WSC
135° W
NAC C NC
JMC EGC EIC
WG C
60
BC
°N
IC
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°N
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45° W
Figure 4 The circulation of the upper layers of the Arctic Mediterranean Sea. Warm Atlantic currents are indicated by red arrows, cold, less-saline polar and arctic currents by blue arrows. Low-salinity transformed currents are shown by green arrows. The maximum ice extent is shown by a blue and the minimum ice extent by red striped line. AC, Anadyr Current; ACC, Alaskan Coastal Current; BC, Baffin Current; BIC, Bear Island Current; BG, Beaufort gyre; EGS, East Greenland Current; EIC, East Iceland Current; ESC, East Spitsbergen Current; IC, Irminger Current; JMC, Jan Mayen Current; MC, Murman Current; NAD, North Atlantic Drift; NAC, Norwegian Atlantic Current; NCC, Norwegian Coastal Current; SB, Siberian branch (of the Transpolar Drift); SCC, Siberian Coastal Current; TPD, Transpolar Drift; WGC, West Greenland Current; WSC, West Spitsbergen Current.
and Yenisey, forming the low-salinity water present on the Kara Sea shelf. Most of this shelf water flows through the Vilkiltskij Strait into the Laptev Sea, where it receives additional freshwater, mostly from the Lena River. The main export of shelf water from the Eurasian shelves to the deep basins occurs across the Laptev Sea shelf break into the Amundsen Basin, and the low-salinity shelf water overruns the boundary current and the two inflow branches. The upper parts of the Fram Strait and the Barents Sea branches become isolated from the ice, the sea surface, and the atmosphere and evolve into halocline waters. The Fram Strait branch supplies the halocline in the Amundsen Basin and the Makarov Basin and the lower halocline, beneath the Pacific water, in the northern part of the Canada Basin. The Barents Sea branch halocline, which remains at the continental slope, is affected by the stronger mixing occurring over the steep topography and becomes warmer by mixing with AW from below. It eventually supplies
the lower halocline in the southern Canada Basin. The shelf water outflow supplies the PML in the Amundsen and Makarov Basins but it has high enough salinity also to contribute to the halocline waters in the Canada Basin, beneath most of the Pacific waters (Figure 6). The properties of the Atlantic and intermediate waters in the Canada Basin are distinctly different from those in the Eurasian Basin, indicating that the waters of the boundary current become transformed along their paths in the Canadian Basin. The temperature maximum in the Atlantic layer becomes colder and less saline, while in the intermediate water range the temperature and the salinity increase and the intermediate uPDW salinity minimum disappears. The yS curves below the temperature maximum approach a straight line toward higher salinities and lower temperatures with increasing depth (Figure 3). The boundary current circulates cyclonically around the Canada Basin and splits at the different
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Pressure (dbar)
ARCTIC OCEAN CIRCULATION
0
0
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.7 0
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.9
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0.5 30.444 1.5 35.142 2.5 39.738 35
27
0.5 0 −0.5 −1.0 34.7
Salinity
34.8
34.9
35
Salinity
Figure 5 Potential temperature and salinity profiles and yS curves showing the interaction and interleaving between the Fram Strait branch and the Barents Sea branch north of Severnaya Zemlya (right column) and the water properties of the Nansen and Amundsen basins offshore of the Fram Strait branch (left column). Red stations: Fram Strait branch, blue station: the Barents Sea branch, black and cyan stations: active mixing between the branches. Interleaving is present not only in the AW but also in the deeper layers. Offshore of the Fram Strait branch, the warm, saline intrusions in the Nansen Basin (black and magenta stations) suggest a close recirculation of the Fram Strait branch in the Nansen Basin, while the colder, less saline intrusions in the Amundsen Basin (green and blue stations) indicate that the intermediate part of the water column here is dominated by Barents Sea branch waters.
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135° E
90° E
180°
45
135° W
°N
60
0°
°N
50
90° W
45° W
Figure 6 Circulation of the Atlantic-derived halocline waters and the distribution of Pacific water (orange) and Eurasian river runoff (green). The proposed source areas for the Fram Strait branch lower halocline water (black) and the Barents Sea branch lower halocline water (blue), and the circulation of these waters in the Arctic Ocean are indicated. RR, river runoff; PW, Pacific water; AW, Atlantic water. The cross indicates possible contribution of Barents Sea branch lower halocline water to the Baffin Bay bottom water. The isoline shown is 500 m. Based on Rudels B, Jones EP, Schauer U, and Eriksson P (2004) Atlantic sources of the Arctic Ocean surface and halocline waters. Polar Research 23: 181–208.
ridges (Figure 7). One loop enters the Makarov Basin along the Mendeleyev Ridge, and one loop penetrates into the northern Canada Basin at the Chukchi Plateau. The boundary current also enters the Canada Basin at the Alpha Ridge and the Makarov Basin at the North American side of the Lomonosov Ridge. These circulation loops and their interactions with the boundary current induce a time lag that makes the Atlantic and intermediate waters of the different basins distinct, even if the waters originate from the same inflow across the Lomonosov Ridge.
Processes The climatic forcing on the Arctic Ocean is strong; the large variation of incoming solar radiation, the severe cooling during the polar night, and the intense weather systems, either locally formed polar lows or cyclones advected from lower latitudes, all contribute in forming the Arctic Ocean characteristics. The strong stability of the deep Arctic Ocean basins confines the effects of these forcing fields to the ice-ocean surface and to the PML, but the shallow Arctic shelves, which
experience the largest impact of the forcing, the strong runoff in summer and excessive ice formation in winter, do not only add low-salinity upper water and ice to the central basins but also create waters dense enough to break through the stratification, transforming the deeper advected layers. Ice formation releases brine, and in lee polynyas, areas of open water where the ice is removed by offshore winds, sufficient ice may form to allow the released brine to overcome the initial low salinity on the shelves and create saline dense bottom waters at freezing temperature. These bottom waters eventually cross the shelf break and descend into the deep basins until their densities match those of the surrounding. They then merge with the ambient water. Less-dense plumes add colder water to the upper layers, supplying water to the halocline. This occurs in the Chukchi Sea, where the water of the Canada Basin upper halocline with salinity 33.1 is formed. In the Nansen Basin no halocline is present and the density range between the upper mixed layer and the AW is small. The shelf outflows then either enter the mixed layer or sink into and cool the thermocline and the Atlantic layer. The formation of
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ARCTIC OCEAN CIRCULATION
135° E
219
90° E
180°
45°
135° W
°N
60
0°
°
50 N
90° W
45° W
Figure 7 Schematics showing the circulation in the subsurface Atlantic and intermediate layers in the Arctic Mediterranean Sea. The interactions between the Barents Sea and Fram Strait inflow branches north of the Kara Sea as well as the recirculation and different inflow streams in Fram Strait and the overflows across the Greenland–Scotland Ridge are indicated. The isoline shown is 500 m. Based on Rudels B, Jones EP, Anderson LG, and Kattner G (1994) On the intermediate depth waters of the Arctic Ocean. In: Johannesen OM, Muench RD, and Overland JE (eds.) AGU Geophysical Monographs 85: The Polar Oceans and Their Role in Shaping the Global Environment, pp. 33–46. Washington, DC: American Geophysical Union.
halocline waters in the Eurasian Basin occurs, when the boundary current is overrun by the less saline shelf water, and the initial mixed layers become transformed into halocline waters. Comparison between the intermediate and deep waters in the Nordic Seas and in the Arctic Ocean shows that the shelf-slope convection not only brings salt but also heat into the deeper layers (Figures 3 and 5). This implies that the denser, more saline plumes are entraining warmer water, as they pass through the intermediate Atlantic layer. Observations of outflowing, cold and dense water from Storfjorden in southern Svalbard have shown that as the outflow reaches deeper than 1500 m, it has become warmer than the ambient water, supporting the idea of entraining boundary plumes. No plumes have yet been followed from their sources down the slope in the Arctic Ocean, but saline, warmer, and denser bottom layers have been observed deeper than 2000 m at the continental slope north of Severnaya Zemlya. The deep waters of the Canadian and Eurasian Basins differ significantly, the Canadian Basin Deep
Water (CBDW) being warmer and more saline than the Eurasian Basin Deep Water (EBDW). The difference in temperature was first believed due to the presence of the Lomonosov Ridge, blocking the cold deep water from the Nordic Seas, but the temperature difference starts well above the sill depth of the Lomonosov Ridge, showing that heat as well as salt is added to the uPDW and the CBDW in the Canadian Basin. The differences between the Amundsen and Nansen Basins are small but clear. The Nansen Basin is slightly less saline than the Amundsen Basin between 1600 and 2600 m due to a stronger presence of NDW and AIW. In the Amundsen Basin a mid-depth (1700 m) salinity maximum or at least a sharp bend in the yS curves is observed. It is strongest closer to Greenland but is present over most of the basin. This maximum derives from the CBDW that crosses the Lomonosov Ridge and then penetrates into the central Amundsen Basin from the Greenland continental slope. In the Nansen, Amundsen, and Canada Basins, the salinity of the deep water increases toward the
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bottom, while the temperature goes through a minimum and then increases slightly before an isothermal and isohaline bottom layer is reached (Figure 3). Shelf-slope convection cannot explain the temperature minimum in the deep waters, but it could be due to advection between the basins. The deep water in the Makarov Basin has at the sill depth (2400 m) of the Alpha Ridge similar yS characteristics as the temperature minimum in the Canada Basin and could supply the minimum. In the Eurasian Basin this explanation does not work. A minimum is present in both the Amundsen and Nansen Basins and these minima lie too deep to derive from the inflow of colder NDW. One possibility is that the St. Anna Trough intermittently conduits colder, denser bottom water, formed in the Barents Sea, into the Nansen Basin. This cold water would enter below the warm layers of the Fram Strait branch and by entraining less heat it could contribute colder water to the deeper part of the water column than the slope convection occurring farther east around Severnaya Zemlya. Geothermal heat flux has been proposed as an energy source for heating and stirring the bottom layers, keeping them homogenous and gradually increasing their temperature, thereby creating the overlying temperature minimum. The high bottom salinity could then be a remnant from convection events that took place several hundred years ago and brought dense, saline, and cold water to the bottom, where it now is isolated and gradually becomes warmer, thicker, and less dense. Whether the bottom layers are kept homogenous by geothermal heating and convection from below, or if the stirring is due to mechanical mixing as the boundary plumes enter the bottom layer is an open question. However, geothermal heating and ventilation by shelf-slope convection are not two mutually exclusive processes and could operate simultaneously. The Makarov Basin is different. The salinity becomes constant with depth, while the temperature still decreases, creating a temperature-stratified layer above the isothermal and isohaline bottom water (Figure 3). This rules out the gradual heating of an old, fossil bottom layer but requires that colder, rather than more saline, water renews the bottom layer. This then excludes shelf-slope convection as a dominant process. One possibility is that colder water from the Amundsen Basin is brought across the Lomonosov Ridge, either intermittently, forced by, for example, topographically trapped waves, or flowing through yet undetected passages in the ridge. The central part of the Lomonosov Ridge, where such an exchange was believed to occur, has recently been surveyed. A gap was found but not as deep as expected, barely 1900 m, and more critically, the
water mass properties observed during the survey indicated that the flow of the densest water was from the Makarov Basin to the Amundsen Basin – in the wrong direction. This inflow of Makarov Basin water followed the Lomonosov Ridge toward Greenland and is likely to be one source, perhaps the most important one, for the mid-depth salinity maximum in the Amundsen Basin. A flow onto the shelves is needed to compensate for the supply to the PML in summer and the outflow of dense water to the deeper layers and the export of ice in winter. This flow could come from the Arctic Ocean, along the bottom in summer and at the surface in winter, or along horizontally separated inflow and outflow paths across the shelf break. It could also be supplied from behind, like the Bering Strait inflow and the inflow to the Barents Sea. These inflow shelves differ from the inner shelves of the Kara, Laptev, and East Siberian Seas, where a two-way exchange across the shelf break might be more likely, even necessary. However, the difference is not as large as it might appear. Most of the Barents Sea inflow actually passes the Kara Sea before it enters the deep basins and much of the runoff from Ob and Yenisey, together with a substantial fraction of the Barents Sea inflow, continues into the Laptev Sea and then enters either the Amundsen Basin or the East Siberian Sea. Such eastward flow is the usual fate of river plumes, which tend to follow the coastline eastward. This also holds for the East Siberian Coastal Current, which carries the runoff from perhaps the most ‘inner’ shelf of the Arctic Ocean, that of the East Siberian Sea, into the Chukchi Sea. The East Siberian Coastal Current occasionally, but rarely, passes south through Bering Strait into the Bering Sea. The circulation pattern in the Arctic Ocean outlined above gives the impression that water masses are advected, with little mixing, in loops through the different basin, replacing one vintage water mass with another. This is not entirely the case. The structure of the salinity and temperature profiles shows sign of intrusive mixing and the presence of lenses or eddies of anomalous water masses in the water column. Eddies were first reported in the Canadian Basin halocline and were there considered related to the outflow of dense water from the Chukchi Sea. These eddies are mostly anticyclonic, 10–20 km in diameter and highly energetic with maximum velocities above 0.3 m s 1. Eddies have subsequently been observed in all water masses of the Arctic Ocean, although the velocities associated with the deeper-lying eddies have so far not been determined. Their water characteristics imply that the eddies have traveled considerable distances as
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ARCTIC OCEAN CIRCULATION
coherent water bodies without their anomalous properties being removed by smaller-scale mixing, for example double-diffusive convection, merging the eddies with the Arctic Ocean water columns in different basins (Figure 8). The intrusive layers are particularly intriguing (Figures 5 and 8). The largest property amplitudes are found at the frontal zones but the layers appear to reach over entire basins, making their extent perhaps the largest one observed in the World Ocean. The inversions allow for the release of the potential energy stored in the unstably stratified component by double-diffusive convection, which can drive the interleaving layers across the basins, and this has been suggested as a mechanism for transfer of heat from the boundary current into the deep basins. For the layers to expand, along-layer gradients in heat and salt are required. In the interior of the basins such gradients are often absent, and an alternative explanation for the wide extent of the intrusive layers is that they are formed in the frontal zones, expand initially, driven by double-diffusive fluxes, until the potential energy, available in the stratification, is removed. After this the layers remain as
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fossil structures that are advected around the main circulation gyres. The intrusions between the Fram Strait and the Barents Sea branch water observed in the Amundsen Basin and over the Gakkel Ridge have been interpreted as fossil structures, initially formed by the interactions between the Fram Strait and the Barents Sea branches north of the Kara Sea and now being advected toward Fram Strait (Figure 5). Interleaving layers are observed in all background stratifications: saltfinger unstable, diffusively unstable, as well as when both components are stably stratified. In the stable–stable situation, disturbances extending across the front are necessary to create the initial inversions. Differential diffusion, taking into account the more rapid diffusion of heat in weakly turbulent surroundings, has been proposed as a mechanism for creating interleaving in a stable–stable stratification. The time required to create layers would then be in the order of years. This appears long, since interleaving layers are found very close to the area where the parent water masses first meet also when both components are stably stratified.
Variability 2.5
Potential temperature (°C)
2.0 1.5 1.0 0.5
7
27.9
0.5 30.444 1.5 35.142 2.5 39.738
0 −0.5 −1.0
34.86
34.88
34.9 Salinity
34.92
34.94
Figure 8 yS diagram showing eddies present in the intermediate and deep layers on a section taken in 1996 across the Eurasian Basin. The red station shows the warm and saline Canadian Basin Deep Water, the blue station indicates an isolated lens of cold, low-salinity Barents Sea branch water. The cyan and black stations show an eddy of Barents Sea branch water (cyan station) with warmer and more saline water in both the slope and the basin directions (black stations) and surrounded by interleaving structures like a ‘meddy’. This is in contrast to the slope to basin decrease in salinity and temperature seen in the interleaving in the Atlantic layer (although here also an eddy (not shown) was detected). Finally an eddy of warm, saline Fram Strait branch water, yellow station, was present in the colder, Barents Sea branch dominated Atlantic layer in the Amundsen Basin.
Until 1990 an underlying assumption has been that the Arctic Ocean, at least in its deeper parts, is reasonably quiet and unchanging and that observations made during a longer period, 10–20 years, could be merged and used to describe the basic hydrographic conditions. The observations in the 1990s proved the Arctic Ocean to be as variable and changing as any other ocean. An inflow of anomalously warm AW was reported in 1990 and has been observed propagating in the boundary current and into the different basins, changing the characteristics of the Atlantic layer in the Arctic Ocean. This warm inflow event persisted for almost a decade. Colder water then entered the Arctic Ocean for a short period after which the inflowing AW again became warmer. These inflow events can be traced upstream and originate from the input of warmer AW across the Greenland–Scotland Ridge. The pulses have also been followed in the Arctic Ocean, tracing several of the suggested circulation loops in the basins, giving timescales for the movements along the different loops. Model work has reproduced many of these events and their pathways around the Arctic Ocean. The first pulse has now left the Eurasian Basin and partly exited the Arctic Ocean through Fram Strait. In the Amundsen Basin it has been replaced by colder water, while the part that entered the Makarov Basin at the Mendeleev Ridge has circulated around the
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Makarov Basin and is now found at the Makarov Basin side of the Lomonosov Ridge, practically removing the temperature front previously present along the ridge. The pulse has also entered the northern Canada Basin at the Chukchi Plateau. Its movements around and south of the Chukchi Plateau have taken comparably long, and it has been proposed that the AW here enters not mainly in the boundary current but directly into the basin as intrusive, doublediffusively driven layers. Older water, previously found in the southern Canada Basin, has shifted northward along the slope and is seen penetrating along the Alpha Ridge into the northern Canada Basin. The 1990 inflow event coincided with a strong, positive state of the North Atlantic Oscillation (NAO) and of the Arctic Oscillation (AO), which also affected the distribution of the runoff from the Siberian rivers. Instead of entering the Amundsen Basin from the Laptev Sea it continued eastward to the Makarov Basin and the northern Canada Basin. The upper Pacific water lens, which in the 1970s extended over the entire Canadian Basin to the Lomonosov Ridge, then contracted to the Canada Basin. Similar scenarios have been reproduced in model studies. The shifting of the Pacific/Atlantic surface front as well as the river water (shelf water) front counterclockwise toward the Canada Basin elevated the effects of the warmer AW. In the Amundsen Basin, and partly in the Makarov Basin, it approached closer to the sea surface, into the levels previously occupied by the halocline, thus magnifying the increase of both temperature and salinity at these levels. The shifting of the river water front also caused an increase in salinity of the surface layers of the Amundsen and Makarov Basins. The area with a deep winter mixed layer, previously confined to the Nansen Basin, expanded into the Amundsen basin, almost reaching the Lomonosov Ridge. The winter convection reached down to 120–130 m and actually caused a temperature decrease immediately above the Atlantic layer. The river water front has during the 2000s moved back into the Amundsen Basin, almost as far as the Gakkel Ridge, indicating that the shelf water outflow from the Laptev Sea again primarily enters the Amundsen Basin, recreating the PML–halocline structure in that basin. The inflow in the Barents Sea branch also appears to vary. The intermediate salinity minimum in the Eurasian Basin has become more pronounced and less saline uPDW has crossed the Lomonosov Ridge in the boundary current along the continental slope and entered the Canadian Basin. This has made the uPDW characteristics in the Makarov Basin more similar to those in the Eurasian Basin, with a curved rather than
a straight yS curve below the temperature maximum. A question is, if these anomalies in the Canadian Basin will stay long enough to be removed by shelfslope convection and interior mixing processes, recreating the older, smooth, yS characteristics, or if this inflow will create entirely new yS structures? If so, does this mean that the gyre circulation has changed and the communication between the Eurasian and Canadian Basins has increased, leading to a more rapid ventilation of the Canadian Basin? Historical data from the Russian archives indicate that sudden changes have occurred before, and the situation now observed may not be unique. The use of CTDs instead of Nansen bottles also reveals structures in the water column previously not resolved. The circulation pattern in the Arctic Ocean responds to long periodic (decadal) variations in the atmospheric circulation, the NAO and the AO. The positive AO state increases the inflow of AW and weakens the Beaufort gyre, while the negative state leads to a well-developed Beaufort gyre and a smaller inflow of AW. The negative state then retains the fresh water and the ice, while the positive state acts to reduce the storage of ice and low-salinity upper waters, which, together with a larger inflow of AW, increases the mean salinity of the Arctic Ocean waters. Perhaps the most prominent change observed in the Arctic Ocean has been the retreat of the ice cover. A reduction of the minimum ice extent of 1.5–2 106 km2 (420%) between 1979 and 2005 has been observed. Comparison between submarine observations of ice thicknesses 20 years apart indicates a thinning of the ice, from 3.1 to 1.8 m. However, the magnitude of this change has been contested. Thickness observations at the same position years apart might not be relevant, since the distribution of the ice will depend upon the forcing of the ice field, which will vary between the different years. Nevertheless, even disregarding changes in ice thickness, the ice cover has become significantly reduced. A thinner and less extensive ice cover indicates a loss of ice storage, which could either be due to reduced formation of ice, or to an increased export of ice out of the Arctic Ocean. The amount of freshwater released by the reduction of the ice cover could have contributed to perhaps the largest fraction of the freshening of the Nordic Seas and to the subpolar gyre reported in recent years, but this does not give any information about where the phase change occurred, in the Arctic Ocean or south of Fram Strait. The time series of the ice export are not long enough to provide an answer and the knowledge of the liquid freshwater export is uncertain, if not completely absent. The presence of the halocline between the PML and the Atlantic layer makes the stirring in the PML
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ARCTIC OCEAN CIRCULATION
entrain cold water from below, and the ice cover is isolated from the heat stored in the Atlantic layer. The situation in the mid-1990s with no river (shelf) water present in the Amundsen Basin could lead to the entrainment of warm AW into the uppermost layer and thus bring heat to the sea/ice surface. This sensible heat rather than latent heat of ice formation could then be supplied to the atmosphere and less ice would form in winter. The heat might also directly melt the ice, but if the stirring of the mixed layer, driving the entrainment of heat from below, is due to haline convection, a change from freezing to melting would stop the convection, and the entrainment as well as the heat transport from below would cease.
Transports The last 10 years have seen large programs focusing on measuring the transports through the key passages of the Arctic Ocean. Much of this activity has been coordinated by the ASOF (Arctic and Subarctic Ocean Fluxes) program but other projects have also been involved. The observations of the Bering Strait inflow have largely confirmed the transport estimates proposed by Russian researchers 50 years ago. A mean inflow of 0.8 Sv of low-salinity (Sr32) water takes place. The seasonal variations are large; the inflow is 1.2 Sv in summer and 0.4 Sv in winter. The recent observations have shown that the freshwater transport through Bering Strait might be substantially (20%) higher than previously estimated, making it twothirds of the river runoff. This reevaluation is due to the inclusion of the transport in the low-salinity Alaskan Coastal Current. The transports through the Canadian Arctic Archipelago, notoriously difficult to measure due to the remoteness, the severe climate, and the proximity to the magnetic north pole that makes direction determinations extremely difficult, have in recent years been measured in the Hell Gate and Cardigan Strait (the Jones Sound) and in Barrow Strait (Lancaster Sound). This gives observed transports through two of the three main passages through the Archipelago. Transport measurements have also been made in Nares Strait, but year-long transport estimates from Nares Strait are not yet available. The fluxes through the narrow Hell Gate are directed out of the Arctic Ocean, barotropic, and almost constant, while in the neighboring Cardigan Strait the flow is weaker and reversals are observed. The combined average transport is estimated to 0.3 Sv. The transports in the wider Barrow Strait are largely barotropic on the southern side of the channel
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and directed eastward, toward Lancaster Sound and Baffin Bay. A weak, baroclinic westward flow is observed on the northern side, indicating transport of runoff and penetration of water from Baffin Bay and Lancaster Sound. The flow is highly variable with an estimated mean transport of 0.75 Sv from the Arctic Ocean to Baffin Bay. Models indicate that the Barrow Strait (Lancaster Sound) might contribute one-third to one-half of the total outflow through the Canadian Arctic Archipelago. Should this be correct the total transport would be 1.5–2 Sv. It is low-salinity, primarily Pacific, water that passes through the archipelago, but the bottom water in Baffin Bay, 0.5 1C and 34.5, likely derives from the Arctic Ocean through Nares Strait and would then be supplied by lower halocline waters. The transport of the AW to the Barents Sea between Bjørnøya and Norway has been measured continuously for 10 years. The mean net transport is into the Barents Sea, but there is a return flow of transformed, colder, and denser water. The mean net transport of AW to the Barents Sea has been estimated to 1.5 Sv. However, short periodic variations are large and in spring, due to changing wind conditions, a whole month of small net inflow, occasionally even outflow, has been observed. There are also indications of variability on longer timescales, 3–4 years, but no trend has been detected. The AW is warm and saline at the entrance, but it loses much of its heat during transit and does not contribute heat to the central Arctic Ocean. To the inflow of AW should be added a transport of 0.7 Sv of less saline (34.4) water from the Norwegian Coastal Current, increasing the net inflow to B2.2 Sv. Fram Strait, which also has been monitored regularly since 1997, has a two-directional flow, and not only polar surface water (PSW), comprising the PML and halocline waters, and AW but also intermediate and deep water masses pass through the strait. The flow is largely barotropic and highly variable in space and time. Eddies are present, both barotropic and baroclinic, which complicates the transport estimates. Most of the steady flow occurs in the northward-flowing West Spitsbergen Current to the east and the southward-flowing East Greenland Current to the west. The inflow comprises warm AW and colder, less-saline AIW and NDW, while the outflow carries sea ice, low-salinity PSW, cool Arctic Atlantic Water (AAW), uPDW with temperature between 0 and 0.5 1C and CBDW, seen as a salinity maximum at 1700 m, and the colder, but also saline EBDW close to the bottom. The observed total northward and southward transports are large, 10–15 Sv. The mean net transport is much smaller and lies between 1.5 and 2.5 Sv
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southward, out of the Arctic Ocean. A large recirculation is present in the strait. This appears partly associated with the barotropic eddies that drift westward along the sill of the strait and carry water from the West Spitsbergen Current to the East Greenland Current. The transports through Fram Strait of largest climate importance are the export of ice and the heat carried by the AW. There is a large interannual variability but no trend has been found. The estimated mean ice export is B0.9 Sv, while the outflow of PSW is B1 Sv. The inflow of AW and heat also varies strongly from year to year but larger transports and higher temperatures have been observed in the last years, giving a flux B50 TW. The average heat transport since 1997 is 35–40 TW. This is estimated relative to an assumed mean temperature of 0.1 1C in the Arctic Ocean. However, since the mass budget is not closed, the heat transport will depend upon the choice of reference temperature. The long-term mean inflow of AW, 42 1C, is c. 3 Sv. Whether an increased heat transport through Fram Strait, connected with warmer AW and perhaps a stronger flow, also leads to more heat being available for the Arctic, for ice melt and for the atmosphere, is not yet clear. The fact that pulses of warm AW can be traced around the Arctic Ocean could imply that most of the heat is not lost but only stored and will eventually leave the Arctic Ocean through Fram Strait. The heat advected into the Arctic Ocean through Bering Strait is located closer to the sea surface and could have a larger impact on the heat flux to the ice cover from below and on the thickness of the ice cover. The large retreat in ice extent reported in the last years has mostly occurred in the southern Canada Basin, close to the Chukchi Sea.
and in the Labrador Sea. However, the fresh water largely remains in the East Greenland Current and mostly bypasses the Greenland Sea and Iceland Sea gyres. In recent years, ice has not been formed in the central Greenland Sea and the Greenland Sea has been dominated by thermal convection. The AAW and the uPDW contribute to the Denmark Strait overflow, while the CBDW and the EBDW mainly enter the Greenland Sea, where they supply the mid-depth (1800 m) temperature maximum and the deep salinity maximum. In recent years, the local convection in the Greenland Sea has not penetrated through the temperature maximum and only less-dense AIW has been formed. The production of AIW has lead to a more direct contribution of the Greenland Sea to the AMOC, especially to the East Greenland Current and the Denmark Strait overflow. The denser Arctic Ocean deep water masses, previously continuing in the East Greenland Current to Denmark Strait are now entering the Greenland Sea and gradually replacing the old, colder, and less saline Greenland Sea deep and bottom waters, making the Greenland Sea water column more ‘Arctic’ in character.
See also Bottom Water Formation. Double-Diffusive Convection. Heat Transport and Climate. Intrusions. Meddies and Sub-Surface Eddies. Non-Rotating Gravity Currents. North Atlantic Oscillation (NAO). Ocean Circulation. Ocean Circulation: Meridional Overturning Circulation. Open Ocean Convection. Overflows and Cascades. Polynyas. Rotating Gravity Currents. Sea Ice: Overview. Upper Ocean Heat and Freshwater Budgets. Upper Ocean Mixing Processes. Water Types and Water Masses.
Significance for Climate The influence of the Arctic Ocean on the circulation at lower latitudes is mainly through the export of freshwater as ice and as low-salinity PSW, and through the export of dense intermediate and deep waters that contribute to the Greenland–Scotland overflow and to the North Atlantic Deep Water, enforcing the thermohaline circulation and the Atlantic meridional overturning circulation (AMOC). Of the around 6-Sv overflow water supplied to the North Atlantic Deep Water by the Arctic Mediterranean about 3 Sv have passed through the Arctic Ocean. The outflow of ice and less-dense surface water could increase the stability of the water column in the convection areas to the south, in the Nordic Seas,
Further Reading Bjo¨rk G, Jakobsson M, Rudels B, et al. (2007) Bathymetry and deep-water exchange across the central Lomonosov Ridge at 88–891 N. Deep-Sea Research I 54: 1197--1208 (doi:10.1016/j.dsr.2007.05.010). Bjo¨rk G, So¨derqvist J, Winsor P, Nikolopoulos A, and Steele M (2002) Return of the cold halocline to the Amundsen Basin of the Arctic Ocean: Implications for the sea ice mass balance. Geophysical Research Letters 29(11): 1513 (doi:10.1029/2001GL014157). Coachman LK and Aagaard K (1974) Physical oceanography of the Arctic and sub-Arctic seas. In: Herman Y (ed.) Marine Geology and Oceanography of the Arctic Ocean, pp. 1--72. New York: Springer.
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Dickson B, Meincke J, and Rhines P (eds.) (2008) Arctic– Subarctic Ocean Fluxes: Defining the Role of the Northern Seas in Climate. New York: Springer. Dickson RR, Rudels B, Dye S, Karcher M, Meincke J, and Yashayaev I (2007) Current estimates of freshwater Arctic and subarctic seas. Progress in Oceanography 73: 210--230 (doi:10.1016/j.pocean.2006. 12.003). Fahrbach E, Meincke J, Østerhus S, et al. (2001) Direct measurements of volume transports through Fram Strait. Polar Research 20: 217--224. Jakobsson M, Cherkis NZ, Woodward J, Macnab R, and Coakley B (2000) New grid of Arctic bathymetry aids scientists and mapmakers. EOS, Transactions of American Geophysical Union 81(9): 89--96. Jakobsson M, Macnab R, Mayer L, et al. (2008) An improved bathymetric portrayal of the Arctic Ocean: Implications for ocean modelling and geological, geophysical and oceanographic analyses. Geophysical Research Letters 35: L07602 (doi:10.1029/2008 GL0335220). Johannesen OM, Muench RD, and Overland JE (eds.) (1994) AGU Geophysical Monographs 85: The Polar Oceans and Their Role in Shaping the Global Environment. Washington, DC: American Geophysical Union. Jones EP, Rudels B, and Anderson LG (1995) Deep water of the Arctic Ocean: Origin and circulation. Deep-Sea Research 42: 737--760. Leppa¨ranta M (ed.) (1998) Physics of Ice-Covered Seas. Helsinki: Helsinki University Press. Lewis EL, Jones EP, Lemke P, Prowse T, and Wadhams P (eds.) (2000) The Freshwater Budget of the Arctic Ocean. Dordrecht: Kluwer Academic Publishers. Merryfield WJ (2002) Intrusions in double-diffusively stable Arctic waters: Evidence for differential mixing? Journal of Physical Oceanography 32: 1452--1459. Nansen F (1902) Oceanography of the North Polar Basin. Scientific Results III(9): The Norwegian North Polar Expedition 1893–96. Oslo: Jacob Dybwad. Peterson BJ, McClelland J, Curry R, Holmes RM, Walsh JE, and Aagaard K (2006) Trajectory shifts in the Arctic and subarctic freshwater cycle. Science 313: 1061--1066. Prinsenberg SJ and Hamilton J (2005) Monitoring the volume, freshwater and heat fluxes passing through Lancaster Sound in the Canadian Arctic Archipelago. Atmosphere-Ocean 43: 1--22. Quadfasel D, Sy A, Wells D, and Tunik A (1991) Warming of the Arctic. Nature 350: 385. Rudels B, Jones EP, Anderson LG, and Kattner G (1994) On the intermediate depth waters of the Arctic Ocean. In: Johannesen OM, Muench RD, and Overland JE (eds.) AGU Geophysical Monographs 85: The Polar Oceans and Their Role in Shaping the Global Environment, pp. 33--46. Washington, DC: American Geophysical Union. Rudels B, Jones EP, Schauer U, and Eriksson P (2004) Atlantic sources of the Arctic Ocean surface and halocline waters. Polar Research 23: 181--208.
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Rudels B, Muench RD, Gunn J, Schauer U, and Friedrich HJ (2000) Evolution of the Arctic Ocean boundary current north of the Siberian shelves. Journal of Marine Systems 25: 77--99. Schauer U, Fahrbach E, Østerhus S, and Rohardt G (2004) Arctic warming through Fram Strait: Oceanic heat transports from 3 years of measurements. Journal of Geophysical Research 109: C06026 (doi:10.1029/2003 JC001823). Serreze MC, Barrett A, Slater AJ, et al. (2006) The largescale freshwater cycle in the Arctic. Journal of Geophysical Research 111: C11010 (doi:10.1029/2005 JC003424). Smith WO, Jr. (ed.) (1990) Polar Oceanography, Part A: Physical Sciences. San Diego: Academic Press. Smith WO, Jr. and Grebmeier JM (eds.) (1995) Coastal and Estuarine Studies: Arctic Oceanography, Marginal Ice Zones and Continental Shelves. Washington, DC: American Geophysical Union. Steele M and Boyd T (1998) Retreat of the cold halocline layer in the Arctic Ocean. Journal of Geophysical Research 100: 881--994. Stein R and MacDonald RW (eds.) (2004) The Organic Carbon Cycle in the Arctic Ocean, 362pp. Berlin: Springer. Swift JH, Aagaard K, Timokhov L, and Nikiforov EG (2005) Long-term variability of Arctic Ocean waters: Evidence from a reanalysis of the EWG data set. Journal of Geophysical Research 110: C03012 (doi:10.1029/ 2004JC002312). Timmermanns M-L, Garrett C, and Carmack E (2003) The thermohaline structure and evolution of the deep water in the Canada Basin, Arctic Ocean. Deep-Sea Research I 50: 1305--1321 (doi:10.1016/S0967-0637(03)00125-0). Untersteiner N (ed.) (1986) The Geophysics of Sea Ice. New York: Plenum. Wadhams P, Gascard J-C, and Miller L (1990) Topical studies in oceanography: The European Subpolar Ocean Programme: ESOP. Deep-Sea Research II 46: 1011--1530. Walsh D and Carmack EC (2002) A note on evanescent behavior of Arctic thermohaline intrusions. Journal of Marine Research 60: 281--310. Wassmann P (ed.) (2006) Special Issue: Structure and Function of Contemporary Food Webs on Arctic Shelves: A Pan-Arctic Comparison. Progress in Oceanography 71(2–4): 123–477. Wheeler PA (1997) Topical studies in oceanography: 1994 Arctic Ocean section. Deep-Sea Research II 44: 1483--1758. Woodgate RA and Aagaard K (2005) Revising the Bering Strait freshwater flux into the Arctic Ocean. Geophysical Research Letters 32: L02602 (doi:1029/2004 GL021747). Woodgate RA, Aagaard K, Muench RD, et al. (2001) The Arctic boundary current along the Eurasian slope and the adjacent Lomonosov Ridge. Deep-Sea Research I 48: 1757--1792.
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ARTIFICIAL REEFS W. Seaman and W. J. Lindberg, University of Florida, Gainesville, FL, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction Artificial reefs are intentionally placed benthic structures built of natural or man-made materials, which are designed to protect, enhance, or restore components of marine ecosystems. Their ecological structure and function, vertical relief, and irregular surfaces vary according to location, construction, and degree to which they mimic natural habitats, such as coral reefs. For ages, humans have taken advantage of the behavior of some organisms to seek shelter at submerged objects by introducing structures into shallow waters, where biological communities could form and fishes could be harvested. Contemporary purposes for artificial reefs include increasing the efficiency of artisanal, commercial, and recreational fisheries, producing new biomass in fisheries and aquaculture, boosting underwater recreation and ecotourism opportunities, preserving and renewing coastal habitats and biodiversity, and advancing research. This article reviews the evolution, spread, and increased scale of such practices globally, discusses their scientific basis, and describes trends concerning artificial reef planning, evaluation, and their appropriate application in natural resource management.
Utilization of Reefs in the World’s Oceans The near-shore ocean ecosystems of all inhabited continents contain artificial reefs. Their scales and purposes vary widely, ranging, for example, from placement in local waters of small structures by individuals in artisanal fishing communities, to deploying complex systems of heavy modules in distant areas by organizations concerned with commercial seafood production, to pilot studies using experimentally based structures for restoration of seagrass, kelp, and coral ecosystems. The origins of artificial reef technology are traced to places as diverse as Japan and Greece. Modern deployment of reefs has the longest history and has been most widespread in Eastern Asia, Australia, Southern Europe, and North America. In recent years, the technology has been introduced or more extensively established in scores of nations worldwide, including in Central and South America and the Indo-Pacific basin. No organization maintains a database for artificial reef development globally, so that assessments of reef-related activity must be derived from the records of economic development, fishery, and environment organizations, as well as from scientific journal articles which serve as a proxy for gauging national efforts. For centuries structures of natural materials have been used to support artisanal and subsistence fishing in coastal communities, particularly in tropical areas (Figure 1). In India, traditionally tree branches have been weighted and sunken. Brush parks made of
Figure 1 In some areas of long-standing artisanal fisheries, structures such as the pesquero (left), a bundle of mangrove tree branches, are used for benthic structure, as in Cuba. More recently, small fabricated modules, such as this example from India (right), are used to complement or replace natural materials.
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branches stuck into the substrates have been used in estuaries of several African countries, Sri Lanka, and Mexico. In the Caribbean basin, casitas of wood logs provide benthic shelters from which spiny lobsters (Panulirus argus) are harvested. Indigenous knowledge of local fisherfolk is important in sustainably managing these reef systems, even when in their initial stages they may function principally to aggregate larger mobile fishes, such as around bundles of mangrove branches, pesqueros, in Cuba. In recent decades, some artisanal fishing communities have deployed newer designs of artificial reefs, and at larger scales. In part, this has been in response to damage of habitat and fisheries by coastal land-use practices and more intense fishing practices, such as trawling. In India, for example, a national effort has deployed reef modules of steel plates, to increase the time fishing cooperatives actually devote to harvest as a means of promoting social and economic well-being. The coastal waters of Thailand received concrete modules as part of a plan to balance small- and large-scale fisheries, whereby areas of 14– 22 km2 received between 2400 and 3300 units. The most extensive deployment of artificial reefs, for any purpose, is in Japan. There, fishing enhancement is the goal of a national plan begun in 1952 and since greatly expanded. By 1989, 9% of the coastal shelf (o200 m depth) had been affected by reef development. Early emphasis was on engineering and construction of reefs to withstand rigorous high seas environmental conditions. Industrial manufacturers have used materials such as steel, fiberglass, and concrete in structures that are among the largest units in the world, attaining heights of 9 m, widths of 27 m, and volumes of 3600 m3 (Figure 2). Research at sea has been augmented by laboratory studies in which scale models of reefs have been analyzed for effects including deflection of
Figure 2 The largest reef modules have been fabricated and deployed in Japan.
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bottom water currents to create nutrient upwellings for enhancing primary production. Observations of fish behavior led to a classification of the affinity of fishes to associate with physical structure. The Japanese aims have been to develop nursery, reproduction, and fishing grounds to supply seafood including abalone, clams, sea urchins, and finfishes. This partially was in response to the closure of distant waters to fishing by other nations. In contrast to Japan (and many other nations more recently involved), the United States, while also an early entrant in this field, built thousands of artificial reefs to enhance recreational fishing through numerous, independent small-scale efforts organized by local interest groups or governmental organizations at the state level. Until recently, materials of opportunity predominated, including heavier and more durable materials such as concrete rubble from bridge and building demolition sites. Designed structures are becoming more common, especially as reefs gain acceptance in habitat restoration. One state, Florida, has about half of the nation’s total number of reefs, and in one four-county area annual expenditures by nonresidents and visitors on fishing and diving at dozens of artificial reefs were US $1.7 billion (2001 values). Recreational fishing has also been the focus of widespread reef building in Australia, and lesser efforts in areas such as Northern Ireland and Brazil. Other regions more recently deploying artificial reefs to enhance or simply maintain fisheries to supply human foodstuffs include southern Europe and Southeast Asia. From inception, their approach has been to design, experimentally test, and document reefs. As early as the 1970s, 8 m3, 13 t concrete blocks (and other designs) were deployed in the Adriatic Sea of Italy, forming reefs of volumes up to 13 000 m3 and coverage up to 2.4 ha. They serve multiple, fishery-related objectives that are sought in many nations: shelter for juvenile and adult organisms; reproduction sites; capture nutrients in shellfish biomass; protect fish spawning and nursery areas from illegal trawling; and protect artisanal set fishing gear from illegal trawling damage. Higher catches and profits for small-scale fisheries were the result. This is especially true in Spain, where heavy concrete and steel rod reef designs have been used effectively to protect seagrass beds from illegal trawling (Figure 3). Artificial reefs are being used as part of a larger strategy to implement marine ranching programs. In Korea, for example, the decline of distant water fisheries led to a transition from capture-based to culturebased fisheries. From 1971 to 1999, artificial reef placements have affected 143 000 ha of submerged
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Figure 3 Concrete reef modules with projecting steel beams such as railroad tracks are deployed in protected marine areas of Spain to discourage illegal trawling, so as to promote restoration of seagrass habitats and fish populations. Photograph Courtesy of Juan Goutayer.
coastal areas, out of the total 307 000 ha planned. At certain reef sites, hatchery-reared juvenile invertebrates and fishes have been released, and enhanced fishery yield reported, as part of long-term experiments. Artificial reefs are also used exclusively for aquaculture, such as in settings where excess nutrients stimulate primary production that is transferred into biomass of harvestable species, such as mussels (Mytilus galloprovincialis) and oysters (Ostrea edulis and Crassostrea gigas) in Italy. Restoration of marine ecosystems using artificial reefs has focused on plant and coral communities. On the Pacific coast of the United States, for example, a project to mitigate for electric powerplant impacts is developing a 61 ha artificial reef for kelp (Macrocystis pyrifera) colonization, survival, and growth. Along the coast of the Northwest Mediterranean Sea, artificial reefs are deployed in protected areas to preserve and promote colonization of seagrasses (Posidonia). In numerous tropical areas, artificial reefs have been used in repairs to damaged coral reefs (Figure 4), or to hasten replacement of dead or removed coral. In the Philippines, for example, hollow concrete cubes are deployed as sites for coral colonization, sometimes being located in marine reserves. Increasing popularity of artificial reefs to promote tourism is seen in the development of diving opportunities, both for people in submersible vehicles and for scuba divers. In places such as Mexico, the Bahamas, Monaco, and Hawaii, submarine operators provide trips to artificial reefs. On a larger scale, obsolete ships have been sunk to create recreational dive destinations in areas as diverse as British Columbia, Canada, Oman, New Zealand, Mauritius, and Israel. Purposes include generation of economic revenues in
Figure 4 This larval enclosure tent is for settlement of the coral, Montastraea faveolata, at the site of a ship grounding on Molasses Reef, Florida Keys National Marine Sanctuary. Coral are mass-spawned, larvae cultured through the planktonic phase, and then introduced into the tent. In the laboratory, larvae may settle onto rubble pieces which subsequently are cemented to the reef. Photograph by D. Paul Brown.
local communities and diversion of diver pressure from natural reefs. As strictly works of engineering, structures such as rock jetties for shore protection also serve as de facto reefs with biological communities. Submerged breakwaters have been used to create waves for recreational surfboarding.
Progress toward Scientific Understanding Advances in research on artificial reefs include methodologies, subjects and rigor of approaches. Methodological advances include improved underwater biological census techniques, and the application of remote sensing. Investigations have become more ecological process-oriented in explaining biological phenomena such as sheltering, reproduction, recruitment, and feeding, better able to explain physical and hydrodynamic behavior of reefs, and have led to specifying reef designs consistent with the ecology of organisms. Newer approaches include an expansion of hypothesis-driven and experimental research, scaling up of pilot designs into larger reefs, ecological modeling and forecasting, and interdisciplinary studies with various combinations of physical, social, life, and mathematical sciences. Comparative studies of natural and artificial habitats have advanced understanding of both systems. The earlier history of research on artificial reefs may be observed in the subjects, methods, and findings reported at eight international conferences held from
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1974 to 2005. Subsequent to a peak number of peerreviewed articles contained in a proceedings from the 1991 conference (84), fewer resulted from the latest conferences (1999 (56) and 2005 (20)). Scientists studying artificial reefs increasingly are reporting in journals devoted to fish biology, fisheries, and marine ecology. Colonization of artificial reefs has the oldest and most extensive record of scientific inquiry. The appearance of large fishes at some reef sites almost immediately after placement of structure is the most visible form of colonization (and also a factor in one of the most controversial artificial reef issues, discussed below). However, microbes, plants, and invertebrates also colonize reefs, and patterns of succession occur that converge toward naturally occurring assemblages of species (Figure 5). Augmenting earlier studies describing diversity and abundance of artificial reef flora and fauna, later research addressed the influence of environmental variables upon reef ecology, interactions of individual species with the artificial reef structure and biota, and long-term life history studies of selected species. Comparison of patch reefs of stacked concrete blocks with translocated coral reefs in the Bahamas (Figure 6) determined that the reefs had similar fish species composition, while the natural reefs supported more individuals and species, owing to their greater structural complexity (i.e., variety of hole sizes) and associated forage base. Meanwhile, long-term studies of the spiny lobster experimentally established that artificial reef structures, with dark recessed crevices approximately one body-length deep and multiple entrances, indeed augmented recruitment in a situation of habitat limitation in the Florida Keys, USA. While these situations allowed observation by scuba diving, in deeper waters hydroacoustic and video-recording instruments have been used to document fish species, such as at petroleum production platforms in the North Sea and Gulf of Mexico. Prominent among the scientific issues regarding the efficacy of artificial reefs for fisheries production and management is the ‘attraction-production question’: Do artificial reefs merely attract fishes thereby improving fishing efficiency, or do they contribute to biological production so as to enhance fisheries stocks? While this popularized dichotomy is an oversimplification, and an expectation of one allencompassing answer is unrealistic, this question has had great heuristic value for the advance of reef science and responsible reef development. The origins of this issue derive from early observations and assumptions about artificial reefs. High catch rates and densities of fish at artificial reefs were popularly
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Figure 5 Aspects of attachment surfaces for marine plants and shelter and feeding sites for invertebrates and fishes provided by artificial reefs. Drawing courtesy of S. Riggio.
Figure 6 Experimental modules such as the one pictured on the right from the Bahamas exemplify a global trend toward greater manipulation of artificial reefs as tools for testing of hypotheses concerning ecological processes in the ocean, and comparison with systems such as coral reefs, pictured left. Drawing courtesy of American Fisheries Society.
taken as proof that artificial reefs benefited fisheries stocks, based on assumptions that hard substrata were limiting and reef fish standing stocks were dependent on food webs that have hard substrata as their foundations. However, a few fisheries scientists challenged this reasoning in the 1980s. They noted that before heavy exploitation the existing natural habitat supported abundant reef fishes, presumably at or near carrying capacity. Fishing then reduced populations to some lower level, yet the amount of natural habitat remained the same, still capable of supporting higher population numbers. The fisheries scientists reasoned that with fish stocks substantially below carrying capacity, the amount of hard-bottom habitat could not be the factor limiting population size, so the addition of artificial reefs would not benefit fish stocks. Thus, observed high densities of
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fish and high catch rates at artificial reefs were considered an artifact of fish behavioral preferences, which simply concentrated them at reef sites and intensified fishing mortality, what conservationists today call an ecological trap. The fisheries scientists’ formal argument was more thorough and complex than the popularized version, as they recognized a continuum from solely attraction, with little if any direct biological benefit, to potentially high levels of added biological production from each unit of additional reef structure. Species fit along that continuum according to their life history characteristics and use of reef habitat ecologically. Thus small, highly site-attached fishes (e.g., blennies and gobies) that derive all resource needs (i.e., food, shelter, and mates) directly from the occupied reef structure, and that complete all but the planktonic phase of their life cycle in one place, would have the greatest net production potential from artificial reef development. Conversely, large, transient reef-visiting fishes (e.g., jacks and mackerels) were expected to have the least production potential from artificial reef development. In between these extremes were the majority of exploited reef fishes for which one would expect some combination of attraction and production affected by reef ecological setting and overall stock abundance. (Implicit in the argument by the fisheries scientists was the assumption that fishing pressure would be intensely focused at the artificial reefs, which has more to do with the management of human activities than the intrinsic ecological characteristics of the reefs themselves.) Differing professional positions concerning attraction-production still remain, and the debate continues to stimulate research questions and artificial reef management issues. In part, this is because the complexity of reef ecological functions and the effects of spatially explicit fishing mortality are beyond simple answers to the seemingly dichotomous question that has been popularized among lay audiences. And in part, this is because researchers necessarily address this issue from their own disciplinary perspectives and study systems, generally not integrating the multiple levels of biological organization or spatial-temporal domains over which this question can be legitimately addressed. Pertinent studies can focus on life history stages of populations, communities, or ecosystems and encompass individual reefs, broad landscapes, or geographic ranges of fisheries stocks, while covering seasonal, interannual, or long-term time frames. Practical solutions to the fisheries management implications will require more specific attention to the desired species and assemblages, and the processes that sustain them.
Now a pivotal question is how the manipulation of habitat affects demographic rates and exploitation rates across spatial and temporal scales. An example of a way to integrate large sets of ecological data to analyze different scenarios of artificial reef development is through ecosystem simulations such as for Hong Kong where fishing levels in and out of protected areas have been predicted as part of a fishery recovery program.
Advances in Planning, Design, and Construction The increase of artificial reefs in the world’s coastal seas, as measured by the growing number, cost, size, intricacy, and footprint of structures deployed, has prompted an emphasis on their planning. This is to promote more efficient and cost-effective structures, enable work at larger scales and with more precision, satisfy regulatory requirements, reduce conflicts with other natural and human aspects of ecosystems, and minimize the prospects of unintended consequences from improperly constructed reefs. The scientific basis for reefs has been strengthened by research efforts in dozens of countries, coupled with the practical experiences of numerous individuals and groups that have worked independently to build, manage, and, increasingly, evaluate reefs. Independently in different geographic areas, common approaches and some published guidelines for developing artificial reefs have emerged. The earliest handbooks focused on physical and oceanographic conditions as first encountered by the Coastal Fishing Ground Enhancement and Development Project of Japan. (Translations of certain documents into English in the 1980s by the United States National Marine Fisheries Service disseminated information.) The United States produced a national plan in 1984 (in revision in 2006) which spurred promulgation of plans among the coastal states. In some nations, planning is done through a coordinated reef program run by a federal or provincial marine or fishery resources organization, as in Italy or China. Finally, where reef deployment is directed to selected localities or areas, but not nationwide, master plans or other guidelines for siting, materials, and other aspects have been produced, such as for the Aegean Sea of Turkey and northern Taiwan. With a legitimate intent assured, the initial component of reef planning is to frame measurable objectives, define expectations for success, and forecast interactions of that reef with the ecosystem. Subsequently, a valid design and site plan for the reef and procedures for its construction must be developed
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and meet regulatory requirements. The context for management of the reef must be defined. For example, an artificial reef program in Hong Kong was based on an extensive consultation with community interests, leading to stakeholder support for regulatory practices in the local fisheries and protection of marine reserve areas containing some of the reefs. Finally, protocols for evaluation of reef performance and management of assessment data must be established, with communication of results to all interested parties. Construction is a three-phase process: fabrication, transportation, and placement. Representative costs of larger reefs include US $16 000 000 for a 56 ha reef in California, USA, and between US $38 and $57 million annually during 1995–99 for a national program in Korea. The ecological and economic impacts of poorly designed artificial reefs are exemplified by an effort in the United States – projected to cost US $5 000 000 – to remove 2 000 000 automobile tires that had been deployed in the 1970s, but which were drifting onto adjacent live coral reefs and beaches. The design phase of reef development has changed dramatically due to growing emphasis on making the structural attributes of intentional, fabricated reef materials conform and contribute to the biological life history requirements of organisms that are particularly desired as part of the reef assemblage. Reef design should be dictated by: (1) a set of measurable and justified objectives for the reef, (2) the ecology of species of concern associated with the reef, and (3) predicted and understandable environmental and socioeconomic consequences of introduction of the reef into the aquatic ecosystem. This requires a multidisciplinary approach using expertise of biologists, engineers, economists, planners, sociologists, and others. From a physical standpoint, key aspects of design and construction include stability of a reef site, durability of the reef configuration, and potential adverse impacts in the environment. Also, the use of physical processes to enhance reef performance is a desirable aspect. To build a quantitative understanding of the environment into which a reef is to be placed, large-scale oceanographic processes and local conditions must be determined by site surveys early in reef planning, including water circulation driven by tides, wind and baroclinic/density fields, locally generated wind waves, swells propagated from distance, sediment/substrate composition, distribution and transport, and depth. These factors are then coupled with the attributes of the reef material, such as weight, density, dimensions, and strength, in order to forecast reef physical performance.
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From an ecological standpoint, abiotic and biotic influences considered in design include geographic location, type and quality of substrate surrounding the reef site, isolation, depth, currents, seasonality, temperature regime, salinity, turbidity, nutrients, and productivity. Substrate attributes affecting reef ecology include its composition and surface texture, shape, height, profile, surface area, volume and hole size, which taken together contribute to the structural complexity of the reef. Spacing of individual and groups of reef structure is important. A plethora of designs for artificial reefs exist (Figure 7). For a mitigation reef aiming to create new kelp beds off California, USA, biologists and engineers concluded that the most effective design should place boulders and concrete rubble in low-relief piles (o1 m height), at depths of 12–14.5 m on a sand layer of 30–50 cm overlaying hard substrate; success criteria include (1) support of four adult plants per 100 m2 and (2) invertebrate and fish populations similar to natural reefs. As part of a system of artificial reefs in Portugal, experiments quantified production of sessile invertebrates (upon which fishes could feed) according to location on different facets of cubic settlement structures placed on reefs. For fishes, meanwhile, one of five designs used in Korea is the box reef, a 3 3 3 m concrete cube (Figure 8), targeted to two species: small, dark spaces in the lower two-thirds of the reef are provided for rockfish (Sebastes schlegeli), while the upper third is more open to satisfy behavioral preferences of porgy (Pagrus major). Repeatedly, authorities cite complexity of structure as a primary factor in design. As large individual and sets of reefs are planned, more organizations are using pilot projects to determine physical, biological, economic, and even
Figure 7 Artificial reefs of concrete modules are used worldwide, with designs intended to meet ecological requirements of designated species and habitats.
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Figure 8 The box reef design used in Korea resulted from a collaboration of engineers and biologists. Modules have been placed in research plots within Tongyong Marine Reserve for use in experiments with stocking juvenile fishes. Photograph by Kim Chang-gil.
political feasibility. The 42 000 t Loch Linne Artificial Reef in Scotland was started in 2001 as a platform for scientific investigation of the performance of different structures, ultimately to establish fisheries for target species, specifically lobster (Homarus gammarus). This effort also is representative of a trend toward increased predeployment research – in this case a 4-year study of seabed, water-column, and biological parameters – intended to enable better forecasting of reef impacts and measurement of results. A special case of working with reef materials concerns the deployment of obsolete ships and petroleum/gas production platforms, due to the need to handle, prepare, and place them in environmentally compatible ways. In Canada, for example, decommissioned naval vessels require extensive removal of electrical wiring and other components to eliminate release of pollutants into the sea, in conformance with strict federal rules. In the northern Gulf of Mexico (where over 4000 platforms provide a considerable area of hard surface for sessile organism attachment), eastern Pacific, North Sea, and Adriatic Sea offshore platforms either act as de facto reefs or in a limited number of cases are being toppled in place or transported to new locations to serve as dedicated reefs, which costs less than removal to land.
have led to more realistic expectations bolstered by two decades of scientific advances. National plans (e.g., Japan, Korea, and United States), regional programs (e.g., Hong Kong, Singapore, and Turkey), and large-scale pilot projects (e.g., Loch Linne, Scotland; San Onofre, USA) concerning artificial reefs are better defining their role in ecosystem and fishery management. International scientific bodies such as the North Pacific Marine Sciences Organization (PICES) have addressed the relevance of artificial reefs to core fisheries issues, including stock enhancement, fishing regulations, and conservation. Integrated coastal management in the Philippines, India, Spain, and elsewhere now includes artificial reefs in multifaceted responses to issues of habitat destruction, fishery decline, and socioeconomic development. Finally, private consultants and businesses have found markets for their services and products, with one patented design being deployed in over 40 countries, and ecological engineers have recognized reefs as constructed ecosystems for use in restoration ecology. The evaluation of reef performance is fostering an increased acceptance and utilization of artificial reef technology by resource management organizations. Quantitative evaluation is increasingly driven by agency concerns for demonstrating positive returns on their investments, and the increasing scale of artificial reef projects. Evaluations vary in intensity and complexity, ranging from descriptive studies of short duration (e.g., pre- and postdeployment) to extensive studies of ecological processes and dynamics, to the synthesis of complex databases through quantitative modeling. One response of the scientific community was through formation of the European Artificial Reef Research Network, which promulgated priorities and protocols for research across its membership. A significant trend is for artificial reef projects to be planned and evaluated in an adaptive management framework, in which expectations are more explicitly stated, the projects implemented and rigorously evaluated, and then adjustments made to the management practices based on findings from the evaluations. When applied consistently, this cycle will continue to evolve the application of artificial reef technologies toward an ever-increasing standard of practical effectiveness.
Integration of Reefs in Ecological and See also Human Systems Increasingly, artificial reef technology is being applied globally in fisheries and ecosystem management. Early disappointments and healthy skepticism
Coastal Zone Management. Cold-Water Coral Reefs. Coral Reef and Other Tropical Fisheries. Coral Reef Fishes. Coral Reefs. Fiordic Ecosystems. Fisheries: Multispecies Dynamics. Fisheries
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Overview. Fishery Management, Human Dimension. Fishery Manipulation through Stock Enhancement or Restoration. Fishing Methods and Fishing Fleets. Large Marine Ecosystems. Mariculture Diseases and Health. Mariculture, Economic and Social Impacts. Mariculture of Aquarium Fishes. Mariculture of Mediterranean Species. Mariculture Overview. Ocean Gyre Ecosystems. Upwelling Ecosystems.
Further Reading American Fisheries Society (1997) Special Issue on Artificial Reef Management. Fisheries 22: 4--36. International Council for the Exploration of the Sea (2002) ICES Journal of Marine Science 59(supplement S15363).
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Jensen AC, Collins KJ, and Lockwood APM (eds.) (2000) Artificial Reefs in European Seas. Dordrecht: Kluwer. Lindberg WJ, Frazer TK, Portier KM, et al. (2006) Density-dependent habitat selection and performance by a large mobile reef fish. Ecological Applications 16(2): 731--746. Love MS, Schroeder DM, and Nishimoto MM (2003) The ecological role of oil and gas production platforms and natural outcrops on fishes in southern and central California: A synthesis of information. OCS Study MMS 2003-032. Seattle. WA: US Department of the Interior, US Geological Survey, Biological Resources Division. Seaman W (ed.) (2000) Artificial Reef Evaluation. Boca Raton, FL: CRC Press. Svane I and Petersen JK (2001) On the problems of epibioses, fouling and artificial reefs: A review. Marine Ecology 22: 169--188.
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ATLANTIC OCEAN EQUATORIAL CURRENTS 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 1, pp 188–191, & 2001, Elsevier Ltd.
because they provide invaluable checks on the theories and models that explain and simulate oceanic currents. Those currents play a central role in the Earth’s climate, by influencing sea surface temperature patterns for example.
Introduction
Time-averaged Currents
The circulations of the tropical Atlantic and Pacific Oceans have much in common because similar trade winds, with similar seasonal fluctuations, prevail over both oceans. The salient features of these circulations are alternating bands of eastward- and westward-flowing currents in the surface layers (see Figure 1). Fluctuations of the currents in the two oceans have similarities not only on seasonal but even on interannual timescales; the Atlantic has a phenomenon that is the counterpart of El Nin˜o in the Pacific. The two oceans also have significant differences. The Atlantic, but not the Pacific, has a net transport of heat from the southern into the northern hemisphere, mainly because of an intense, crossequatorial coastal current in the Atlantic, the North Brazil Current. The similarities and differences between the tropical Atlantic and Pacific (and also the Indian Ocean) are of enormous interest to modelers
Although the trade winds that prevail over the tropical Atlantic Ocean have a westward component, the currents driven by those winds include the eastward North Equatorial Countercurrent, between the latitudes 31 and 101N approximately. Sverdrup, in one of the early triumphs of dynamical oceanography, first pointed out that this current is attributable to the curl of the wind. Flanking this eastward current are westward currents to its north, the North Equatorial Current, and to its south, the South Equatorial Current. The latter current is particularly intense at the equator, where it can attain speeds in excess of 1 m s 1. Figure 1, a schematic map of the various currents, actually depicts conditions between July and September when the south-east trades are particularly intense and penetrate into the northern hemisphere. Centered on the equator, and below the westward surface flow, is an intense eastward jet known as the
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Figure 1 Schematic map showing the major surface currents of the tropical Atlantic Ocean between July and September when the North Equatorial Countercurrent (NECC) flows eastward into the Guinea Current in the Gulf of Guinea. From January to May the North Equatorial Countercurrent disappears and the surface flow is westward everywhere in the western tropical Atlantic.
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Seasonal Variations of the Currents The seasonal variations of the winds are associated with the north–south movements of the Intertropical Convergence Zone (ITCZ), the band of cloudiness
and heavy rains where the south-east and north-east trades meet. The south-east trades are most intense and penetrate into the northern hemisphere during the northern summer when the ITCZ is between 101 and 151N. During those months the surface currents are particularly strong. The North Brazil Current, after crossing the equator, veers sharply eastward to feed the North Equatorial Countercurrent. The Equatorial Undercurrent is also strongest during this season when the east–west slope of the equatorial thermocline is at a maximum. During the summer of the southern hemisphere, the zone where the north-east and south-east trades meet (the ITCZ) shifts equatorward so that the winds are relaxed along the equator. The North Brazil Current no longer veers offshore after crossing the equator, but continues to flow along the coast into the Gulf of Mexico. It is fed by surface flow that is westward at practically all latitudes in the tropics because, during this season, the eastward North Equatorial Countercurrent disappears from the surface layers, as is evident in Figure 2. At this time, the northward heat transport across 101N is huge – on the order of a peta-watt; during the northern summer it is practically zero.
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Equatorial Undercurrent which amounts to a narrow ribbon that precisely marks the location of the equator. The undercurrent attains speeds on the order of 1 m s 1 has a half-width of approximately 100 km; its core, in the thermocline, is at a depth of approximately 100 m in the west, and shoals towards the east. The current exists because the westward trade winds, in addition to driving divergent westward surface flow (upwelling is most intense at the equator), also maintain an eastward pressure force by piling up the warm surface waters in the western side of the ocean basin. That pressure force is associated with equatorward flow in the thermocline because of the Coriolis force. At the equator, where the Coriolis force vanishes, the pressure force is the source of momentum for the eastward Equatorial Undercurrent which, in a downstream direction, continually loses water because of intense equatorial upwelling which sustains the divergent, poleward Ekman flow in the surface layers. Along the African coast, cold equatorward coastal currents, the Canary Current off north-west Africa, and the Benguela Current off south-west Africa, are driven by the components of the winds parallel to the coast. These currents, which are associated with intense coastal upwelling and low sea surface temperatures, feed the westward North and South Equatorial Currents respectively. Along the coast of South America, the most prominent current is the North Brazil Current, which carries very warm water from about 51N across the equator. Some of that water feeds the Equatorial Undercurrent, but much of it continues into the northern hemisphere. Further south along the coast of Brazil, the flow is southward. The net north–south circulation associated with the various currents is a northward flow of warm surface waters, and a southward return flow of cold water at depth, resulting in a transport of heat from the southern into the northern Atlantic. The southward flow below the thermocline is part of the global thermohaline circulation, which involves the sinking of cold, saline waters in the northern Atlantic. The absence of such formation of deep water in the northern Pacific – that ocean is less saline than the northern Atlantic – is part of the reason why there is a northward transport of heat across the equator in the Atlantic but not the Pacific.
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Figure 2 The seasonal disappearance of the North Equatorial Countercurrent from the western tropical Atlantic. The eastward velocity in cm s 1 (negative values correspond to westward flow) is shown as a function of latitude and month, starting in January. The data, which have been averaged over a band of longitudes in the western equatorial Atlantic from 231W to 331W, are from shipdrift records.
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ATLANTIC OCEAN EQUATORIAL CURRENTS
The upwelling along the west African coast, and the coastal currents too, are subject to large seasonal fluctuations in response to the variations in the local winds. Thus upwelling is most intense off south-western Africa, and surface temperatures there are at a minimum, during the late northern summer when the local alongshore winds are most intense. Off northwestern Africa the season for such conditions is the late northern winter. The northern coast of the Gulf of Guinea (along 51N approximately) also has seasonal upwelling, with lowest temperatures during the northern summer, even though the local winds along that coast have almost no seasonal cycle. In that region, changes in the depth of the thermocline (which separates warm surface waters from the cold water at depth) depend on winds everywhere in the equatorial Atlantic, even the winds off Brazil which are most intense during the northern summer when they cause a shoaling of the thermocline throughout the Gulf of Guinea. If the winds over the ocean were suddenly to stop blowing, how long would it be before the currents in Figure 1 disappear? The answer to this question (which is the same as asking how long it would take for the currents to be generated from a state of rest) is of central importance in climate studies because, associated with the currents, are sea surface temperature patterns that profoundly affect climate. (From a strictly atmospheric perspective, the cause of El Nin˜o is a change in the surface temperature pattern of the tropical Pacific.) The Indian Ocean is ideal for studying these matters because there the abrupt onset of the south-west monsoons in May quickly generates the intense Somali Current along the eastern coast of Africa. Theoretical studies indicate that the generation of such currents, and more generally the adjustment of the ocean to a change in the winds, depend critically on waves (known as Rossby waves) that propagate across the ocean basin along the thermocline. The speed of those waves increases with decreasing latitude, reaching a maximum at the equator, which serves as a guide for the fastest waves – there they travel westward at about 50 cm s 1. The equator is also a guide for a very rapid eastward traveling wave, a Kelvin wave with a speed on the order of 150 cm s 1. The Somali Current near the equator can therefore be generated far more rapidly than can the Gulf Stream in mid-latitudes. The time it takes for the ocean to adjust (for the currents to be generated) depends not only on the speed of certain oceanic waves, but also on the width of the ocean basin. Hence it takes longer to generate the Kuroshio Current in the very wide Pacific, than the Gulf Stream in the smaller Atlantic. If the winds change gradually rather than abruptly, then the timescale of the gradual changes relative to
the time it takes the ocean to adjust determines the nature of the oceanic response. Thus winds that fluctuate on a timescale much longer than the adjustment time of the ocean will force an equilibrium response in which the ocean, at any given time, is in equilibrium with the winds at that time. (The currents and winds fluctuate essentially in phase.) From results such as these it can be inferred that the seasonally varying trade winds over the tropical Atlantic and Pacific Oceans should force an equilibrium response near the equator in the case of the small ocean basin, the Atlantic, but not in the case of the much larger Pacific. The measurements confirm this theoretical result: the seasonal variations of the currents and of the thermocline slope are in phase with the variations of the winds in the equatorial Atlantic, but not in the equatorial Pacific.
Interannual Variations Given the similarities between the climates of the tropical Atlantic and Pacific – arid, cool conditions on the eastern sides, along the shores of Peru and south-western Africa, and warm moist conditions on the western sides – it should come as no surprise that the climate fluctuation known as El Nin˜o has an Atlantic counterpart. As in the Pacific, such events involve a relaxation of the trades so that the warm waters that are usually confined to the western side of the basin flow eastward, causing a rise in sea surface temperatures off the south-west African coast where rainfall can increase significantly. To attribute this phenomenon to a relaxation of the trades is of course an oceanographic perspective. From a meteorological point of view, the warming of the eastern tropical Atlantic is the reason for the weakening of the winds and for several other changes in atmospheric conditions. This circular argument – changes in sea surface temperature are both the cause and consequence of changes in the winds – implies that interactions between the ocean and atmosphere are at the heart of the matter. Those interactions give rise to a variety of natural modes of oscillation which, in the Pacific, appear to be neutrally stable so that random atmospheric disturbances are able to sustain a continual oscillation, the Southern Oscillation, with a distinctive timescale on the order of 4 years. In the Atlantic the possible natural modes appear to be strongly damped and hence are far more sporadic than in the Pacific; there is no distinctive timescale for interannual fluctuations in the Atlantic. The main reason for this difference is the modest dimensions of the Atlantic relative to those of the Pacific. Some of the natural modes attributable to
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ATLANTIC OCEAN EQUATORIAL CURRENTS
ocean–atmosphere interactions depend on the delayed response of the ocean to changes in the winds. If that delay is small, that is the case in an ocean basin of modest size – then the natural modes tend to be damped. Another factor that can inhibit interannual fluctuations is a particularly strong seasonal cycle. That cycle has a larger amplitude in the equatorial Atlantic than Pacific, because the influence of continents on the seasonal changes in the winds can exceed those of ocean–atmosphere interactions in a basin of small dimensions. For a damped mode of oscillation to appear, a suitable perturbation is necessary. The occurrence of El Nin˜o in the Pacific provides such a perturbation in the Atlantic by causing an intensification of the trade winds, and unusually low surface temperatures, in the Atlantic. (This is the impact of the presence of deep atmospheric convection over the eastern tropical Pacific during El Nin˜o.) Apparently El Nin˜o in the Pacific can amount to a preconditioning of the Atlantic because, on several occasions, El Nin˜o in the Pacific was followed a year later by a similar phenomenon in the Atlantic. The amplitude of El Nin˜o is generally much larger in the Pacific than Atlantic – the reason why the Pacific but not the Atlantic phenomenon is capable of a global impact. El Nin˜o, in the Atlantic and Pacific, has a structure that is essentially symmetrical about the equator. The Atlantic has an additional climate fluctuation that is anti-symmetrical relative to the equator, with sea surface temperatures that are high on one side of that line, low on the other side. The cross-equatorial winds then blow towards the warm side with exceptional intensity. If the higher temperatures are to the north, then the zonal band of heavy rains, the ITCZ, persists in a northerly position, bringing drought to north-eastern Brazil, and good rains to the Sahel, the region to the south of the Sahara desert in west Africa. The reverse happens when the ocean temperatures are high south of the equator, cool to the north.
Stability of the Currents During the northern summer, the currents in the western equatorial Atlantic are so intense that they become unstable. One important factor is the enormous latitudinal shear between the eastward North Equatorial Countercurrent and the adjacent westward South Equatorial Current. The instabilities
237
result in meanders that drift westward at a speed near 50 cm s 1, that have a wavelength on the order of a 1000 km, and a period of approximately 1 month. The unstable conditions are confined to the western equatorial region where there is room for two or three waves at most – they sometimes appear in satellite photographs of sea surface temperature. The waves persist for a few months at most so that approximately three oscillations appear during the summer. The counterparts of these unstable waves in the eastern equatorial Pacific have a shorter period (close to 3 weeks) than in the Atlantic, cover a much larger region, and persist far longer. In the Pacific it is possible to observe very long wave trains – they can extend from the Galapagos Islands in the east to the dateline – that persist for many months.
See also Brazil and Falklands (Malvinas) Currents. Current Systems in the Atlantic Ocean. Coastal Trapped Waves. El Nin˜o Southern Oscillation (ENSO). El Nin˜o Southern Oscillation (ENSO) Models. Florida Current, Gulf Stream and Labrador Current. Rossby Waves. Satellite Remote Sensing of Sea Surface Temperatures.
Further Reading The Journal of Geophysical Research Volume 103 (1998) is devoted to a series of excellent and detailed review articles on tropical ocean–atmosphere interactions, including an article on oceanic currents. Carton J and Huang B (1994) Warm events in the tropical Atlantic. Journal of Physical Oceanography 24: 888--903. Chang P, Ji L, and Li H (1997) A decadal climate variation in the tropical Atlantic ocean from thermodynamic air– sea interaction. Nature 385: 516--518. Merle J, Fieux M, and Hisard P (1980) Annual signal and interannual anomalies of sea surface temperature in the eastern equatorial Atlantic. Deep Sea Research 26: 77--101. Philander SGH (1990) El Nin˜o, La Nin˜a and the Southern Oscillation. New York: Academic Press. Richardson PL and Walsh DW (1986) Mapping climatological seasonal variations of surface currents in the tropical Atlantic using ship drifts. Journal of Geophysical Research 91: 10537--10550.
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ATMOSPHERIC INPUT OF POLLUTANTS R. A. Duce, 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 1, pp 192–201, & 2001, Elsevier Ltd.
Introduction For about a century oceanographers have tried to understand the budgets and processes associated with both natural and human-derived substances entering the ocean. Much of the early work focused on the most obvious inputs – those carried by rivers and streams. Later studies investigated sewage outfalls, dumping, and other direct input pathways for pollutants. Over the past decade or two, however, it has become apparent that the atmosphere is also not only a significant, but in some cases dominant, pathway by which both natural materials and contaminants are transported from the continents to both the coastal and open oceans. These substances include mineral dust and plant residues, metals, nitrogen compounds from combustion processes and fertilizers, and pesticides and a wide range of other synthetic organic compounds from industrial and domestic sources. Some of these substances carried into the ocean by the atmosphere, such as lead and some chlorinated hydrocarbons, are potentially harmful to marine biological systems. Other substances, such as nitrogen compounds, phosphorus, and iron, are nutrients and may enhance marine productivity. For some substances, such as aluminum and some rare earth elements, the atmospheric input has an important impact on their natural chemical cycle in the sea. In subsequent sections there will be discussions of the input of specific chemicals via the atmosphere to estuarine and coastal waters. This will be followed by considerations of the atmospheric input to open ocean regions and its potential importance. The atmospheric estimates will be compared with the input via other pathways when possible. Note that there are still very large uncertainties in all of the fluxes presented, both those from the atmosphere and those from other sources. Unless otherwise indicated, it should be assumed that the atmospheric input rates have uncertainties ranging from a factor of 2 to 4, sometimes even larger.
238
Estimating Atmospheric Contaminant Deposition Contaminants present as gases in the atmosphere can exchange directly across the air/sea boundary or they may be scavenged by rain and snow. Pollutants present on particles (aerosols) may deposit on the ocean either by direct (dry) deposition or they may also be scavenged by precipitation. The removal of gases and/or particles by rain and snow is termed wet deposition. Direct Deposition of Gases
Actual measurement of the fluxes of gases to a water surface is possible for only a very few chemicals at the present time, although extensive research is underway in this area, and analytical capabilities for fast response measurements of some trace gases are becoming available. Modeling the flux of gaseous compounds to the sea surface or to rain droplets requires a knowledge of the Henry’s law constants and air/sea exchange coefficients as well as atmospheric and oceanic concentrations of the chemicals of interest. For many chemicals this information is not available. Discussions of the details of these processes of air/sea gas exchange can be found in other articles in this volume. Particle Dry Deposition
Reliable methods do not currently exist to measure directly the dry deposition of the full size range of aerosol particles to a water surface. Thus, dry deposition of aerosols is often estimated using the dry deposition velocity, vd. For dry deposition, the flux is then given by: Fd ¼ vd Ca
½1
where Fd is the dry deposition flux (e.g., in g m2 s1), vd is the dry deposition velocity (e.g., in m s1), and Ca is the concentration of the substance on the aerosol particles in the atmosphere (e.g., in g m3). In this formulation vd incorporates all the processes of dry deposition, including diffusion, impaction, and gravitational settling of the particles to a water surface. It is very difficult to parameterize accurately the dry deposition velocity since each of these processes is acting on a particle population, and they are each dependent upon a number of factors, including wind speed, particle size, relative humidity, etc. The following are dry deposition velocities that have been
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ATMOSPHERIC INPUT OF POLLUTANTS
used in some studies of atmospheric deposition of particles to the ocean:
• • •
Submicrometer aerosol particles, 0.001 m s17 a factor of three Supermicrometer crustal particles not associated with sea salt, 0.01 m s17 a factor of three Giant sea-salt particles and materials carried by them, 0.03 m s17 a factor of two
Proper use of eqn [1] requires that information be available on the size distribution of the aerosol particles and the material present in them. Particle and Gas Wet Deposition
The direct measurement of contaminants in precipitation samples is certainly the best approach for determining wet deposition, but problems with rain sampling, contamination, and the natural variability of the concentration of trace substances in precipitation often make representative flux estimates difficult using this approach. Studies have shown that the concentration of a substance in rain is related to the concentration of that substance in the atmosphere. This relationship can be expressed in terms of a scavenging ratio, S: S ¼ Cr r Ca=g 1
½2
where Cr is the concentration of the substance in rain (e.g., in g kg1), r is the density of air (B1.2 kg m3), Ca/g is the aerosol or gas phase concentration in the atmosphere (e.g., in g m3), and S is dimensionless. Values of S for substances present in aerosol particles range from a few hundred to a few thousand, which roughly means that 1 g (or 1 ml) of rain scavenges D1 m3 of air. For aerosols, S is dependent upon such factors as particle size and chemical composition. For gases, S can vary over many orders of magnitude depending on the specific gas, its Henry’s law constant, and its gas/water exchange coefficient. For both aerosols and gases, S is also dependent upon the vertical concentration distribution and vertical extent of the precipitating cloud, so the use of scavenging ratios requires great care, and the results have significant uncertainties. However, if the concentration of an atmospheric substance and its scavenging ratio are known, the scavenging ratio approach can be used to estimate wet deposition fluxes as follows: Fr ¼ P Cr ¼ P S Ca=g r1
½3
where Fr is the wet deposition flux (e.g., in g m2 year1) and P is the precipitation rate (e.g., in m year1), with appropriate conversion factors to translate rainfall
239
depth to mass of water per unit area. Note that P S r1 is equivalent to a wet deposition velocity.
Atmospheric Deposition to Estuaries and the Coastal Ocean Metals
The atmospheric deposition of certain metals to coastal and estuarine regions has been studied more than that for any other chemicals. These metals are generally present on particles in the atmosphere. Chesapeake Bay is among the most thoroughly studied regions in North America in this regard. Table 1 provides a comparison of the atmospheric and riverine deposition of a number of metals to Chesapeake Bay. The atmospheric numbers represent a combination of wet plus dry deposition directly onto the Bay surface. Note that the atmospheric input ranges from as low as 1% of the total input for manganese to as high as 82% for aluminum. With the exception of Al and Fe, which are largely derived from natural weathering processes (e.g., mineral matter or soil), most of the input of the other metals is from human-derived sources. For metals with anthropogenic sources the atmosphere is most important for lead (32%). There have also been a number of investigations of the input of metals to the North Sea, Baltic Sea, and Mediterranean Sea. Some modeling studies of the North Sea considered not only the direct input pathway represented by the figures in Table 1, but also considered Baltic Sea inflow, Atlantic Ocean inflow and outflow, and exchange of metals with the Table 1 Estimates of the riverine and atmospheric input of some metals to Chesapeake Bay Metal
Riverine input (106 g year1)
Atmospheric input (106 g year1)
% Atmospheric input
Aluminum Iron Manganese Zinc Copper Nickel Lead Chromium Arsenic Cadmium
160 600 1300 50 59 100 15 15 5 2.6
700 400 13 18 3.5 4 7 1.5 0.8 0.4
81 40 1 26 6 4 32 10 14 13
Data reproduced with permission from Scudlark JR, Conko KM and Church TM (1994) Atmospheric wet desposition of trace elements to Chesapezke Bay: (CBAD) study year 1 results. Atmospheric Environment 28: 1487–1498.
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ATMOSPHERIC INPUT OF POLLUTANTS
sediments, as well as the atmospheric contribution to all of these inputs. Figure 1 shows schematically some modeling results for lead, copper, and cadmium. Note that for copper, atmospheric input is relatively unimportant in this larger context, while atmospheric input is somewhat more important for cadmium, and it is quite important for lead, being approximately equal to the inflow from the Atlantic Ocean, although still less than that entering the North Sea from dumping. As regards lead, note that approximately 20% of the inflow from the Atlantic to the North Sea is also derived from the atmosphere. This type of approach gives perhaps the most accurate and in-depth analysis of the importance of
atmospheric input relative to all other sources of a chemical in a water mass. Nitrogen Species
The input of nitrogen species from the atmosphere is of particular interest because nitrogen is a necessary nutrient for biological production and growth in the ocean. There has been an increasing number of studies of the atmospheric input of nitrogen to estuaries and the coastal ocean. Perhaps the area most intensively studied is once again Chesapeake Bay. Table 2 shows that approximately 40% of all the nitrogen contributed by human activity to Chesapeake Bay enters via
Direct atmospheric deposition 1172 (1172)
Direct atmospheric deposition 79 (79) Copper Rivers and direct discharge 2230 (15)
Atlantic outflow 6166 (102)
Dumping 1561 (0)
North Sea
Lead Rivers and direct discharge 1597 (288) Dumping 3015 (0)
Atlantic outflow 4122 (1167) North Sea
Baltic inflow 84 (17)
Baltic inflow 454 (5)
Atlantic inflow 1262 (252)
Atlantic inflow 3785 (38) Sedimentation 2543 (35)
Sedimentation 3248 (562)
Erosion 600 (0) Sediments
Erosion 240 (0) Sediments
(Tons/ year)
Direct atmospheric deposition 20 (20)
(Tons/ year)
Direct atmospheric deposition 13.3 (13.3) Cadmium Rivers and direct discharge 115 (6)
Atlantic outflow 461 (49)
Dumping 72 (0)
North Sea
Sedimentation 39 (5)
Rivers and direct discharge 3.9 (3.5)
Atlantic outflow 15.4 (15.3)
Dumping 0.0 (0)
North Sea
Baltic inflow 30 (3)
Baltic inflow 0.2 (0.2)
Atlantic inflow 251 (25)
Atlantic inflow 5.0 (5.0)
Erosion 12 (0) Sediments
Lindane
Sedimentation 7.0 (6.7)
(Tons/ year)
Erosion 0.0 (0) Sediments
(Tons/ year)
Figure 1 Input of copper, lead, cadmium, and lindane to the North Sea. Values in parentheses denote atmospheric contribution. For example, for copper the atmospheric contribution to rivers and direct discharges is 15 tons per year. (Figure reproduced with permission from Duce, 1998. Data adapted with permission from van den Hout, 1994.)
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ATMOSPHERIC INPUT OF POLLUTANTS
Table 2
241
Estimates of the input of nitrogen to Chesapeake Bay
Source
Total input (109g year1)
Areal input rate (g m2year1)
% of the total
Animal waste Fertilizers Point sources Atmospheric precipitation nitrate ammonium Total
5 48 33
0.4 4.2 2.9
3 34 24
35 19 140
3.1 1.7 12.3
25 14 100
Data reproduced with permission from Fisher D, Ceroso T, Mathew T and Oppenheimer M (1988) Polluted Coastal Waters: The Role of Acid Rain. New York: Environmental Defense Fund.
precipitation falling directly on the Bay or its watershed. These studies were different from most earlier studies because the atmospheric contributions were considered not only to be direct deposition on the water surface, but also to include that coming in via the atmosphere but falling on the watershed and then entering the Bay. Note from Table 2 that atmospheric input of nitrogen exceeded that from animal waste, fertilizers, and point sources. In the case of nitrate, about 23% falls directly on the Bay, with the remaining 77% falling on the watershed. These results suggest that studies that consider only the direct deposition on a water surface (e.g., the results shown in Table 1) may significantly underestimate the true contribution of atmospheric input. The total nitrogen fertilizer applied to croplands in the Chesapeake Bay region is B5.4 g m2 year1, while the atmospheric nitrate and ammonium nitrogen entering the Bay is B4.8 g m2 year1. Chesapeake Bay is almost as heavily fertilized from atmospheric nitrogen, largely anthropogenic, as the croplands are by fertilizer in that watershed!
Table 3
Results from studies investigating nitrogen input to some other estuarine and coastal regions are summarized in Table 3. In this table atmospheric sources for nitrogen are compared with all other sources, where possible. The atmospheric input ranges from 10% to almost 70% of the total. Note that some estimates compare only direct atmospheric deposition with all other sources and some include as part of the atmospheric input the portion of the deposition to the watershed that reaches the estuary or coast. Synthetic Organic Compounds
Concern is growing about the input of a wide range of synthetic organic compounds to the coastal ocean. To date there have been relatively few estimates of the atmospheric fluxes of synthetic organic compounds to the ocean, and these estimates have significant uncertainties. These compounds are often both persistent and toxic pollutants, and many have relatively high molecular weights. The calculation of the atmospheric input of these compounds to the
Estimates of the input of nitrogen to some coastal areasa
Region
Total atmospheric inputb (109g year1)
Total input all sources (109g year1)
% Atmospheric input
North Sea Western Mediterranean Sea Baltic Sea Chesapeake Bay New York Bight Long Island Sound Neuse River Estuary, NC
400c 400c 500 54 – 11 1.7
1500 577d B1200 140 – 49 7.5
27c 69c 42 39 13c 22 23
a
Data from several sources in the literature. Total from direct atmospheric deposition and runoff of atmospheric material from the watershed. c Direct atmospheric deposition to the water only. d Total from atmospheric and riverine input only. b
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ATMOSPHERIC INPUT OF POLLUTANTS
Table 4
Estimates of the input of synthetic organic compounds to the North Sea
Organic compound
Atmospheric input (106 g year1)
Input from other sources (106 g year1)
% Atmospheric input
PCB Lindane Polycyclic aromatic hydrocarbons Benzene Trichloroethene Trichloroethane Tetrachloroethene Carbon tetrachloride
40 36 80 400 300 90 100 6
3 3 90 500 80 60 10 40
93 92 47 44 80 94 91 13
Data reproduced with permission from Warmerhoven JP, Duiser JA, de Leu LT and Veldt C (1989) The Contribution of the Input from the Atmosphere to the Contamination of the North Sea and the Dutch Wadden Sea. Delft, The Netherlands: TNO Institute of Environmental Sciences.
coastal ocean is complicated by the fact that many of them are found primarily in the gas phase in the atmosphere, and most of the deposition is related to the wet and dry removal of that phase. The atmospheric residence times of most of these compounds are long compared with those of metals and nitrogen species. Thus the potential source regions for these compounds entering coastal waters can be distant and widely dispersed. Figure 1 shows the input of the pesticide lindane to the North Sea. Note that the atmospheric input of lindane dominates that from all other sources. Table 4 compares the atmospheric input to the North Sea with that of other transport paths for a number of other synthetic organic compounds. In almost every case atmospheric input dominates the other sources combined.
Atmospheric Deposition to the Open Ocean Studies of the atmospheric input of chemicals to the open ocean have also been increasing lately. For many substances a relatively small fraction of the material delivered to estuaries and the coastal zone by rivers and streams makes its way through the near shore environment to open ocean regions. Most of this material is lost via flocculation and sedimentation to the sediments as it passes from the freshwater environment to open sea water. Since aerosol particles in the size range of a few micrometers or less have atmospheric residence times of one to several days, depending upon their size distribution and local precipitation patterns, and most substances of interest in the gas phase have similar or even longer atmospheric residence times, there is ample opportunity
for these atmospheric materials to be carried hundreds to thousands of kilometers before being deposited on the ocean surface. Metals
Table 5 presents estimates of the natural and anthropogenic emission of several metals to the global atmosphere. Note that ranges of estimates and the best estimate are given. It appears from Table 5 that anthropogenic sources dominate for lead, cadmium, and zinc, with essentially equal contributions for copper, nickel, and arsenic. Clearly a significant fraction of the input of these metals from the atmosphere to the ocean could be derived largely from anthropogenic sources. Table 6 provides an estimate of the global input of several metals from the atmosphere to the ocean and compares these fluxes with those from rivers. Estimates are given for both the dissolved and particulate forms of the metals. These estimates suggest that rivers are generally the primary source of particulate Table 5 Metal
Lead Cadmium Zinc Copper Arsenic Nickel
Emissions of some metals to the global atmosphere Anthropogenic emissions (109 g year1)
Natural emissions (109g year1)
Range
Best estimate
Range
Best estimate
289–376 3.1–12 70–194 20–51 12–26 24–87
332 7.6 132 35 18 56
1–23 0.15–2.6 4–86 2.3–54 0.9–23 3–57
12 1.3 45 28 12 30
Data reproduced with permission from Duce et al., 1991.
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ATMOSPHERIC INPUT OF POLLUTANTS
243
Estimates of the input of some metals to the global ocean
Table 6 Metal
Atmospheric input
Iron Copper Nickel Zinc Arsenic Cadmium Lead
Riverine input
Dissolved (109 g year1)
Particulate (109 g year1)
Dissolved (109 g year1)
Particulate (109 g year1)
1600–4800 14–45 8–11 33–170 2.3–5 1.9–3.3 50–100
14 000–42 000 2–7 14–17 11–55 1.3–3 0.4–0.7 6–12
1100 10 11 6 10 0.3 2
110 000 1 500 1 400 3 900 80 15 1 600
Data reproduced with permission from Duce et al., 1991.
Pb concentration in seawater, pmol/kg and _ 10 3 Pb concentration in the atmosphere, 10 g / m
metals in the ocean, although again a significant fraction of this material may not get past the coastal zone. For the dissolved phase atmospheric and riverine inputs are roughly equal for metals such as iron, copper, and nickel; while for zinc, cadmium, and particularly lead atmospheric inputs appear to dominate. These estimates were made based on data collected in the mid-1980s. Extensive efforts to control the release of atmospheric lead, which has been primarily from the combustion of leaded gasoline, are now resulting in considerably lower concentrations of lead in many areas of the open ocean. For example, Figure 2 shows that the concentration of dissolved lead in surface sea water near Bermuda has been decreasing regularly over the past 15–20
years, as has the atmospheric lead concentration in that region. This indicates clearly that at least for very particle-reactive metals such as lead, which has a short lifetime in the ocean (several years), even the open ocean can recover rather rapidly when the anthropogenic input of such metals is reduced or ended. Unfortunately, many of the other metals of most concern have much longer residence times in the ocean (thousands to tens of thousands of years). Figure 3 presents the calculated fluxes of several metals from the atmosphere to the ocean surface and from the ocean to the seafloor in the 1980s in the tropical central North Pacific. Note that for most metals the two fluxes are quite similar, suggesting the potential importance of atmospheric input to the
180 Marine dissolved Pb Atmospheric Pb
160 140 120 100 80 60 40 20 0 1970
1975
1980
1985 Year
1990
1995
2000
Figure 2 Changes in concentration of atmospheric lead at Bermuda and dissolved surface oceanic lead near Bermuda from the mid1970s to the mid-1990s. (Data reproduced with permission from Wu J and Boyle EA (1997) Lead in the western North Atlantic Ocean: Completed response to leaded gasoline phaseout. Geochimica et Cosmochimica Acta 61: 3279–3283; and from Huang S, Arimoto R and Rahn KA (1996) Changes in atmospheric lead and other pollution elements at Bermuda. Journal of Geophysical Research 101: 21 033–21 040.)
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ATMOSPHERIC INPUT OF POLLUTANTS
Air/sea exchange
1200
560
0.67
7.8
8.9
67
4.5
7
Al
Fe
Th
V
Cu
Zn
Se
Pb
4700
2600
0.55
5
10
12
0.007
0.3
Flux to sediments
Figure 3 A comparison of the calculated fluxes of aluminum (Al), iron (Fe), thorium (Th), vanadium (V), copper (Cu), zinc (Zn), selenium (Se), and lead (Pb) (in 109 g cm2 year1) from the atmosphere to the ocean and from the ocean to the sediments in the central tropical North Pacific. For each metal note the relative similarity in the two fluxes, except for lead and selenium. (Reproduced with permission from Duce, 1998.)
marine sedimentation of these metals in this region. Lead and selenium are exceptions, however, as the atmospheric flux is much greater than the flux to the seafloor. The fluxes to the seafloor represent average fluxes over the past several thousand years, whereas the atmospheric fluxes are roughly for the present time. The atmospheric lead flux is apparently much larger than the flux of lead to the sediments, primarily because of the high flux of anthropogenic lead from the atmosphere to the ocean since the introduction of tetraethyllead in gasoline in the 1920s. (The atmospheric flux is much lower now than in the 1980s, as discussed above.) However, in the case of selenium the apparently higher atmospheric flux is an artifact, because most of the flux of selenium from the atmosphere to the ocean is simply marine-derived selenium that has been emitted from the ocean to the atmosphere as gases, such as dimethyl selenide (DMSe). DMSe is oxidized in the atmosphere and returned to the ocean, i.e., the selenium input is simply a recycled marine flux. Thus, care must be taken when making comparisons of this type.
surface waters. Table 7 presents a recent estimate of the current input of fixed nitrogen to the global ocean from rivers, the atmosphere, and nitrogen fixation. From the numbers given it is apparent that all three sources are likely important, and within the uncertainties of the estimates they are roughly equal. In the case of rivers, about half of the nitrogen input is anthropogenic for atmospheric input perhaps the most important information in Table 7 is that the organic nitrogen flux appears to be equal to or perhaps significantly greater than the inorganic (i.e., ammonium and nitrate) nitrogen flux. The source of the organic nitrogen is not known, but there are indications that a large fraction of it is anthropogenic in origin. This is a form of atmospheric nitrogen input to the ocean that had not been considered until very recently, as there had been few measurements of organic nitrogen input to the ocean before the mid1990s. The chemical forms of this organic nitrogen are still largely unknown. Of particular concern are potential changes to the input of atmospheric nitrogen to the open ocean in
Nitrogen Species
Table 7 Estimates of the current input of reactive nitrogen to the global ocean
There is growing concern about the input of anthropogenic nitrogen species to the global ocean. This issue is of particular importance in regions where nitrogen is the limiting nutrient, e.g., the oligotrophic waters of the central oceanic gyres. Estimates to date suggest that in such regions atmospheric nitrogen will in general account for only a few percent of the total ‘new’ nitrogen delivered to the photic zone, with most of the ‘new’ nutrient nitrogen derived from the upwelling of nutrient-rich deeper waters and from nitrogen fixation in the sea. It is recognized, however, that the atmospheric input is highly episodic, and at times it may play a much more important role as a source for nitrogen in
Source
From the atmosphere Dissolved inorganic nitrogen Dissolved organic nitrogen From rivers (dissolved inorganic þ organic nitrogen) Natural Anthropogenic From nitrogen fixation within the ocean
Nitrogen input (1012 g year1)
28–70 28–84
14–35 7–35 14–42
Data reproduced with permission from Cornell S, Rendell A and Jickells T (1995) Atmospheric inputs of dissolved organic nitrogen to the oceans. Nature 376: 243–246.
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ATMOSPHERIC INPUT OF POLLUTANTS
Table 8
245
Estimates of anthropogenic reactive nitrogen production, 1990 and 2020
Region
Energy (NOx)
Fertilizer
1990 2020 (1012g N year1)
D
Factor
% of total increase
1990 2020 (1012g N year1)
D
Factor
% of total increase
USA/Canada Europe Australia Japan
7.6 4.9 0.3 0.8
10.1 5.2 0.4 0.8
2.5 0.3 0.1 0
1.3 1.1 1.3 1.0
10 1 0.4 0
13.3 15.4 — —
14.2 15.4 — —
0.9 0 — —
1.1 1.0 — —
1.6 0 — —
Asia Central/South America Africa Former Soviet Union Total
3.5 1.5 0.7 2.2 21
13.2 5.9 4.2 5.7 45
9.7 4.4 3.5 3.5 24
3.8 3.9 6.0 2.5 2.1
39 18 15 15 100
36 1.8 2.1 10 79
85 4.5 5.2 10 134
49 2.7 3.1 0 55
2.4 2.5 2.5 1.0 1.7
88 5 6 0 100
Data adapted with permission from Galloway et al., 1995.
the future as a result of increasing human activities. The amount of nitrogen fixation (formation of reactive nitrogen) produced from energy sources (primarily as NOx, nitrogen oxides), fertilizers, and legumes in 1990 and in 2020 as a result of human activities as well as the current and predicted future geographic distribution of the atmospheric deposition of reactive nitrogen to the continents and ocean have been evaluated recently. Table 8 presents estimates of the formation of fixed nitrogen from energy use and production and from fertilizers, the two processes which would lead to the most important fluxes of reactive nitrogen to the atmosphere. Note that the most highly developed regions in the world, represented by the first four regions in the table, are predicted to show relatively little increase in the formation of fixed nitrogen, with none of these areas having a predicted increase by 2020 of more than a factor of 1.3 nor a contribution to the overall global increase in reactive nitrogen exceeding 10%. However, the regions in the lower part of Table 8 will probably contribute very significantly to increased anthropogenic reactive nitrogen formation in 2020. For example, it is predicted that the production of reactive nitrogen in Asia from energy sources will increase r fourfold, and that Asia will account for almost 40% of the global increase, while Africa will have a sixfold increase and will account for 15% of the global increase in energy-derived fixed nitrogen. It is predicted that production of reactive nitrogen from the use of fertilizers in Asia will increase by a factor of 2.4, and Asia will account for B88% of the global increase from this source. Since both energy sources (NOx, and ultimately nitrate) and fertilizer (ammonia and nitrate) result in the extensive release of reactive nitrogen to the atmosphere, the predictions above indicate that there
should be very significant increases in the atmospheric deposition to the ocean of nutrient nitrogen species downwind of such regions as Asia, Central and South America, Africa, and the former Soviet Union. This prediction has been supported by numerical modeling studies. These studies have resulted in the generation of maps of the 1980 and expected 2020 annual deposition of reactive nitrogen to the global ocean. Figure 4 shows the expected significant increase in reactive nitrogen deposition from fossil fuel combustion to the ocean to the east of all of Asia, from Southeast Asia to the Asian portion of the former Soviet Union; to the east of South Africa, northeast Africa and the Mideast and Central America and southern South America; and to the west of northwest Africa. This increased reactive nitrogen transport and deposition to the ocean will provide new sources of nutrient nitrogen to some regions of the ocean where biological production is currently nitrogen-limited. There is thus the possibility of significant impacts on regional biological primary production, at least episodically, in these regions of the open ocean. Synthetic Organic Compounds
The atmospheric residence times of many synthetic organic compounds are relatively long compared with those of the metals and nitrogen species, as mentioned previously. Many of these substances are found primarily in the gas phase in the atmosphere, and they are thus very effectively mobilized into the atmosphere during their production and use. Their long atmospheric residence times of weeks to months leads to atmospheric transport that can often be hemispheric or near hemispheric in scale. Thus
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246
ATMOSPHERIC INPUT OF POLLUTANTS
1 1.5 2
3 4
Figure 4 The ratio of the estimated deposition of reactive nitrogen to ocean and land surfaces in 2020 relative to 1980. (Figure reproduced with permission from Watson AJ (1997) Surface Ocean–Lower Atmosphere Study (SOLAS). Global Change Newsletter, IGBP no. 31, 9–12; data in figure adapted with permission from Galloway JN, Levy H and Kasibhatla PS (1994) Year 2020: Consequences of population growth and development on deposition of oxidized nitrogen. Ambio 23: 120–123.).
Table 9
Estimates of the atmospheric input of organochlorine compounds to the global ocean
Ocean
SHCH (106g year1)
SDDT (106g year1)
SPCB (106g year1)
HCB (106g year1)
Dieldrin (106g year1)
Chlordane (106g year1)
North Atlantic South Atlantic North Pacific South Pacific Indian Global input via the atmosphere Global input via rivers % Atmospheric input
850 97 2600 470 700 B4700 B60 B99%
16 14 66 26 43 B170 B4 B98%
100 14 36 29 52 B230 B60 B80%
17 10 20 19 11 B80 B4 B95%
17 2.0 8.9 9.5 6.0 B40 B4 B91%
8.7 1.0 8.3 1.9 2.4 B22 B4 B85%
(Data reproduced with permission from Duce et al., 1991.)
atmospheric transport and deposition in general dominates all other sources for these chemicals in sea water in open ocean regions. Table 9 compares the atmospheric and riverine inputs to the world ocean for a number of synthetic organochlorine compounds. Note that the atmosphere in most cases accounts for 90% or more of the input of these compounds to the ocean. Table 9 also presents estimates of the input of these same organochlorine compounds to the major ocean basins. Since most of these synthetic organic compounds are produced and used in the northern hemisphere, it is not surprising that the flux into the northern hemisphere ocean is greater than that to southern hemisphere marine regions. There are some differences for specific compounds in different ocean basins. For example, HCH (hexachlorocyclohexane) and DDT have a higher input rate to the North Pacific than the North Atlantic, largely because of the greater use of these compounds in Asia than in North America or
Europe. On the other hand, the input of PCBs (polychlorinated biphenyls) and dieldrin is higher to the North Atlantic than the North Pacific, primarily because of their greater use in the continental regions adjacent to the North Atlantic.
Conclusions The atmosphere transports materials to the ocean that are both harmful to marine life and that are essential for marine biological productivity. It is now apparent that atmospheric transport and deposition of some metals, nitrogen species, and synthetic organic compounds can be a significant and in some cases dominant pathway for these substances entering both estuarine and coastal waters as well as some open ocean regions. Atmospheric input clearly must be considered in any evaluation of material fluxes to marine ecosystems. However, the uncertainties in the
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ATMOSPHERIC INPUT OF POLLUTANTS
atmospheric fluxes of these materials to the ocean are large. The primary reasons for these large uncertainties are:
• • • • •
The lack of atmospheric concentration data over vast regions of the coastal and open ocean, particularly over extended periods of time and under varying meteorological conditions; The episodic nature of the atmospheric deposition to the ocean; The lack of accurate models of air/sea exchange, particularly for gases; The inability to measure accurately the dry deposition of particles; and The inability to measure accurately the air/sea exchange of gases.
See also Chlorinated Hydrocarbons. Metal Pollution. Refractory Metals. Transition Metals and Heavy Metal Speciation
Further Reading
247
Marine Meteorology Technical Conference on Marine Pollution, World Meteorological Organization TD-No. 890, Geneva, Switzerland. Duce RA, Liss PS, Merrill JT, et al. (1991) The atmospheric input of trace species to the world ocean. Global Biogeochemical Cycles 5: 193--259. Galloway JN, Schlesinger WH, Levy H, Michaels A, and Schnoor JL (1995) Nitrogen fixation: anthropogenic enhancement – environmental reponse. Global Biogeochemical Cycles 9: 235--252. Jickells TD (1995) Atmospheric inputs of metals and nutrients to the oceans: their magnitude and effects. Marine Chemistry 48: 199--214. Liss PS and Duce RA (1997) The Sea Surface and Global Change. Cambridge: Cambridge University Press. Paerl HW and Whitall DR (1999) Anthropogenicallyderived atmospheric nitrogen deposition, marine eutrophication and harmful algal bloom expansion: Is there a link? Ambio 28: 307--311. Prospero JM, Barrett K, Church T, et al. (1996) Nitrogen dynamics of the North Atlantic Ocean – Atmospheric deposition of nutrients to the North Atlantic Ocean. Biogeochemistry 35: 27--73. van den Hout KD (ed.) (1994) The Impact of Atmospheric Deposition of Non-Acidifying Pollutants on the Quality of European Forest Soils and the North Sea. Report of the ESQUAD Project, IMW-TNO Report No. R 93/ 329.
Duce RA (1998) Atmospheric Input of Pollution to the Oceans, pp. 9–26. Proceedings of the Commission for
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ATMOSPHERIC TRANSPORT AND DEPOSITION OF PARTICULATE MATERIAL TO THE OCEANS J. M. Prospero, University of Miami, Miami, FL, USA R. Arimoto, New Mexico State University, Carlsbad, NM, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction The atmosphere is the primary pathway for the transport of many geochemically important substances to the oceans. Although the magnitude of these wind-borne transports is not accurately known, there is growing evidence that atmospheric deposition significantly impacts chemical and biological processes in the oceans. It is only over the past several decades that marine scientists have come to appreciate the importance of atmospheric transport. Historically it had been assumed that the fluxes of continental materials to the oceans were dominated by rivers. But over time it was recognized that much of the riverine load was deposited in estuaries or on the continental shelves. In contrast, winds can rapidly span great distances to reach even the most remote ocean regions. The transport and deposition of particulate matter (PM) to the oceans depends on many factors including the distribution of sources, physical and chemical properties of the particles, meteorological conditions, and removal mechanisms. Our interest here focuses on particles between about 0.1 and 10 mm in diameter which, because of their small size, have atmospheric lifetimes ranging from days to several weeks. These are commonly referred to as aerosol particles or aerosols. Larger particles are deposited close to their sources and do not contribute substantially to ocean deposition except in some coastal regions. Smaller aerosols carry little mass and, while they are important for other atmospheric issues, they are not particularly important for air/sea chemical exchange. Winds carry billions of tons of PM to the ocean. Some of the PM is emitted by natural processes and some is produced anthropogenically, that is, as a result of human activities. Much PM is emitted directly as primary particles; this includes mineral (soil) dust, organic particles from plants, and emissions from anthropogenic combustion processes (e.g., from industry, homes, and vehicles) and biomass burning,
248
which can be natural (e.g., started by lightning) or anthropogenic (e.g., in clearing land, burning agricultural waste). An important PM fraction – secondary PM – is that produced from gases, natural and anthropogenic, that react in the atmosphere to form particles. One goal of marine scientists is to characterize atmospheric transport and chemical deposition to the ocean and to assess the impact of the air/sea exchange process. This is a difficult task which can only be achieved when we know the kinds of materials deposited and their temporal and spatial variability. Because of the patchy distribution of sources and the relatively short tropospheric residence times of aerosols, PM concentrations over the oceans vary by orders of magnitude in time and space. Here we review the sources and composition of aerosols and the removal mechanisms relevant to deposition issues. We then present estimates of deposition rates of some PM classes to the oceans.
Aerosol Sources, Composition, and Concentrations The composition of PM over the oceans varies greatly due to the myriad sources and variations in their strengths. Soils emit fine mineral particles. Plants produce a wide range of organic particles, ranging from decayed leaf matter, to plant waxes, and condensed organic compounds. Volcanoes sporadically inject many tons of material into the atmosphere, and much of this is deposited in the oceans; but the total amount of PM deposited over time is relatively small compared with other sources. Smelters, power plants, and incinerators emit PM highly enriched with trace metal pollutants. Combustion sources, both natural (wild fires, biomass burning) and anthropogenic (power plants, vehicles, home heating) emit thousands of organic compounds. Combustion processes and the use of fertilizers contribute to the production of nitrogen-rich particles. Pesticides and other synthetic organic compounds are emitted from industrial and domestic sources. Typically the dominant marine aerosol species by mass are: (a) sea salt, produced by breaking waves and bursting bubbles; (b) sulfate, including that from sea salt aerosol and non-sea-salt sulfate (nss-SO4 2 ), the latter of which is both transported from pollution sources on the continents and produced from
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ATMOSPHERIC TRANSPORT AND DEPOSITION OF PARTICULATE MATERIAL TO THE OCEANS
gaseous precursors such as dimethyl sulfide (DMS) emitted from the oceans; (c) nitrate, originating from pollution sources on the continents and produced by lightning; (d) ammonium, derived mostly from continental sources but in some areas from ocean sources; (e) mineral dust, from arid lands; (f) organic carbon (OC), largely from anthropogenic and natural sources on the continents; and (g) black carbon (BC), from biomass burning and anthropogenic sources. PM composition is strongly size dependent not only as a result of various production mechanisms but also because size-selective removal occurs during transport. Physical production mechanisms (grinding of rocks, bursting bubbles) normally produce large particles with most of the mass in PM greater than 1-mm diameter (coarse particles). For example, the mass median diameter (MMD: 50% of the mass is greater than the MMD and 50% is less) of dust particles over deserts can be extremely large, many tens or hundreds of micrometers, but over the oceans, it is typically only several micrometers. The MMD of sea salt particles is generally in the range of about 5–10 mm, depending on wind conditions. Other primary particles including soot emitted from smoke stacks, diesel exhaust, and particles shed by plants (e.g., plant waxes, fibers), tending to be in the size range of B0.1–1.0 mm.
Table 1
Gas-phase reactions produce secondary PM ranging in size range of 0.001–0.1 mm diameter. Examples are sulfate particles produced from SO2 emitted from power plants or from the oxidation of DMS emitted from the oceans. Particles in this very fine particle mode are highly mobile. They can rapidly coagulate to form larger particles (typically 0.1–1 mm) or they can diffuse to the surface of cloud or fog droplets or to larger particles (e.g., sea salt, mineral dust). Table 1 presents concentration data for the aerosols that make up most of the PM mass over the oceans; it also includes data for vanadium, which is included as an example of an element strongly affected by pollution sources. The column on the extreme right shows the total aerosol concentration less that of sea salt, so as to better illustrate the impact of transported PM. These data are the product of longterm measurements obtained from a global ocean network. In general, PM concentrations are much higher in the Northern Hemisphere relative to the Southern Hemisphere. Mineral dust shows an extremely wide range of concentrations over the oceans; the highest are over the tropical North Atlantic and the western North Pacific. These reflect the impact of dust transport from North Africa and China, respectively. Dust concentrations in the southern oceans tend to be extremely low due to the
Annual mean aerosol concentrations measured at remote ocean stations Station locationa Lat 1N
North Atlantic Mace Head Bermuda Barbados
Sea salt (mg m 3)
NO3 (mg m 3)
nss-SO4 (mg m 3)
NH4 (mg m 3)
Dust V (mg m 3) (mg m 3)
Total b (mg m 3)
Total SSc (mg m 3)
19.8
4.1
7.2
3.0
15.5
4.1
49.5
29.8
13.8 15.1
0.3 0.4
0.5 0.5
0.8 0.0
0.7 0.7
0.2 0.3
15.4 16.7
1.6 1.6
Lon 1E
North Pacific Western Pacific Cheju, 33.5 126.5 Korea Central Pacific Midway 28.2 177.4 Oahu 21.36 157.7 53.5
9.9
14.1
1.5
2.0
0.9
0.5
0.9
19.0
4.9
32.3 13.2
64.9 59.4
13.7 16.5
1.1 0.5
2.2 0.8
0.3 0.1
5.6 14.6
1.3 1.9
22.8 32.5
9.2 16.0
South Pacific American 14.3 Samoa
170.6
16.7
0.1
0.4
0.0
0.1
17.2
0.5
62.5
0.3
0.0
0.1
0.5
0.1
Antarctic Mawson
249
67.6
0.0
a
Station Location: negative latitudes – Southern Hemisphere: negative longitudes – Western Hemisphere. Total aerosol: the sum of the major aerosol components – sea salt, soil dust, nss-SO4, NO3, and NH4. c Total SS: Total aerosol minus sea salt. b
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ATMOSPHERIC TRANSPORT AND DEPOSITION OF PARTICULATE MATERIAL TO THE OCEANS
dearth of strong dust sources combined with the great distances to central ocean regions. The impact of air pollution is evident over much of the Northern Hemisphere. Extremely high NO3 and nss-SO4 2 concentrations are seen in the western Pacific near Asia; these are attributable to continental outflow and exacerbated by limited emission controls. Moderately high pollutant levels are seen over the North Atlantic as well, a result of emissions from North America and Europe. In contrast, the concentrations of NO3 and nss-SO4 2 at American Samoa and the Antarctic stations Mawson and Palmer are extremely low; these represent conditions that one might expect when pollution impacts are minimal. The larger-scale picture of PM distributions is provided by sensors such as the advanced very high resolution radiometer (AVHRR), which measures solar radiation backscattered to space by PM to estimate aerosol optical thickness (AOT) (Figure 1). There are three characteristics of the global distributions of AOT, all consistent with the data in Table 1. First, the highest AOT (i.e., the greatest column loadings of PM) is found close to the continents. This distribution affirms the fact that over most of the ocean PM is largely the result of material transported from the continents. Second, there are large seasonal differences in PM concentrations due to the seasonality of emissions and meteorology. Third, some continents emit more PM than others, illustrating the large-scale differences in production and transport. Especially notable in satellite images is a large plume of AOT over the tropical Atlantic, extending from the coast of Africa to South America (December– February) and to the Caribbean (June–August). This plume is mainly African dust. A large region of high AOT over the Arabian Sea (June–August) is attributed to dust from Africa and the Middle East. In this same season, a large PM plume seen off the west coast of southern Africa is attributed to smoke from intense biomass burning. Substantial aerosol plumes are also seen over the North Atlantic; these are caused by pollutants from North America and Europe. Large regions of high AOT are seen along the coast of Asia; but the Asian plume is most prominent in spring when large quantities of soil dust mix with pollution aerosols. The attribution of these plumes to these dominant aerosol types is supported by evidence from field studies.
Aerosol Removal Mechanisms Estimating Wet and Dry Deposition
PM is deposited to the ocean by two broadly characterized mechanisms: (1) dry deposition, in which a
particle is transferred directly from the atmosphere to the surface; and (2) wet deposition, in which a particle is first incorporated into a cloud or rain droplet that subsequently falls to the surface. The relative efficiency of the removal processes is dependent on a number of factors, especially the particle-size distribution and the hygroscopic properties of the aerosol. In most ocean regions, wet removal is thought to dominate for most aerosol species. Wet deposition is relatively easy to measure using precipitation collectors, such as automatic bucket systems that open only when precipitation falls. Dry deposition, on the other hand, is much more difficult to collect because this process is affected by a variety of factors, all highly variable: the properties of the aerosol and the water surface, atmospheric stability, relative humidity, wind velocity, etc. Furthermore, while there is a vast quantity of data on wet deposition to continental areas, there is very little for ocean environments. There are some long-term records for selected species in precipitation at a few island stations but there are no matching data for dry deposition. Wet Deposition
Long-term studies show that on average the wet deposition rates of many species are related to their concentrations in the atmospheric aerosol and to rainfall rates. This relationship is expressed in terms of a dimensionless scavenging ratio, S: S ¼ Cp rC1 a
½1
where Cp is the concentration of the substance in precipitation (g kg 1), r the density of air (B1.2 kg m 3), and Ca the aerosol concentration of the species of interest (g m 3). Wet deposition rates depend on the vertical distribution of PM and the type of precipitation event (e.g., frontal, cumulus, and stratus). In practice, comprehensive, long-term, aerosol data are only available from surface sites; consequently one must assume that over the longer term the PM concentrations in surface-level air are correlated with their vertical distributions. Therefore, S is appropriately calculated only when data have been obtained over periods of a year or more. That is, the variability in the aerosol and precipitation events must be smoothed by the averaging process. Typically used values for S fall in the range of 200–1000. Scavenging ratios can be applied to regions where no precipitation data exist using the following
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ATMOSPHERIC TRANSPORT AND DEPOSITION OF PARTICULATE MATERIAL TO THE OCEANS
251
(a) Dec., Jan., Feb.
EAOT 0.5 0.45 0.4 0.35 0.3 0.25 0.2 0.15 0.1 0.05
Husar, Stove and Prospero, 1996
(b) Jun., Jul., Aug.
EAOT 0.5 0.45 0.4 0.35 0.3 0.25 0.2 0.15 0.1 0.05
Husar, Stove and Prospero, 1996
Figure 1 The distribution of aerosols over the oceans inferred from aerosol optical thickness estimated (EAOT) by National Oceanic and Atmospheric Administration (NOAA) AVHRR. Aerosol optical thickness is a measure of the attenuation of direct solar radiation at a specific wavelength due to the scattering and absorption caused by aerosols. Large values of optical thickness suggest high concentrations of aerosols. Distributions are shown for the months (a) Dec.–Feb. and (b) Jun.–Aug. Adapted by permission of American Geophysical Union from Husar RB, Prospero JM, and Stowe LL, Characterization of tropospheric aerosols over the oceans with the NOAA advanced very high resolution radiometer optical thickness operational product, Journal of Geophysical Research, vol. 102(D14), pp. 16889–16909, 1997. Copyright 1997 American Geophysical Union.
expression: Fp ¼ PCp ¼ PSr1 Ca
½2
where Fp is the wet deposition flux (g m 2 yr 1) and P is the precipitation rate (m yr 1), using conversion factors to translate rainfall amounts to the mass of water deposited per unit area. Note
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ATMOSPHERIC TRANSPORT AND DEPOSITION OF PARTICULATE MATERIAL TO THE OCEANS
that the combined terms PSr 1Ca have a unit of velocity. Dry Deposition
There are no widely accepted methods for directly measuring PM dry deposition to water surfaces. In practice, dry deposition is almost always calculated by assuming that the deposition rate is proportional to PM concentration times a ‘deposition velocity’, vd. The dry PM flux, Fd (g m 2s 1), is given by Fd ¼ vd Ca
½3
where vd is the dry deposition velocity (m s 1) and Ca is the mass concentration of the substance in the atmosphere (g m 3). Deposition velocities have been modeled based on physical principles, and they have been empirically derived by concurrently measuring the concentration of PM species in the atmosphere and the amount deposited to a surrogate surface (typically a flat plate). While there have been some determinations of deposition velocities at continental sites, the data for ocean regions are scant. In eqn [3], vd incorporates all the processes of dry deposition, but it is difficult to accurately parametrize vd for the ambient aerosol because the importance of these processes varies with particle size. For PM between 0.1 and 1.0 mm, dry deposition is inefficient, and wet removal is normally the major sink. Gravitational settling and surface impaction control the dry removal of PM larger than about 1 mm while below about 0.1 mm, Brownian diffusion dominates. Each of these mechanisms depends on wind speed, aerosol hygroscopicity, relative humidity near the surface, and other factors which are poorly characterized. Unfortunately, there is no general agreement on how to resolve these uncertainties. Nonetheless, many estimates of dry deposition make the following assumptions about the dependence on particle size and the uncertainties in the resulting estimated deposition rate:
• • •
submicrometer aerosol particles: 0.001 m s 17 a factor of 3, supramicrometer crustal particles not associated with sea salt: 0.01 m s 17 a factor of 3, giant sea salt particles and materials carried by them: 0.03 m s 17 a factor of 2.
Despite the widespread use of these values, it should be emphasized that they are only estimates and that the error range is, if anything, probably optimistic. For example, wind speed has a great influence on deposition velocities: for PM B0.1–1.0 mm in diameter, the deposition velocity ranges from c. 0.005 cm s 1 at 5 m s 1 to c. 0.1 cm s 1 at 15 m s 1.
Deposition of Aerosols to the Oceans In this section, we present estimates of the deposition of a number of PM species that are potentially important for biogeochemical processes in the oceans. Estimates of deposition to specific locations can be made using the above relationships, assuming that the necessary aerosol concentration data are available. Larger-scale estimates of deposition are best obtained with atmospheric chemical transport models as discussed below. These models are subject to large uncertainties because they generally rely on estimates of aerosol properties and concentrations over the oceans and they incorporate highly parametrized removal schemes. For illustrative purposes, we present results for mineral dust, selected trace elements, and a group of nitrogen-containing species. A wide range of natural and anthropogenic organic species are also transported to the oceans and deposited there. Of these, certain persistent organic pollutants are known to have a harmful impact on marine biological systems. There are, however, relatively little data on the large-scale transport of organics that would enable us to address this issue on a global scale. Consequently, we do not include organic species in this report.
Deposition of Mineral Dust and Eolian Iron
In many ocean regions, primary (photosynthetic) biological production is limited by the classical nutrients such as nitrate and phosphate. In nutrient-rich surface waters, biological activity is usually high which results in high chlorophyll concentrations. But in large areas of the world’s oceans, nutrient concentrations are high, yet chlorophyll remains low which suggests low productivity. These are termed high-nutrient, low-chlorophyll (HNLC) waters; prominent examples include the equatorial Pacific and much of the high-latitude southern oceans. In the 1980s, it was found that primary production in some HNLC regions was limited by the availability of iron, an essential micronutrient in certain enzymes involved in photosynthesis. Remote ocean regions are largely dependent on atmospheric dust for the input of iron. The deposition of this eolian iron and its impact on productivity has important implications for the global CO2 budget and, hence, climate. Increased iron fluxes could conceivably fertilize the oceans, thereby increasing productivity and drawing down atmospheric CO2. In addition, certain nitrogen fixers (e.g., Trichodesmium sp.) have a high iron requirement; an increased eolian iron flux could stimulate the growth
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ATMOSPHERIC TRANSPORT AND DEPOSITION OF PARTICULATE MATERIAL TO THE OCEANS
of nitrogen fixers, thereby increasing nitrate levels and further contributing to the CO2 drawdown. Much effort has gone into estimating the temporal and spatial patterns of dust deposition to the oceans. Some studies have used satellite aerosol measurements coupled with network measurements of dust to calculate wet and dry deposition fluxes using scavenging ratios and deposition velocities. Recently, regional- and global-scale models have been developed to provide estimates of dust emissions, transport, and deposition. Dust is generally included as a passive tracer, and its removal is highly parametrized. Dry deposition is calculated using the model’s dust size distribution and size-dependent deposition (see section ‘Dry deposition’ above). Wet removal is also modeled, but the interaction of dust with clouds is not well constrained, partly because aerosol–cloud interactions in general are not well understood, and also because there are few measurements of cloud microphysical measurements in dusty regions. Mineral dust is not readily soluble in water; so some models assume that mineral aerosols do not interact with clouds directly, but rather are scavenged via subcloud removal – hence, simple scavenging ratios are used (see section ‘Wet deposition’ above). As a consequence of these uncertainties, current models show large differences in dust wet deposition lifetimes, ranging from 10 to 56 days. A typical model estimate of dust deposition rates to the oceans is shown in Figure 2. In the tropical North Atlantic, rates typically range from 2 to 10 g m 2 yr 1; over the Arabian Sea, they are as
0.000
0.2
0.5
1
2
253
high as 20 g m 2 yr 1. Over much of the North Pacific, rates are in the range 0.5–1 g m 2 yr 1, increasing to 1–2 g m 2 yr 1 closer to the coast of Asia. Dust deposition rates in Figure 2 tend to mirror the PM distribution shown in Figure 1, which, as previously stated, is largely linked to the presence of dust and, in some regions, smoke from biomass burning. Table 2 shows estimates of deposition rates to the major ocean basins produced by eight commonly used models. There is considerable agreement for the North Atlantic which is heavily impacted by African dust. In contrast, there is a considerable spread in the estimates for other regions, especially the Indian Ocean and South Pacific. Despite these differences, current models yield a reasonable, albeit broad, match with sediment trap measurements in the oceans. These various studies show that North Africa is clearly the world’s most active dust source followed by the Middle East and Central Asia. In effect, these combine to form a global dust belt that dominates transport to the oceans. These sources account for the much greater deposition rates to the northern oceans compared with southern oceans. Nonetheless, there are some substantial and important dust sources in the Southern Hemisphere in Australia, southern Africa, and southern South America. Within these continental regions, certain specific environments are particularly active dust sources, and they are sensitive to changes in climate, especially rainfall and wind speed. The presence of large, deep, alluvial deposits, usually deposited in the Pleistocene or
5
10
20
50
Figure 2 Model estimates of dust deposition rates (units: g m 2 yr 1) to the continents and the oceans. Reproduced by permission of American Geophysical Union from Mahowald NM, Baker AR, Bergametti G, et al., Atmospheric global dust cycle and iron inputs to the ocean. Global Biogeochemical Cycles, vol. 19, GB4025, 2005. Copyright 2005 American Geophysical Union.
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ATMOSPHERIC TRANSPORT AND DEPOSITION OF PARTICULATE MATERIAL TO THE OCEANS
Table 2
Estimate of mean annual dust deposition to the global ocean and to various ocean basins Annual dust deposition ratea (1012 g yr 1)
Reference
Duce et al. (1991) Prospero (1996) Ginoux et al. (2001) Zender et al. (2003) Luo et al. (2003) Ginoux et al. (2004) Tegen et al. (2004) Kaufman et al. (2005)
GO
NAO
SAO
NPO
SPO
NIO
SIO
910 358 478 314 428 505 422
220 220 184 178 230 161 259 140
24 5 20 29 30 20 35
480 96 92 31 35 117 56
39 8 28 8 20 28 11
100 20 154 36 113 164 61
44 9 12 15
a
GO, Global Oceans; NAO, North Atlantic Ocean; SAO, South Atlantic Ocean; NPO, North Pacific Ocean; SPO, South Pacific Ocean; NIO, North Indian Ocean; SIO, South Indian Ocean. Adapted from Engelstaedter S, Tegen I, and Washington R (2006) North African dust emissions and transport. Earth-Science Reviews 79(1–2): 73–100.
Table 3 Metal
Iron Copper Nickel Zinc Arsenic Cadium Lead
Atmospheric and riverine fluxes of dissolved and particulate trace metals to the oceana Atmospheric transports
Riverine transports
Dissolved transports
AtmDiss
AtmPart
Ratio Diss/Part
RivDiss
RivPart
Ratio Diss/Part
Ratio AtmDiss/RivDiss
3200 30 10 102 4 3 75
28 000 5 16 33 2 1 9
0.1 6.6 0.6 3.1 1.7 4.7 8.3
1100 10 11 6 10 0.3 2
11 000 1 500 1 400 3 900 80 15 1 600
0.100 0.007 0.008 0.002 0.125 0.020 0.001
2.9 3.0 0.9 16.9 0.4 8.7 37.5
Units: 109 g yr 1. Atm, atmospheric; Riv, riverine; Diss, Dissolved; Part, particulate. Adapted from Duce RA, Liss PS, Merrill JT, et al. (1991) The atmospheric input of trace species to the world ocean. Global Biogeochemical Cycles 5: 193–259.
a
Holocene, is a common prerequisite for strong dust sources. Trace Element Deposition
Some oceanographers study the biogeochemical cycling of trace elements and seek to quantify the elements’ oceanic sources and sinks. In regions dominated by mineral dust, the ratios of many trace elements (e.g., Al, Ba, Ca, Cs, Fe, Hf, Mn, Rb, Sc, Ta, Th, and Yb) are similar to those in geological materials such as soils, thus implicating mineral dust as their main source. Several elements (Co, Cr, Eu, Mg, and Na) show slight enrichments over crustal values while others such as As, Cd, Cu, Ni, Pb, Sb, Se, V, and Zn are strongly enriched. Such large enrichments are typically associated with pollution impacts, but emissions from natural sources such as volcanoes can also be responsible.
The impact of trace element deposition on ocean biogeochemistry depends to a great extent on the degree to which the elements are soluble in seawater. Extensive studies of trace element solubilities in various natural aqueous media or in aqueous solutions of similar composition yield a wide range of solubilities; these depend on the types of aerosols used, the exposure times, and other experimental variables, especially pH. Thus it is difficult to convert the estimated air/sea exchange rates into an effective or bioavailable flux of trace elements, and therefore it is difficult to accurately assess the impact of PM deposition on ocean processes. Comparison of Trace Element Transport by Rivers and by the Atmosphere
Rivers carry large quantities of dissolved and suspended materials to the oceans. Table 3 compares the
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ATMOSPHERIC TRANSPORT AND DEPOSITION OF PARTICULATE MATERIAL TO THE OCEANS
amounts of selected trace elements carried by rivers with that carried by winds and also the relative amounts of particulate and dissolved or soluble phases in the transported material. Riverine transports of trace elements are overwhelmingly in the particulate phases and the ratio of dissolved to particulate phases ranges from about 0.001 to 0.13. In comparison, the corresponding ratio for atmospheric transport is much larger, ranging from 0.1 to 8.3. Table 3 also shows that the dissolved or soluble inputs to the ocean from the atmosphere exceed those from rivers; in almost all cases the ratio is greater than unity, in some cases much larger. The comparison of river versus atmospheric inputs is based on the measured concentrations in rivers before they reach the oceans. However, much of the material carried by rivers is rapidly deposited when the rivers reach the sea; therefore, the impact of air/ sea exchange on ocean systems is in reality much greater than suggested by Table 3. We emphasize, however, that the data in Table 3 are rather old. Recent work shows that the concentrations of some trace elements have changed significantly over time. For example, during the mid-1900s, aerosol lead greatly increased in response to increasing pollution emissions but in later years concentrations decreased as controls were implemented. In recent decades, other trace elements have increased due to growing industrialization in developing nations. Also, recent research suggests that there is considerably more uncertainty in PM trace metal solubility than shown in Table 3. Nonetheless, one would still expect that the impact of atmospheric transport is much greater than river transport, especially for the open ocean. Nitrogen Deposition
Anthropogenic activities have greatly increased the amounts of nitrogenous materials that enter the atmosphere and find their way into the rivers (Table 4). There is interest in the possible impacts of these materials on the marine environment, especially about chemicals such as nitrate that can affect primary production. This issue is of particular importance in regions where nitrogen is the limiting nutrient, for example, the oligotrophic central oceanic gyres where an enhancement in productivity would increase the drawdown of CO2 and hence affect climate. In coastal waters, atmospheric inputs could contribute to eutrophication although one would expect the inputs from rivers to dominate. There are two broad classes of nitrogen compounds of interest: oxidized and reduced. The most important oxidized species are NO and NO2
Table 4
255
Atmospheric emissions of fixed nitrogen, 1993a
Sources
NOx
NH3
Anthropogenic Biomass burning Agricultural activity Fossil fuel combustion Industry Total anthropogenic
6.4 2.6 20.9 6.4 36.3
4.6 39.7 0.1 2.8 47.2
2.9 5.4 0.8 0.6 0.0 6.8
4.6 0.0 0.8 0.0 5.6 11.0
43.1 5.3
58.2 4.3
Natural Soils, vegetation, and animals Lightning Natural fires Stratosphere exchange Ocean exchange Total natural Grand total Ratio: anthropogenic/natural
Units Tg/g(N) yr 1. Adapted from Jickells TD (2006) The role of air–sea exchange in the marine nitrogen cycle. Biogeosciences 3: 271–280. a
(collectively referred to as NOx) and NOy (termed reactive odd nitrogen) which is comprised of NOx plus the compounds produced from its oxidation, including HNO3 and other compounds. NOx is rapidly oxidized in the atmosphere to a wide range of compounds, many of which are ultimately converted to HNO3 and aerosol NO3 . In the marine boundary layer, HNO3 reacts rapidly with sea salt particles and promptly deposits on the sea surface. Indeed, NO3 is the N-containing compound of greatest interest in terms of impact on the oceans, and it is the N compound most commonly measured and modeled. Table 4 lists the major sources of oxidized and reduced N emitted to the atmosphere. The primary natural sources of NOx are biological fixation and lightning, the latter being rather minor. In modern times, fossil-fuel combustion along with industry and biomass burning dominate the oxidized N cycle. The ratio of anthropogenic NOx emissions to that of natural sources is 2.5 and continues to increase. While most research has focused on inorganic N (IN), there is evidence that organic nitrogen (ON) also may play an important role in marine biogeochemical cycling. However, there are relatively few data on ON compounds and most focus only on dissolved ON. The major reduced nitrogen species (NHx) are aerosol NH4 þ and NH3, the latter being the only gas-phase species that significantly titrates atmospheric acidity. The major natural sources of NH3 (Table 4) include soils, vegetation, and excreta from wild animals. However, the emissions of NH3 to the atmosphere are now dominated by fertilizers and the
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ATMOSPHERIC TRANSPORT AND DEPOSITION OF PARTICULATE MATERIAL TO THE OCEANS
excreta from dairy and beef cattle. Indeed, the ratio of anthropogenic and natural NH3 emissions is 9. The oceans can also be a source of NH3 under certain conditions in some regions, but the continental sources clearly dominate. Models provide estimates of the present-day atmospheric N fluxes to the oceans and their spatial distribution. Figure 3 presents the deposition rate of
reactive nitrogen NOy þ NHx for the year 2000. As was the case for dust, emissions and deposition in the Northern Hemisphere are much greater than those of the Southern Hemisphere. Deposition rates are extremely high adjacent to the continental coastlines; thus one would expect that the adjacent water bodies would be most heavily impacted. The total N flux to the ocean (NOy and NHx but not including ON) in mg N m−2 6000 3000
60° N
2000 1000 900
30° N
800
Latitude (deg)
700 600 EQ.
500 400 300 200
30° S
100 75 60° S
50 25 10
180° W 150° W 120° W
90° W
60° W
30° W
0° E
30° E
60° E
90° E
120° E 150° E 180° E
Longitude (deg)
Figure 3 Model estimates of the deposition rate of total reactive nitrogen (NOy þ NHx) (units: mg N m 2 yr 1) in the year 2000. Reprinted with permission from Dentener F, Stevenson D, Ellingsen K, et al. (2006) The global atmospheric environment for the next generation. Environmental Science and Technology 40(11): 3586–3594 (doi:10.1021/es0523845). Copyright (2006) American Chemical Society.
Table 5
NOy and NHx deposition for the year 2000
Deposition region
Ocean Coastal ocean NH SH World Ratio: ocean/ world
NOy þ NHx
NHx
NOy Total (Tg(N) yr 1)
Mean rate (mg(N) m2 yr 1)
Total (Tg(N) yr 1)
Mean rate (mg(N) m2 yr 1)
Total (Tg(N) yr 1)
23 4 38 14 52 0.43
61 192 150 54 102
24 4 48 16 65 0.36
63 206
47 8 87 30 117 0.40
126
Adapted from Dentener F, Drevet J, Lamarque JF, et al. (2006) Nitrogen and sulfur deposition on regional and global scales: A multimodel evaluation. Global Biogeochemical Cycles 20: GB4003 (doi:10.1029/2005GB002672).
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ATMOSPHERIC TRANSPORT AND DEPOSITION OF PARTICULATE MATERIAL TO THE OCEANS
2000 was 46 Tg N yr 1 of which 8 Tg N yr 1 is deposited to the coastal ocean (Table 5). The deposition of oxidized forms (NOy) is essentially equal to that of reduced forms (NHx). The ocean N deposition amounts to c. 40% of global emissions. The total reactive N transport by rivers is about 48 Tg N yr 1, essentially the same as the atmospheric source. However, there is evidence that fluvial nitrogen inputs to the oceans are denitrified on the shelf and that the shelf region is a sink rather than a source of nitrogen for the open oceans. Thus it appears that air/sea transfer is the major source of N transported to the open ocean. ON compounds could also be playing a significant role in total N fluxes. Studies from many different environments suggest that ON constitutes about a third of the total atmospheric reactive nitrogen. Thus, ON could add significantly to the total global flux to the oceans, conceivably raising the total to about 69 Tg N yr 1.
Conclusions It is now recognized that atmospheric transport plays a central role in ocean biogeochemical processes. There is increased interest in the chemically coupled ocean/atmosphere system, how this system has changed over time, and how it might respond to global change. Although many models are currently focusing on this question, the development of these models is handicapped by the dearth of measurements over many ocean regions. It remains a formidable challenge to the community to carry out the necessary measurements over such large ocean areas.
See also Anthropogenic Trace Elements in the Ocean. Atmospheric Input of Pollutants. Iron Fertilization. Metal Pollution. Nitrogen Cycle. Biological Pump and Particle Fluxes. Trace Element Nutrients.
Further Reading Arimoto R, Kim YJ, Kim YP, et al. (2006) Characterization of Asian dust during ACE-Asia, global and planetary change. Monitoring and Modelling of Asian Dust Storms 52(1–4): 23--56.
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Arimoto R, Ray BJ, Lewis NF, Tomza U, and Duce RA (1991) Mass-particle size distributions of atmospheric dust and the dry deposition of dust to the remote ocean. Journal of Geophysical Research – Atmospheres 102(D13): 15867--15874. Boyd PW, Watson A, Law CS, et al. (2000) A mesoscale phytoplankton bloom in the polar Southern Ocean stimulated by iron fertilization of waters. Nature 407: 695--702. Dentener F, Drevet J, Lamarque JF, et al. (2006) Nitrogen and sulfur deposition on regional and global scales: A multimodel evaluation. Global Biogeochemical Cycles 20: GB4003 (doi:10.1029/2005GB002672). Dentener F, Stevenson D, Ellingsen K, et al. (2006) The global atmospheric environment for the next generation. Environmental Science and Technology 40(11): 3586--3594 (doi:10.1021/es0523845). Duce RA, Liss PS, Merrill JT, et al. (1991) The atmospheric input of trace species to the world ocean. Global Biogeochemical Cycles 5: 193--259. Engelstaedter S, Tegen I, and Washington R (2006) North African dust emissions and transport. Earth-Science Reviews 79(1–2): 73--100. Harrison SP, Kohfeld KE, Roeland C, and Claquin T (2001) The role of dust in climate today, at the Last Glacial Maximum and in the future. Earth-Science Reviews 54: 43--80. Husar RB, Prospero JM, and Stowe LL (1997) Characterization of tropospheric aerosols over the oceans with the NOAA advanced very high resolution radiometer optical thickness operational product. Journal of Geophysical Research 102(D14): 16889--16909. Jickells TD (2006) The role of air–sea exchange in the marine nitrogen cycle. Biogeosciences 3: 271--280. Jurado E, Jaward F, Lohmann R, et al. (2005) Wet deposition of persistent organic pollutants to the global oceans. Environmental Science and Technology 39(8): 2426--2435 (doi:10.1021/es048599 g). Mahowald NM, Baker AR, Bergametti G, et al. (2005) Atmospheric global dust cycle and iron inputs to the ocean. Global Biogeochemical Cycles 19: GB4025 (doi:10.1029/2004GB002402). Parekh P, Follows MJ, and Boyle EA (2005) Decoupling of iron and phosphate in the global ocean. Global Biogeochemical Cycles 19: GB2020 (doi:10.1029/ 2004GB002280). Prospero JM (1996) The atmospheric transport of particles to the ocean. In: Ittekkot V, Scha¨fer P, Honjo S and Depetris PJ (eds.) Particle Flux in the Ocean. SCOPE Report 57, pp. 19--52. Chichester: Wiley. Wesely ML and Hicks BB (2000) A review of the current status of knowledge on dry deposition. Atmospheric Environment 34(12–14): 2261--2282.
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AUTHIGENIC DEPOSITS G. M. McMurtry, University of Hawaii at Manoa, Honolulu, HI, USA Copyright & 2001 Elsevier Ltd.
usually lithified sediment or basalt substrate of a seamount or island slope to create ‘pseudo’ or ‘seamount nodules’ that are otherwise indistinguishable from crustal pavements.
This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 1, pp 201–212, & 2001, Elsevier Ltd.
Ferromanganese Nodules
Introduction There are many kinds and forms of authigenic mineral deposits in the deep-sea. Here we concentrate on the more abundant and, in some cases, potentially economic types, namely ferromanganese deposits, phosphorites, marine barite, and the authigenic silicate minerals in deep-sea sediments. Hydrothermal deposits and a class of sedimentary iron and manganese oxides associated with seafloor hydrothermal activity – the metalliferous sediments found as a basal sequence in the deep-sea sedimentary column and in deep basins adjacent to active spreading centers – are described elsewhere in the encyclopedia. Ferromanganese deposits are here divided into nodules and crusts on the basis of morphology and environment of formation. The sources of the major and minor metal enrichments in these deposits are contrasted between continental weathering and submarine volcanism, with important enrichment influences from the marine biosphere. Marine barite has a similar mixed origin with strong biological influence. Knowledge of the origins of submarine phosphorites has been greatly aided by studies of recent seafloor deposits. There are a variety of authigenic silicate minerals in deep-sea sediments. The more abundant zeolites and clays such as smectite form mainly from diagenetic alteration of metastable volcanic glasses, but many other reaction paths are known and suggested.
Ferromanganese Deposits Marine ferromanganese deposits are complex assemblages of authigenic minerals and detrital components. They are broadly characterized by shape and by environment of formation as deep-sea nodules and crusts. Ferromanganese nodules are commonly found on the abyssal seafloor at and near the top of the underlying pelagic sediment cover. Crusts are usually found on the steep slopes of islands and seamounts but are also found with the pelagic sediment on the deep-sea floor where sedimentation is extremely low. Occasionally, nodules may grow to coalesce into a pavement or crust of ferromanganese deposits. Conversely, crusts can grow around small fragments of the
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Deep-sea ferromanganese nodules are classified as three major types on the basis of their origin: hydrogenetic, early diagenetic, and mixed hydrogenetic–early diagenetic. In this same scheme, crusts are described as primarily hydrogenetic, with comparatively minor diagenetic and hydrothermal occurrences (Figure 1). Hydrogenetic nodules obtain their ferromanganese oxides as precipitates from sea water. However, they are rarely if ever pure and also contain a variety of detrital minerals and biogenic debris from the pelagic sediment and fallout from the overlying water column and atmosphere. Early diagenetic nodules obtain their ferromanganese oxides from metals dissolved from the sediment and transported within the pore waters of the upper, peneliquid sediment layer (Figure 1). Because manganese and minor metals such as nickel and copper are more readily mobile in pore waters under suboxic seafloor conditions, diagenetic nodules are enriched in these metals relative to concentrations of iron and cobalt. The hybrid or mixed type nodule has both hydrogenetic and diagenetic components. These nodules are usually found at the sediment–water interface where their outer layers record their most recent position. The upper, relatively smooth hydrogenetic top is composed of iron- and cobalt-rich ferromanganese oxides from the bottom water and the lower, relatively bumpy or ‘botroyidal’ diagenetic bottom is composed of manganese-, nickel-, and copper-rich ferromanganese oxides from the sediment. Attempts have been made to classify nodules on the basis of morphology and to relate their occurrence to the composition of the underlying sediment. Such field classifications were made to predict metal enrichment and genesis, but reliance upon morphological data for other than abundance information from either seafloor photography or remote acoustic imaging has not yet been widely practiced in exploration. Nodule abundance appears to correlate more strongly with siliceous ooze than with red clays, and both siliceous ooze and red clays have higher nodule abundance than does carbonate ooze. Ferromanganese nodules and crusts are mainly composed of three types of manganese oxide minerals: vernadite or d-MnO2; birnessite; and todorokite. These
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AUTHIGENIC DEPOSITS
259
Ocean water
Hydrogenetic crusts
Hydrogenetic nodules Range of calcite compensation depth Mixed-type nodules
Early diagenetic nodules
Precipitation of colloidal particles from near-bottom sea water Upward migration of dissolved 2+
2+
metals such as Mn , Cu , 2+ 2+ 2+ Ni , Zn , Mg under pore water conditions
Basaltic substrate rocks
Peneliquid sediment layer
Consolidated pelagic sediment
Figure 1 Schematic figure of the principal types of ferromanganese nodules (diameters vary from 1 to 15 cm) and crusts (thicknesses vary from o1 to 10 cm) distributed in pelagic sediments and on seamounts and island slopes. (After Halbach P (1986) Processes controlling the heavy metal distribution in Pacific ferromanganese nodules and crusts. Geologische Rundschau 75: 235–247.)
minerals are listed in order of decreasing Eh or oxidation potential of the formation environment, and increasing crystal order. In vernadite, all manganese occurs as Mn4þ, which allows for crystal lattice substitution of metals such as Co3þ and Pb4þ. Todorokite contains some manganese as Mn2þ, which allows for crystal lattice substitution of metals such as Cu2þ and Ni2þ. Birnessite is an intermediate form. Iron oxides are primarily found as closely intermingled X-rayamorphous FeOOH and occasionally as goethite or lepidocrocite. Other minerals include detrital quartz, plagioclase feldspar, clay minerals such as smectites, authigenic zeolites, apatite, barite, and calcite. The global distribution of ferromanganese nodules on the seafloor (Figure 2) largely reflects the inverse distribution of thickest sediment cover. Nodules, which accumulate at rates of several millimeters per million years, cannot grow where sediments are accumulating at rates that are greater than 103 times faster (several centimeters per thousand years). Otherwise, rapid burial and dissolution would be their fate. Paradoxically, nodules are found to lie atop sediments that are accumulating at rates up to 103 times faster; this phenomenon has been explained by bottom current winnowing and gentle
bumping or tilting of nodules during the burrowing activities of benthic organisms. Nodules are truly pelagic deposits and are rarely found near the continents or near islands where rapid hemipelagic sedimentation occurs. Exceptions are found on the Blake Plateau off eastern Florida, within the Drake Passage off South America, and near the Cape of Good Hope off Southern Africa, where bottom currents are strong and sweep sediments from these areas (Figure 2). Highly concentrated nodule fields occur between the Clarion and Clipperton fracture zones just north of the equatorial high productivity zone in the Eastern Pacific, just south of this zone in the Indian Ocean, and near the Antarctic Convergence in the Southeast Pacific. Source of major metals The sources of the major metals in ferromanganese nodules and crusts – iron and manganese – are commonly described as products of the weathering of the continents, seafloor hydrothermal activity, and sediment diagenesis. Other than a small but distinct cosmogenic component (see Ferromanganese Crusts below), the iron and manganese must ultimately come from the Earth’s crust and mantle. The
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AUTHIGENIC DEPOSITS
Areas of nodule coverage
Areas where nodules are particularly abundant
Figure 2 Distribution of ferromanganese nodule abundance on the world ocean floor. (After Cronan, 1980.)
fraction coming from the continents is more geologically recycled than that from the ocean crust, and diagenesis in this regard is clearly not an ultimate source but a process of further recycling of these metals within the sediment column. The relative importance of continental weathering versus seafloor hydrothermal activity as a source of iron and manganese to pelagic sediments and marine ferromanganese deposits has been debated since the 1890s. Early work by able geochemists on both sides of this issue was hampered by a lack of convincing evidence. For example, on the basis of early distribution maps of manganese and other trace metals in surface Atlantic sediments and evidence for iron and manganese enrichments in atmospheric dust, it was suggested that continental weathering and atmospheric transport was an important pathway of these metals to the ocean. However, the manganese distribution in Atlantic sediments can more readily be explained by hydrothermal enrichments from hot springs along the Mid-Atlantic Ridge, unknown at that time but well-known today. Likewise, early arguments for extensive hydrothermal iron and manganese enrichments from volcanic hot springs located along the mid-ocean ridges were compelling, but were not confirmed until the discovery of hot springs there in 1977. Water column trace-metal analytical techniques have improved to the point that a very convincing argument can be made for continental weathering and atmospheric transport as a path for at least some fraction of the iron and manganese in marine ferromanganese deposits. Another approach, which additionally answers the old
question of the relative contributions of these two sources to ferromanganese deposits at any given location of the seafloor, was made for the Pacific (Figure 3) using a ‘Co-chronometer’. This dating method is based upon the inverse relationship to growth rate of some minor metals’ concentrations in hydrogenetic ferromanganese crust deposits, more painstakingly measured by radiometric techniques such as U-series (excess 230Th) or 10Be dating. Using the relatively easy Co-chronometer, hundreds of ferromanganese deposits were measured to construct the detailed map of ferromanganese crust growth rates for the Pacific in Figure 3. The highest growth rates are associated with active spreading along the East Pacific Rise, Juan de Fuca Ridge, and Galapagos Rift, and with submarine arc volcanism in the Mariana Island arc. The deposits in these areas therefore receive most of their iron and manganese from hydrothermal sources. The slowest crust growth rates and highest cobalt enrichments are located near the Mid-Pacific Mountains seamount province west of Hawaii, far from active hot springs and the continents. Source of minor metals Relative to sea water and the Earth’s crust, ferromanganese nodules are enriched in nickel, copper, cobalt, and a host of other minor metals such as the platinum-group elements (PGE) and the rare-earth elements (REE). Concentrations of nickel and copper reach up to 2 wt% in ‘high-grade’ nodules from fields between the Clarion and Clipperton fracture zones, with cobalt concentrations approaching 0.5 wt%. Broadly, nodules from the Pacific appear to be more enriched
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AUTHIGENIC DEPOSITS
120˚
140˚
160˚
180˚
160˚
140˚
120˚
100˚
80˚
R (mm/Myr) 60˚
60˚
40˚
40˚
20˚
20˚
0˚
0˚
20˚
20˚
40˚
40˚
60˚
60˚
261
inverse relationship of depth of copper and nickel concentration in Pacific nodules with the sedimentary concentration of calcium carbonate tests from planktonic organisms, with deeper sediment deposits containing lesser CaCO3 because of increased dissolution. Comparative studies of copper, nickel, and cobalt in plankton and the labile or easily mobilized fraction of pelagic sediment and associated deep-sea nodules has also suggested that diagenesis of organic matter in the sediments leads to an enrichment of these metals in the nodules. All told, scavenging by organisms in the surface waters enriches the particle rain in minor metals that originally enter the world’s oceans by winds and rivers, and further concentration of these metals occurs within the surface sediments during suboxic diagenesis, with ferromanganese nodules as the final metals depository. Ferromanganese Crusts
120˚
140˚
160˚
180˚
160˚
140˚
120˚
100˚
80˚
Figure 3 Distribution of ferromanganese crust accumulation rates in the Pacific Ocean. (After Manheim FT and Lane-Bostwick CM (1988) Cobalt in ferromanganese crusts as a monitor of hydrothermal discharge on the Pacific sea floor. Nature 335, 59–62.)
in these metals than those from the Indian and Atlantic Oceans, but metal concentrations can vary widely within basins and nodule fields. For at least the better-studied Pacific, concentrations of nickel and copper show good correlation to the nodule Mn/Fe ratio, suggesting that manganese-rich, todorokitecontaining diagenetic nodules are the major depositories of these metals (Figure 4A, B). There is also a regional trend of nickel and copper enrichments and Mn/Fe ratios in the Pacific that is highest in the equatorial NE Pacific, the Peru Basin, and the SE Pacific near the Antarctic Convergence. The distribution of nodule cobalt concentrations shows little correlation to the Mn/Fe ratio, however, being enriched instead near the Mid-Pacific Mountains and in the South Central Pacific (Figure 4C). The ultimate source for the minor metals in ferromanganese nodules is again the Earth’s crust and mantle, but with minor metals the case for continental weathering versus seafloor hydrothermal activity is stronger. First, hydrothermal ferromanganese crusts, like many rapidly accumulating diagenetic crusts, are known to have very low concentrations of minor metals. Second, studies of copper and nickel distributions in pelagic surface sediments suggest that the highest concentrations are nearest to regions of high surface productivity, implying that plankton are involved in the enrichment process. There is an
Crust deposits differ from nodules in form, occurrence, and composition. Hydrogenetic crusts are principally composed of vernadite with iron oxides and minor detrital mineral and carbonate fluorapatite (CFA) contaminates. They grow at extremely slow rates from less than 1 mm to tens of millimeters per million years. Because the cobalt flux is for the most part invariant throughout the world’s oceans, cobalt concentrations approach 2 wt% on comparatively shallow seamount slopes between 800 and 2000 m water depth within the Western Equatorial Pacific (Figure 3 – compare with Figure 4C) where crust growth rates are slowest. Above approximately 800 m dilution and coverage of crustal pavements by principally carbonate-rich sediments limits their growth, and below 2000 m incorporation of increasing fluxes of seamount and wind-blown detritus and ferromanganese oxides causes crust growth rates to increase, effectively decreasing the cobalt concentration. The oxygen minimum zone (OMZ) presently intersects seamounts and island slopes in the Central Pacific at between 500 and 1500 m water depth. Although not presently anoxic enough to inhibit manganese oxide precipitation in this area, the OMZ was probably more intense in the past, thereby providing a mechanism for both transport of dissolved Mn2þ and inhibition of crust growth. Cobalt concentrations also show a general trend of decreasing values within older crustal layers, suggesting that past seafloor conditions favored increased manganese fluxes and growth. Relative to nodules, crusts are on average enriched in Fe, Ca, P, Ti, Pb, Ce, As, and Pt, as well as cobalt, and are depleted in Si, Al, Ni, Cu, and Zn. Manganese concentrations are similar. The relative enrichments of calcium and phosphorus in crusts reflect the more widespread incorporation of CFA in seamount crusts
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AUTHIGENIC DEPOSITS
(B)
(A)
>5 2_ 5
> 1% 0.5−1%
1_ 2
0.25−0.5%
<1
< 0.25%
(C)
> 0.50% 0.25−0.50% < 0.25% Figure 4 (A) Distribution of Mn/Fe ratios in ferromanganese nodules from the Pacific. (B) Distribution of nickel concentrations in ferromanganese nodules from the Pacific. Copper shows a similar distribution. (C) Distribution of cobalt concentrations in ferromanganese nodules from the Pacific. (After Calvert SE (1978) In: Sea Floor Development: Moving into Deeper Water. London: The Royal Society.)
(see Phosphorites below), whereas the higher silicon and aluminum contents of nodules reflect their origin within loose, aluminosilicate-rich sediments of the abyssal seafloor. The mechanism for cobalt enrichment is oxidative scavenging, whereby dissolved Co2þ is oxidized to Co3þ on the manganese oxide surface. This mechanism also explains the relative enrichments of cerium, and possibly some of the lead and titanium enrichments in crusts. Lower concentrations of iron (and arsenic, an oxyanion in sea water scavenged by positively charged FeOOH; MnO2 surfaces are negatively charged at sea water pH) in nodules reflect a
greater diagenetic component that favors mobilization of manganese over iron. Because of their extremely slow growth in areas relatively free of detrital input, crusts also accumulate large amounts of cosmogenic debris. Enrichments of platinum and PGE metals can be partially explained by this source, but several lines of geochemical evidence indicate that the majority of platinum and the other PGE (Ir, Os, Pd, Rh, Ru) are scavenged from sea water. Oxidation and reduction mechanisms have been proposed for both incorporation and postdepositional remobilization of the PGE in crusts and nodules.
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AUTHIGENIC DEPOSITS
Phosphorites Submarine phosphorites are principally composed of CFA or francolite with usually large admixtures of detrital impurities from the sediments in which they form or replace during authigenesis. Concentrations of P2O5 commonly range from 5 to 28 wt%, reaching values of 35 wt%, with minor element enrichments of strontium, yttrium, and the REE. There are three principal types of phosphorite deposits in the world’s oceans: offshore or continental shelf-slope nodules and concretions; massive insular deposits originating from sea bird guano; and ‘seamount’ phosphorites that often associate with ferromanganese crusts as layers, veins, impregnations, and substrate material (Figure 5). Submarine phosphorites were originally thought to be fossil deposits no younger than Miocene age, like those marine phosphorites mined on land, but were subsequently found to be actively forming in shelf-slope areas off Western South America, Southwest Africa, and Australia. Studies of the deposits off Peru and Chile have demonstrated their contemporaneous origin near anoxic sediments where the OMZ impacts the continental shelf-slope there at between 200 and 400 m (Figure 6). Upwelling currents bring waters rich in dissolved PO4 onto the continental slope and shelf, increasing surface productivity and enhancing the reduction potential in sediments underlying the OMZ. Bacterial consumption of organic matter within the sediment releases HPO4 2 that can combine with dissolved Ca2þ, F and CO3 2 in pore waters to form authigenic precipitates of CFA. This process is thought to occur in sediments near the boundaries of the OMZ where the hydrogen ion
Figure 5 Photograph of a ferromanganese crust on a phosphorite substrate recovered from the summit of Schumann seamount, north of Kauai, Hawaii. Note the inclusions of altered basalt clasts from the seamount within the phosphorite, which has likely replaced limestone. (From G. McMurtry and D. L. VonderHaar, unpublished data.)
263
concentrations (low pH) produced during bacterial consumption and in CFA precipitation are diluted by relatively alkaline (high pH) sea waters. Associated sediment diagenetic reactions involving dolomite formation and Mg2þ exchange in clays reduce the concentration of apatite precipitation inhibiting Mg2þ in pore waters, aiding CFA formation. The distribution of phosphorites throughout the Phanerozoic rock record is highly irregular, correlating broadly with periods of warm climate and higher sea level. Off South America and South-west Africa, Useries dating of the deposits has demonstrated their correlation with warm, interglacial periods back to 150 000 years before present. Increased chemical weathering of the continents during warm periods may provide enhanced PO4 fixation into hemipelagic
Figure 6 Schematic figure of the western South American shelf and slope, showing contours of dissolved phosphate and oxygen in the bottom waters and the locations of aerobic and anaerobic sediments, phosphorite and Recent phosphorite deposits. (After Burnett WC, Veeh HH and Soutar A (1980) U-series, oceanographic and sedimentary evidence in support of recent formation of phosphate nodules off Peru. In: Marine Phosphorites, SEPM Special Publ. no. 29, pp. 61–72.)
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sediments. Concurrent encroachment of the OMZ onto the continental shelf-slope during marine transgressions increases the extent of organic matter preservation. When combined, these effects are thought to promote initial phosphorite formation. Subsequent marine regressions may enhance the initial enrichment by winnowing detritus from older phosphatic sediments, producing deposits of economic value. Insular phosphorites are relatively easily understood deposits from guano in the nesting areas of ancient sea birds that created the ore deposit by transporting and further biologically concentrating scarce dissolved phosphate in surface sea waters within the flight radius of the island. The island of Nauru in the central Equatorial Pacific is a prime example. Although some are presently submerged, these deposits have all formed subaerially and are no older than Tertiary. Comparatively less is known about the formation of seamount phosphorites. Substrate samples show cryptocrystalline CFA matrix replacement of preexisting limestones that often contain altered basalt breccia and other evidence of high-energy deposition upon the summits and upper slopes of Cretaceous seamounts in the Pacific (Figure 5). Within ferromanganese crusts, these deposits show the effects of lowered redox conditions (e.g., platinum remobilization) with impregnation and replacement of older ferromanganese oxides. Strontium and oxygen isotope-derived ages of CFA formation appear to center on the Eocene–Oligocene (36 Ma) and Oligocene– Miocene (24 Ma) boundaries, with some evidence for a minor Middle Miocene event at about 15 Ma. Paleotracking of the Pacific Cretaceous seamounts shows that many were close to the equatorial high productivity belt during the late Cretaceous and Paleocene–Eocene periods, where their then-shallower summits and slopes could have intersected the OMZ in a region of equatorial upwelling. Later depositional episodes would require a greatly expanded and more intense OMZ than the present one.
mineral in hydrothermal deposits on land and has more recently been found as a principal component of moderate-temperature hydrothermal chimneys on midocean ridges and volcanically active seamounts. Early work on the distribution of marine barite in deep-sea sediments had difficulty distinguishing the relative importance of hydrothermal sources from continental weathering as cycled through the marine biosphere. Distribution maps of marine barite in Pacific deep-sea sediments and the more quantitative barium accumulation rate (Figure 7) show both an association with the equatorial high productivity zone and the East Pacific Rise–Bauer Basin areas that are heavily influenced by metalliferous hydrothermal deposition. The key question is how much of the relatively dispersed, finegrained (usuallyo2 mm) marine barite particles in deep-sea sediments result from hydrothermal plume fallout and bottom current redistribution of ridge-crest metalliferous sediments versus those carried to the seafloor in the particle rain from surface productivity. The association of marine barite with organic matter complicates the interpretation of occurrence of increased barite deposition found along midocean ridges because greater preservation of organic matter also occurs with increased carbonate sedimentation above the calcite compensation depth. Additionally, basin-scale dispersal of hydrothermal particles appears limited, especially for the relatively dense barite. Studies of marine barite saturation show that barite is below saturation in the water column but rapidly approaches saturation in the pore waters of deep-sea sediments. Discrete barite particles are nevertheless found in the microenvironments of suspended
Marine Barite and Authigenic Silicates Marine Barite
Barite (BaSO4) is a widespread mineral in deep-sea sediments, varying between 1 and 10 wt% on a carbonate-free basis. It is the predominant barium phase in the ocean. BaSO4 is known to compose a solid solution series with SrSO4 as celestobarite in the skeletal portions of some marine organisms (i.e., the Xenophyophoria) and is often found in association with marine organic matter, such as in suspended particles and fecal pellets. Barite is also a well-known gangue
Figure 7 Distribution of barium accumulation rates (units of mg cm2 per 1000 years) in the Pacific sediments. (After Bostro¨m K, Joensuu O, Moore C et al. (1973), Geochemistry of barium in pelagic sediments. Lithos 6: 159–174.)
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Figure 8 Photograph of barite crystals inside a ferromanganese nodule. (A) Scanning electron micrograph of foraminifera shell walls that have been replaced by barite. (B) Barium elemental X-ray map produced by EMP. (C) Magnification of (A) showing individual barite crystals. (D) Elemental spectrum of the barite crystals: S Ka at 2300 eV; Ba La, b, g at 4460 eV. (After Lalou C, Brichet E, Poupeau G, Romany P and Jehanno C (1979) Growth rates and possible age of a North Pacific manganese nodule. In Bischoff JL and Piper DZ (eds) Marine Geology and Oceanography of the Central Pacific Manganese Nodule Province, pp. 815–834. New York and London: Plenum Press.)
biogenic matter and in fecal pellets that rapidly fall to the seafloor; their dissolution on the seafloor and in transit is likely the dominant control on the barium concentration in the deep ocean. Within the sediment and in ferromanganese deposits, barite can form authigenically as discrete particles and diagenetically as replacement of biogenic skeletal remains (Figure 8). Authigenic Silicates
Zeolites Low-temperature alteration of basalt and metastable volcanic glasses on pillow rims, in layers, and dispersed throughout deep-sea sediments can produce a variety of diagenetic alteration phases. (Low temperature is defined as ranging from modern ambient seafloor conditions, or near 01C to up to
1501C, the low-temperature metamorphism limit.) These include X-ray-amorphous palagonite, the zeolites phillipsite, clinoptilolite, analcite (plus several others of rarer occurrence), smectites, and authigenic K-feldspar (Table 1). Phillipsite is a hydrated potassium- and sodium-rich aluminosilicate that forms elongated and sometimes twinned crystals from 8 to 250 mm length with numerous inclusions that indicate rapid growth. Oxygen isotope values of 34% for phillipsite indicate formation in the marine environment at modern seafloor temperatures. Phillipsite appears metastable on geological timescales, becoming rarer beyond Cenozoic age. The mineral is found globally at the sediment–sea water interface and continues to grow within the sediment column until it dissolves and disappears in the deepest
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Table 1
Summary of observed and suggested diagenetic reactions in deep-sea sediments
Opal-A-quartz Basaltic glass-palagonite Volcanic glass-smectite minerals Volcanic glass-zeolites Basalt-smectite minerals Basalt þ H4SiO4-smectite minerals Basaltic glass þ H4SiO4-phillipsite Poorly cyrstalline smectite-phillipsite Andesitic and rhyolitic glass þ H4SiO4-clinoptilolite Andesitic and rhyolitic glass-smectite minerals Phillipsite þ H4SiO4-clinoptilolite Phillipsite þ smectite þ H4SiO4-clinoptilolite þ palygorskite þ Opal-A þ Al(OH) 4 þ K -clinoptilolite Opal-A þ Al(OH)4 þ Kþ-opal-CT þ clinoptilolite Plagioclase þ Kþ þ H4SiO4-K-feldspar þ Naþ þ Ca2þ -K-feldspar ?-albite Clinoptilolite þ Naþ-analcite þ Kþ þ quartz Clinoptilolite- K-feldspar þ quartz Clinoptilolite þ Naþ-analcite þ K-feldspar þ quartz H4SiO4 þ Mg2þ þ Al(OH) 4 - palygorskite H4SiO4 þ Mg2þ-sepiolite Volcanic glass þ Mg2þ þ H4SiO4-palygorskite Smectite þ Mg2þ þ H4SiO4- palygorskite Clinoptilolite þ palygorskite þ calcite- K-feldspar þ dolomite þ quartz Clinoptilolite þ sepiolite þ calcite- K-feldspar þ dolomite þ quartz Amorphous hydroxides (mainly Fe) þ H4SiO4 þ Mg2þ-Fe-montmorillonite Nontronite þ Mg2þ þ reduced sulfur-saponite þ FeS2 Dissolved silica adsorption by clay minerals Amorphous aluminosilicate reconstitution? Smectite-mixed-layer illite–smectite? After Kastner (1981).
drill holes (usually below 500 m depth). Marine clinoptilolite forms finer grained (o45 mm), platy crystals of relatively silicon-rich hydrated potassiumand sodium-rich aluminosilicate (high Si/Al ratio). Clinoptilolite is more frequently encountered at depths 4100 m in the sediment column and persists to great depths and geological age, suggesting that it is thermodynamically stable in the deep-sea. Analcite is a rarer, sodium-rich zeolite in deep-sea sediments that displays a general trend of increasing abundance with geological age which parallels that of clinoptilolite. Both clinoptilolite and analcite can form directly from alteration of more siliceous volcanic glass (andesite, rhyolite, from active island arcs and explosive continental volcanism), by diagenetic reaction of phillipsite and dissolved silica (clinoptilolite) or Na/K exchange with clinoptilolite (analcite). Palygorskite and sepiolite Authigenic clay minerals (o2 mm particle size) in the deep-sea include the fibrous minerals palygorskite and sepiolite and the smectite family of expandable phyllosilicates. Both palygorskite and sepiolite are rare in recent marine
sediments, occuring more often in Eocene and older sediments. Palygorskite is a hydrous silicate containing Mg, Al and Fe3þ, whereas sepiolite is almost a pure hydrous magnesium silicate. Pore water solutions of alkaline pH with high concentrations of dissolved silica and magnesium favor the formation of both minerals. Fine fibrous textures and overgrowths of siliceous tests and opal-CT attest to their authigenesis. Reactions range from the diagenetic alteration of (mainly silicic and intermediate) volcanic ash, either directly or indirectly via smectite with biogenic silica, formation upon magnesium release after conversion of biogenic opal-A to opal-CT, and reaction of biogenic silica tests with marine pore waters, including hypersaline brines enriched in magnesium (Table 1). Smectites Most clay minerals in deep-sea sediments are detrital phases from continental weathering. These minerals compose the bulk of the nonbiogenic sediments in the o2 mm fraction and include clay-sized quartz and the phyllosilicates kaolinite, illite, chlorite, and smectite. The smectite
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267
Figure 9 Distribution of montmorillonite (smectite) abundance in the o2 mm fraction of sediments in the world ocean, carbonate-free basis. (After Griffin JJ, Windom H and Goldberg ED (1968) The distribution of clay minerals in the world ocean. Deep-Sea Research 15: 433–459.)
group of expandable clay minerals includes these common end-members: magnesium-rich saponite, found mostly as a diagenetic or low-temperature metamorphic product of basalt; iron-rich nontronite, often found as a low-temperature (o1001C) hydrothermal deposit; and aluminumrich beidellite predominantly derived from volcanic ash alteration on land. The surface distribution of marine smectite in the world’s oceans shows broad areas, such as in the south-eastern Pacific where 470% of the mineral’s abundance is far from land and blankets the active East Pacific Rise spreading
center (Figure 9). Most of this smectite is iron-rich montmorillonite. Up to 50% of this mineral has been described as authigenically formed from the seafloor temperature alteration of volcanic glass, with the remaining 50% from detrital sources. Early oxygen isotope studies of deep-sea montmorillonite suggested that much of it formed pedogenically on land, but iron-rich montmorillonite is not typically found in windblown dust. More recent oxygen isotope work has suggested either a lowtemperature submarine hydrothermal origin for the iron-montmorillonite or formation from low-
Fe2O3
Galapagos nontronite
East Pacific Rise Fe-mont NE Pacific Fe-mont OCP ridge Fe-mont
Galapagos Fe-mont
Detrital Al-mont /beidellite
MgO
Al2O3
Figure 10 Triangle plot of octahedral Mg–Al–Fe composition for marine smectites. (After McMurtry GM, Wang CH and Yeh HW (1983) Chemical and isotopic investigations into the origin of clay minerals from the Galapagos hydrothermal mounds field. Geochimica et Cosmochimica Acta 47: 475–489.)
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temperature hydrothermal alteration of ash within subaerial volcanoes that is periodically deposited in the ocean by explosive volcanism. Nontronite and iron-montmorillinite have also been suggested to form at seafloor temperatures by reaction of hydrothermally derived iron oxides with biogenic silica. One of the difficulties in assigning a precise origin to this mineral class is the possibility of physical admixtures of truly authigenic iron-rich smectites with more aluminum-rich smectites derived from the continents (Figure 10).
Glossary Aluminosilicate A silicate containing aluminum in coordination with oxyhydroxides and/or in substitution for silicon in SiO4 tetrahedra. Authigenesis New origin. Process of formation of new minerals in place. Detrital Formed from detritus, usually from rocks, minerals, or sediments from elsewhere than the depositional site. Diagenesis Changed origin. Recombination or rearrangement of a mineral that results in a new mineral, usually postdepositionally. Hemipelagic Deep-sea sediment that accumulates near the continental margin, so that the sediment contains abundant continentally derived material and rates of sedimentation are high. Hypersaline Excessive salinity, much greater than the normal salinity of sea water. Metalliferous Metal bearing, usually enriched toward economic extraction of the metals. Metastable State of a phase that is stable toward small disturbance, but is capable of reaction if sufficiently disturbed. Pedogenesis Soil origin. Mineral formation within the soil. Pelagic Open ocean environment. A marine sediment with that fraction derived from the continents indicating deposition from a dilute suspension distributed throughout deep-sea water. Phyllosilicate Layered or sheet silicate mineral, formed by sharing three of the four oxygens in neighboring silicon tetrahedra. Plankton Aquatic organisms that drift, or swim weakly. Can be either plants (phytoplankton) or animals (zooplankton). Redox Abbreviation for reduction–oxidation, usually expressed as a potential. Seamount Underwater mountain, 1000 m or higher elevation from seafloor base. Morphology may be peaked or flat-topped, with the latter called guyot.
Suboxic Condition lacking free oxygen, but not extremely reducing. Zeolite Any of the minerals of the zeolite group. Aluminosilicate minerals with an open framework structure that allows for easily reversible hydration, gas adsorption, and either cation or anion exchange.
See also Aeolian Inputs. Clay Mineralogy. Hydrothermal Vent Deposits. Manganese Nodules. Mineral Extraction, Authigenic Minerals. Platinum Group Elements and their Isotopes in the Ocean. Pore Water Chemistry. Rare Earth Elements and their Isotopes in the Ocean. River Inputs. Sediment Chronologies. Sedimentary Record, Reconstruction of Productivity from the. Tracers of Ocean Productivity. Transition Metals and Heavy Metal Speciation. Uranium-Thorium Decay Series in the Oceans Overview.
Further Reading Bentor YK (ed.) (1980) Marine Phosphorites; a Symposium. Oklahoma: SEPM Special Publication no. 29. Burns RG and Burns VM (1981) Authigenic oxides. The Sea, vol. 7, pp. 875--914. New York: Wiley. Chamley H (ed.) (1989) Clay Sedimentology. Berlin: Springer-Verlag. Cronan DS (1974) Authigenic minerals in deep-sea sediments. In: Goldberg ED (ed.) The Sea, vol. 5, pp. 491--525. New York: Wiley. Cronan DS (ed.) (1980) Underwater Minerals. London: Academic Press. Cronan DS (ed.) (2000) Handbook of Marine Mineral Deposits. Boca Raton, FL: CRC Press. Glasby GP (ed.) (1977) Marine Manganese Deposits. Elsevier Oceanography Series. Amsterdam: Elsevier. Glenn CR, Pre´vot-Lucas L, and Lucas J (eds.) (2000) Marine Authigenesis: from Global to Microbial. Oklahoma: SEPM Special publication no. 66. Halbach P, Friedrich G, and von Stackelberg U (eds.) (1988) The Manganese Nodule Belt of the Pacific Ocean: Geological Environment, Nodule Formation, and Mining Aspects. Stuttgart: F. Enke Verlag. Kastner M (1981) Authigenic silicates in deep-sea sediments: formation and diagenesis. In: Emiliani C (ed.) The Sea, vol. 7, pp. 915--980. New York: Wiley. Manheim FT (1986) Marine cobalt resources. Science 232: 600--608. Margolis SV and Burns RG (1976) Pacific deep-sea manganese nodules: their distribution, composition, and origin. Annual Review of Earth and Planetary Science 4: 229--263.
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BACTERIOPLANKTON H. W. Ducklow, The College of William and Mary, Gloucester Point, VA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 1, pp 217–224, & 2001, Elsevier Ltd.
Introduction Marine bacteria, unicellular prokaryotic plankton usually less than 0.5–1 mm in their longest dimension, are the smallest autonomous organisms in the sea – or perhaps in the biosphere. The nature of their roles in marine food webs and the difficulty of studying them both stem from their small size. A modern paradigm for bacterioplankton ecology was integrated into oceanography only following development of modern epifluorescence microscopy and the application of new radioisotopic tracer techniques in the late 1970s. It was not until a decade later, with the use of modern genomic techniques, that their identity and taxonomy began to be understood at all. Thus we are still in the process of constructing a realistic picture of marine bacterial ecology, consistent with knowledge of evolution, plankton dynamics, food web theory, and biogeochemistry. The lack of bacterioplankton compartments in most numerical models of plankton ecology testifies to out current level of ignorance. Nevertheless, much is now well known that was just beginning to be guessed in the 1980s. Bacterioplankton are important in marine food webs and biogeochemical cycles because they are the principal agents of dissolved organic matter (DOM) utilization and oxidation in the sea. All organisms liberate DOM through a variety of physiological processes, and additional DOM is released when zooplankton fecal pellets and other forms of organic detritus dissolve and decay. By recovering the released DOM, which would otherwise accumulate, bacterioplankton initiate the microbial loop, a complicated suite of organisms and processes based on the flow of detrital-based energy through the food web. The flows of energy and materials through the microbial loop can rival or surpass those flows passing through traditional phytoplankton-grazerbased food chains. For further information on the topics summarized here, the reader may consult the Further Reading, especially the recent book edited by Kirchman.
Identity and Taxonomy Most bacterial species cannot be cultivated in the laboratory and, until the development of cultureindependent genomic methods, the identity of over 90% of bacterial cells enumerated under the microscope was unknown. Only those few cells capable of forming colonies on solid media (agar plates) could be identified by classical bacteriological techniques. However, since the application of molecular genomic methods to sea water samples in the mid-1980s, our understanding of marine bacterial systematics and evolution has undergone a profound revolution. In this approach, plankton samples including bacterioplankton cells are collected and lysed to yield a mixture of DNA strands reflecting the genetic composition of the original assemblage. Then individual genes on the DNA molecules can be cloned and amplified via the polymerase chain reaction (PCR) for further analysis. Theoretically, any gene complex can be cloned, and several major groups of genes have been studied to date – for example, genes controlling specific biogeochemical transformations like ammonium oxidation, nitrogen fixation, sulfate reduction, and even oxidation of xenobiotic pollutant molecules. The most useful and widely studied genes for elucidating evolutionary relationships among bacterioplankton have been the genes coding for small subunit ribosomal RNA (SSU rRNA), because they evolve relatively slowly and their characters have been conserved across all life forms during the course of evolution. By sequencing the base pairs making up individual SSU rRNA molecules, the similarity of different genes can be established with great sensitivity. To date, nearly 1000 individual microbial SSU rRNA genes have been cloned and sequenced, yielding an entirely new picture of the composition of marine communities. The most important aspect of our understanding is that what we term ‘bacterioplankton’ really consists of two of the fundamental domains of life: the Bacteria and the Archaea (Figure 1). Domain Archaea is a group of microbial organisms with unique genetic, ultrastructural, and physiological characters that are about as different, genetically, from the Bacteria as either group is from higher life forms. Members of the Archaea may be typified by organisms from extreme habitats including anaerobic environments, hot springs, and salt lakes, but marine archaeal groups I and II are common in sea water. They make up about 10% of the microbial plankton in the
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Cytophaga group
lopha
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iae on m c io eu um su rad os hilus pn n i er us a t il sp nop ac cc ph m ob oco ydo obiu s lim r r e b n m Fi Dei hla omic myc leyi pallida C ruc cto sta era n a r a Ve Plairellul osph P Is m lla" inutu rix parvice m m h icrot obiu Atop idatus "M Cand Marine Actinobacteria Rubrobacter xylanophilus Nitrospira marina SAR2 Therm 02 omicro bium r Chlo oseum r o H f l The e erp eto xus aura rm s oto Geo ntiac ga tog iphon us aur Aq ma a su a uife ritim bte ntia rran xp cus yro a ea ph ilus n ci
n ra du
0.10 substitutions per nucleotide position
Mag
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spir
illum
mag
neto
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16 icum Azospirillu m brasilen se
tal lire m du dof ethylo c era x fe troph ens rme us nta ns SAR86 Bathymodiolus thermophilus gi ll sym. Oceanospirillum linum macleodii Alteromonas um SAR11 linar o sa us i r b i t dov sula Rho cap R83 us r e A ng act / S lo dob ter ter Rho c c a ob oba se r o h R yt Er Rho
Proteobacteria
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Cyanobacteria
sm a oc Arc ido ha ea ph Gr ilu Hal ou m pI oba I cte rium hal Halofe obi u rax vo lc anii m Methanosp irillum hun gatei Methanosarcina barkeri
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ii asch jann tae ler s u l c e m o ococ s v s c cu m han ccu occu mici hicu Met o c no o c or p tha erm m f otro Me h T teriu aut o ac ob erm n h ha m t et iu M er ct a ob an h et M
Archaea
Figure 1 Dendrogram showing relationships among the most widespread SSU rRNA gene clusters among the marine prokaryotes (the ‘bacterioplankton’). (Modified after Giovannoni SJ, in Kirchman (2000).)
surface waters of the oceans, and are relatively more numerous at greater depths, where they approach about half the total abundance. Since most of these organisms are known only from their RNA genes
and have never been cultured, their physiology and roles in the plankton are almost entirely unknown. Domain Bacteria contains all the familiar, culturable eubacterial groups and also a large number of
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BACTERIOPLANKTON
1000 _1
Knowledge of the nutrition and physiology of naturally occurring bacterioplankton as a functional group in the sea is based partly on laboratory study of individual species in pure culture, but mostly on sea water culture experiments. Traditional laboratory investigations show that bacteria can only utilize small-molecular-weight compounds less than B500 Daltons. Larger polymeric substances and particles must first be hydrolyzed by extracellular enzymes. In the sea water culture approach, samples with natural bacterioplankton assemblages are incubated for suitable periods (usually hours to a few days) while bacterial abundance is monitored, the utilization of various compounds with 14C- or 3Hlabelled radiotracers is estimated, and the net production or loss of metabolites like oxygen, CO2, and inorganic nutrients is measured. Such experiments, combined with size-fractionation using polycarbonate filters with precise and uniform pores of various diameter (0.2–10 mm), revealed that over 90% of added organic radiotracers are utilized by the smallest size fractions (o1 mm). Bacteria are overwhelmingly the sink for DOM in all habitats studied to date. Nutrient limitation of bacterial growth can
_2
Nutrition and Physiology
be identified by adding various compounds (e.g., ammonium, phosphate, or iron salts; monosaccharides and amino acids) singly or in combination to experimental treatments and comparing growth responses to controls. Using this approach, it has been learned that bacteria are effective competitors with phytoplankton for inorganic nutrients, including iron, which bacteria can mobilize by producing ironbinding organic complexes called siderophores. In general, bacterial growth in the sea, from estuaries to the central gyres, tends to be limited by organic matter. Sea water cultures most often respond to additions of sugars and amino acids, with the response sometimes enhanced if inorganic nutrients (including iron) are also added. At larger scales, the ultimate dependence of bacteria on organic matter supply is indicated by significant correlations between bacterial standing stocks or production (see below) and primary production (PP) across habitats (Figure 2). At withinhabitat scales and shorter timescales, significant relationships are less common, indicating time lags between organic matter production and its conversion by bacteria. Such uncoupling of organic matter production and consumption is also shown by transient accumulations of DOM in the upper ocean, where production processes tend to exceed utilization. It is not yet understood why DOM accumulates. Some fraction might be inherently refractory or rendered so by ultraviolet radiation or chemical condensation reactions in sea water. Deep ocean
Bacterial production (mg C m d )
unculturable, previously unknown groups. The main culturable groups include members of the Proteobacteria, marine oxygenic, phototrophic Cyanobacteria, and several other major groups including methylotrophs, planctomycetes, and the Cytophaga– Flavobacterium–Bacteroides group. But the most abundant genes recovered so far are not similar to those of the known culturable species. These include the most ubiquitous of all groups yet recovered, the SAR-11 cluster of the alpha Proteobacteria, which have been recovered from every bacterial clone library yet isolated. It appears to be the most widely distributed and successful of the Bacteria. The photosynthetic Cyanobacteria, including Synecchococcus spp. and the unicellular prochlorophytes, are functionally phytoplankton and they dominate the primary producer populations in the open sea, and at times in coastal and even estuarine regimes. They are treated elsewhere in this encyclopedia, so our discussion here is limited to heterotrophic forms of Bacteria and to the planktonic Archaea, although we cannot specify what many (or most) of them do. Genomic techniques are now being used to investigate bacterial and archaeal species succession during oceanographic events over various timescales, much as phytoplankton and higher organism successions have been observed for a century or more.
271
100
10
1 100
1000
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_1
Primary production (mg C m d ) Figure 2 Bacterial production plotted against primary production for the euphotic zone in several major ocean regimes or provinces. The overall data set has a significant regression, but the individual regions do not. J, Sargasso Sea; K, Arabian Sea; B, equatorial Pacific; n, equatorial Pacific.
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DOM has a turnover time of centuries to millennia, and seems to become labile (vulnerable to bacterial attack) when the ocean thermohaline circulation returns it to the illuminated surface layer. Alternatively, bacterial utilization of marine DOM, which generally has a high C:N ratio, might be limited by availability of inorganic nutrients. The latter hypothesis is supported by observations that DOM accumulation tends to be greater in the tropics and subtropics, where nitrate and phosphate are depleted in surface waters. The efficiency with which bacteria convert organic matter (usually expressed in carbon units) into biomass can be estimated by comparing the apparent utilization of individual compounds or bulk DOM with increases in biomass or with respiration. Respiration is usually measured by oxygen utilization but precise new analytical techniques for measuring carbon dioxide make CO2 production a preferable approach. Bacterial respiration (BR) is difficult to measure because water samples must first be passed through filters to remove other, larger respiring organisms, and because the resulting respiration rates are low, near the limits of detection of oxygen and CO2 analyses. It is also not easy to estimate bacterial biomass precisely (see below). The conversion efficiency or bacterial growth efficiency (BGE) is the quotient of net bacterial production (BP) and the DOM utilization: BGE ¼
BP BP ¼ DDOM BP þ BR
½1
Bacteria have rather uniform biomass C:N composition ratios of 4–6. Intuitively, it seems reasonable to expect that they would utilize substrates with high C:N ratios at lower efficiency. Enrichment cultures initiated from natural bacterial assemblages grow in sea water culture in the laboratory on added substances with efficiencies of 30–90%. The BGE is inversely related to the C:N ratio of the organic substrate if just a single compound is being utilized, but when a mixture of compounds is present, as is probably always the case in the environment, there is no discernible relationship between the chemical composition of the materials being used and the BGE. In the open ocean, BGE averages about 10–30%, a relatively low value that has important implications for our understanding and modeling of organic matter turnover and ocean metabolism. At larger scales, BGE appears to increase from B10% to 50% along an offshore-to-onshore gradient of increasing primary productivity, probably reflecting greater organic matter availability. This pattern has been used
to support an argument suggesting that in lake and oceanic systems with the lowest primary productivity, respiration exceeds production; that is, such oligotrophic systems might be net heterotrophic. This possibility has also been supported by results from careful light–dark bottle studies in which oxygen consumption exceeds production. This finding, however, is inconsistent with a large amount of geochemical evidence, for instance showing net oxygen production at the basin and seasonal to annual scale. Resolution of this debate probably rests on improved estimates of BGE. Pure culture, sea water culture, and the latest genomic studies indicate fundamental metabolic and genetic differences among different bacterial populations, which can generally be grouped into two broad classes based on organic matter utilization. Native marine bacteria capable of utilizing DOM at concentrations below 100 nmol l1, termed oligotrophs, cannot survive when DOM is greater than about 0.1–1 mmol l1. Copiotrophic bacteria found in some habitats with higher ambient DOM levels thrive on concentrations far exceeding this threshold. Observations that copiotrophs shrink and have impressive survival capability under severe starvation conditions (thousands of days to, apparently, centuries) led some investigators to suggest that the dominant native marine bacteria are starving (nongrowing) copiotrophs in a survival mode, awaiting episodes of nutrient enrichment. A variable fraction of the total population usually does appear to be dormant, as indicated by autoradiography, vital staining, and RNA probes, but the timescales of the transition from active growth to dormancy and back again are not well defined. Maintenance of dormant cells in a population depends on strong predator preferences for actively growing cells and prey selection against the nongrowing cells. Most oligotrophs so far isolated in the laboratory under stringent low-DOM conditions appear to be unrelated to known bacterial groups.
Bacterial Biomass, Growth, and Production The standing stock of bacteria is still most commonly assessed by epifluorescence microscopy, following staining of the cells with a fluorochrome dye. Flow cytometric determination is gradually taking over, and has several key advantages over microscopy: faster sample processing, improved precision, and discrimination of heterotrophic and phototrophic bacteria. There is a gradient in bacterial abundance proceeding from B1010 cells l1 in estuaries to 109
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BACTERIOPLANKTON
cells l1 in productive ocean regimes and 108 cells l1 in the oligotrophic gyres (Figure 3). These horizontal gradients parallel gradients in primary production and organic matter fluxes, suggesting the overall importance of bottom-up controls on bacterial abundance. Chlorophyll a concentrations, indicative of phytoplankton biomass, vary somewhat more widely than bacterial abundance over basin to global scales, but within habitats, the variability of bacterial and phytoplankton biomass is about equal, reflecting the generally close coupling between the two groups and the similarity of removal processes (grazing, viral lysis, unspecified mortality) acting on them. It is more difficult to estimate bacterial biomass, because we cannot measure the mass (e.g., as carbon) directly, and have to convert estimates of cell volumes to carbon instead. The best estimates now indicate 7–15 1015 g C cell1 for oceanic cells and 15–25 1015 g C cell1 for the slightly larger cells found in coastal and estuarine habitats. Thus the biomass gradient is steeper than the abundance gradient because the cells are larger inshore. Table 1 shows data compiled from Chesapeake Bay and the Sargasso Sea off Bermuda, two well-studied sites that illustrate the contrasts in phytoplankton and bacterioplankton from a nutrient-rich estuary to the oligotrophic ocean gyres. Bacterial and phytoplankton biomass are much greater in the estuary, as
10
Cells l
_1
10
9
10
expected. Interestingly, assuming a mean euphotic zone depth of 1 m in the Bay and 140 m off Bermuda, we find that the standing stocks of bacteria in these euphotic zones are B10 and 50 mmol C m2 in the estuaries and open sea, respectively. The oceanic euphotic zone is somewhat more enriched in bacteria than the more productive estuaries. Bacterial and phytoplankton stocks are nearly equal in the open sea, but phytoplankton exceeds bacterial biomass inshore. Carbon from primary producers appears to be more efficiently stored in bacteria in oceanic systems compared to estuarine ones. Bacterial stocks in different environments can be assessed using the relationship Bmax ¼ F=m
10
94
136
30
37
27
162
H
aw ai i Sa rg as so R os s Se G a ul fS Eq tre am ua to ria lP ac N . C A he tla sa nt ic pe ak e Ba y
66
Figure 3 Bacterial abundance in the euphotic zone of several major ocean provinces. The box plots show the median, 10th, 25th, 75th and 90th centiles of the data. The number of samples is listed for each region. There is no statistical difference among the regions except for Chesapeake Bay.
½2
where Bmax is the carrying capacity in the absence of removal, F is the flux of utilizable organic matter to the bacteria, and m is their maintenance efficiency (the specific rate of utilization when all of F is used to meet cellular maintenance costs, with nothing left for growth). The problem is specifying values for F and m. The DOM flux can be evaluated by flow analysis and is about 20–50% of the net primary production (NPP) in most systems. Maintenance costs are poorly constrained and possibly very low if most cells are near a starvation state, but 0.01 d1 is a reasonable value for actively growing cells. Thus for the oligotrophic gyres where the latest NPP estimates are about 200–400 mg C m2 d1, we can calculate that Bmax should be about 4–8 109 cells l1, an order of magnitude greater than observed. Removal processes must maintain bacterial stocks considerably below their maximum carrying capacity. Bacteria convert preformed organic matter into biomass. This process is bacterial production, which can be expressed as the product of the biomass and the specific growth rate (m) BP ¼ dB=dt ¼ mB
8
273
½3
Like biomass, BP cannot be measured directly in mass units. Instead, metabolic processes closely coupled to growth are measured and BP is derived using conversion factors. The two most common methods follow DNA and protein synthesis using (3H)thymidine and (3H)leucine incorporation rates, respectively. The values for the conversion factors are poorly constrained and hard to measure, leading to uncertainty of at least a factor of two in the BP estimates. Few measurements were performed in the open sea before the 1990s. The Joint Global Ocean Flux Study (JGOFS) time-series station at Bermuda is perhaps the best-studied site in the ocean (Table 1).
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BACTERIOPLANKTON
Table 1
The biomass (B) and production rates (P) of bacterioplankton and phytoplankton at estuarine and open ocean locationsa Biomass (mmol m3)
Production rate (mmol m3 d1)
P/B (d 1)
Location
Phytoplankton
Bacteria
Phytoplankton
Bacteria
Phytoplankton
Bacteria
Chesapeake Bay
5–400 (56)
1–80 (11)
20–47 (33)
0.1–50 (4)
0.07–1.9
0.01–2 (0.34)
Sargasso Sea
0.3–3.2 (1.0)
0.2–0.6 (0.4)
0.06–0.9 (0.3)
0.002–0.07 (0.02)
0.1–1 (0.3)
0.01–0.16 (0.06)
a The values are annual, euphotic zone averages derived from published reports. P/B is the specific turnover rate for the population. The data are presented as ranges with the mean of various estimates in parentheses. Ranges encompass observations and assumptions about conversion factors for deriving values from measurements (see text).
In the open sea, far removed from allochthonous inputs of organic matter, we can compare BP and PP directly, since all the organic matter ultimately derives from the PP. One difficulty is that BP itself is not constrained by PP, since if the recycling efficiency of DOM and the BGE are sufficiently high, BP can exceed PP. BP also commonly exceeds local PP in estuaries, where inputs of terrestrial organic matter are consumed by bacteria. Bacterial respiration, however, cannot exceed the organic matter supply and serves as an absolute constraint on estimates of BP. But as noted above, bacterial respiration is very hard to measure and there are many fewer reliable measurements than for BP itself. BR is usually estimated from the BGE. Rearranging eqn [1], BR ¼
ð1 BGEÞBP BGE
½4
Most commonly, variations of eqn [1] have been used to estimate the total bacterial carbon utilization or demand (BCDBR þ BP) from estimates or assumptions about BP and BGE. Earlier estimates and literature surveys suggested that BP was as high as 30% of PP. Combining this value with a BGE of 20% yields a BCD of 1.5 times the PP. This estimate in itself is possibly acceptable, if recycling of DOM is high, but then eqn [4] yields a BR of 1.2 times the total PP – an impossibility. More recent estimates of BP, typified by the Sargasso Sea data, suggest BP is about 10% of PP in the open sea. Applying this value and the mean BGE for the region (0.14), we find that BR consumes about 55% of the primary production in the Sargasso Sea, still a substantial figure. Similar calculations for other well-studied ocean areas suggest that zooplankton (including protozoans and microzooplankton) and bacteria consume nearly equal amounts of the total primary productivity. These estimates illustrate the biogeochemical importance of bacterioplankton in the ocean carbon cycle: although their growth efficiency is low,
bacteria process large amounts of DOM. DOM produced by a myriad of ecological and physiological processes must escape bacterial metabolism to enter long-term storage in the oceanic reservoir.
Role in Food Webs and Biogeochemical Cycles The process of bacterivory (consumption of bacteria by bacteriovores) completes the microbial loop. Bacterioplankton cells are ingested by a great diversity of predators, but, because of the small size of the prey, most bacteriovores are small protozoans, typically o5 mm nanoflagellates and small ciliates. Bacterial cells only occupy about 107 of the volume of the upper ocean, indicating the difficulty of encountering these small prey. Larger flagellates, small ciliates, and some specialized larger predators can also ingest bacterial prey. The most important of the larger predators are gelatinous zooplankton like larvaceans, which use mucus nets to capture bacterial cells sieved from suspension. But most bacteriovores are also very small. Nanoflagellates can clear up to 105 body volumes per hour, thus making a living from harvesting small, rare bacterial prey, and generally dominating bacterivory in the sea. Protozoan bacterivory closely balances BP in lessproductive oceanic regimes. Most crustacean zooplankton cannot efficiently harvest bacterioplankton unless the latter are attached to particles, effectively increasing their size. Bacterial prey enter marine food webs following ingestion by flagellates, and ingestion of the flagellates by other flagellates, ciliates, and copepods. This means that bacteria usually enter the higher trophic levels after several cycles of ingestion by consumers of increasing size, with attendant metabolic losses at each stage. The microbial loop and its characteristic long, inefficient food chains can be short-circuited by the gelatinous bacteriovores, which package bacterial cells into larger prey.
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BACTERIOPLANKTON
Compared to phytoplankton and to bacteriovores, bacteria are enriched relative to body carbon in nitrogen, phosphorus, protein, nucleic acids, and iron. Their excess nutritional content, coupled with the many trophic exchanges that bacterial biomass passes through as it moves in food webs, means that the microbial loop is primarily a vehicle for nutrient regeneration in the sea rather than an important source of nutrition for the upper trophic levels. The main function of bacteria in the microbial loop is to recover ‘lost’ DOM, enrich it with macro- and micronutrients, and make it available for regeneration and resupply to primary producers. Lower bacterial production estimates (see above) would also tend to decrease the importance of bacteria as a subsidy for higher consumers. In estuaries and other shallow near-shore habitats, BP is not as closely balanced by planktonic bacteriovores as in ocean systems. In these productive habitats, bacteria are larger and more often associated with particles, so they are vulnerable to a wider range of grazers. Bacteria can also be consumed by mussels, clams, and other benthic suspension feeders. External subsidies of organic matter mean that BP is higher inshore, so bacteria are a more important food source in coastal and estuarine food webs than in oceanic waters. In these productive systems where bacterial abundance is greater, more of the bacterial stock is also attacked and lysed by viruses, resulting in release of DOM and nutrients instead of entry into food webs. The relative importance of viruses and bacteriovores in removing bacteria is not yet well known, but has important implications for food web structure. Bacteria are the major engines of biogeochemical cycling on the planet, and serve to catalyze major transformations of nitrogen and sulfur as well as of carbon. They participate in the carbon cycle in several ways. Their principal role is to serve as a sink for DOM, and thus regulate the export of DOM from the productive layer. Bacteria also have intensive hydrolytic capability and participate in decomposition and mineralization of particles and aggregates. Bacteria rapidly colonize fresh particulate matter in the sea, and elaborate polymeric material that helps to cement particles together, so they both reduce particle mass by enzymatic hydrolysis and promote particle formation by fostering aggregation. The balance of bacterial activity for forming particles and accelerating particle sedimentation or, in contrast, decomposing particles and reducing it, is not clear. Larvaceans and other giant, specialized bacteriovores, centimeters to meters in size, can repackage tiny bacterial cells into large, rapidly sinking
275
fecal aggregates, thus feeding the ocean’s smallest organisms into the biological carbon pump.
Conclusions Knowledge of the dynamics of bacterioplankton, their identity, roles in food webs and biogeochemical cycles is now becoming better known and integrated in a general theory of plankton dynamics, but these aspects are not yet common in plankton ecosystem models. The differential importance of bacteria in plankton food webs in coastal and oceanic systems might serve as a good test of our understanding in models. The dynamics of DOM are only crudely parametrized in most models, and explicit formulation of bacterial DOM utilization may help in better characterizing DOM accumulation and export. Other interesting problems such as the effects of size-selective predation, bacterial community structure, and species succession are just beginning to be explored. Exploration of marine bacterial communities together with molecular probes and numerical models should lead to a new revolution in plankton ecology.
See also Carbon Cycle. Copepods. Hydrothermal Vent Deposits. Hydrothermal Vent Fauna, Physiology of. Marine Mammal Trophic Levels and Interactions. Microbial Loops. Network Analysis of Food Webs. Primary Production Methods. Primary Production Processes. Protozoa, Planktonic Foraminifera. Protozoa, Radiolarians. Small-Scale Physical Processes and Plankton Biology.
Further Reading Azam F (1998) Microbial control of oceanic carbon flux: the plot thickens. Science 280: 694--696. Carlson C, Ducklow HW, and Sleeter TD (1996) Stocks and dynamics of bacterioplankton in the northwestern Sargasso Sea. Deep-Sea Research II 43: 491--516. del Giorgio P and Cole JJ (1998) Bacterial growth efficiency in natural aquatic systems. Annual Review of Ecological Systems 29: 503--541. Ducklow HW and Carlson CA (1992) Oceanic bacterial productivity. Advances in Microbial Ecology 12: 113--181. Kemp PF, Sherr B, Sherr E, and Cole JJ (1993) Handbook of Methods in Aquatic Microbial Ecology. Boca Raton, FL: Lewis Publishers. Kirchman DL (2000) Microbial Ecology of the Oceans. New York: Wiley. Pomeroy LR (1974) The ocean’s food web, a changing paradigm. BioScience 24: 499--504.
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BALEEN WHALES J. L. Bannister, The Western Australian Museum, Perth, Western Australia Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 1, pp 224–235, & 2001, Elsevier Ltd.
Diagnostic Characters and Taxonomy Baleen or whalebone whales (Mysticeti) comprise one of the two recent (nonfossil) cetacean suborders. They differ from the other suborder (toothed whales, Odontoceti), particularly in their lack of functional teeth. Instead they feed on relatively very small marine organisms by means of a highly specialized filter-feeding apparatus made up of baleen plates ‘whalebone’ attached to the gum of the upper jaw. Other differences from toothed whales include the baleen whales’ paired blowhole, symmetrical skull, and absence of ribs articulating with the sternum. Baleen whales are generally huge. In the blue whale they include the largest known animal, growing to more than 30 m long and weighing more than 170 tonnes. Like all other cetaceans, baleen whales are totally aquatic. Like most of the toothed whales, they are all marine. Many undertake very long migrations, and some are fast swimming. A few species come close to the coast at some part of their life cycle and may be seen from shore; however, much of their lives is spent remote from land in the deep oceans. Baleen whale females grow slightly larger than the males. Animals of the same species tend to be larger in the Southern than in the Northern Hemisphere. Within the mysticetes are four families: right whales (Balaenidae, balaenids), pygmy right whales (Neobalaenidae, neobalaenids), gray whales (Eschrichtiidae, eschrichtiids); and ‘rorquals’ (Balaenopteridae, balaenopterids). Within the suborder, 12 species are now generally recognized (Table 1). Right whales are distinguished from the other three families by their long and narrow baleen plates and arched upper jaw. Other balaenid features include, externally, a disproportionately large head (c. one-third of the body length), long thin rostrum, and huge bowed lower lips; they lack multiple ventral grooves. Internally, there is no coronoid process on the lower jaw and cervical vertebrae are fused together. Within the family are two distinct groups: the bowhead (Balaena mysticetus) of northern polar waters (formerly known as the ‘Greenland’ right
276
whale) and the three ‘black’ right whales (Eubalaena spp.) of more temperate seas (so called to distinguish them from the ‘Greenland’ right whale). (Figure 1). All balaenids are robust. Pygmy right whales (Capnea marginata) have some features of both right whales and balaenopterids. The head is short (c. one-quarter of the body length), although with an arched upper jaw and bowed lower lips, and there is a dorsal fin. The relatively long and narrow baleen plates are yellowish-white, with a dark outer border, quite different from the all-black balaenid baleen plates. Internally, pygmy right whales have numerous broadened and flattened ribs. Gray whales (Eschrichtius robustus) are also somewhat intermediate in appearance between right whales and balaenopterids. They have short narrow heads, a slightly arched rostrum, and between two and five deep creases on the throat instead of the balaenopterid ventral grooves. There is no dorsal fin, but a series of 6 to 12 small ‘knuckles’ along the tail stock. The yellowish-white baleen plates are relatively small. Balaenopterids comprise the five whales of the genus Balaenoptera blue, B. musculus; fin, B. physalus; sei, B. borealis; Bryde’s, and minke, B. acutorostrata & B. bonaerensis; whales and the humpback whale (Megaptera novaeangliae). All have relatively short heads, less than a quarter of the body length. In comparison with right whales, the baleen plates are short and wide. Numerous ventral grooves are present, and there is a dorsal fin, sometimes rather small. Internally, the upper jaw is relatively long and unarched, the mandibles are bowed outward, and a coronoid process is present; cervical vertebrae are generally free. All six balaenopterids are often known as ‘rorquals’ (said to come from the Norse ‘whale with pleats in its throat’). Strictly speaking, the term should probably be applied to the five Balaenoptera species, recognizing the rather different humpback in its separate genus, but many authors now use it for all six balaenopterids. Baleen whales are sometimes called ‘great whales.’ Despite their generally huge size, some of the species are relatively small, and it seems preferable to restrict the term to the larger mysticetes (blue, fin, sei, Bryde’s, humpback) together with the largest odontocete (the sperm whale, Physeter macrocephalus) (Figure 2). In a recent review of the systematics and distribution of the world’s marine mammals, Rice (1998)
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BALEEN WHALES
Table 1
277
Mysticetes (baleen whales)
Family
Genus
Species
Balaena Eubalaena
B. mysticetus E. glacialis
Balaenidae
E. australis E. japonica Neobalaenidae Neobalaena Caperea marginata Eschrichtiidae Eschrichtius E. robustus Balaenopteridae Megaptera M. novaeangliae BalaenopteraB. acutorostrata
B. bonaerensis B. edeni B. borealis B. physalus B. musculus
Subspecies
Common name
Maximum length (m)
Right whales Bowhead whale 19.8 North Atlantic right 17.0 whale Southern right 17.0 whale North Pacific right 17.0 whale Pygmy right whales Pygmy right whale 6.4 Gray whales Gray whales 14.1 ‘Rorquals’ Humpback whale 16.0 Common minke whale B. a. N. Atlantic minke 9.2 acutorostrata whale B. a. N. Pacific minke ? scammoni whale B. a. subsp. Dwarf minke whale? Antarctic minke whale Bryde’s whale Sei whale Fin whale Blue whale Blue whale
10.7 14.0 17.7 26.8
B. m. 26.0 musculus B. m. indica Great Indian ? rorqual B. m. Pygmy blue whale 24.4 brevicauda B. m. ‘True’ blue whale 30.5 intermedia
has drawn attention to the derivation of the Latin word Mysticeti and clarified the status of a variant, Mystacoceti. He describes the former as coming from Aristotle’s original Greek mustoketos, meaning ‘the mouse, the whale so-called’ or ‘the mousewhale’ (said to be an ironic reference to the animals’ generally vast size). Mystacoceti means ‘moustache whales,’ and although used occasionally in the past (and more obviously appropriate for whales with baleen in their mouths), it has been superseded by Mysticeti. The 12 species in Table 1 differ somewhat from those listed by Rice. Some authors disagree with his use of the genus Balaena for Eubalaena and his preference for the single species glacialis rather than
Generalized distribution
Circumpolar in the Arctic Temperate–Arctic Temperate–N. Atlantic Temperate N. Pacific
Temperate, Southern Hemisphere only
North Pacific–Arctic Worldwide Worldwide Temperate–Arctic Temperate–Arctic Temperate–subantarctic, Southern Hemisphere only Temperate–Antarctic Circumglobal, tropical–subtropical Worldwide, largely temperate Worldwide Worldwide N. Atlantic, N. Pacific N. Indian Ocean Southern Hemisphere, temperate– subAntarctic Southern Hemisphere, temperate– Antarctic
the three species, Eubalaena glacialis, E. japonica and E. australis (the North Atlantic, North Pacific and southern right whales). While acknowledging the need for further investigation, they refer to present-day biologists’ usage, and genetic information, in preferring a separation between Northern and Southern Hemisphere animals and in recognizing two species in the Northern Hemisphere; Eubalaena is, however, the only mysticete genus where one or more separate species is recognized in each hemisphere. Rice also distinguishes between two species of Bryde’s whale: Balaenoptera edeni and B. brydei. The taxonomic status of these ‘inshore, smaller’ and ‘offshore, larger’ forms has yet to be determined and here they are subsumed within B. edeni. In the case
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278
BALEEN WHALES
Bowhead
Right
Pygmy right
Gray
Blue
Minke
Humpback
Figure 1 Lateral profiles of representative baleen whales, with a human figure, to scale.
of the blue whale, Rice’s inclusion of a northern Indian Ocean form (B. m. indica, referred to by Rice as ‘the great Indian rorqual’) has been followed. Similarly, his listing of three subspecies of minke whale, including the Southern Hemisphere dwarf minke, which has yet to be formally described, has been retained. However, other subspecies, e.g., two sei whales and two fin whales, have not been included.
Baleen plates
Tongue
Distribution and Ecology Habitat
Baleen fringes Figure 2 Head of a right whale, showing the arrangement of the filter-feeding apparatus (from Bonner, 1980).
In addition the subspecies listed in the previous section, many stocks or populations have been
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BALEEN WHALES
recognized, some mainly for management purposes, based on more or less valid biological grounds. Some significant examples include the following. Bowhead whales In addition to the currently most abundant population (the Bering-Chukchi-Beaufort Seas stock), four others are recognized: Baffin Bay/ Davis Strait, Hudson Bay, Spitzbergen, and Okhotsk Sea. Right whales In the North Atlantic species, two populations are currently recognized, western and eastern, with calving grounds off the southeastern United States and northwestern Africa. The latter may now represent only a relict population(s). In the North Pacific species, the current view is that there well may once have been two or more stocks, based on feeding ground information: at least one now centered in summer on the Sea of Okhotsk and another, although possibly not now a functioning unit, summering in the Gulf of Alaska. In the southern right whale, there are several populations, defined by currently occupied calving grounds, but these cover only a proportion of the many areas known from historical whaling records to have once been occupied by right whales. Up-todate information is available on presumed discrete populations off eastern South America, South Africa, southern Australia, and sub-Antarctic New Zealand. Gray whales A western North Atlantic population may have persisted until the 17th or 18th centuries, but is now extinct. The species now survives only in the North Pacific, where, in addition to a flourishing ‘Californian’ stock, wintering on the coast of Baja California, and summering in the Bering Sea, animals are now being reported from a remnant western stock, summering in the northern Okhotsk Sea. Humpback whales In the North Atlantic, two major populations are recognized: one based on animals wintering in the West Indies and the other, now possibly only a relict population, wintering around the Cape Verde Islands. In the North Pacific, three discrete wintering grounds have been recorded: around the Bonin, Mariana, and Marshall Islands in the west; around the Hawaian Islands in the center; and off Mexico in the east. In the Southern Hemisphere, seven populations have been postulated. Six are well defined, based on calving (wintering) grounds either side of each continent (one off eastern Australia is closely related to animals wintering off Fiji and Tonga), and a possible seventh in the central Pacific. In the northwest Indian
279
Ocean, there seems to be a separate population where animals have been reported present throughout the year. Baleen whales thus occupy a wide variety of habitats, from open oceans to continental shelves and coastal waters, from the coldest waters of the Arctic and Antarctic, through temperate waters of both hemispheres and into the tropics. Most specialized is the bowhead, Balaena, restricted to the harsh cold and shallow seas of the Arctic and sub-Arctic. The black right whales (Eubalaena) are more oceanic and prefer generally temperate waters, but come very close to coasts in winter to give birth, particularly in the Southern Hemisphere. Once believed not to penetrate much further south than the Antarctic convergence (c. 50– 551S), there have been recent records in the Antarctic proper, south of 601S. Whether this is a new phenomenon is unclear: a report by Sir James Clark Ross of many ‘common black’ (i.e., right) whales in the Ross Sea (eastern Antarctic) at 631S in December 1840 was discounted when their presence there later that century could not be confirmed. It has been suggested that the currently greatly reduced population of the western North Atlantic right whale, now wintering off the south eastern United States and summering in coastal waters north to the Bay of Fundy (c. 451N), may represent the peripheral remnant of a more widely distributed stock, formerly summering north to Labrador and southern Greenland, i.e., to at least 601N (Figure 3). The pygmy right whale (Caperea) is restricted to Southern Hemisphere temperate waters, between about 301 and 521S; it can be found coastally in winter in some areas, and all-year round in others. Gray whales (Eschrichtius) are the most obviously coastal baleen whales. The long coastal migration of the ‘Californian’ stock, from Mexico to Alaska, supports a major whale-watching industry from December to April. In spring the animals migrate through the Bering Strait into the more open waters of the Bering Sea, but still favoring more shallow waters. Among the balaenopterids, fin and sei whales are probably the most oceanic, with the former penetrating into colder waters than the latter in summer. Blue whales can be found closer inshore, but are often associated with deep coastal canyons, e.g., off central and southern California. The Southern Hemisphere pygmy blue whale (subspecies B. m. brevicauda) has been regarded as restricted to more temperate waters than the ‘true’ blue whale (B. m. intermedia), not often being found much beyond 551S. The most coastal balaenopterid is the humpback (Megaptera), with long migrations between
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280
BALEEN WHALES
Breeding cycle Month
Zone of migrations
Antarctic feeding zone
20˚ Mean latitude south
Probable range of breeding grounds
1st year A M
2nd year
Impregnation
Calving Pregnent
C
I 20˚
Pregnant in opposite phase
30˚ 40˚ 50˚
Mother and calf
Mother and calf in opposite phase
30˚
Weaning Resting
60˚
2-year cycle of adult female
60˚ Latitude of pack ice
Gestation
Lactation
Resting
Fetal growth
Neonatal growth (suckling)
Independent feeding
Birth Conception
70˚ Feeding
Feeding
14 12 10 8 6 4 2 0
40˚ 50˚
70˚
Length of fetus and calf (m)
New cycle
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 Winter Summer Winter Winter Summer
Weaning
Accelerated growth
14 12 10 8 6 4 2 0
Grow th of fetus and calf
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
Figure 3 ‘Typical’ life cycle of a southern baleen whale (as modified by Bonner, 1980, from Mackintosh, 1965).
temperate/tropical breeding grounds and cold-water feeding grounds. In the Southern Hemisphere, much of its journey occurs along the east and west coasts of the three continents. In the Northern Hemisphere, humpbacks are rather more oceanic, but still coastal at some stage in their migration: in the North Pacific they can be found wintering off the Hawaiin Islands and summering off Alaska and in the western North Atlantic they winter in the Caribbean and summer between New England, the west coast of Greenland, and Iceland. Minke whales are wide ranging, from polar to tropical waters in both hemispheres. In the Southern Hemisphere they can, with blue whales, be found closest to the ice edge in summer. Elsewhere they can often occur near shore, in bays and inlets. Their migrations are less well defined and predictable than other migratory balaenopterids; in some regions they are present year-round. The most localized balaenopterid is Bryde’s whale. It is the only species restricted entirely to tropical/ warm temperate waters and probably does not undertake long migrations. The two forms – inshore and offshore, in several areas – can differ in their movements. Off South Africa, for example, the inshore form is thought to be present throughout the year, whereas the offshore form appears and disappears seasonally, presumably in association with movements of its food, shoaling fish. Food and Feeding
Although they include the largest living animals, baleen whales feed mainly on very small organisms
and are strictly carnivorous, feeding on zooplankton or small fish. In ‘filter-feeding’ – sieving the sea – baleen whales are quite different from toothed whales, where the prey is captured individually. Filter feeding has been described as requiring, in addition to a supply of food in the water, three basic features: a flow of water to bring prey near the mouth; a filter to collect the food but allow water to pass through; and a means of removing the filtered food and conveying it to the stomach for digestion. Baleen whales meet those requirements by (a) seeking out areas where their food concentrates, (b) either swimming open-mouthed through food or gulping it in, (c) possessing a highly efficient filter formed by the baleen plates, and (d) forcing the water containing the food out through the baleen plates and then transfering the trapped food back to the gullet and hence to the stomach. In the latter the tongue is presumed to be involved; in balaenopterids the process is aided by the distensible throat. While all baleen whales possess a filter based on baleen plates, two rather different systems – essentially ‘skimming’and ‘gulping’ – have evolved to filter a large volume of water containing food. Each relies on a series of triangular baleen plates, borne transversely on each upper jaw. The inner, longer border (hypotenuse) of each plate bears a fringe of fine hairs, forming a kind of filtering ‘doormat.’ Quite unrelated to teeth (which appear as early rudiments in the gums of fetal baleen whales), baleen is closest in structure to mammalian hair and human fingernails. In the right whales, filtration is achieved with very long and narrow plates in the very large mouth, itself carried in the very large head. The plates, up to 4 m long in bowheads and
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BALEEN WHALES
2.7 m in the right whales, are accommodated in the mouth by an arched upper jaw and are enclosed in massively enlarged and upwardly bowed lower lips. There is a gap between the rows at the front of the mouth, and the whole arrangement allows the whale to scoop up a great quantity of water while swimming slowly forward. In balaenopterids, with their much smaller heads, the baleen plates are shorter and broader and the rows are continuous at the front. Taking in a large volume of water and food is usually achieved by swimming through a food swarm and gulping, while simultaneously enlarging the capacity of the mouth greatly by extending the ventral grooves and depressing the tongue. The two systems allow, on the one hand, the relatively slow-swimming balaenids to concentrate their rather sparse slow-swimming food over a period, and on the other, the faster-swimming balaenopterids to take in large amounts of their highly concentrated fast-swimming prey over a shorter time. Typically, baleen whales feed on zooplankton, mainly euphausiids or copepods, swarming in polar
Table 2
Baleen Whale Food Items
Species
Subspecies
or subpolar regions in summer. That is particularly so in the Southern Hemisphere, where the summer distributions of several balaenopterids depend on the presence of Euphausia superba known to whalers by the Norwegian word ‘krill’ (see Krill) in huge concentrations in the Antarctic. In the Northern Hemisphere, with a more variable availability of food, balaenopterids are more catholic in their feeding. Humpbacks and fin whales, for example, feeding almost exclusively on krill in the south, commonly take various species of schooling fish in the north. The variety of organisms taken by the various species in different regions is listed in Table 2. While most feeding occurs in colder waters, baleen whales may feed opportunistically elsewhere. All baleen whales but one, the gray whale, feed generally within 100 m of the surface and, consequently, unlike many toothed whales, do not dive very deep or for long periods. Gray whales feed primarily on bottom-living organisms, almost exclusively amphipods, in shallow waters.
Common name
Food items Northern Hemisphere
B. mysticetus
Southern Hemisphere
Bowhead whale
Mainly calanoid copepods; euphausiids; occasional mysids, amphipods, isopods, small fish North Atlantic right whale Calanoid copepods; euphausiids Southern right whale Copepods; postlarval Munida gregaria; Euphausia superba Pygmy right whale Calanoid copepods
E. glacialis E. australis Caperea marginata E. robustus
Gray whale
M. novaeangliae
Humpback whale
B. acutorostrata B. a. acutorostrata B. a. scammoni
N. Atlantic minke N. Pacific minke
B. bonaerensis B. edeni
B. a. subsp.
Dwarf minke Antarctic minke Bryde’s whale
B. borealis
Sei whale
B. physalus
Fin whale
B. musculus
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B. m. musculus B. m. indica B. m. intermedia
Blue whale Great Indian rorqual ‘True’ blue
B. m. brevicauda
Pygmy blue
Gammarid amphipods; occasional polychaetes Schooling fish; euphausiids
E. superba (Antarctic); euphausiids, postlarval M. gregaria, occasional fish (ex-Antarctic)
Schooling fish; euphausiids Euphausiids; copepods; schooling fish
Pelagic crustaceans, including euphausiids Schooling fish Schooling fish; squid; euphausiids; copepods Euphausiids ?Euphausiids; copepods
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?Euphausiids, schooling fish E. superba Schooling fish; euphausiids Copepods, including Calanus and E. superba E. superba (Antarctic); other euphausiids (ex-Antarctic)
E. superba (Antarctic); other euphausiids (ex-Antarctic) Euphausiids, mainly E. vallentini
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BALEEN WHALES
The baleen plate structure, particularly the inner fringing hairs, to some extent mirrors the food organisms taken or (in the case of E. superba) different size classes. Thus there is some correlation between decreasing size of prey and fineness of baleen by species, i.e. gray, blue, fin, humpback, minke, sei, and right whales. Where food stocks are very dense, e.g., around sub-Antarctic South Georgia, fin, blue, and sei whales may all overlap in their feeding on E. superba. Baleen whale food consumption per day has been calculated as some 1.5–2.0% of body weight, averaged over the year. Given that feeding occurs mainly over about 4 months in the summer in the larger species, the food intake during the feeding season has been calculated at some 4% of body weight per day, c. 4000 kg per day for a large blue whale. To survive the enormous drain of pregnancy and lactation, it has been calculated that a pregnant female baleen whale needs to increase its body weight by up to 65%. The ability to achieve such an increase in only a few months’ feeding indicates the great efficiency of the baleen whales’ feeding system. Predators and Parasites
Apart from humans, the most notable baleen whale predator is the killer whale (Orcinus orca). Minke whales have been identified as a major diet item of some killer whales in the Antarctic. Killer whale attacks have been reported on blue, sei, bowhead, and gray whales, although their frequency and success are unknown. Humpbacks often have killer whale tooth marks on their bodies and tail flukes. Humpback and right whale calves in warm coastal waters are susceptible to attack by sharks. There are anecdotal reports of calving ground attacks on humpbacks by false killer whales (Pseudorca crassidens). A form of harassment, only recently described, occurs on right whales on calving grounds off Peninsula Valdes, Argentina. Kelp gulls have developed the habit of feeding on skin and blubber gouged from adult southern right whales’ backs as they lie at the surface. Large white lesions can result. The whales react adversely to such gull-induced disturbance and calf development may be affected. External parasites, particularly ‘whale lice’ (cyamid crustaceans) and barnacles (both acorn and stalked) are common on the slower-swimming more coastal baleen whales, such as gray, humpback, and right whales. In the latter, aggregations of light-colored cyamids on warty head callosities have facilitated research using callosity-pattern photographs for individual identification. External parasites are much less common on the faster swimming species,
although whale lice have been reported on minke whales (in and around the ventral grooves and umbilicus); the highly modified copepod Penella occurs particularly on fin and sei whales in warmer waters. The commensal copepod Balaenophilus unisetus often infests baleen plates in such waters, especially on sei and pygmy blue whales. A variety of internal parasites have been recorded, although some baleen whales seem less prone to infection than others. They appear, for example, to be less common in blue whales, but prevalent in sei whales. Records include stomach worms (Anisakis spp.), cestodes, kidney nematodes, liver flukes, and acanthocephalans (‘thorny-headed’ worms) of the small intestine. The cold water diatom Cocconeis ceticola often forms a brownish-yellow film on the skin of blue and other baleen whales in the Antarctic. Because the film takes about a month to develop, its extent can be used to judge the length of time an animal has been there. Its presence led to an early common name for the blue whale ‘sulphur bottom.’ For many years the origin of small scoop-shaped bites on baleen whale bodies in warmer waters remained a mystery until they were found to be caused by the small ‘cookie-cutter’ shark, Isistius brasiliensis. Some species are highly prone to such attacks. In Southern Hemisphere sei whales the overlapping healing scars can impart a galvanized-iron sheen to the body.
Life History Behavior
Sound production Unlike toothed whales, baleen whales are not generally believed to use sound for echo location, although bowheads, for example, are thought to use sound reflected from the undersides of ice floes to navigate through ice fields. However, sound production for communication, for display, establishment of territory, or other behavior, is well developed in the suborder. Blue whales produce the loudest sustained sounds of any living animal. At up to nearly 190 decibels, their long (half-minute or more), very low frequency (o20 Hz) moans may carry for hundreds of kilometers or more in special conditions. Fin whales produce similarly low (20Hz) pulsed sounds. Minke whales also produce a variety of loud sounds. Right whales produce long low moans; bowhead sounds, recorded on migration past hydrophone arrays in nearshore leads, have been used in conjunction with sightings to estimate population size off northern Alaska.
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BALEEN WHALES
Southern right whales, at least, seem to use sound to communicate with their calves. Humpbacks produce the longest, most complex sound sequences in ‘songs’, described as an array of moans, groans, roars, and sighs to high-pitched squeaks and chirps, lasting 10 or more minutes before repetition, sometimes over hours. It seems that only the adult males sing, generally only in or close to the breeding season. In any one breeding season, all the males sing the same song, changing slightly over successive seasons. Different populations have different songs; so much so, for example, that those off western Australia have a distinctly different song – less complex, less ‘chirpy’ – than that heard on breeding grounds separated by the Australian continent, off the east coast. ‘Songs’ may also be heard in migrating humpbacks, but less so on the cold-water feeding grounds, where if they occur at all, they appear generally only as ‘snatches’ or isolated segments. Swimming and migration With their stream-lined bodies, rorquals include the fastest-swimming baleen whales. Sei whales have been recorded at around 35 knots (more than 60 km/h) in short bursts; minke and fin whales are also known as fast swimmers, the latter up to 20 knots (37 km/h). Blue whales are among the most powerful swimmers, able to sustain speeds of over 15 knots (28 km/h) for several hours. On migration, humpbacks and gray whales average about 4 knots (8–9 km/h) and bowheads only about 2.7 knots (5 km/h). Migration speeds for southern right whales are not known, but medium range coastal movements off southern Australia indicate 1.5–2.3 knots (2.7–4.2 km/h) over 24 h, for cow/calf pairs. Baleen whales undertake some of the longest migrations known. Gray whales cover some 10 000 nautical miles (18 000 km) on the round trip between the Baja California breeding grounds and the Alaskan feeding grounds, among the longest migrations of any mammal. Southern Hemisphere humpbacks may cover as much as 501 of latitude either way between breeding and feeding grounds, a round trip of some 6000 nautical miles (11 000 km). Not all baleen whale migrations are so well marked. The biannual movements of Bering Sea bowheads are governed by the seasonal advance and retreat and of the sea ice, which varies from year to year. Although Southern Hemisphere blue and fin whales all feed extensively in the Antarctic in summer, the locations of their calving grounds are not known. Sei whale migrations are relatively diffuse and can vary from year to year in response to changing environmental conditions. By comparison, Bryde’s whales hardly
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migrate at all, presumably being able to satisfy both reproductive and nutritional needs in tropical/warm temperate waters. Even among such migratory animals as humpbacks, it may be that not all animals migrate every year; studies off eastern Australia indicate that a proportion of adult females may not return to the calving grounds each year, and individuals have even been reported in summer further north. However, Southern Hemisphere migrating humpbacks show segregation in the migrating stream: immatures and females accompanied by yearling calves are in the van of the northward migration, followed by adult males and nonpregnant mature females; pregnant females bring up the rear. A similar pattern occurs on the southward journey, with cow/calf pairs traveling last. Very similar segregation is recorded among migrating gray whales. Baleen whale migrations have generally been regarded as taking place in response to the need to feed in colder waters and reproduce in warmer waters. Explanations for such long-range movements have included direct benefits to the calf (better able to survive in calm, warm waters), evolutionary ‘tradition’ (a leftover from times when continents were closer together), and the possible ability of some species to supplement their food supply from plankton encountered on migration or on the calving grounds. Corkeron and Connor (1999) have rejected these explanations, suggesting that there may be a major advantage to migrating pregnant female baleen whales in reducing the risk of killer whale predation on newborn calves in low latitudes. It cites in its favor the greater abundance of killer whales in higher latitudes, that their major prey (pinniped seals) is more abundant there, and that killer whales do not seem to follow the migrating animals. Social activity Large aggregations of baleen whales are generally uncommon. Even on migration, in those species where well-defined migration paths are followed (e.g., gray whales and humpbacks), individual migrating groups are generally small, numbering only a few individuals. It has been stated that predation is a main factor in the formation of large groups of cetaceans, e.g., open-ocean dolphins. Given the large size of most adult baleen whales, predation pressure is low and group size is correspondingly small. Blue whales are usually solitary or in small groups of two to three. Fin whales can be single or in pairs; on feeding grounds they may form larger groupings, up to 100 or more. Similarly, sei whales can be found in large feeding concentrations, but in groups of up to only about six elsewhere. The same is true for minke whales, found in concentrations on the
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BALEEN WHALES
feeding grounds, but singly or in groups of two or three elsewhere. Social behavior has been studied most intensively in coastal humpbacks, e.g., on calving grounds. Male humpbacks compete for access to females by singing and fighting. The songs seem to act as a kind of courtship display. Males congregate near a single adult female, fighting for position. Such aggression can involve lunging at each other with ventral grooves extended, hitting with the tail flukes, raising the head while swimming, fluke and flipper slapping, and releasing streams of bubbles from the blowhole. As a result of such encounters, individuals can be left with raw and bleeding wounds caused by the sharp barnacles. Among southern right whales, similar ‘interactive’ groups are often observed on the coastal calving grounds in winter, involving a tight group with up to seven males pursuing an adult female, but not generally resulting in wounded animals. As for humpbacks, it is not yet certain whether such behavior results in successful mating, although, at least in such right whale groups, intromission is often observed. Feeding balaenopterids have often been reported as circling on their sides through swarms of plankton or fish. It has been suggested that gray whales feed on their right sides, as those baleen plates are more worn down, presumably through contact with the seabed. The most remarkable behavior, however, is reported from humpbacks. In the Southern Hemisphere, on swarms of krill, they may feed in the same ‘gulping’ way as other balaenopterids. In the Northern Hemisphere, two methods are commonly reported: ‘lunging’ and ‘bubble netting.’ In the former, individuals emerge almost vertically at the surface with their mouths partly open, closing them to force the enclosed water out through the baleen. In the latter, an animal circles below the food swarm; as it swims upward, it exhales a series of bubbles, forming a ‘net’ encircling the prey. It then swims upward through the prey with its mouth open, as in lunging.
22 meters, and live for possibly 80–90 years. Adult female blue whales give birth every 2–3 years, with pregnancy lasting some 10–11 months. Other balaenopterids follow the same general pattern. Mating takes place in warm waters in winter, with birth following some 11-months later. A 7to 11-month lactation period may be followed by a year ‘resting’, or almost immediately by another pregnancy. Most adults are able to reproduce from between 5 and 10 years of age and reach maximum growth after 15 or more years. The smallest balaenopterid, the common minke whale, is born after a pregnancy of some 10 months, at a length of just under 3 meters. Weaning occurs at just under 6 meters, after 3–6 months. The adult female can become pregnant again immediately following birth, but the resulting short calving interval is generally uncommon in baleen whales: 2–3 years is the norm, although female humpbacks can achieve a similar birth rate, enabling their stocks to recover rapidly after depletion (see Further Reading section). Right whales follow a similar general pattern, but there are some differences. In right whales, gestation lasts about 11 months and weaning for about another year. Females are able to reproduce successfully from about 8 years (there are records of successful first pregnancies from 6 years), but the calving interval is usually a relatively regular 3 years. For bowheads, it has been reported, rather surprisingly, that while growth is very rapid during the first year of life (from c. 4.5 meters), it may be followed by a period of several years with little or no growth. Sexual maturity occurs at 13–14 meters; at the reduced growth rate, that would not be reached until 17–20 years. Similarly, evidence shows considerable longevity in this species: stone harpoon heads found in harvested whales and last known to be used off Alaska early this century suggest that individual animals can be at least 100 years old.
Population Status Growth and Reproduction
Young baleen whales, particularly the fetus and the calf, grow at an extraordinary rate. In the largest species, the blue whale, fetal weight increases at a rate of some 100 kg/day toward the end of pregnancy. The calf’s weight increases at a rate of about 80 kg/day during suckling. During that 7-month period of dependence on the cow’s milk, the blue whale calf will have increased its weight by some 17 ton and increased in length from around 7 to 17 meters. Blue whales attain sexual maturity at between 5 and 10 years, at a length of around
For centuries, baleen whales have borne the brunt of human greed, for products and profit. Only the sperm whale, largest of the toothed whales, has rivaled them as a whaling target. Black right whales (Eubalaena) were taken in the Bay of Biscay from the 12th century, with the fishery extending across the North Atlantic by the 16th century. Attention then shifted to the Greenland whale (Balaena) near Spitzbergen and later off southern and western Greenland. Both species’ numbers were reduced to only small remnants, and in several areas (e.g., Spitzbergen and Greenland for Balaena and the
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BALEEN WHALES
northeast Atlantic and the North Pacific for Eubalaena) the stocks were virtually exterminated. That destruction was undertaken using the ‘old’ whaling method, with open boats and hand harpoons, on the ‘right’ species – ‘right’ because they were relatively easy to approach, floated when dead, and provided huge quantities of products [oil for lighting, lubrication, and soap and baleen (‘whalebone’) for articles combining flexibility with strength, such as corset stays, umbrella spokes, fishing rods]. Development of the harpoon gun and steam catcher, from 1864, increased the rate of catching greatly, but also allowed attention to turn to the largest baleen whales, the blue and fin whales, whose size, speed, and tendency to sink when dead had prevented capture by the old methods. From its beginning in the North Atlantic, then, by the end of the century, in the North Pacific, ‘modern’ whaling’s next and most intensive phase moved south, first in 1904 at South Georgia in the South Atlantic, just within the Antarctic zone. Initially on humpbacks [up to 12 000 were taken in one year (1912), leading to very rapid stock decline] and then on blue and fin whales, southern whaling based on such land stations – in the Antarctic in summer and the tropics in winter – was overtaken from the late 1920s by the great development of pelagic whaling using floating factory ships. Huge annual Southern Hemisphere catches resulted – a maximum of over 40 000 in 1931 – averaging around 30 000 animals per year in the later 1930s and again after World War II until 1965. Whereas blue whales had been the preferred target in the 1930s, their great reduction in numbers led to a shift in attention progressively over the years to fin whales, to sei whales in the 1940s, and finally to minke. With depletion of stocks and more stringent conservation measures (killing of humpbacks, blue, and fin whales was banned from the mid-1960s, even though some illegal catching continued until the early 1970s or even later), catches fell to between 10 000 and 15 000 per year in the 1970s. The ‘old’ whaling story had virtually repeated itself – enormous reductions through overfishing of one species or stock leading to exploitation of other species and stocks until, apart from minke whales, only remnants were left. Since 1989, a moratorium on all commercial whaling has eliminated that pressure, with the exception of limited whaling carried out under exemption for scientific research, and since 1993, limited commercial catching of minke whales in the eastern North Atlantic. Some ‘aboriginal’ whaling has also continued in the Northern Hemisphere, on bowheads off northern Alaska, on gray whales in the Bering Sea, on fin and minke whales off Greenland, and on humpbacks in the Caribbean.
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Despite the great scale of the kill in ‘old’ and ‘modern’ whaling, no whale species has become extinct through whaling, although a number of individual stocks have been reduced greatly; at least one, the North Atlantic gray whale, has disappeared within the past 200–300 years. In its most recent (1996) ‘Red List’ of threatened animals, the World Conservation Union (IUCN) (Table 3) includes no baleen whale species or stocks as either extinct or critically threatened (the latter within the threatened category). Within the threatened category, seven taxa – three species, one subspecies and three stocks – are listed as endangered (E), four taxa – one species and three stocks – are vulnerable (V). Six taxa – two species, one subspecies, and three stocks – are listed as at lower risk (LR), and two taxa – one species and one subspecies – as data deficient (DD). Those species under greatest current threat (E) are the Northern Atlantic and North Pacific right, sei, and fin whales, together with the ‘true’ blue subspecies, two of the five bowhead stocks (Okhotsk Sea, Spitzbergen), and the northwest Pacific gray. Next most threatened (V) are the humpback, two bowhead stocks (Hudson Bay, Baffin Bay/Davis Strait), and the North Atlantic blue. At lower risk (LR) are the southern right and Antarctic minke, one bowhead stock (Bering-Chukchi-Beaufort Seas), the North Atlantic minke, northeast Pacific gray, and North Pacific blue; all but one are further qualified as conservation dependent (cd, not vulnerable because of specific conservation efforts). The exception is the North Atlantic minke, listed as near threatened (nt, not conservation dependent but almost qualifying as vulnerable). The two taxa for which insufficient information is currently available (DD) are Bryde’s whale and the pygmy blue. The International Whaling Commission’s Scientific Committee, responsible for the assessments of such stocks’ current status, has reported encouraging recent reversals of stock decline for some stocks of some species. One, the northeast Pacific gray whale, has recovered under protection from commercial whaling (but with aboriginal catches up to some 150 per year) to at or near its ‘original’ (prewhaling) state (c. 26 000 animals). Similarly, the northwest Atlantic humpback and several Southern Hemisphere humpback populations have been showing marked increases. The latest estimate of the North Atlantic stock, some 10 600 animals in 1993 (cf. 5500 in 1986), must reflect some population growth in the intervening period, whereas two Southern Hemisphere stocks (off eastern and western Australia) have been increasing steadily, at 10% or more per year since the early 1980s. Indeed, in all areas where surveys have been undertaken recently on Southern Hemisphere humpback populations they
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Table 3
IUCN ‘Red List’ Categories for Baleen Whales (1996)
Species
Subspecies
Common name
Category EX
EW
Threatened CR
B. mysticetus E. glacialis* E. australis Caperea marginatae E. robustus M. novaeangliae B. acutorostrata
B. B. B. B. B.
bonaerensis edeni borealis physalus musculus
B. a. acutorostrata B. a. scammoni B. a. subsp.
B. B. B. B.
m. m. m. m.
musculus indica intermedia brevicauda
Bowhead whale North Atlantic right whale Southern right whale Pygmy right whale Gray whale Humpback whale N. Atlantic minke N. Pacific minke Dwarf minke Antarctic minke Bryde’s whale Sei whale Fin whale Blue whale Great Indian rorqual ‘True’ blue Pygmy blue
LR
EN
VU
*a
*b
*
(cd)c
*
(cd)
*
(cd)g
*
(nt)
DD
NE
*d
*f
*
(cd) * * * *h
*
(cd)i
* *
*
Includes E. japonica. Okhotsk Sea bowhead whale, Spitzbergen bowhead whale. b Hudson Bay bowhead whale, Baffin Bay/Davis Strait bowhead whale. c Bering-Beaufort-Chuckchi Seas bowhead whale. d North Atlantic and North Pacific northern right whales. e Pygmy right whale removed from 1996 Red List. f Northwest Pacific gray whale. g Northeast Pacific gray whale. h North Atlantic blue whale. i North Pacific blue whale. Categories: Ex, extract; EW, extinct in the wild; CR, current risk; EN, endangered; VU, vulnerable; LR, lower risk; DD, data deficient; NE, not evaluated; cd, conservation dependent; nt, not threatened. a
have been shown to be undergoing some recovery. Three southern right whale stocks (off eastern South America, South Africa, and southern Australia) have been increasing since the late 1970s at around 7–8% per year, although at some 3000, 3000, and 1200 animals, respectively, all are still well below their ‘original’ stock size. Even the ‘true’ blue whale, whose future has been of considerable concern, with estimates for the late 1980s at fewer than 500 animals for the whole Antarctic, has shown recent encouraging signs. Based on a series of Antarctic sightings cruises, mainly for minke whales but including other large whales, the most recent calculations (admittedly using only small absolute numbers sighted) show that the population must have been increasing since the last estimate (1991). As yet the analyses do not permit a firm conclusion on the number now present, although it seems likely to be more than 1000. The one species or stock for which there is now very great concern is the North Atlantic right whale. At very low absolute abundance (only some 300
animals), not recovering despite protection from whaling in the 1930s, now even decreasing through a reduced survival rate and an increase in calving interval, and subject to increasing removals from ship strikes and fishing gear entanglement, the only way to ensure the species survival is to reduce such anthropogenic mortality (ship strike and entanglement) to zero. While research on mortality reduction measures should be pursued, immediate management action is urgently needed. It has been calculated that the great reduction of baleen whales by whaling, for the Antarctic to around one-third of original numbers and one-sixth in biomass, must have left a large surplus of food – some 150 million tonnes per year – available for other consumers, such as seals, penguins, and fish. (In a different way, earlier whaling in the North Atlantic, particularly on right whales, is believed to have influenced the spread of one sea bird – the fulmar – by providing food in the form of discarded whale carcasses.) In response to an increase in
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BALEEN WHALES
available food, there may well have been increases in growth rates, earlier ages at maturity, and higher rates of pregnancy in some baleen whale species. However, the evidence is equivocal, as it is for competition between individual whale species. For some, e.g., right whales and sei, it has been suggested that an increase in one (right whales) could be inhibited by competition with another (sei whales). In the North Pacific, both sei and right whales can feed on the same prey – copepods – and sei whales can at times be ‘skimming’ feeders, like right whales. However, evidence that they actually compete on the same prey, in the same area, at the same time, and even on the same prey patch is lacking. Similarly, there has been much debate and speculation on whether the recovery of the Southern Hemisphere ‘true’ blue whale has been inhibited by an apparent increase in Antarctic minke whales. In that case, there may in fact be very little direct competition for food where the common prey is not limited in abundance (as in the Antarctic) and is available in large patches. The well-authenticated increases in the substantial annual rates for several stocks of Southern Hemisphere humpbacks and right whales and the possibility of at least a limited increase in numbers for the ‘true’ blue whale suggest that such competition is unlikely, at least where, as in the Antarctic, food supplies are abundant.
Glossary Baleen Plates of keratin hanging transversely in the roof of the mouth of baleen whales, forming the ‘baleen apparatus’ for filter feeding on surface plankton; formerly known as ‘whalebone’ but bearing no resemblance to true bone. Ventral grooves A series of parallel grooves or plates running longitudinally on the undersurface of the throat and chest region in balaenopterid
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whales, allowing great expansion of the mouth during feeding. Krill Planktonic shrimp-like crustaceans of the genus Euphausia, particularly the Antractic Euphautia superba.
See also Bioacoustics. Copepods. Krill. Marine Mammal Migrations and Movement Patterns. Marine Mammal Overview. Marine Mammal Social Organization and Communication. Marine Mammal Trophic Levels and Interactions. Marine Mammals, History of Exploitation.
Further Reading Bonner WN (1980) Whales. Poole, Dorset; Blandford Press. Corkeron PJ and Connor RC (1999) Why do baleen whales migrate? Marine Mammal Science 15: 1228--1245. Harrison R and Bryden MM (eds.) (1988) Whales, Dolphins and Porpoises. Hong Kong: International Publishing Corporation. IUCN (1996) 1996 IUCN Red List of Threatened Animals (plus annexes). Gland, Switzerland: IUCN. IUCN (1996) 1996 IUCN Red List of Threatened Animals (plus annexes). Gland, Switzerland: IUCN. Laws RM (1977) Seals and Whales of the Southern Ocean. Philosophical Transactions of the Royal Society of London, Series B 279, 81--96. Leatherwood S and Reeves RR (1983) The Sierra Club Handbook of Whales and Dolphins. San Francisco: Sierra Club Books. Rice DW (1998) Marine Mammals of the World:: Systematics and Distribution, Special Publication Number 4. Lawrence, Kansas: Society for Marine Mammalogy.
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BALTIC SEA CIRCULATION W. Krauss, Institut fu¨r Meereskunde an der Universita¨t Kiel, Kiel, Germany Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 1, pp 236–244, & 2001, Elsevier Ltd.
Introduction The Baltic Sea consists of a number of sub-basins connected by straits and deep channels. They have an important influence on the current system. A map of the Baltic Sea with topographic subdivision and bottom topography is shown in Figure 1. The North Sea, via Skagerrak and Kattegat, is connected to the Baltic Proper by three pathways, called the Danish Straits (Figure 2). The eastern route, the Sound, is only 2 km wide at its narrowest location, shallow (sill depth 8 m) and only about 55 km long. It is the shortest inflow route for saline water. The central pathway is about 180 km long and about 13 m deep on the average, 25–30 m along the axis. It consists of the Great Belt and the Fehmarnbelt and is terminated at the east by the Darss sill (18 m deep), the shallowest sill for the main inflow. The third connection is the Little Belt, having a cross section of only 16 000 m2 compared to 255 000 m2 of the Great Belt and 80 000 m2 of the Sound. Therefore, about 70% of the in- and outflows occur through the Great Belt; the Little Belt is negligible. Moving from west to east, the Arkona Basin (45 m) is connected to the Bornholm Basin (95 m) by the Bornholm Channel. The Stolpe Channel (20 km wide by the 60 m isobath) allows inflow into the Gdansk Basin (110 m) and the Eastern Gotland Basin with a maximum depth of 250 m. Further to the north the Faro¨ Deep is followed by the Northern Basin (200 m) which extends to the east towards the Gulf of Finland as a deep channel with decreasing width and depth. By contrast, the Gulf of Bothnia is separated from the deeper layers of the Baltic Proper by the Aland Sea with its numerous islands. The Gulf of Bothnia consists of two basins, the Bothnian Sea and the Bothnian Bay. The deep basins and the connecting channels are important for the water exchange between the North Sea and the Baltic Sea and determine the current structure. Due to the large river runoff, the Baltic Sea is the largest brackish water body of the world with a volume of 21 000 km3. The mean annual total contribution of all rivers for the period 1950–1990 was
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446 km3 year1. This would correspond to a sealevel rise of 1.18 m year1 if the Baltic Sea were not connected to the North Sea. An additional surplus of fresh water results from the difference of precipitation minus evaporation, which amounts to about 60 km3 year1. The river runoff shows both a seasonal cycle and interannual variations. The lowest and highest annual values differ from the mean value by 27% and þ 22%, respectively. The seasonal runoff varies between 25 000 m3 s1 in spring and 12 000 m3 s1 in winter. As a consequence of this freshwater supply the salinity in the eastern and northern parts of the Baltic Sea is reduced to about 4 PSU (practical salinity units), increasing to about 8 PSU in the central and western Baltic Sea. Due to the associated density difference saline water penetrates from the Kattegat (30–34 PSU) through the Danish Straits and the channels into the Baltic Sea and yields a strong halocline, which separates the water of the deeper layers from the brackish upper water masses (Figure 3). As a consequence vertical mixing is strongly reduced. The Baltic Sea extends from about 541N to 661N thus ranging from mild and humid to a subarctic climate. Frequent and complex synoptic-scale cyclonic activity and subsynoptic-scale depressions are characteristic, leading to a high variability of the prevailing westerly winds. The wind fields produce a highly variable current system, especially in the upper layers, superimposed on the weak baroclinic flow field induced by the salinity differences. Beneath the wind-mixed layer bottom topography has a strong steering influence. The wind-induced currents and the associated sea-level variations may drastically change, when parts of the Baltic Sea are icecovered. Sea ice occurs every year in the Baltic Sea. Under normal winter conditions about 45% of the Baltic Sea is covered with ice and the ice season lasts about 6 months in the northern parts. Severe winters may lead to almost total ice coverage. Ice first appears in the innermost parts of the Bothnian Bay during mid-November. In a normal winter, the entire Bothnian Sea, Aland Sea, Gulf of Finland and the northernmost part of the Baltic Proper are covered by ice. In severe ice-winters the Kattegat, the Belt Sea, the Sound and large parts of the Baltic Proper are also ice covered. Wind and thermohaline forcing determine the currents of the Baltic Sea, strongly influenced by bottom topography and ice coverage. A large
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BALTIC SEA CIRCULATION
66°
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Figure 1 Skagerrak (1), Kattegat (2) and the subareas of the Baltic Sea: Beltsea (3), Arkona Basin (4), Bornholm Basin (5), Gdansk Basin (6), Eastern (7) and Western Gotland Basin (8), Gulf of Riga (9), Northern Basin (10), Gulf of Finland (11), Aland Sea (12), Bothnian Sea (13) and Bay of Bothnia (14). Depth contours: thin broken line 40 m, heavy dotted line 100 m, heavy full line 200 m. The location of the section of Figure 3 is also shown (full line with open dots).
number of current measurements has been made in the past decades. However, due to the high variability of the wind-induced currents and the extensive fishing activities, which make it impossible to install observational systems in some areas, it was not possible in the past to derive a consistent circulation pattern from observations. In recent years, three-dimensional models, combined with data assimilation, have improved in such a way that both the mean circulation and its variability can now be described with sufficient accuracy.
The Estuarine Circulation In a stratified sea it is generally not possible to separate wind-induced currents properly from
thermohaline ones. Along the coasts and the bottom slopes wind produces up- and downwelling and thus inclinations of the density surfaces which may largely amplify the wind effects. However, some insight is gained by considering the horizontal pressure gradients, which result from the mean surface inclination and the horizontal density differences. They produce an estuarine type of circulation. Figure 3 shows the mean salinity section from the Kattegat along the main axis of the Baltic Sea into the Gulf of Finland. Except at the level of the summer thermocline (about 25–30 m depth) the density distribution is similar. Near the bottom the density decreases from about 25 kg m3 in the Kattegat to 9 kg m3 in the Gotland Basin and less than 4 kg m3 in the interior of the Gulf of Finland.
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57° 06'
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E
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Hälsingborg
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t Be
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Kopenhagen
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Arkona basin 40 m
Fe h
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DS
20 m
Be
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Germany 54°
54° 9°
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Figure 2 The Danish Straits, Kattegat and Arkona Basin. SS, Samsoe Sill, DS, Darss Sill. The Belt Sea is the area between the Kattegat and Arkona Basin.
Because of the river runoff and the low salinity in the eastern and northern Baltic Sea the sea level is higher by 0.3 m in the interior of the Gulf of Finland and the Gulf of Bothnia compared to that in the Belt Sea. This sea level inclination produces a pressure gradient which drives the upper layers out of the Baltic Sea. Due to the Coriolis force this outflow is Great Belt
Kattegat
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concentrated on the Swedish coast. On average about 1250 km3 year1 leave the Baltic Sea through the Danish Straits. As compensation about 740 km3 year1 of saline, dense and oxygen-rich water penetrates in the deeper layers through the Danish Straits into the Arkona Basin and from there through the Bornholm Channel into the Bornholm Basin.
>10
Mean salinity section August
120 140
12
160 >
200
180 200
Figure 3 Mean salinity (practical salinity units) distribution (1902–1956) along a section from the Kattegat to the Gulf of Finland (for location see Figure 1).
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9˚
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Figure 4 Mean salinity (practical salinity units; PSU) distribution at the bottom of the Baltic Sea (1902–1956).
Horizontal pressure gradients due to the density gradients, which overcompensate the sea level gradient, act as forcing term. The path of this inflow further to the north is shown in Figure 4 by the salinity distribution close to the bottom. Water of 12 PSU flows through the Stolpe Channel into the Gdansk Basin and further into the Eastern Gotland Basin. Part of this water continues into the Northern
Basin, and another part flows around Gotland and south into the Western Gotland Basin. However, these observations also represent the cyclonic windproduced circulation in the Baltic Proper and are only partially due to thermohaline driving (see next section). Only in the Danish Straits, where the salinity differences are large and the flow is channeled, are
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baroclinic currents as strong as 0.1–0.2 m s1 observed. During light and moderate winds, a twolayer system of currents prevails in the Belt, with outflow from the Baltic Sea in the upper and inflow in the deeper layers. This bidirectional flow is driven by the contrast in density and the surface inclination and includes the outflow from the Baltic Sea due to the fresh water surplus. In the Fehmarnbelt this flow system is observed in 50% of all cases in summer and 29% in winter. In the Great Belt the corresponding numbers are 95% in summer and 65% in winter. In contrast to the Belt with its typical bidirectional current system, outflow into the Kattegat extends over the entire Sound, reaching maximum values of 0.2 m s1 near the surface. This flow forms the Baltic Current along the Swedish coast in the Kattegat and Skagerrak which becomes the Norwegian Coastal Current further to the north.
Wind-driven and Thermohaline Circulation As mentioned above the wind field over the Baltic Sea is highly variable with a well pronounced annual cycle. Maximum speed is reached during November– March with an absolute maximum in December. The minimum occurs in May/June. In December the maximum mean speed is 7–9 m s1, the minimum in May/June is 4–5 m s1. The mean wind direction is from south west/west towards north east/east, i.e., nearly along the line from the western Baltic to the Gulf of Finland. The current system produced by wind blowing in the direction of an infinitely long basin with constant cross-section and horizontally uniform stratification is depicted in Figure 5. 1. In the surface layers there results an Ekman current which is deflected to the right of the wind direction by the Coriolis force, yielding an Ekman transport in cross direction. 2. This Ekman transport produces a sea-level rise on the right-hand coast (looking in wind direction) and a fall on the left-hand side. Furthermore, downwelling occurs on the right-hand side and upwelling on the left-hand side, resulting in baroclinicity of the same sign at both coasts. Note the different length scales of deflection of the sea surface and the isopycnals, which are governed by the external and internal Rossby radius, a measure which describes the adjustment of the deflection under gravity of a rotating fluid. 3. Consequently, geostrophically balanced coastal jets are produced by the inclination of the sea surface along both coasts in wind direction and a
Wind (1) Coastal jet
Transport
Ekman Upwelling
Coastal jet
(2) Downwelling Return flow
Figure 5 Schematic of wind-produced currents in an infinitely long basin. Sea surface (1) and density surfaces (2) are inclined due to the Ekman transport perpendicular to the coast.
slow return flow compensates this transport in the central area of the basin. The inclination of the density surfaces enhances these currents. If the depth contours are closed, as in the deep basins of the Baltic Sea, the return flow in the center is prohibited. Instead topographic effects become dominant. Outside of the up- and downwelling regions along the coasts the flow in the basins is determined by the sea level inclination of Figure 5. For these regions Rossby’s potential vorticity theorem holds, d=dtððz þ f Þ=HÞ ¼ 0
½1
where z is the relative vorticity, f is the Coriolis parameter and H is depth. In basins like the Baltic, which are large enough for rotation to play an important role in the dynamics, but not so large that variations of the Coriolis parameter must be taken into account, and where the mean currents are weak (x5f ), the vorticity theorem for the mean flow, v, can be reduced to v grad H ¼ 0
½2
i.e. the vertically averaged currents below the Ekman layer try to follow the depth contours. The wind-driven currents of the Baltic Sea show much of this typical response besides the strong vertical stratification, e.g. coastal jets are often observed along the eastern slopes with current speeds up to 1 m s1. Numerical models reveal clear evidence of characteristic persistent circulation patterns which comprise mostly the sub-basins with less transport between the basins. Figure 6 shows the annual average of the vertically integrated transport
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BALTIC SEA CIRCULATION
10°
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66°
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Figure 6 Annual average of the depth-integrated transports (m3 s1) for 1992 (model results). Maximum vector 0.934 105 m3 s1.
for the year 1992. As can be seen from this figure the most pronounced structure with the highest stability is a cyclonic circulation cell comprising the Eastern Gotland basin. Most of the water is recirculating in the Eastern Gotland Basin, but part of it flows back along the western slope of the Western Gotland Basin, continues towards the Bornholm Basin and closes the return flow through the Stolpe Channel. On average the stability of this circulation pattern is higher than 50%. The northward flow along the eastern Baltic slope is much more pronounced than the southward flow along the Swedish slope. This can be interpreted according to Figure 5: the currents along the eastern slope are intensified by the coastal jet and correspondingly reduced along the Swedish coast. Similar patterns as in the Gotland Basin are observed in all other basins. The transport associated with these wind-induced patterns is about an order of magnitude higher in the
Stolpe Channel than the fresh water supply by the river runoff. Although highly variable between the years, this area shows the most intense transport rates of the entire Baltic Sea. The bottom currents are only slightly different from the vertically averaged ones. Bottom currents towards the east are found in the Bornholm and in the Stolpe Channel with stabilities higher than 50%. The annual average of the bottom currents in the Bornholm Channel reaches about 0.1 m s1 and in the Stolpe Channel 0.05–0.08 m s1. This demonstrates the importance of these channels for the inflow of saline water into the eastern basins. However, the strongest bottom currents are in most cases not coincident with the highest stability. This indicates that these bottom currents are not as permanent as the vertically averaged currents. As expected from Figure 5 up- and downwelling occurs along the coasts. Figure 7 shows the vertical
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10°
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66°
66°
N
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62°
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Figure 7 Upwelling and downwelling in the Baltic Sea based on four-year vertical averages of vertical velocities (model results). Grey areas, upwelling 4 4 106 m s1; black areas, downwelling o 4 106 m s1.
average of the vertical velocities for a four-year modeling period. Along the western coast of the Baltic upwelling occurs at a distance of about 20– 30 km offshore, and along the eastern coast a continuous band of downwelling can be seen. Larger islands and topographic features have similar effects. The areas can be regarded as the most active areas in the vertical exchange of water masses.
Currents in the Danish Straits and Major Inflow Events The currents in the Danish Straits are highly variable, mainly determined by the sea-level differences between the Skagerrak and the Arkona Basin. Westerly winds, for example, pile up water in the Skagerrak– Kattegat area and simultaneously lower the sea level
in the Arkona Basin. Large sea-level differences are therefore common along the Danish Straits. The flow in these straits responds like the flow in a manometer between two large basins. Table 1 depicts the mean daily outflow rates and the standard deviations through the straits for some years (model results). Typically the standard deviation is 30–40 times larger than the mean flow. Actual mass transport across Darss Sill calculated from observations is depicted in Figure 8 for 1994 and 1995, showing this high variability. In the Great Belt the speed reaches more than 71 m s1, and in the ¨ resund as much as 71.5 m s1. At Darss Sill, where O the central channel widens, 70.5 m s1 is observed. The inflow of saline and oxygen-rich water along the bottom of the Baltic Sea is essential for bottomliving animals in the Baltic Proper. Oxygen depletion in the deep basins causes oxygen deficiencies and the
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BALTIC SEA CIRCULATION
Table 1 Mean daily outflow rates and standard deviation through the Danish Straits (model results) (1 km3 day1 ¼ 11.6 103 Sv) Year
Mean outflow (km3day 1)
Standard deviation (km3day 1)
1986 1988 1993 1994
1.28 1.39 2.26 1.02
41.0 54.4 63.6 59.1
production of hydrogen sulfide after long periods of stagnation. Inflow of water of intermediate salinity occurs more or less continuously, as mentioned above, but this water is normally not dense enough to replace the bottom water in the deep basins. Renewal occurs only intermittently after periods of strong and persistent westerly winds. The greatest renewal during the last 50 years occurred in 1951, followed by a stagnation period until 1963. By then about onethird of the area of the Gotland Basin was covered with hydrogen sulfide-enriched water near the bottom. The inflows in 1963 and 1969 only partially replaced these water masses. The situation became
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worse during the next decades; the last major inflow occurred in 1993, which is the best documented and modeled inflow event. The following conditions are most favorable for a major inflow, which usually occurs between November and January. 1. An extended period of easterly winds must drive water out of the Baltic and lower the sea level in the entire Baltic Sea. Maximum observed sea-level difference between Kattegat and Baltic Sea then amounts to about 0.7 m on average (maximum value 1.66 m). During this period the Kattegat front, which separates the saline Kattegat water from the outflowing Baltic Sea water, is pushed far to the north. 2. This period must be followed by a phase of strong westerly winds over the North Sea and the Baltic, such that the sea level and barocline gradients drive the water back into the Baltic. This period lasts typically 2–3 weeks, during which the brackish water of the Straits is flowing into the Arkona Basin and allows the saline water of the Kattegat to penetrate towards Darss and Drogden Sill. Due to this inflow, sea level rises in the Baltic up to normal or even above normal.
Figure 8 Mass transport over Darss Sill during June 1994–April 1995 based on two stations at the sill. ‘27.6’ on the time axis means 27 June.
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3. A major inflow of saline water starts when the Kattegat water, usually reduced in salinity to 22 PSU by vertical mixing, reaches Darss and Drogden Sill. It then penetrates along the bottom into the Arkona Basin. During major inflows the water transports 2 109–5 109 tonnes of salt into the Baltic Sea. The water takes about 22 days on average (0.11 m s1) to pass the Arkona Basin. The reason for this slow propagation is doming of the inflowing water. The heavy water forms a large cyclonic eddy (geostrophic flow around the dense water) which prevents the water from flowing rapidly through the basin. Thus the heavy water remains there as a pool over extended periods, its salinity may be further reduced by windinduced mixing and the water in the upper layers may even flow back into the Danish Straits. After passing Bornholm Channel the same occurs in the Bornholm Basin, where the water takes about 40 days to reach the center. In many cases it does not fill the basin up to 60 m (the sill depth of Stolpe Channel) and its further penetration into the Gotland Basin depends on the wind conditions.
which allow water exchange up to 4 105 m3 s1 (Figure 8). Typically, we find only two or three oscillations. Observations indicate that the dominant oscillation belongs to the western Baltic–Gulf of Finland system, which represents the second mode of the entire Baltic. The period varies between 26 and 27 h (theoretically 26.4 h). The Gulf of Bothnia is little influenced by this mode, in contrast to the first mode (period 31 h). Both the first and second modes show amphidromic points north of Gotland, where the vertical displacement vanishes. This allows the tide gauge of Landsort (south of Stockholm) to be used as reference station for the mean sea level of the Baltic Sea. Currents due to seiches are most pronounced in the shallow areas. Tidal currents are negligible in the Baltic Sea east of Darss Sill.
See also Ekman Transport and Pumping. Estuarine Circulation. Rossby Waves. Wave Generation by Wind.
Oscillatory Currents A pronounced phenomenon in the current variations is the appearance of inertial oscillations, a response of the Ekman currents to the varying wind conditions. They include about 30% of the observed kinetic energy within the wind-mixed layer. Their frequency is determined by the Coriolis parameter, slightly modified by stratification. Furthermore, seiches (standing waves of the entire Baltic Sea, transformed to amphidromic systems due to the Earth’s rotation) are a dominant feature in all tide gauge records, especially at the Finnish coast. These free oscillations were first recognized by F. A. Farel in Lake Geneva, and the local name ‘seiches’ has generally been adopted as a term for free oscillations of more or less enclosed water basins. They occur when the wind decreases after piling-up of water at one end of the basin. In the Baltic Sea the seiches are heavily damped due to the complicated configuration of the sea, the division into several subbasins and the open connection to the North Sea,
Further Reading Bock KH (1971) Monatskarten des Salzgehaltes der Ostsee, dargestelt fu¨r verschiedene Tiefenhorizonte. Deutsche Hydrographische Zeitschrift, Erga¨nzungsheft Reihe B 12: 1--147. Lehman A and Hinrichsen H-H (2000) On the wind driven and thermohaline circulation of the Baltic Sea. Physics and Chemistry of the Earth B 25(2): 183--189. Ma¨lkki P and Tamsalu R (1985) Physical Features of the Baltic Sea. Finnish Marine Research No. 252, 110 pp. Helsinki. Rheinheimer G (ed.) (1996) Meereskunde der Ostsee, 2nd edn. Berlin: Springer. Simons TJ (1980) Circulation Models of Lakes and Inland Seas. Canadian Bulletin of Fisheries and Aquatic Sciences, Bulletin 203. Ottawa: Department of Fisheries and Oceans. Wu¨bber Ch and Krauss W (1979) The two-dimensional seiches of the Baltic Sea. Oceanologica Acta 2: 435--446.
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BATHYMETRY D. Monahan, University of New Hampshire, Durham, NH, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction Bathymetry is the art and science of mapping the surface of the seafloor. Unlike subaerial land surfaces which are exposed in their entirety and can be captured using a variety of optical instruments, the surface of the seafloor is invisible beyond a few meters depth, or beyond the light beams of human divers and remote vehicles. Mapping it consists of making measurements between it and a reference surface, plotting those measurements in their correct geographic locations, then interpreting the combined measurements into a mapped surface or a grid of calculated depths, both representatives of the actual seafloor. Most human activity in the marine environment has a spatial or locational component. Geographic coordinate systems are enormously useful, but are more easily perceived by machines than humans. Humans can more quickly and easily visualize maps, with a shoreline (the edge of the seafloor) and depth portrayed by isobaths (contour lines) or color graduation: consequently, published bathymetry maps are used in all branches of marine science. They are the most common base maps, and are readily available in a number of series. Digital grids of depths are also calculated and published. Despite their value and ubiquitousness, bathymetry maps are often not understood and frequently taken for granted. There is a general belief that they are accurate and complete, and can be used unquestioningly. Unfortunately, this is seldom the case. The surface of the seafloor is a major component of the Earth, covering as it does some 70% of the globe. It may be the place where sediments (and pollutants) accumulate, or may be composed of bare rock, some dating to the early formation of the Earth’s crust, some formed only yesterday. It acts as the deeper wave guide for water motion, regulating the flow and velocity of tides and currents, including tsunamis. It is the home of myriad species, some who live their entire lives within it, some who have only sporadic contact with it. Some parts of it support rich biological habitats, some parts are deserts. Elements of it will form international boundaries
under the United Nations Convention on the Law of the Sea (UNCLOS). Where it nears the ocean/atmosphere surface, it can become a hazard to navigation or provide protection from storm waves. Despite the seafloor’s importance, it is poorly known, being mapped at a resolution that is 1 or 2 orders of magnitude less than some of the planets and the Moon.
Uses of Bathymetry Bathymetry is used in most human activities in the marine environment, with uses ranging from providing a geographic reference frame to being a major component of a process. Economic and social uses include, but are not limited to:
• • •
• •
sustainable resource management: comprising fisheries, fisheries enhancement, aquaculture, and petroleum and mineral extraction; environmental stewardship: comprising habitat monitoring, land-use planning, environmental quality, national heritage, marine protected areas; health and safety: comprising sewage sites, dredge spoils disposal, risk reduction, oil spill readiness plans, coastal planning/global warming sea level rise, hazard mitigation (earthquakes, storm surge and tsunami, and seafloor hazards), accident investigation; infrastructure: comprising cables and pipeline, offshore exploration and production platforms, coastal structures, dredging, opening Northwest Passage; sovereignty: comprising boundaries definition, juridical continental shelf under UNCLOS, seabed surveillance and intervention, contraband and drugs.
In addition, bathymetry contributes to most other marine sciences. Physical oceanographers study how ocean water circulation is controlled by the shape of the seafloor, while to biologists the seafloor is an important component of the natural habitat of bottom-dwelling creatures, so that better resolution of bathymetry leads to better understanding of these phenomena. The shape of the seafloor can tell geologists the history and composition of the rocks and sediments that compose it, give clues to where seafloor hazards like landslides are likely to occur, and trace depositional patterns of material brought from land, including pollutants. Despite these many
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uses and its importance, the World Ocean is very poorly mapped.
What Bathymetry Is and Is Not Bathymetric Maps Differ from Navigation Charts
Two types of maps showing ‘depth’ are in common use. Both bathymetric maps and navigation charts are highly specialized products, which are designed for different purposes and show different things. Since they both show ‘depths’, they are sometimes substituted one for the other, often due to the fact that in some areas only one of them existed at or near the desired scale. Navigation charts are carefully constructed instruments designed to provide a basis for charting at sea. To do so, a chart must (1) form the base for the graphical exercise of navigation; (2) provide information on the identification and characteristics of navigational aides; and (3) provide information on the nature and position of navigational hazards, which in the case of bathymetry means providing shoal-biased depths. By contrast, a bathymetry map is not constrained by the need to protect mariners and their vessels. Like its land-based equivalent, the topographic map, it shows the relief of the seafloor, without any constraints imposed by the demands of safety. Every contour is shown in its best-known location; there is no bias toward depicting shallower rather than deeper depths – isolated deeps are included. In fact, the aim of the bathymetric map is to show every ‘nook and cranny’ of the seafloor in the best possible way permitted by the data available and the horizontal scale of the map. It may be useful to think in terms of the surfaces that bathymetry maps and navigation charts show. Imagine a layer model, in which the deepest layer is the actual seafloor, the surface that would be seen if the water were removed. It undulates smoothly in places, changing texture to rough abruptly, the creation of the interplay of different geologic forces acting over different timescales. The next lowest layer is the bathymetry map. It is a representation of the lower layer; its creators would like it to conform exactly to the deepest layer. Although in a few areas of very densely spaced measurements it does, mostly the bathymetry layer conforms closely to the real surface in only a few places. There are two prime reasons for this. First, the real surface has only been measured in a few locations, and these are usually insufficient to allow a complete one-to-one correspondence between the bathymetry map surface and the real surface. Second, where the real surface has been measured by single-beam echo sounders,
their beams tend to smooth the real surface, filling in holes and broadening peaks. The upper layer is the navigation chart. Because the navigation chart’s role is to help guide mariners away from areas where a ship could run aground, the navigation chart layer touches only the shallowest portions of the real surface, and magnifies the size of dangers to navigation. Only about 8% of the ocean is shallower than 200 m, and depths that impact navigation are less than half that depth. Bathymetric Surfaces Differ from Potential Field Surfaces
Maps of potential field (magnetics, gravity) are made from data collected along lines as is bathymetry. The mathematical approaches to producing contours and grids of potential field data have sometimes been applied to bathymetric acoustic measurements, but doing so can lead to erroneous results. Potential fields are usually (much) smoother than seafloor, since potential fields radiate in all directions, declining according to the inverse distance squared law. This means that potential field can only be singlevalued, and has limits to its gradients; bathymetry represents a real world surface that in places includes overhangs and can have faces that are absolutely vertical. Assumptions made when processing potential field data may or may not apply to bathymetry.
Bathymetry of a Small Area It is easy to begin with a model for constructing a bathymetry map or grid of a restricted area, one in which all the measurements are made using the same equipment. Local bathymetry maps are made almost exclusively from echo soundings, carried out by surface ships, or operated near the bottom in remote vehicles. This means that they do not travel very quickly! Essentially, a measurement is made at one location, the echo sounder moved, another measurement taken, the echo sounder moved again, and so on until resources are used up or weather forces termination of activities. From the resulting set of measurements, bathymetry is created. Acoustics
Most of the data accepted for use in bathymetry come from acoustic measurements. In these measurements, sound is propagated into the water from a transducer, reflected from the seafloor and its arrival recorded at the point of origin. From the travel time, distance is calculated using a measured or assumed speed of sound.
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Acoustic systems can be classified into: 1. single-beam echo sounders, with one sound beam oriented vertically, which produces a profile across the seafloor (these were the only method available until c. 1985 and were used in all applications; they are being supplanted by (2) and (3)); 2. multibeam echo sounders (MBESs) with many beams radiating across the ship; and 3. side-scan sonars with one broad beam on each side; these are sometimes configured to produce measurements of depth, but more often give only a visual depiction. These instruments are most frequently hull-mounted, and are also shallow-towed, or carried close to the seafloor by deep-towing or remote vehicles for highresolution mapping. Depth measurements are converted to data by being combined with navigation and attitude measurements. Navigation determines the sensor’s position, while attitude determines where the sensor is pointed, from a combination of heading, speed, roll, pitch, and yaw. Nonacoustic Methods
There are other ways of mapping the ocean floor besides using acoustics-based instruments. It has long been the dream of some bathymetrists to develop instruments that would not have to operate in the water, so that the progress of ocean mapping would not be tied to the slow speed and limited number of surface ships and their constrained area of operation. One nonacoustic instrument that is being used in ocean mapping is LIDAR (light detection and ranging), in which a laser beam is projected from an aircraft to the seafloor. While this approach does benefit from the greater speed of aircraft, it is limited to a depth of 50 m or so, and consequently can only be used in shallow coastal waters.
express precision of depths as the standard deviation of differences at crossovers.
Scale and Resolution A map is a picture representing the Earth’s continuous surface drawn at a greatly reduced scale. Conventionally the horizontal scale of a map is the ratio of a unit of length on the map to a unit of length on the Earth (e.g., a scale of 1:1 000 000 means that 1 cm on the map surface represents 1 000 000 cm on the Earth’s surface). Determining a suitable scale at which to produce a paper map requires balancing a number of general factors like intended use, cost, method of reproduction, and accuracy required (see Table 1). Although the display of digital maps is not constrained by a fixed scale, they cannot portray the seafloor to a higher resolution than the data captured. A grid is a numerical model of the Earth’s continuous surface comprising numerical values at discrete intervals. It too has a scale. Resolution of a bathymetry map or grid specifies the size of the features it can resolve, with smaller indicating more detail. Scale imposes a limit on resolution, so the trick is to ensure that map portrays features down to the resolution limit. There are a number of stages in the production of a bathymetry map which contribute to the final resolution of the map. Table 1 Summary of horizontal and vertical resolution requirements for various oceanographic activities by area Topic
Area
Horizontal resolution
Vertical resolution
Ocean models
Open ocean Sills Continental shelf o2.5 m deep Continental shelf 2.5–10 m deep Continental shelf 10–250 m deep Canyons and ridges Open ocean Open ocean Continental slope Continental shelf
10 km 1–5 km 50 m
50 m 20 m
Mid-ocean ridges
Rift valleys
100 m
Deep seafloor fabric Paleoceanography Subsurface geology
Abyssal hills
250 m
Few meters Few meters 10 m
Continental shelf Continental margin
500 m 5 km
10–25 m
Tsunami models
Uncertainty
Bathymetric uncertainty can most easily be minimized before undertaking a single survey by designing the data collection to cover the entire area to be mapped in a cohesive fashion, using the same positioning and sounding instruments. Once a survey has been completed, the reliability with which we can determine uncertainty is completely dependent upon the degree to which the bathymetric data are redundant (repeated measurements at track crossovers). With sufficient crossovers, working only from the measurements themselves, it is possible to
299
Tides
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100 m 100–500 m 500 m 42 km 5 km 1 km 100 m
10 m
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First, resolution of bottom features by acoustic measurements varies with the beam width of the echo sounder. As the beam spreads out away from the instrument, an increasingly large area of the seafloor (footprint) is ensonified, meaning that larger and larger features are masked as depth increases. (This is why deep tow is finer resolution.) Thus, even within the same map area, measurement resolution will not be the same if depth varies by hundreds of meters. Second, the spacing and spatial arrangement of the depth measurements determines resolution. This takes a bit of explanation and a unit of measurement is needed. Horizontal resolution may be expressed by considering the surface to be made up of a series of superimposed waves. All Earth surfaces, be they the seafloor or landscapes, are made up of features of many different sizes. For example, mid-ocean ridges are large, and contain smaller mountains and foothills, which can have smaller valleys and ridges on them, which in turn may have ledges, spurs, steps. A ‘wavelength’ is the horizontal distance from the peak of one wave to the peak of the next (so that mountains might have a wavelength of tens of kilometers while ledges have one measured in tens of meters). Where the spacing between measurements is greater than any of the wavelengths that comprise the seafloor, those wavelengths cannot be captured; consequently, horizontal resolution is limited to twice the measurement spacing. Wavelengths that can be reconstructed from the measurements are closer to the original seafloor when the sampling interval is less than the original wavelength, but only when the sampling interval is less than one half the original wavelength are the regenerated wavelengths of a size similar to the original. Consequently, evenly spaced measurements can only capture features whose half wavelength is less than twice the distance between measurements. This is called the Nyquist frequency. Spacing between depth measurements takes several forms. Although single-beam echo sounders record continuously along-track, the depths they record are usually stored at discrete intervals, with twice the spacing being the Nyquist frequency in the direction of the track. The spacing between one track and another may be different than the spacing within a track, and where it is the Nyquist frequency will be twice the line spacing. Consequently, there can be different resolutions in different directions within the same map if made from single-beam soundings. However, in the last few years, the introduction and deployment of the MBES is changing bathymetric measurements, having an impact similar to the one that the invention of the telescope had on astronomy. Rather than collecting a profile, MBES measures the
(almost) complete surface of the seafloor on both sides of the instrument out to a distance of 2–5 times the water depth. With the proper data processing, it produces maps very similar to the familiar topographic maps on land. Within the area ensonified, MBES produces a grid of depths, whose spacing yields a Nyquist frequency that is much smaller than that normally obtainable by single-beam sounders. It sometimes may happen that the measurements create wavelengths on the map that do not exist on the seafloor: this is termed aliasing. All too often overlooked is the fact that maps also have a vertical scale. On bathymetric maps, vertical scale is expressed in the contour interval, since the minimum height of features that the map can resolve is twice the contour interval. For features of larger amplitude than this, contours portray the undulations of the land surface quite well, but smaller features are masked.
Making a Bathymetry Map of a Large Area Data Assembly
Bathymetry for large areas includes portions of the deep ocean, where there have been very few routine surveys. Instead, there are usually an assortment of tracks from research cruises on which the data were measured from a variety of platforms, using different positioning and sounding systems, using (or not using) different sound velocities, units and plotting methods, collected over a time span of many years. Although much of the data will be held by the National Geophysical Data Center and other data centers worldwide, many are scattered in other data centers and research institutions round the world. Assembling these data usually includes identification and accumulation, assessment of quality conversion of analog records to digital form, incorporation of digital data sets into an efficient data management system, construction of plots to portray the distribution of observations within the study area, reduction of all observations to a consistent set of processing parameters (velocity of sound in seawater, seasonal variations of sound velocity), the correction of obvious data errors (blunders), and elimination of bad data sets. The data so assembled usually have random orientation, widely variable spacing of tracks and within tracks, widely variable uncertainty in position and depth corrections and some redundancy. Increasingly, the data will consist of a mixture of MBES and single-beam data. Making maps from this data assemblage is more difficult than working from single data sets; hence, this can be
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Table 2
301
Summary of characteristics of the principal types of bathymetric data
Data type
Spot SDG
Single beam E/S profile
Swath/multibeam
Satellite altimetry
Age Distribution
Old, aging Limited, remote areas Twice spacing
1950s to present Most of world
Recent Limited areas High
Small scale None None Very low
High along track, 2 line spacing across All scales Good along track, poor across track Variable with layout Good along track, poor across track
Recent World except Poles See text
Sketchy
Needs high level
Resolution, wavelength Shown on Detail Redundancy Coverage Interpretation
Large scale High High High within survey area Automatic
Small scale Poor High High Automatic
Spot or single soundings are included as a base level. Single-beam echo sounding is being replaced by multibeam, a trend which will continue. These characteristics must be kept in mind when working with the different types of data together.
treated as numerical exercise (algorithm) only up to a point and requires interpretation and consideration of other types of data (see Table 2).
Interpretation
Mapping of the land topography is largely an automatic process since the land surface can be seen by the human eye and can be photographed and imaged. However, we have never seen the seafloor, over areas larger than tens of square meters, and have bounced sound off only a small percentage of it, so that producing maps and grids is far from automatic. There are, however, two approaches to how to proceed from the incomplete and possibly inefficiently arranged data available. One is based on the assumption that only the depth data should contribute to the creation of the bathymetry, while the other makes use of other related data. The first approach fits a surface to the data or calculates individual depth values at grid nodes using mathematical techniques. Different workers use different variations of mathematical contouring, triangulating, trend surfaces, and more, with varying and arguable results. In the second approach, ancillary geological, geophysical, and physical oceanographic information, as well as satellite altimetry (see next section) and sidescan sonar images, may be used in support of the acoustic data. For instance, the center lines of canyons may be established and the mapped surface generated to conform to them. Another example is that of seafloors made of bedrock having different shapes than those that are sediments, and sediments that are accumulating in a still area having different shapes to those emanating from a canyon. In areas of
sparse acoustic data, satellite altimetry gives the main direction and extent of larger features. The introduction of MBES has added another dimension to interpretation. Within the area of seafloor ensonified during a multibeam survey, there is so much data that there is no need to interpret the shape of the seafloor as part of mapping it. Furthermore, the MBES swathes include more information than just depth. They show characteristics of the surface like texture, roughness, graininess, smoothness, nature, as well as backscatter. These can be extrapolated into the areas not surveyed by MBES. Satellite Altimetry
In addition to its secondary role of guiding the interpretation of scarce acoustic data, satellite altimetry stands on its own as the most consistent and cohesive data set covering the world oceans (except for the Arctic). The beautiful world-scale maps produced by satellite altimetry techniques are widely distributed. Satellite altimetry has the advantages of producing data of uniform quality at regular and close line spacing. Despite having these advantages, satellite altimetry maps cannot be directly substituted for acoustic maps for two reasons. First, they do not measure water depth directly but calculate it from the displacement of the sea’s surface due to variations in mass on the seafloor. Because of the inverse distance squared law inherent in potential fields, the bathymetry produced is much smoother than the actual seafloor. Consequently accuracy is diminished, although this can be improved by incorporating echo sounder data into the satellite data. Second, they resolve only long wavelength (12–20 km) features of the seafloor morphology.
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Metadata
Since bathymetry data are scarce and valuable, they should be shared by submitting them to a data center. Submitted data are much more valuable when accompanied by metadata, but the preparation of metadata is often viewed as an onerous task. Nevertheless, metadata is useful to describe data sets’ attributes to enable accurate searching and allow the data set to be fully understood and used. It also builds a trail to the originator to allow proper acknowledgement of the contribution. Different data centers publish their recommended standards for metadata on their websites: although no common agreed standard exists at the time of writing, data center specialists from several countries are attempting to construct one. Uncertainty
Uncertainty over large areas is highly variable and difficult to determine, and it is almost always worse in this case than over small areas. Where measurements are very scarce, rather than trying to calculate a numerical value for uncertainty, standard practice has been to show the locations of the ships’ tracks and let the user of the map make their own judgments about the reliability. This cannot be applied to grids, however. Rates of Progress
The rate at which mapping of any natural surface progresses depends on the interrelationship of:
• • • •
the size of the surface to be mapped; the speed of the platforms that carry those tools, their amount of use, and the areas that they are deployed in; the efficiency of the mapping tools that can be applied; and the organizational structures in place to assimilate the raw data and transform it into maps.
The seafloor is twice as large as all the land surface on Earth, research ships travel at about the same speed as a rider taking a leisurely bicycle ride, and MBES is the most efficient tool yet devised to map the seafloor as it maps a swath equal in width to 2 or 3 times the water depth. Unfortunately, research ships have never been deployed in a program to systematically ensonify every part of the ocean. If they were, it would take at least 800 ship-years. Rather than a systematic and complete data set, there is a pattern of tracks that is analogous to the pattern of roads and streets on land. If data were collected along roads, in the cities, data would be closely
spaced and arranged with some regularity, but in the rural areas data would be sparse and arranged randomly. Maps of world bathymetry are thus very uneven in quality, accuracy, and resolution. Despite the absence of a systematic survey program, over the next 10 years data will be collected primarily on continental slopes to satisfy UNCLOS Article 76 requirements to map the foot of the slope and the 2500 m contour, tsunami-affected areas, primarily the shallow water run-up zone, and in areas of specialized scientific interest, for example, the Ridge Program (which focuses on mid-ocean ridges). ‘Random’ tracks that collect data will decrease due to the shift from the ‘expeditionary’ style of at-sea oceanographic data collection to repetitive measurements of the same point or small area to collect time series. Every track, especially those outside well-surveyed areas, is valuable and it is essential that all data be assembled for the common good. Need for Cooperation
Because the ocean is so large and the portion of it that any one ship can survey so relatively small, it has long been recognized that only international cooperation will accomplish the mapping of ocean basins. The first international chart of the world ocean was produced in 1903 as the General Bathymetric Chart of the Oceans (GEBCO), a series that has been continually improved and updated since then. Supporting this production is the IHO Data Center for Digital Bathymetry (IHO DCDB) operated by the National Geophysical Data Center in Boulder, CO, USA. The DCDB and a number of national data centers assemble contributions of depth data from all possible sources, perform fundamental quality control, maintain inventories of digital bathymetric data and their metadata, and make the collected data available to the world. The data centers collaborate with one another in establishing standards and sharing data. A path therefore exists for all research cruises to contribute bathymetry data so that existing bathymetry maps and grids can be continuously upgraded. Collecting and sharing of data by all research cruises is the only way to advance in the absence of a systematic data collection program.
Assembly, Interpretation, Distribution, and Updating Bathymetry maps and grids are produced at scales ranging from local to world-encompassing. The results of a single research cruise may be portrayed as a figure in a published paper or a poster at a
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conference. Some countries produce national series of bathymetry maps at a variety of scales. Larger regions are often mapped by International Oceanographic Commission’s (IOC’s) International Bathymetric Charts (IBC) projects. Maps and grids covering the entire World Ocean are produced by the GEBCO. In the maps not produced from a single cruise, maps covering smaller areas are incorporated into those covering larger areas, along with other depth data. Bathymetry normally needs updating as soon as new measurements are made and submitted, and the method production and distribution must be designed to accommodate continual updating. Early approaches of printing successive editions are being replaced with continuous updating of maps and grids using websites.
See also Biogeochemical Data Assimilation. Geomorphology. Satellite Altimetry.
Further Reading Carron MJ, Vogt PR, and Jung W-Y (2001) A proposed international long-term project to systematically map the world’s ocean floors from beach to trench: GOMaP (Global Ocean Mapping Program). International Hydrographic Review 2: 49--55. Clarke JEH (1996) Are you really getting ‘full bottom coverage’? A collection of thoughts and images. http:// www.omg.unb.ca/%7Ejhc/coverage_paper.html (modified 8 Nov. 1996). Clarke JEH (2002) The challenge of technology: Improving sea-floor mapping methodologies (part of ‘‘Visualizing the sea floor: Mapping submarine landscapes’’). Symposium on Visualizing and Looking Beyond Earth. Washington, DC: American Association for the Advancement of Science. http://www.omg.unb.ca/AAAS/ UNB_Seafloor_Mapping.html (modified 12 Feb. 2002). Clarke JH (2000) Present-day methods of depth measurement. In: Cook PJ and Carleton CM (eds.) Continental Shelf Limits: The Scientific and Legal Interface, pp. 139--159. Oxford, UK: Oxford University Press. Committee on National Needs for Coastal Mapping and Charting Ocean Division of Earth and Life Studies (2003) Voyage into the Unknown, 213pp. Washington, DC: The National Academies Press. Cormier M-H, De Moustier C, Hall JK, Mayer L, Monahan D, and Vogt P (2004) The Global Ocean Mapping Project (GOMaP): Promoting international collaboration for a systematic, high-resolution mapping of the world’s oceans. 32nd International Geological Conference, Florence, Italy, 24–27 August.
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Dunn RA, Scheirer DS, and Forsyth DW (2001) A detailed comparison of repeated bathymetric surveys along a 300-km-long section of the southern East Pacific Rise. Journal of Geophysical Research 106: 463--471. Dunne D and Sutton G (2006) Integrating marine data into Google Earth. e-Newsletter Hydro International, no. 17. http://cmrc.ucc.ie/publications/journals/Reprint_ HI_0706_Dunne.pdf (accessed Mar. 2008). Hare R (2002) Bathymetry error modelling: Approaches, advances and applications. Canadian Hydrographic Conference, Toronto, Canada, 28–31 May (CD-ROM). Holcombe TL and Moore C (2000) Data sources, management and presentation. In: Cook P and Carleton C (eds.) Sovereign Limits beneath the Oceans: The Scientific and Technological Aspects of the Definition of the Continental Shelf, pp. 230--249. Oxford, UK: Oxford University Press. IOC and UNESCO (2002) Improved global bathymetry. IOC Technical Series: Final Report of SCOR Working Group 107, p. 111. http://unesdoc.unesco.org/images/ 0012/001272/127238e.pdf (accessed Mar. 2008). Marks KM (2002) Acoustic bathymetry for altimetric bathymetry calibration studies, NOAA Laboratory for Satellite Altimetry. http://ibis.grdl.noaa.gov/SAT/Bathy. intro.html (accessed 22 Apr. 2008). Monahan D (2000) Interpretation of bathymetry. In: Cook PJ and Carleton CM (eds.) Continental Shelf Limits: The Scientific and Legal Interface, pp. 160--176. Oxford, UK: Oxford University Press. Monahan D (2004a) Altimetry applications to continental shelf delineation under the United Nations Convention on the Law of the Sea. Oceanography 17: 71--78. Monahan D (2004b) GEBCO: The second century. Looking towards a general bathymetric chart. Hydro International 8: 45--47. Sandwell DT and Smith WHF (1996) Global Bathymetric Prediction for Ocean Modelling and Marine Geophysics. http://topex.ucsd.edu/marine_topo/text/topo. html (accessed Mar. 2008). Smith WH (1998) Seafloor tectonic fabric from satellite altimetry. Annual Review Earth Planetary Science Letters 26: 697--738. Wells D, Kleusberg A, and Vanicek P (1996) A seamless vertical-reference surface for acquisition, management and display of ECDIS hydrographic data. Proceedings of the Canadian Hydrographic Conference ’96, p. 64. Fredericton, NB: Department of Geodesy and Geomatics Engineering, University of New Brunswick.
Relevant Websites http://www.awi.de – Alfred Wegener Institute: Bathymetry data and maps of polar regions. http://www.bodc.ac.uk – British Oceanographic Data Center: Producers of GEBCO Digital Atlas (GDA), free downloadable grid, information on how to use the GDA.
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http://www.ccom.unh.edu – Center for Coastal and Ocean Mapping, University of New Hampshire: Education and research in all elements of ocean mapping, with many examples. http://gebco.net – General Bathymetric Chart of the Oceans (GEBCO): Free 1 min grid, information and publications on all things bathymetric. http://mp-www.nrl.navy.mil – Global Ocean Mapping Project (US Naval Research Laboratory Marine Physics Branch): Mostly inactive site, valuable figures on deep-water data collection. http://www.ifremer.fr – IFREMER: French marine data sets. http://www.ngdc.noaa.gov – IHO Data Center for Digital Bathymetry (IHO DCDB), NOAA’s National Geophysical Data Center (NGDC): Data base of depth data submitted by National Hydrographic offices worldwide; various bathymetry data sets at Marine Geology and Geophysics Division of NGDC; World Data Center (WDC) System: NOAA links to major data centers worldwide, many with bathymetry. http://www.ices.dk – International Council for the Exploration of the Sea (ICES): Bathymetry and its use in fisheries.
http://www.imca-int.com – International Marine Contractors Association: Graphical Guide to the Offshore Survey Industry, an excellent set of interrelated diagrams showing industrial bathymetric data collection. http://www.omg.unb.ca – Ocean Mapping Group, University of New Brunswick: Animations of multibeam interaction with the seafloor. http://www.pangaea.de – PANGAEA Publishing Network for Geoscientific and Environmental Data: Listing of maps, grids and cruise data sets. http://ocean-ridge.ldeo.columbia.edu – RIDGE Multibeam Synthesis Project (ocean floor databases): Data base of multibeam data over the midocean ridges. http://topex.ucsd.edu – Satellite Geodesy at the Scripps Institution of Oceanography, University of California San Diego: Measured and estimated seafloor topography, bathymetry from space. http://www.scor-int.org – Scientific Committee on Oceanic Research (SCOR): Interrelationships of all oceanography branches.
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BEACHES, PHYSICAL PROCESSES AFFECTING A. D. Short, University of Sydney, Sydney, Australia Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 1, pp 245–255, & 2001, Elsevier Ltd.
Introduction Beach systems occur on shorelines throughout the world and are a product of three basic factors – wave, sediment, and substrate. They are, however, also influenced by tides and impacted by a range of geological, biotic, chemical, and temperature factors. Beach systems therefore occupy a considerable range of environments, from the picturesque tropical beaches, to those exposed to the full forces of North Atlantic storms, to the usually frozen shores of the Arctic Ocean. In addition formative waves can range from low to extreme, and sediment from fine sand to boulders. This article looks at how the two main processes – waves and tides – interact with sediment to form three types of beach systems: wave-dominated, tide-modified, and tide-dominated, the three representing all the world’s beach types. It then considers some of the major factors that modify beach systems on a global and regional scale, including geology, temperature, biota, and chemical processes.
Beaches Beaches are a wave-deposited accumulation of sediment lying between wave base and the limit of wave uprush or swash. The wave base is the seaward limit or depth from which average waves can entrain and transport sediment shoreward. The essential requirements for a beach to form are therefore an underlying substrate, sediment, and waves. In twodimensions as waves shoal across the nearshore zone, between wave base and the breaker zone or shoreline, they generate a concave upward profile. As the waves transform at breaking to swash, the swash builds a steep beach face or swash zone, which caps the beach system. In reality beaches are usually more complex. Waves can range in size from a few centimeters to several meters, and period from 1 to 20 s, and additional processes, particularly tides influence both the waves and the resulting beach morphology. Beaches can also be composed of sediments ranging from fine sand to boulders. In addition, the underlying
substrate, known as the geological inheritance provides both the two-dimensional foundations, as well as potential two- and three-dimensional irregularities such as reefs and headlands, all of which influence wave transformation, sediment transport, and twoand three-dimensional beach morphology. Finally when waves break across a surf zone, they generate three-dimensional flows that produce a more complex three-dimensional surf zone and beach morphology, that lies between the nearshore and swash zones. Beaches, therefore, all contain a nearshore zone dominated by wave shoaling processes; they may contain a surf zone dominated by wave breaking and surf zone currents, and finally a swash zone dominated by wave uprush and backwash (Figure 1). Beaches also range from relatively simple lowenergy, two-dimensional, tideless systems with waves surging against a uniform shoreline, to more complex higher energy systems, incorporating surf zones with three-dimensional circulation, such as rip currents, associated with highly rhythmic bars and shorelines. In addition geological inheritance can introduce a wide spectrum of bedrock and sedimentology inputs that cause additional wave transformation through variable wave refraction and attenuation, as well as influence sediment type and mobility. This article will cover the full range of physical beach systems from the low to high wave energy wave-dominated beaches. It will then address the influence of increasing tide range in the tide-modified and tide-dominated beaches. Finally it will briefly review the role of geology (e.g., headlands and reefs) and other physical, chemical, and biotic factors in modifying beach systems.
Wave-dominated Beaches Wave-dominated beaches occur in environments where waves are high relative to the tide range. This can be defined quantitatively by the relative tide range (RTR) RTR ¼ TR=Hb
½1
where TR is the spring tide range and Hb the average breaker wave height. When RTRo3 beaches are wave dominated, when 3oRTRo15 they are tidemodified and when the RTR415 they become tidedominated.
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Beach system Morphology
Subaerial beach
Inner continental shelf
Nearshore zone
Surf zone
Wave Swash Wave breaking Wave shoaling process lian Runnel Aeo Berm crest Berm Beachface Dune Swash zone Breakpoint Swash limit
Channel
Bar
Shoreline Bars: Swell coasts (0 _ 2 bars) Sea coasts (0 _ 5 bars)
1_ 30 m depth (depends on modal wave height)
Nearshore: The more protected a beach and lower the waves, the more narrow and shallow this zone.
Wave base 50 _ 100 m
~1_ 2 km
~100 m+
Figure 1 Definition sketch of a high-energy wave-dominated beach system. The beach contains a subaerial beach dominated by swash, a surf zone dominated by breaking waves and the nearshore zone across which waves shoal. (Reproduced from Short, 1999.)
Within any beach environment, wave height and period and sediment size can all vary substantially. In order to accommodate the potentially wide range of wave height–period–sediment combinations Figure 2 presents a sensitivity plot which defines the three wave-dominated beach types, reflective, intermediate and dissipative, using the dimensionless fall velocity O ¼ Hb=TWs
½2
where T is wave period (s) and Ws is sediment fall velocity (ms1). Reflective Beaches
Reflective beaches occur when Oo1, which requires a combination of lower waves, longer periods and particularly coarser sands. They occur on sandy open swell coasts when waves average less than 0.5 m, and on all coasts when beach sediments are coarse sand or coarser, including all gravel through boulder beaches. On gravel–boulder reflective beaches waves may exceed a few meters. They are, however, all characterized by a nearshore zone of wave shoaling that extends to the shoreline. Waves then break by plunging and/or surging across the base of the beach face. The ensuing strong swash rushes up the beach, combining with the coarse sediments to build a steep beach face, commonly capped by well-developed beach cusps and/or a berm (Figures 3 lower and 4). When the sediment consists of a range of grain sizes, the coarser grains accumulate as a coarser steep step
below the zone of wave breaking, at the base of the beach face. Reflective beaches therefore consist of a concave upward nearshore zone composed of wave ripples and relatively fine sand across which waves shoal. As the waves break at the base of the beach there may be a coarser step up to a few decimeters high, then a relatively steep (4–101) beach face, possibly capped with cusps. The cusps are a product of cellular circulation on the high tide beach resulting from subharmonic edge waves produced from the interaction of the incoming and reflected backwash. The high degree of incident wave reflection off the beach face is responsible for naming of this beach type, i.e., reflective. Apart from the cosmetic beach cusps and swash circulation, if present, these are essentially two-dimensional beaches with no alongshore variation in either processes or morphology. On sand beaches they also represent the lower energy end of the beach spectrum and as such are relatively stable systems, only experiencing beach change during periods of higher waves which tend to erode the berm and deposit it as an attached low tide bar or terrace (Figure 3 lower). During following lower waves the terrace quickly migrates back up onto the beach as a new berm or cusps. Intermediate Beaches
Intermediate beaches are called such as they represent a suite of beach types between the lower energy reflective and higher energy dissipative. They are the
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BEACHES, PHYSICAL PROCESSES AFFECTING
307
4
8
14 T= 12 T= 10 T=
T=
6
T=
3
T=
High
Dissipative Fine sand Higher waves Shorter periods
T
4 =1 2 1 = T 10 =
T
Moderate
2
=6
T =
8
Breaker wave height (Hb) meters
4 Intermediate Ω = 1_ 6 Fine _ medium sand Moderate _ high w aves Short_ long periods
T = 6
T
1
=
4
Low
Reflective Ω<1 Coarser sand Lower waves Longer periods
=1
0 4
2 Coarse
1
0.500 Medium
0.250
0.125 Fine
0.0405
0.0240
0.067 WS
Sand size (mm) 0.3000
0.1820
0.1100
0.0660
_1
0.01450 m s
Figure 2 Sensitivity plot of the contribution of wave height, sediment size and wave period to O and beach type. To use the chart determine the breaker height, period (T) and grain size/fall velocity (phi or m s1). Read off the wave height and grain size, then use the period to determine where the boundary of reflective/intermediate, or intermediate/dissipative beaches lie. O ¼ 1 along solid T lines, and O ¼ 6 along dashed T lines. Below the solid lines Oo1 the beach is reflective, above the dashed lines O 4 6 the beach is dissipative, between the solid and dashed lines O is between 1 and 6 and the beach is intermediate. (Reproduced from Short, 1999.)
beaches that form under moderate (Hb 4 0.5 m) to high waves, on swell and seacoasts, in fine to medium sand (Figure 2). The two most distinguishing characteristics of intermediate beaches is first a surf zone, and second cellular rip circulation, usually associated with rhythmic beach topography. Since intermediate beaches can occur across a wide range of wave conditions, they consist of four beach states ranging from the lower energy low tide terrace to the high energy longshore bar and trough (Figure 3). Intermediate beaches are controlled by processes related to wave dissipation across the surf zone which transfers energy from incident waves with periods of 2–20 s, to longer infragravity waves with periods 4 30 s. As a consequence, incoming long waves associated with wave groupiness, increase in
energy and amplitude across the surf zone and are manifest at the shoreline as wave set up (crest) and set down (trough). They then reflect off the beach leading to an interaction between the incoming and outgoing waves to produce a standing wave across the surf zone. It is believed that standing edge waves trapped in the surf zone are responsible for the cellular circulation that develops into rip current circulation, that in turn is responsible for the high degree of spatial and temporal variability in intermediate beach morphodynamics. Low tide terrace Low tide terrace (or ridge and runnel) beaches are characterized by a continuous attached bar or terrace located at low tide (Figure 5). Usually smaller waves (0.5–1 m) break
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BEACHES, PHYSICAL PROCESSES AFFECTING
Figure 4 Pearl Beach near Sydney, Australia, is a modally reflective beach composed of medium to coarse sand and receiving usually lower swell waves. Note the very narrow zone of wave breaking, well-developed beach cusps and cusp-controlled segregated swash circulation (A.D. Short).
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Figure 6 Well-developed transverse bar and rip beach with high rhythmic shoreline (megacusp horns and embayments), attached bars and well-developed rip channels, Johana Beach, Victoria (A.D. Short).
which reworks the beach to form prominent megacusp horns, in lee of bars, and embayments, in lee of channels (Figures 3 and 6). Water tends to flow shoreward over the bars, then into the rip feeder channels. The flow moves close to the shoreline and converges laterally in the rip embayment. It then moves seaward in the rip channel as a relatively narrow (a few meters), strong flow (0.5–1 m s1), called a rip current. The current usually moves through the surf zone and beyond the breakers as a rip head. This beach state has extreme spatial–longshore variation in wave breaking, surf zone and swash circulation, and beach and surf zone topography, leading to a high unstable and variable beach system. Figure 5 A wide, well-developed low tide terrace at Crowdy Head, New South Wales. Note the transition from transverse bar and rips (foreground) to mini-rips, then continuous low tide terrace (A.D. Short).
across the bar at low tide, whereas at high tide they may remain unbroken and surge up the reflective high tide beach face. At mid to low tide cellular circulation in the surf zone usually produces weak rip current flows, which may be transitory across a planar bar, or may reside in shallow rip channels (mini-rips) (Figure 3). Transverse bar and rip Transverse bar and rip beaches form under moderate waves (1–1.5 m) on swell coasts. They consist of well-developed rip channels, which are separated by shallow bars, which are attached and perpendicular or transverse to the beach. Variable wave breaking and refraction across the shallow bars and deeper rip channels leads to a longshore variation in swash height and approach,
Rhythmic bar and beach The rhythmic bar and beach state forms during periods of moderate to high waves (1.5–2 m) on swell coasts. The high waves lead to greater surf zone discharges that require deeper and wider rip feeder channels and rip channels to accommodate the flows. Rips flow in well-developed rip channels, separated by transverse bars, but the bars are detached from the beach by the wider feeder channels (Figure 7). The bar is highly rhythmic moving seaward where rips exit the surf, and shoreward over the bars, whereas the shoreline is also highly rhythmic for the reason stated above. It is not uncommon for the more energetic swash in the rip embayments to scarp the beach face, while the adjoining horns are accreting (Figure 3). Longshore bar and trough The longshore bar and trough systems are a product of periods of higher waves which excavate a continuous longshore trough between the bar and the beach. Waves break heavily on the outer bar, reform in the trough and
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BEACHES, PHYSICAL PROCESSES AFFECTING
Figure 8 Aracaju Beach in north-east Brazil is a well-developed double bar dissipative beach produced by persistent trade wind waves breaking across a fine sand beach. Note the straight beach scarp, a characteristic of dissipative beach erosion. Photograph by AD Short.
Figure 7 Rhythmic bar and beach at Netherlands. Note the shoreline rhythms crescentic bar (A.D. Short).
Egmond, The and detached
then break again at the shoreline, often producing a steep reflective beach face (coarser sand, Figure 3) or a low tide terrace (finer sand). On swell coasts waves must exceed 2 m to produce this intermediate beach state. The high level of surf zone discharge excavates the deep trough to accommodate the flow. Surf zone circulation is both cellular rip flows, as well as increasingly shore normal bed return flows (see below). Dissipative Beaches
Dissipative beaches represent the high-energy end of the beach spectrum. They occur in areas of high waves, prefer short wave periods, and must have fine sand. They are relatively common in exposed sea environments where occasional periods of high, short storm waves produce multibarred dissipative beach systems, as in the North and Baltic seas. They also occur on high-energy midlatitude swell and storm wave coasts as in northwest USA, southern Africa, Australia, and New Zealand. On swell coasts waves must exceed 2–3 m for long periods to generate fully dissipative beaches. They are characterized by a wide long gradient beach face and surf zone,
with two and more shore parallel bars forming across the surf zone (Figures 3 upper and 8). The low gradient is a product of both the fine sand, and the dominance of lower frequency infragravity swash and surf zone circulation, which act to plane down the beach. Their name comes from the fact that waves dissipate their energy across the many bars and wide surf zone. The dissipation of the incident wave energy leads to a growth in the longer infragravity energy, which becomes manifest as a strong set up and set down at the shoreline. The standing wave
