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 Volume 3: G - M 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
Permissions may be sought directly from Elsevier’s Science & Technology Rights Department in Oxford, UK: phone (+44)(0) 1865 843830; fax (+44)(0) 1865 853333; email:
[email protected]. Alternatively you can submit your request online by visiting the Elsevier website at (http://elsevier.com/locate/permissions), and selecting Obtaining permissions to use Elsevier material
Notice No responsibility is assumed by the publisher for any injury and/or damage to persons or property as a matter of products liability, negligence or otherwise, or from any use or operation of any methods, products, instructions or ideas contained in the material herein, Because of rapid advances in the medical sciences, in particular, independent verification of diagnoses and drug dosages should be made
British Library Cataloguing in Publication Data A catalogue record for this book is available from the British Library
Library of Congress Control Number: 2009935908
ISBN: 978-0-12-375044-0
For information on all Elsevier publications visit our website at www.elsevierdirect.com
PRINTED AND BOUND IN ITALY 09 10 11 12 13 10 9 8 7 6 5 4 3 2 1
(c) 2011 Elsevier Inc. All Rights Reserved.
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.
v
Editorial Advisory Board John H. S. Blaxter Scottish Association for Marine Science Dunstaffnage Marine Laboratory Oban Argyll, UK Quentin Bone The Marine Biological Association of the United Kingdom Plymouth, UK Kenneth H. Brink Woods Hole Oceanographic Institution Woods Hole MA, USA Harry L. Bryden School of Ocean and Earth Science James Rennell Division University of Southampton Empress Dock Southampton, UK Robert Clark University of Newcastle upon Tyne Marine Sciences and Coastal Management Newcastle upon Tyne, UK J. Kirk Cochran State University of New York at Stony Brook Marine Sciences Research Center Stony Brook NY, USA Jeremy S. Collie Coastal Institute Graduate School of Oceanography University of Rhode Island South Ferry Road Narragansett RI, USA
Paul G. Falkowski Departments of Geological Sciences & Marine & Coastal Sciences Institute of Marine & Coastal Sciences School of Environmental & Biological Sciences Rutgers University New Brunswick NJ, USA Mike Fashamw Southampton Oceanography Centre University of Southampton Southampton UK John G. Field MArine REsearch (MA-RE) Institute University of Cape Town Rondebosch South Africa Michael Fogarty NOAA, National Marine Fisheries Service Woods Hole MA, USA Wilford D. Gardner Department of Oceanography Texas A&M University College Station TX, USA Ann Gargett Old Dominion University Center for Coastal Physical Oceanography Crittenton Hall Norfolk VA, USA
Peter J. Cook Australian Petroleum Cooperative Research Centre Canberra, Australia
vi
Robert A. Duce Departments of Oceanography and Atmospheric Sciences Texas A&M University College Station TX, USA
w
deceased
(c) 2011 Elsevier Inc. All Rights Reserved.
Editorial Advisory Board
Christopher Garrett University of Victoria Department of Physics Victoria British Columbia, Canada
Lindsay Lairdw Aberdeen University Zoology Department Aberdeen UK
W. John Gould Southampton Oceanography Centre University of Southampton Southampton UK
Peter S. Liss University of East Anglia School of Environmental Sciences Norwich, UK
John S. Grayw Institute of Marine Biology and Limnology University of Oslo Blindern Oslo, Norway
Ken Macdonald University of California Department of Geological Sciences Santa Barbara CA, USA
Gwyn Griffiths Southampton Oceanography Centre University of Southampton Southampton UK
Dennis McGillicuddy Woods Hole Oceanographic Institution Woods Hole MA, USA Alasdair McIntyre University of Aberdeen Department of Zoology Aberdeen UK
Stephen J. Hall World Fish Center Penang Malaysia Roger Harris Plymouth Marine Laboratory West Hoe Plymouth, UK Porter Hoagland Woods Hole Oceanographic Institution Woods Hole MA, USA George L. Hunt Jr. University of California, Irvine Department of Ecology and Evolutionary Biology Irvine CA, USA William J. Jenkins Woods Hole Oceanographic Institution Woods Hole MA, USA
w
deceased
W. Kendall Melville Scripps Institution of Oceanography UC San Diego La Jolla CA, USA John Milliman College of William and Mary School of Marine Sciences Gloucester Point VA, USA James N. Moum College of Oceanic and Atmospheric Sciences Oregon State University Corvallis OR, USA Michael M. Mullinw Scripps Institution of Oceanography Marine Life Research Group University of California San Diego La Jolla CA, USA
(c) 2011 Elsevier Inc. All Rights Reserved.
vii
viii
Editorial Advisory Board
Yoshiyuki Nozakiw University of Tokyo The Ocean Research Institute Nakano-ku Tokyo Japan
Ellen Thomas Yale University Department of Geology and Geophysics New Haven CT, USA
John Orcutt Scripps Institution of Oceanography Institute of Geophysics and Planetary Physics La Jolla CA, USA Richard F. Pittenger Woods Hole Oceanographic Institution Woods Hole MA, USA Gerold Siedler Universita¨t Kiel Institut fua¨r Meereskunde Kiel Germany
Peter L. Tyack Woods Hole Oceanographic Institution Woods Hole MA, USA Bruce A. Warren Woods Hole Oceanographic Institution Woods Hole MA, USA Wilford F. Weeks University of Alaska Fairbanks Department of Geology and Geophysics Fairbanks AK, USA
Robert C. Spindel University of Washington Applied Physics Laboratory Seattle WA, USA
Robert A. Weller Woods Hole Oceanographic Institution Woods Hole MA, USA
Colin P. Summerhayes Scientific Committee on Antarctic Research (SCAR) Scott Polar Institute Cambridge, UK
w
Stewart Turner Australian National University Research School of Earth Sciences Canberra Australia
James A. Yoder Woods Hole Oceanographic Institution Woods Hole MA, USA
deceased
(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.
(c) 2011 Elsevier Inc. All Rights Reserved.
ix
x
Preface to Second Print Edition
In Memoriam During the creation of the Encyclopedia of Ocean Sciences and also in several cases prior to the publication of the electronic Second Edition, several Associate Editors or designated Associate Editors died. We specifically acknowledge their role in making this work an effective publication. They are Mike Fasham, John S. Gray, Lindsay Laird, Michael Mullin and Yoshiyuki Nozaki. J. H. Steele, S. A. Thorpe, and K. K. Turekian Editors
(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.
xi
xii
Guide to Use of the Encyclopedia
iii. To indicate material that broadens the discussion.
ACOUSTICS, ARCTIC See also: Acoustics in Marine Sediments. Acoustic Noise. Acoustics, Shallow Water. Arctic Ocean Circulation. Bioacoustics. Ice–ocean interaction. Nepheloid Layers. North Atlantic Oscillation (NAO). Ocean Circulation: Meridional Overturning Circulation. Platforms: Autonomous Underwater Vehicles. Satellite Passive-Microwave Measurements of Sea Ice. Sea Ice. Sea Ice: Overview. Seals. Seismic Structure. Tomography. Under-Ice Boundary Layer. Water Types and Water Masses.
3. Index The index will provide you with the volume and page number where the material is to be located, and the index entries differentiate between material that is a whole article, is part of an article or is data presented in a table or figure. On the opening page of the index detailed notes are provided.
4. Appendices In addition to the articles that form the main body of the encyclopedia, there are a number of appendices which provide bathymetric charts and lists of data used throughout the encyclopedia. The appendices are located in volume 6, before the index.
5. Contributors A full list of contributors appears at the beginning of volume 1.
(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.
xiii
xiv
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
xv
D H Shull Western Washington University, Bellingham, WA, USA D K Steinberg College of William and Mary, Gloucester Pt, VA, USA L Stramma University of Kiel, Kiel, Germany R N Swift NASA Goddard Space Flight Center, Wallops Island, VA, USA T Takahashi Lamont Doherty Earth Observatory, Columbia University, Palisades, NY, USA P D Thorne Proudman Oceanographic Laboratory, Liverpool, UK
A R Piola Universidad de Buenos Aires, Buenos Aires, Argentina J M Prospero University of Miami, Miami, FL, USA S Rahmstorf Potsdam Institute for Climate Impact Research, Potsdam, Germany P C Reid SAHFOS, Plymouth, UK G Reverdin LEGOS, Toulouse Cedex, France S R Rintoul CSIRO Antarctic Climate and Ecosystems Cooperative Research Centre, Hobart, TAS, Australia J M Roberts Scottish Association for Marine Science, Oban, UK P A Rona Rutgers University, New Brunswick, NJ, USA T C Royer Old Dominion University, Norfolk, VA, USA B Rudels Finnish Institute of Marine Research, Helsinki, Finland
P L Tyack Woods Hole Oceanographic Institution, Woods Hole, USA T Tyrrell National Oceanography Centre, Southampton, UK F E Werner The University of North Carolina at Chapel Hill, Chapel Hill, NC, USA E A Widder Harbor Branch Oceanographic Institution, Fort Pierce, FL, USA D J Wildish Fisheries and Oceans Canada, St. Andrews, NB, Canada A J Williams, III Woods Hole Oceanographic Institution, Woods Hole, MA, USA D K Woolf Southampton Oceanography Centre, Southampton, UK
W Seaman University of Florida, Gainesville, FL, USA
C W Wright NASA Goddard Space Flight Center, Wallops Island, VA, USA
F Sevilla, III, University of Santo Tomas, Manila,The Philippines
J D Wright Rutgers University, Piscataway, NJ, USA
L V Shannon University of Cape Town, Cape Town, South Africa
J R Young The Natural History Museum, London, UK
G I Shapiro University of Plymouth, Plymouth, UK
H J Zemmelink University of Groningen, Haren, The Netherlands
A D Short University of Sydney, Sydney, Australia
W Zenk Universita¨t Kiel, Kiel, Germany
(c) 2011 Elsevier Inc. All Rights Reserved.
xvi
Contributors
Volume 2 G P Arnold Centre for Environment, Fisheries & Aquaculture Science, Suffolk, UK
K Dyer University of Plymouth, Plymouth, UK
K M Bailey Alaska Fisheries Science Center, Seattle, WA, USA
M Elliott Institute of Estuarine and Coastal Studies, University of Hull, Hull, UK
J G Baldauf Texas A&M University, College Station, TX, USA
D M Farmer Institute of Ocean Sciences, Sidney, BC, Canada
J Bascompte CSIC, Seville, Spain
A V Fedorov Yale University, New Haven, CT, USA
A Belgrano Institute of Marine Research, Lysekil, Sweden
M J Fogarty Northeast Fisheries Science Center, National Marine Fisheries Service, Woods Hole, MA, USA
O A Bergstad Institute of Marine Research, Flødevigen His, Norway J H S Blaxter Scottish Association for Marine Science, Argyll, UK
R Fonteyne Agricultural Research Centre, Ghent, Oostende, Belgium
Q Bone The Marine Biological Association of the United Kingdom, Plymouth, UK
D J Fornari Woods Hole Oceanographic Institution, Woods Hole, USA
I Boyd University of St. Andrews, St. Andrews, UK
A E Gargett Old Dominion University, Norfolk, VA, USA
K M Brander DTU Aqua, Charlottenlund, Denmark and International Council for the Exploration of the Sea (ICES), Copenhagen, Denmark
C H Gibson University of California, San Diego, La Jolla, CA, USA
J N Brown Yale University, New Haven, CT, USA T K Chereskin University of California San Diego, La Jolla, CA, USA J S Collie Danish Institute for Fisheries Research, Charlottenlund, Denmark and University of Rhode Island, Narragansett, RI, USA G Cresswell CSIRO Marine Research, Tasmania, Australia
J D M Gordon Scottish Association for Marine Science, Argyll, UK J F Grassle Rutgers University, New Brunswick, New Jersey, USA S J Hall Flinders University, Adelaide, SA, Australia N Hanson University of St. Andrews, St. Andrews, UK P J B Hart University of Leicester, Leicester, UK
J Davenport University College Cork, Cork, Ireland
K R Helfrich Woods Hole Oceanographic Institution, Woods Hole, MA, USA
R H Douglas City University, London, UK
D M Higgs University of Windsor, Windsor, ON, Canada
S Draxler Karl-Franzens-Universita¨t Graz, Graz, Austria
N G Hogg Woods Hole Oceanographic Institution, Woods Hole, MA, USA
J T Duffy-Anderson Alaska Fisheries Science Center, Seattle, WA, USA J A Dunne Santa Fe Institute, Santa Fe, NM, USA and Pacific Ecoinformatics and Computational Ecology Lab, Berkely, CA, USA
E D Houde University of Maryland, Solomons, MD, USA V N de Jonge Department of Marine Biology, Groningen University, Haren, The Netherlands
(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
xvii
T J Pitcher University of British Columbia, Vancouver, Canada A N Popper University of Maryland, College Park, MD, USA J F Price Woods Hole Oceanographic Institution, Woods Hole, Massachusetts, USA R D Prien Southampton Oceanography Centre, Southampton, UK A-L Reysenbach Portland State University, Portland, OR, USA P L Richardson Woods Hole Oceanographic Institution, Woods Hole, MA, USA A R Robinson Harvard University, Cambridge, MA, USA M D J Sayer Dunstaffnage Marine Laboratory, Oban, Argyll, UK
B J McCay Rutgers University, New Brunswick, NJ, USA
R W Schmitt Woods Hole Oceanographic Institution, Woods Hole, MA, USA
J D McCleave University of Maine, Orono, ME, USA
J Scott DERA Winfrith, Dorchester, Dorset, UK
D Minchin Marine Organism Investigations, Killaloe, Republic of Ireland
M P Sissenwine Northeast Fisheries Science Center, Woods Hole, MA, USA
C M Moore University of Essex, Colchester, UK K Moran University of Rhode Island, Narragansett, RI, USA G R Munro University of British Columbia, Vancouver, BC, Canada J D Nash Oregon State University, Corvallis, Oregon, OR, USA A C Naveira Garabato University of Southampton, Southampton, UK
T P Smith Northeast Fisheries Science Center, Woods Hole, MA, USA P V R Snelgrove Memorial University of Newfoundland, St John’s, NL, Canada M A Spall Woods Hole Oceanographic Institution, Woods Hole, MA, USA A Stigebrandt University of Gothenburg, Gothenburg, Sweden D A V Stow University of Southampton, Southampton, UK
J D Neilson Department of Fisheries and Oceans, New Brunswick, Canada
D J Suggett University of Essex, Colchester, UK
Y Nozakiw University of Tokyo, Tokyo, Japan
U R Sumaila University of British Columbia, Vancouver, BC, Canada
R I Perry Department of Fisheries and Oceans, British Columbia, Canada S G Philander Princeton University, Princeton, NJ, USA w
Deceased.
K S Tande Norwegian College of Fishery Science, Tromsø, Norway S A Thorpe National Oceanography Centre, Southampton, UK R S J Tol Economic and Social Research Institute, Dublin, Republic of Ireland
(c) 2011 Elsevier Inc. All Rights Reserved.
xviii
Contributors
K E Trenberth National Center for Atmospheric Research, Boulder, CO, USA J J Videler Groningen University, Haren, The Netherlands
R S Wells Chicago Zoological Society, Sarasota, FL, USA D C Wilson Institute for Fisheries Management and Coastal Community Development, Hirtshals, Denmark
Volume 3 S Ali Plymouth Marine Laboratory, Plymouth, UK
K H Coale Moss Landing Marine Laboratories, CA, USA
J T Andrews University of Colorado, Boulder, CO, USA
M F Coffins University of Texas at Austin, Austin, TX, USA
M A de Angelis Humboldt State University, Arcata, CA, USA
P J Corkeron James Cook University, Townsville, Australia
A J Arp Romberg Tiburon Center for Environment Studies, Tiburon, CA, USA
B C Coull University of South Carolina, Columbia, SC, USA
T Askew Harbor Branch Oceanographic Institute, Ft Pierce, FL, USA
R Cowen University of Miami, Miami, FL, USA
R D Ballard Institute for Exploration, Mystic, CT, USA
G Cresswell CSIRO Marine and Atmospheric Research, Hobart, TAS, Australia
G Barnabe´ Universite´ de Montpellier II, France
D S Cronan Royal School of Mines, London, UK
R S K Barnes University of Cambridge, Cambridge, UK
J Csirke Food and Agriculture Organization of the United Nations, Rome, Italy
E D Barton University of Wales, Bangor, Menai Bridge, Anglesey, UK
G A Cutter Old Dominion University, Norfolk, VA, USA
D Bhattacharya University of Iowa, Iowa City, IA, USA
D J DeMaster North Carolina State University, Raleigh, NC, USA
F von Blanckenburg Universita¨t Bern, Bern, Switzerland
T D Dickey University of California, Santa Barbara, CA, USA
D R Bohnenstiehl North Carolina State University, Raleigh, NC, USA
D Diemand Coriolis, Shoreham, VT, USA
H L Bryden University of Southampton, Southampton, UK J Burger Rutgers University, Piscataway, NJ, USA S M Carbotte Lamont-Doherty Earth Observatory of Columbia University, Palisades, NY, USA G T Chandler University of South Carolina, Columbia, SC, USA M A Charette Woods Hole Oceanographic Institution, Woods Hole, MA, USA
C S M Doake British Antarctic Survey, Cambridge, UK C M Domingues CSIRO Marine and Atmospheric Research, Hobart, TAS, Australia C J Donlon Space Applications Institute, Ispra, Italy F Doumenge 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
xix
S K Hooker University of St. Andrews, St. Andrews, UK H Hotta Japan Marine Science & Technology Center, Japan G R Ierley University of California San Diego, La Jolla, CA, USA
A E Ellis Marine Laboratory, Aberdeen, Scotland, UK C R Engle University of Arkansas at Pine Bluff, Pine Bluff, AR, USA C C Eriksen University of Washington, Seattle, WA, USA V Ettwein University College London, London, UK S Farrow Carnegie Mellon University, Pittsburgh, PA, USA M Fieux Universite´ Pierre et Marie Curie, Paris Cedex, France N Forteath Inspection Head Wharf, TAS, Australia J D Gage Scottish Association for Marine Science, Oban, UK S M Garcia Food and Agriculture Organization of the United Nations, Rome, Italy
G Ito University of Hawaii at Manoa, Honolulu, HI, USA J Jacoby Woods Hole Oceanographic Institution, Woods Hole, MA, USA M J Kaiser Bangor University, Bangor, UK A E S Kemp University of Southampton, Southampton Oceanography Centre, Southampton, UK W M Kemp University of Maryland Center for Environmental Science, Cambridge, MD, USA V S Kennedy University of Maryland, Cambridge, MD, USA P F Kingston Heriot-Watt University, Edinburgh, UK G L Kooyman University of California San Diego, CA, USA
C Garrett University of Victoria, VIC, Canada
W Krijgsman University of Utrecht, Utrecht, The Netherlands
R N Gibson Scottish Association for Marine Science, Argyll, Scotland
J B Kristoffersen University of Bergen, Bergen, Norway
M Gochfeld Environmental and Community Medicine, Piscataway, NJ, USA
K Lambeck Australian National University, Canberra, ACT, Australia
H O Halvorson University of Massachusetts Boston, Boston, MA, USA
R S Lampitt University of Southampton, Southampton, UK
B U Haq Vendome Court, Bethesda, MD, USA
M Landry University of Hawaii at Manoa, Department of Oceanography, Honolulu, HI, USA
G R Harbison Woods Hole Oceanographic Institution, Woods Hole, MA, USA
C G Langereis University of Utrecht, Utrecht, The Netherlands A Lascaratos University of Athens, Athens, Greece
R M Haymon University of California, CA, USA
S Leibovich Cornell University, Ithaca, NY, USA
D L Hebert University of Rhode Island, RI, USA J E Heyning The Natural History Museum of Los Angeles County, Los Angeles, CA, USA P Hoagland Woods Hole Oceanographic Institution, Woods Hole, MA, USA
W G Leslie Harvard University, Cambridge, MA, USA C Llewellyn Plymouth Marine Laboratory, Plymouth, UK R A Lutz Rutgers University, New Brunswick, NJ, USA
(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.
GAS EXCHANGE IN ESTUARIES M. I. Scranton, State University of New York, Stony Brook, NY, USA M. A. de Angelis, Humboldt State University, Arcata, CA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1113–1119, & 2001, Elsevier Ltd.
the gas in either the gas or liquid phase.) For gases that make up a large fraction of the atmosphere (O2, N2, Ar), pgas does not vary temporally or spatially. For trace atmospheric gases (carbon dioxide (CO2), methane (CH4), hydrogen (H2), nitrous oxide (N2O), and others), pgas may vary considerably geographically or seasonally, and may be affected by anthropogenic activity or local natural sources.
Introduction Many atmospherically important gases are present in estuarine waters in excess over levels that would be predicted from simple equilibrium between the atmosphere and surface waters. Since estuaries are defined as semi-enclosed coastal bodies of water that have free connections with the open sea and within which sea water is measurably diluted with fresh water derived from land drainage, they tend to be supplied with much larger amounts of organic matter and other compounds than other coastal areas. Thus, production of many gases is enhanced in estuaries relative to the rest of the ocean. The geometry of estuaries, which typically have relatively large surface areas compared to their depths, is such that flux of material (including gases) from the sediments, and fluxes of gases across the air/water interface, can have a much greater impact on the water composition than would be the case in the open ocean. Riverine and tidal currents are often quite marked, which also can greatly affect concentrations of biogenic gases.
Gas Solubility The direction and magnitude of the exchange of gases across an air/water interface are determined by the difference between the surface-water concentration of a given gas and its equilibrium concentration or gas solubility with respect to the atmosphere. The concentration of a specific gas in equilibrium with the atmosphere (Ceq) is given by Henry’s Law: Ceq ¼ pgas =KH
½1
where pgas is the partial pressure of the gas in the atmosphere, and KH is the Henry’s Law constant for the gas. Typically as temperature and salinity increase, gas solubility decreases. (Note that Henry’s Law and the Henry’s Law constant also may be commonly expressed in terms of the mole fraction of
Gas Exchange (Flux) Across the Air/ Water Interface The rate of gas exchange across the air/water interface for a specific gas is determined by the degree of disequilibrium between the actual surface concentration of a gas (Csurf) and its equilibrium concentration (Ceq), commonly expressed as R: R ¼ Csurf =Ceq
½2
If R ¼ 1, the dissolved gas is in equilibrium with the atmosphere and no net flux or exchange with the atmosphere occurs. For gases with Ro1, the dissolved gas is undersaturated with respect to the atmosphere and there is a net flux of the gas from the atmosphere to the water. For gases with R41, a net flux of the gas from the water to the atmosphere occurs.
Models of Gas Exchange The magnitude of the flux (F) in units of mass of gas per unit area per unit time across the air/water interface is a function of the magnitude of the difference between the dissolved gas concentration and its equilibrium concentration as given by Fick’s First Law of Diffusion: F ¼ kðCsurf Ceq Þ
½3
where k is a first order rate constant, which is a function of the specific gas and surface water conditions. The rate constant, k, also known as the transfer coefficient, has units of velocity and is frequently given as k ¼ D=z
½4
where D is the molecular diffusivity (in units of cm2 s1), and z is the thickness of the laminar layer at the
(c) 2011 Elsevier Inc. All Rights Reserved.
1
2
GAS EXCHANGE IN ESTUARIES
air/water interface, which limits the diffusion of gas across the interface. In aquatic systems, Csurf is easily measured by gas chromatographic analysis, and Ceq may be calculated readily if the temperature and salinity of the water are known. In order to determine the flux of gas (F) in or out of the water, the transfer coefficient (k) needs to be determined. The value of k is a function of the surface roughness of the water. In open bodies of water, wind speed is the main determinant of surface roughness. A number of studies have established a relationship between wind speed and either the transfer coefficient, k, or the liquid laminar layer thickness, z (Figure 1). The transfer coefficient is also related to the Schmidt number, Sc, defined as: Sc ¼ n=D
½5
where n is the kinematic viscosity of the water. In calmer waters, corresponding to wind speeds of o5 m s1, k is proportional to Sc2/3. At higher wind speeds, but where breaking waves are rare, k is proportional to S1/2. Therefore, if the transfer coefficient of one gas is known, the k value for any other gas can be determined as: Scn1 =Scn2
k1 =k2 ¼ Smooth surface regime
150
Rough surface regime
_ 2/ 3
Breaking wave (bubble) regime
_ 1/ 2
Kw Sc
Kw Sc
125
½6
where n ¼ the exponent. For short-term steady winds, the transfer coefficient for CO2 has been derived as kCO2 ¼ 0:31ðU10 Þ2 ðSc=600Þ0:5
½7
where U10 is the wind speed at a height of 10 m above the water surface. Eqns [6] and [7] can be used to estimate k for gases other than CO2. In most estuarine studies wind speeds are measured closer to the water surface. In such cases, the wind speed measured at 2 cm above the water surface can be approximated as 0.5U10. In restricted estuaries and tidally influenced rivers, wind speed may not be a good predictor of wind speed due to limited fetch or blockage of prevailing winds by shore vegetation. Instead, streambed-generated turbulence is likely to be more important than wind stress in determining water surface roughness. In such circumstances, the large eddy model may be used to approximate k as: k ¼ 1:46ðDul1 Þ1=2
½8
where u is the current velocity, and l is equivalent to the mean depth in shallow turbulent systems. Much of the reported uncertainty (and study to study variability in fluxes) is caused by differences in assumptions related to the transfer coefficient rather than to large changes in concentration of the gas in the estuary.
Less soluble gas (e.g O2)
Transfer velocity cm h
_1
Direct Gas Exchange Measurements 100
75
More soluble gas (e.g. CO2)
50
25
0.25 0
0
5
0.5 10
_1
0.75 U ms 15 U ms
_1
1 20
Figure 1 Idealized plot of transfer coefficient (Kw) as a function of wind speed (u) and friction velocity (u*). (Adapted with permission from Liss PS and Merlivat L (1985) Air–sea gas exchange rates. In: Buat-Menard P (ed.) The Role of Air–sea Exchange in Geochemical Cycling, p. 117. NATO ASI Series C, vol. 185. Dordrecht: Reidel.)
Gas exchange with the atmosphere for gases for which water column consumption and production processes are known can be estimated using a dissolved gas budget. Through time-series measurements of biological and chemical cycling, gas loss or gain across the air/water interface can be determined by difference. For example, in the case of dissolved O2, the total change in dissolved O2 concentration over time can be attributed to air/water exchange and biological processes. The contribution of biological processes to temporal changes in dissolved O2 may be estimated from concurrent measurements of phosphate and an assumed Redfield stoichiometry, and subtracted from the total change to yield an estimate of air/water O2 exchange. Gas fluxes also may be measured directly using a flux chamber that floats on the water surface. The headspace of the chamber is collected and analyzed over several time points to obtain an estimate of the net amount of gas crossing the air/water interface enclosed by the chamber. If the surface area enclosed
(c) 2011 Elsevier Inc. All Rights Reserved.
GAS EXCHANGE IN ESTUARIES
by the chamber is known, a net gas flux can be determined. Some flux chambers are equipped with small fans that simulate ambient wind conditions. However, most chambers do not use fans and so do not take account of the effects of wind-induced turbulence on gas exchange. Despite this limitation, flux chambers are important tools for measuring gas exchange in environments (such as estuaries or streams) where limited fetch or wind breaks produced by shoreline vegetation make wind less important than current-induced turbulence in shallow systems. While flux chambers may alter the surface roughness and, hence, gas/exchange rates via diffusion, flux chambers or other enclosed gas capturing devices also are the best method for determining loss of gases across the air/water interface due to ebullition of gas bubbles from the sediment. Measurement of radon (Rn) deficiencies in the upper water column can be used to determine gas exchange coefficients and laminar layer thickness, which can then be applied to other gases using eqn [3]. In this method, gaseous 222Rn, produced by radioactive decay of 226Ra, is assumed to be in secular equilibrium within the water column. 222Rn is relatively short-lived and has an atmospheric concentration of essentially zero. Therefore, in nearsurface waters, a 222Rn deficiency is observed, due to flux of 222Rn across the air/water interface. The flux of 222Rn across the air/water interface can be determined by the depth-integrated difference in measured 222Rn and that which should occur based on the 226 Ra inventory. From this flux, the liquid laminar layer thickness, z, can be calculated. Other volatile tracers have been used in estuaries to determine gas exchange coefficients. These tracers, such as chlorofluorocarbons (CFCs) or sulfur hexafluoride (SF6), are synthetic compounds with no known natural source. Unlike 222Rn, these gases are stable in solution. These tracers may be added to the aquatic system and the decrease of the gas due to flux across the air/water interface monitored over time. In some estuaries, point sources of these compounds may exist and the decrease of the tracer with distance downstream may be used to determine k or z values for the estuary.
Individual Gases
plays an important role controlling atmospheric chemistry, including serving as a regulator of tropospheric ozone concentrations and a major sink for hydroxyl radicals in the stratosphere. Methane concentrations have been increasing annually at the rate of approximately 1–2% over the last two centuries. While the contribution of estuaries to the global atmospheric methane budget is small because of the relatively small estuarine global surface area, estuaries have been identified as sources of methane to the atmosphere and coastal ocean and contribute a significant fraction of the marine methane emissions to the atmosphere. Surface methane concentrations reported primarily from estuaries in North America and Europe range from 1 to 42000 nM throughout the tidal portion of the estuaries. Methane in estuarine surface waters is generally observed to be supersaturated (100% saturation B2–3 nM CH4) with R-values ranging from 0.7 to 1600 (Table 1). In general, estuarine methane concentrations are highest at the freshwater end of the estuary and decrease with salinity. This trend reflects riverine input as the major source of methane to most estuaries, with reported riverine methane concentrations ranging from 5 to 10 000 nM. Estuaries with large plumes have been observed to cause elevated methane concentrations in adjacent coastal oceans. In addition to riverine input, sources of methane to estuaries include intertidal flats and marshes, ground-water input, runoff
Table 1 Methane saturation values (R) and estimated fluxes to the atmosphere for US and European estuaries Geographical regiona
Rb
Flux CH4 (mmol m2 h1)
North Pacific coast, USA North Pacific coast, USA Columbia River, USA Hudson River, USA Tomales Bay, CA, USA Baltic Sea, Germany European Atlantic coast Atlantic coast, USA Pettaquamscutt Estuary, USA
3–290 1–550 78 18–376 2–37 10.5–1550 0.7–1580 n.a. 81–111
6.2–41.7c 3.6–8.3c 26.0c 4.7–40.4c 17.4–26.3c 9.4–15.6c 5.5c 102–1107c 0.8–14.2c 541–3375d
a
Methane (CH4)
Atmospheric methane plays an important role in the Earth’s radiative budget as a potent greenhouse gas, which is 3.7 times more effective than carbon dioxide in absorbing infrared radiation. Despite being present in trace quantities, atmospheric methane
3
For studies that report values for a single estuary, the major river feeding the estuary is provided. For studies that report values for more than one estuary, the oceanic area being fed by the estuaries is given. b R ¼ degree of saturation ¼ measured concentration/atmospheric equilibrium concentration. n.a. indicates values not available in reference. c Diffusive flux. d Ebullition (gas bubble) flux.
(c) 2011 Elsevier Inc. All Rights Reserved.
4
GAS EXCHANGE IN ESTUARIES
from agricultural and pasture land, petroleum pollution, lateral input from exposed bank soils, wastewater discharge and emission from organic-rich anaerobic sediments, either diffusively or via ebullition (transport of gas from sediments as bubbles) and subsequent dissolution within the water column. Anthropogenically impacted estuaries or estuaries supplied from impacted rivers tend to be characterized by higher water-column methane concentrations relative to pristine estuarine systems. Seasonally, methane levels in estuaries are higher in summer compared with winter, primarily due to increased bacterial methane production (methanogenesis) in estuarine and riverine sediments. Methane can be removed from estuarine waters by microbial methane oxidation and emission to the atmosphere. Methane oxidation within estuaries can be quite rapid, with methane turnover times of o2 h to several days. Methane oxidation appears to be most rapid at salinities of less than about 6 (on the practical salinity units scale) and is strongly dependent on temperature, with highest oxidation rates occurring during the summer, when water temperatures are highest. Methane oxidation rates decrease rapidly with higher salinities. Methane diffusive fluxes to the atmosphere reported for estuaries (Table 1) fall within a narrow range of 3.6–41.7 mmol m2 h1 (2–16 mg CH4 m2 day1). Using a global surface area for estuaries of 1.4 106 km2 yields an annual emission of methane to the atmosphere from estuaries of 1–8 Tg y1. Because the higher flux estimates given in Table 1 generally were obtained close to the freshwater endmember of the estuary, the global methane estuarine emission is most likely within the range of 1– 3 Tg y1, corresponding to approximately 10% of the total global oceanic methane flux to the atmosphere, despite the much smaller global surface area of estuaries relative to the open ocean. Methane is also released to the atmosphere directly from anaerobic estuarine sediments via bubble formation and injection into the water column. Although small amounts of methane from bubbles may dissolve within the water column, the relatively shallow nature of the estuarine environment results in the majority of methane in bubbles reaching the atmosphere. The quantitative release of methane via this mechanism is difficult to evaluate due to the irregular and sporadic spatial and temporal extent of ebullition. Where ebullition occurs, the flux of methane to the atmosphere is considerably higher than diffusive flux (Table 1), but the areal extent of bubbling is relatively smaller than that of diffusive flux and, except in organic-rich stagnant areas such as tidal marshes, probably does not contribute
significantly to estuarine methane emissions to the atmosphere. Methane emission via ebullition has been observed to be at least partially controlled by tidal changes in hydrostatic pressure, with release of methane occurring at or near low tide when hydrostatic pressure is at a minimum. Nitrous Oxide (N2O)
Nitrous oxide is another important greenhouse gas that is present in elevated concentrations in estuarine environments. At present, N2O is responsible for about 5–6% of the anthropogenic greenhouse effect and is increasing in the atmosphere at a rate of about 0.25% per year. However, the role of estuaries in the global budget of the gas has only been addressed recently. Nitrous oxide is produced primarily as an intermediate during both nitrification (the oxidation of ammonium to nitrate) and denitrification (the reduction of nitrate, via nitrite and N2O, to nitrogen gas), although production by dissimilatory nitrate reduction to ammonium is also possible. In estuaries, nitrification and denitrification are both thought to be important sources. Factors such as the oxygen level in the estuary and the nitrate and ammonium concentrations of the water can influence which pathway is dominant, with denitrification dominating at very low, but non-zero, oxygen concentrations. Nitrous oxide concentrations are typically highest in the portions of the estuary closest to the rivers, and decrease with distance downstream. A number of workers have reported nitrous oxide maxima in estuarine waters at low salinities (o5–10 on the PSU scale), but this is not always the case. The turbidity maximum has been reported to be the site of maximum nitrification (presumably because of increased residence time for bacteria attached to suspended particulate matter, combined with elevated substrate (oxygen and ammonium)). Table 2 presents a summary of the data published for degree of saturation and air–estuary flux of nitrous oxide from a variety of estuaries, all of which are located in Europe and North America. Concentrations are commonly above that predicted from air–sea equilibrium, and estimates of fluxes range from 0.01 mmol m2 h1 to 5 mmol m2 h1. Ebullition is not important for nitrous oxide because it is much more soluble than methane. Researchers have estimated the size of the global estuarine source for N2O based on fluxes from individual estuaries multiplied by the global area occupied by estuaries to range from 0.22 Tg N2O y1 to 5.7 Tg y1 depending on the characteristics of the rivers studied. Independent estimates based on budgets of nitrogen
(c) 2011 Elsevier Inc. All Rights Reserved.
GAS EXCHANGE IN ESTUARIES
Table 2 Nitrous oxide saturation values (R) and estimated fluxes to the atmosphere for US and European estuaries Estuary
R
Flux a N2O (mmol m2 h1)
Europe Gironde River Gironde River Oder River Elbe Scheldt Scheldt
1.1–1.6 E1.0–3.2 0.9–3.1 2.0–16 E1.0–31 E1.2–30
n.a. n.a. 0.014–0.165 n.a. 1.27–4.77 3.56
UK Colne Tamar Humber Tweed Mediterranean Amvraikos Gulf
0.9–13.6 1–3.3 2–40 0.96–1.1
0.9–1.1
1.3 0.41 1.8 E0
5
Table 3 Fluxes of carbon dioxide from estuaries in Europe and eastern USA Estuary
European rivers Northern Europe Scheldt estuary Portugal UK Clyde estuary East coast USA Hudson River (tidal freshwater) Georgia rivers
R
Fluxa CO2 (mmol m2 h1)
0.7–61.1 0.35–26.2 1.6–15.8 1.1–14.4 E0.7–1.8
1.0–31.7 4.2–50 10–31.7 4.4–10.4 n.a.
1.2–5.4
0.67–1.54
Slight supersaturation to 22.9
1.7–23
a n.a., insufficient data were available to permit calculation of this value.
0.043 7 0.0468
North-west USA Yaquina Bay Alsea River
1.0–4.0 0.9–2.4
0.165–0.699 0.047–0.72
East coast USA Chesapeake Bay Merrimack
0.9–1.4 1.2–4.5
n.a. n.a.
a
All fluxes given are for diffusive flux to the atmosphere. n.a. indicates that insufficient data were given to permit calculation of flux.
input to rivers, assumptions about the fraction of inorganic nitrogen species removed by nitrification or denitrification, and the fractional ‘yield’ of nitrous oxide production during these processes indicate that nitrous oxide fluxes to the atmosphere from estuaries is about 0.06–0.34 Tg N2O y1. Carbon Dioxide (CO2) and Oxygen (O2)
Estuaries are typically heterotrophic systems, which means that the amount of organic matter respired within the estuary exceeds the amount of organic matter fixed by primary producers (phytoplankton and macrophytes). Since production of carbon dioxide then exceeds biological removal of carbon dioxide, it follows that estuaries are likely to be sources of the gas to the atmosphere. At the same time, since oxidation of organic matter to CO2 requires oxygen, the heterotrophic nature of estuaries suggests that they represent sinks for atmospheric oxygen. In many estuaries, primary productivity is severely limited by the amount of light that penetrates into the water due to high particulate loadings in the water. In addition, large amounts of organic matter may be supplied to the estuary by runoff from agricultural and forested land, from ground water, from
sewage effluent, and from organic matter in the river itself. There are many reports of estuarine systems with oxygen saturations below 1 (undersaturated with respect to the atmosphere), but few studies in which oxygen flux to the estuary has been reported. However, estuaries are often dramatically supersaturated with respect to saturation with CO2, especially at low salinities, and a number of workers have reported estimates of carbon dioxide flux from these systems (Table 3).
Dimethylsulfide (DMS)
DMS is an atmospheric trace gas that plays important roles in tropospheric chemistry and climate regulation. In the estuarine environment, DMS is produced primarily from the breakdown of the phytoplankton osmoregulator 3-(dimethylsulfonium)-propionate (DMSP). DMS concentrations reported in estuaries are generally supersaturated, ranging from 0.5 to 22 nM, and increase with increasing salinity. DMS levels in the water column represent a balance between tightly coupled production from DMSP and microbial consumption. Only 10% of DMS produced from DMSP in the estuarine water column is believed to escape to the atmosphere from estuarine surface water, since the biological turnover of DMS (turnover time of 3–7 days) is approximately 10 times faster than DMS exchange across the air/water interface. A large part of the estuarine DMS flux to the atmosphere may occur over short time periods on the order of weeks, corresponding to phytoplankton blooms. The DMS flux for an estuary in Florida, USA, was estimated to be on the order of o1 nmoles m2 h1. Insufficient
(c) 2011 Elsevier Inc. All Rights Reserved.
6
GAS EXCHANGE IN ESTUARIES
data are available to determine reliable global DMS air/water exchanges from estuaries. Hydrogen (H2)
Hydrogen is an important intermediate in many microbial catabolic reactions, and the efficiency of hydrogen transfer among microbial organisms within an environment helps determine the pathways of organic matter decomposition. Hydrogen is generally supersaturated in the surface waters of the few estuaries that have been analyzed for dissolved hydrogen with R-values of 1.5–67. Hydrogen flux to the atmosphere from estuaries has been reported to be on the order of 0.06–0.27 nmol m2 h1. The contribution of estuaries to the global atmospheric H2 flux cannot be determined from the few available data. Carbon Monoxide (CO)
Carbon monoxide in surface waters is produced primarily from the photo-oxidation of dissolved organic matter by UV radiation. Since estuarine waters are characterized by high dissolved organic carbon levels, the surface waters of estuaries are highly supersaturated and are a strong source of CO to the atmosphere. Reported R-values for CO range from approximately 10 to 410 000. Because of the highly variable distributions of dissolved CO within surface waters (primarily as the result of the highly variable production of CO), it is impossible to derive a meaningful value for CO emissions to the atmosphere from estuaries. Carbonyl Sulfide (OCS)
Carbonyl sulfide makes up approximately 80% of the total sulfur content of the atmosphere and is the major source of stratospheric aerosols. Carbonyl sulfide is produced within surface waters by photolysis of dissolved organosulfur compounds. Therefore, surface water OCS levels within estuaries exhibit a strong diel trend. Carbonyl sulfide is also added to the water column by diffusion from anoxic sediments, where its production appears to be coupled to microbial sulfate reduction. Diffusion of OCS from the sediment to the water column accounts for B75% of the OCS supplied to the water column and is responsible for the higher OCS concentrations in estuaries relative to the open ocean. While supersaturations of OCS are observed throughout estuarine surface waters, no trends with salinity have been observed. Atmospheric OCS fluxes to the atmosphere from Chesapeake Bay have been reported to range from 10.4 to 56.2 nmol m2 h1. These areal fluxes are over 50 times greater than those determined for the open ocean.
Elemental Mercury (Hg0)
Elemental mercury is produced in estuarine environments by biologically mediated processes. Both algae and bacteria are able to convert dissolved inorganic mercury to volatile forms, which include organic species (monomethyl- and dimethyl-mercury) and Hg0. Under suboxic conditions, elemental mercury also may be the thermodynamically stable form of the metal. In the Scheldt River estuary, Hg0 correlated well with phytoplankton pigments, suggesting that phytoplankton were the dominant factors, at least in that system. Factors that may affect elemental mercury concentrations include the type of phytoplankton present, photo-catalytic reduction of ionic Hg in surface waters, the extent of bacterial activity that removes oxygen from the estuary, and removal of mercury by particulate scavenging and sulfide precipitation. Fluxes of elemental mercury to the atmosphere have been estimated for the Pettaquamscutt estuary in Rhode Island, USA, and for the Scheldt, and range from 4.2–29 pmol m2 h1, although the values are strongly dependent on the model used to estimate gas exchange coefficients. Volatile Organic Compounds (VOCs)
In addition to gases produced naturally in the environment, estuaries tend to be enriched in byproducts of industry and other human activity. A few studies have investigated volatile organic pollutants such as chlorinated hydrocarbons (chloroform, tetrachloromethane, 1,1-dichloroethane, 1,2-dichloroethane, 1,1,1-trichloroethane, trichloroethylene and tetrachloroethylene) and monocyclic aromatic hydrocarbons (benzene, toluene, ethylbenzene, oxylene and m- and p-xylene). Concentrations of VOCs are controlled primarily by the location of the sources, dilution of river water with clean marine water within the estuary, gas exchange, and in some cases, adsorption onto suspended or settling solids. In some cases (for example, chloroform) there also may be natural biotic sources of the gas. Volatilization to the atmosphere can be an important ‘cleansing’ mechanism for the estuary system. Since the only estuaries studied to date are heavily impacted by human activity (the Elbe and the Scheldt), it is not possible to make generalizations about the importance of these systems on a global scale.
Conclusions Many estuaries are supersaturated with a variety of gases, making them locally, and occasionally
(c) 2011 Elsevier Inc. All Rights Reserved.
GAS EXCHANGE IN ESTUARIES
regionally, important sources to the atmosphere. However, estuarine systems are also highly variable in the amount of gases they contain. Since most estuaries studied to date are in Europe or the North American continent, more data are needed before global budgets can be reliably prepared.
Glossary Air–water interface The boundary between the gaseous phase (the atmosphere) and the liquid phase (the water). Catabolic Biochemical process resulting in breakdown of organic molecules into smaller molecules yielding energy. Denitrification Reduction of nitrate via nitrite to gaseous endproducts (nitrous oxide and dinitrogen gas). Ebullition Gas transport by bubbles, usually from sediments. Estuary Semi-enclosed coastal body of water with free connection to the open sea and within which sea water is measurably diluted with fresh water derived from land drainage. Gas solubility The amount of gas that will dissolve in a liquid when the liquid is in equilibrium with the overlying gas phase. Henry’s Law Constant Proportionality constant relating the vapor pressure of a solute to its mole fraction in solution. Liquid laminar thickness The thickness of a layer at the air/water interface where transport of a dissolved species is controlled by molecular (rather than turbulent) diffusion. Molecular diffusivity (D) The molecular diffusion coefficient. Nitrification Oxidation of ammonium to nitrite and nitrate. Practical salinity scale A dimensionless scale for salinity. Redfield stoichiometry Redfield and colleagues noted that organisms in the sea consistently removed nutrient elements from the water in a fixed ratio (C : N : P=106 : 16 : 1). Subsequent workers have found that nutrient concentrations in the sea typically are present in those same ratios. Transfer coefficient The rate constant which determines the rate of transfer of gas from liquid to gas phase.
See also Air–Sea Gas Exchange. Air–Sea Transfer: Dimethyl Sulfide, COS, CS2, NH4, Non-Methane
7
Hydrocarbons, Organo-Halogens. Air–Sea Transfer: N2O, NO, CH4, CO.
Further Reading Bange HW, Rapsomanikis S, and Andreae MO (1996) Nitrous oxide in coastal waters. Global Biogeochemical Cycles 10: 197--207. Bange HW, Dahlke S, Ramesh R, Meyer-Reil L-A, Rapsomanikis S, and Andreae MO (1998) Seasonal study of methane and nitrous oxide in the coastal waters of the southern Baltic Sea. Estuarine, Coastal and Shelf Science 47: 807--817. Barnes J and Owens NJP (1998) Denitrification and nitrous oxide concentrations in the Humber estuary, UK and adjacent coastal zones. Marine Pollution Bulletin 37: 247--260. Cai W-J and Wang Y (1998) The chemistry, fluxes and sources of carbon dioxide in the estuarine waters of the Satilla and Altamaha Rivers, Georgia. Limnology and Oceanography 43: 657--668. de Angelis MA and Lilley MD (1987) Methane in surface waters of Oregon estuaries and rivers. Limnology and Oceanography 32: 716--722. de Wilde HPJ and de Bie MJM (2000) Nitrous oxide in the Schelde estuary: production by nitrification and emission to the atmosphere. Marine Chemistry 69: 203--216. Elkins JW, Wofsy SC, McElroy MB, Kolb CE, and Kaplan WA (1978) Aquatic sources and sinks for nitrous oxide. Nature 275: 602--606. Frankignoulle M, Abril G, Borges A, et al. (1998) Carbon dioxide emission from European estuaries. Science 282: 434--436. Frost T and Upstill-Goddard RC (1999) Air–sea exchange into the millenium: progress and uncertainties. Oceanography and Marine Biology: An Annual Review 37: 1--45. Law CS, Rees AP, and Owens NJP (1992) Nitrous oxide: estuarine sources and atmospheric flux. Estuarine, Coastal and Shelf Science 35: 301--314. Liss PS and Merlivat L (1986) Air–sea gas exchange rates: introduction and synthesis. In: Baut-Menard P (ed.) The Role of Air–Sea Exchange in Geochemical Cycling, pp. 113--127. Dordrecht: Riedel. Muller FLL, Balls PW, and Tranter M (1995) Processes controlling chemical distributions in the Firth of Clyde (Scotland). Oceanologica Acta 18: 493--509. Raymond PA, Caraco NF, and Cole JJ (1997) Carbon dioxide concentration and atmospheric flux in the Hudson River. Estuaries 20: 381--390. Robinson AD, Nedwell DB, Harrison RM, and Ogilvie BG (1998) Hypernutrified estuaries as sources of N2O emission to the atmosphere: the estuary of the River Colne, Essex, UK. Marine Ecology Progress Series 164: 59--71. Sansone FJ, Rust TM, and Smith SV (1998) Methane distribution and cycling in Tomales Bay, California. Estuaries 21: 66--77.
(c) 2011 Elsevier Inc. All Rights Reserved.
8
GAS EXCHANGE IN ESTUARIES
Sansone FJ, Holmes ME, and Popp BN (1999) Methane stable isotopic ratios and concentrations as indicators of methane dynamics in estuaries. Global Biogeochemical Cycles 463--474. Scranton MI, Crill P, de Angelis MA, Donaghay PL, and Sieburth JM (1993) The importance of episodic events
in controlling the flux of methane from an anoxic basin. Global Biogeochemical Cycles 7: 491--507. Seitzinger SP and Kroeze C (1998) Global distribution of nitrous oxide production and N input in freshwater and coastal marine ecosystems. Global Biogeochemical Cycles 12: 93--113.
(c) 2011 Elsevier Inc. All Rights Reserved.
GELATINOUS ZOOPLANKTON L. P. Madin and G. R. Harbison, Woods Hole Oceanographic Institution, Woods Hole, MA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1120–1130, & 2001, Elsevier Ltd.
Introduction
tissues, which may also deter predation. Large size also permits commensal crustaceans to live on or in the body. Thus, as we look at the diversity of gelatinous zooplankton, we should keep in mind the forces that have led to their remarkable convergence. It is impossible to deal in a short article with the entire range of phyla that have gelatinous representatives, so some of the major groups will be highlighted.
Gelatinous zooplankton comprise a diverse group of organisms with jellylike tissues that contain a high percentage of water. They have representatives from practically all the major, and many of the minor phyla, ranging from protists to chordates. The fact that so many unrelated groups of animals have independently evolved similar body plans suggests that gelatinous organisms reflect the nature of the open ocean environment better than any other group. Whether as predators or grazing herbivores, they seem particularly well adapted to life in the oligotrophic regions of the world oceans, where their diversity and abundance relative to crustacean zooplankton is often greatest. The gelatinous body plan has evolved in a world where physical parameters are relatively constant but food resources are sparse or unpredictable. Gelatinous zooplankton exhibit many common adaptations to this habitat.
Species of polycystine radiolarians form large gelatinous colonies up to several meters in length. Thousands of individual protists are embedded in a common gelatinous matrix from which their pseudopodia extend into the water. The combined efforts of individuals in the colony enable relatively large plankton (such as copepods) to be captured and ingested. In addition to the protistan members of the colony, the matrix also contains symbiotic dinoflagellates (zooxanthellae) that grow on the metabolites of the radiolarians. In turn, the radiolarians digest the zooxanthellae, so that these colonies are planktonic homologues of coral reefs.
•
Medusae
•
•
• •
Transparent tissues provide concealment in the upper layers of the ocean, an environment without physical cover. Transparency is less common below the photic zone. The high water content of gelatinous tissues gives the organisms a density very close to that of sea water. The resulting neutral buoyancy decreases the energy required to maintain depth, but may actually require more energy overall, because of drag. The environment lacks physical barriers, strong turbulence, and current shears, so that gelatinous bodies do not need great structural strength. However, fragility makes many species difficult to sample or handle, and excludes most from more energetic coastal environments. High water content and noncellular jelly permit rapid growth and large body sizes, which can act as, or produce, large surfaces for the collection of food. Relatively large size makes gelatinous animals too big to be attacked by some predators, while their high water content reduces the food value of their
Taxonomic Groups Radiolaria
The phylum Cnidaria has many gelatinous representatives, comprising various groups of medusae and the strictly oceanic siphonophores (see below). What are commonly called jellyfish are medusae belonging to three Classes of the Cnidaria — the Hydrozoa, the Scyphozoa, and the Cubozoa. Since the morphology and life history of all three groups is broadly similar, it is practical to treat them together here. There are perhaps 1000 species of hydro- and scyphomedusae, probably with more to be discovered, especially in deep or polar waters. All are carnivorous, capturing prey with specialized stinging cells, called nematocysts. A wide variety of prey is eaten by different medusae, ranging from larval forms and small crustaceans to other gelatinous animals and large fish. Many epipelagic medusae also harbor zooxanthellae, and presumably they share their resources in the same way as the polycystine radiolarians. Many of these medusae are part of a life history that alternates between a sessile, benthic, asexually reproducing polyp and a sexually
(c) 2011 Elsevier Inc. All Rights Reserved.
9
10
GELATINOUS ZOOPLANKTON
reproducing and dispersing planktonic medusa. However, many oceanic species have lost the polyp stage and evolved instead a variety of sexual and asexual reproductive mechanisms that do not require a benthic habitat. There are several classification schemes for Cnidarians; the group names given here are common usage, but these vary in different taxonomies. Anthomedusae This order of hydromedusae includes small species ranging in size from less than 1 mm to several centimeters. The umbrella is usually shaped like a tall bell, and gonads are almost always found on the sides of the central stomach. There are four radial canals connecting the stomach to a marginal ring canal. Tentacles occur in varying numbers around the umbrella margin and sometimes around the mouth. Anthomedusae alternate with polyp forms, but some also bud medusae directly (Figure 1A). Leptomedusae These medusae are generally flatter than a hemisphere. They usually have four radial canals, but sometimes eight or more, or canals that are branched. Gonads are located on the radial canals, and there may be various sense organs on the margin. The stomach is sometimes flat, and sometimes mounted on a peduncle that can be quite long. There are tentacles around the margin but not the mouth. Leptomedusae also alternate with hydroids, but some species produce new medusae by budding or fission (Figure 1B). Limnomedusae Both high and low umbrella shapes are found in this order. There are usually four radial canals, sometimes branched. Gonads are either on the stomach or on the radial canals, and there is alternation of generations. Species of limnomedusae live in brackish, fresh (one species), or marine environments. Trachymedusae These medusae in the order Trachylina do not alternate generations but develop young medusae directly from planula larvae, or by asexual budding. The umbrella is often high, with stiff mesoglea and well-developed muscle fibers. Most have eight unbranched radial canals and gonads located on them. Many trachymedusae live in deep water and are heavily pigmented (Figure 1D). Narcomedusae Also in the Trachylina, narcomedusae have direct development from planulae, with a larval stage that is often parasitic on other medusae. There are no radial canals, but the flat central stomach is very wide and, in some genera,
extends into radial stomach pouches. Tentacles are solid and stiff, and often extend aborally. Narcomedusae are common in epipelagic and mesopelagic environments; some are strong vertical migrators (Figure 1C). Coronatae This order of scyphomedusae includes mainly deep-water species. The umbrella is divided into a high central part and a thinner marginal part by a coronal groove. The margin of the bell is divided into lappets; sense organs and solid tentacles arise from the cleft between lappets. The mouth has simple lips and the gastrovascular cavity is often deeply pigmented (Figure 1G). Semaeostomae The familiar large jellyfish are mainly in this order of the Scyphozoa. The umbrella margin is divided into lappets, and bears sense organs and hollow tentacles. There is no coronal groove around the umbrella. The mouth opening is surrounded by four long oral arms, often frilled. Gonads are in folds of the subumbrella (Figure 1E). Rhizostomae Medusae in this order of the Scyphozoa are mainly coastal species and can attain large size. They lack tentacles for prey capture, and instead ingest small particles carried into numerous small mouth openings by water currents. Some species in tropical waters host intracellular symbiotic algae. Cubomedusae Medusae in the class Cubozoa also alternate with a benthic polyp form, although details of their life cycles are poorly known. Cubomedusae can be quite large, and have the most virulently toxic nematocysts of any Cnidarians. Some species are responsible for human fatalities. Cubomedusae are also unusual in possessing complex, image-forming eyes, which are not as well developed in other medusae (Figure 1F). Siphonophores
The Order Siphonophora comprises a large and diverse group of predatory Cnidarians in the Class Hydrozoa. Their complex life cycles and colonial morphology are very different from the relatively simple hydromedusae and it is practical to consider the siphonophores as a separate group. The colonial, or polygastric, phase of the life cycle is the largest and most familiar. In this stage, siphonophores consist of an assemblage of medusoid and polypoid zooids, which are budded asexually from a founding larval polyp. The colony may include a gas float, nectophores or swimming bells, and a series of stem
(c) 2011 Elsevier Inc. All Rights Reserved.
GELATINOUS ZOOPLANKTON
11
Figure 1 Medusae. (A) Pandea conica, an anthomedusa about 2 cm high. (B) Aequorea macrodactyla, a leptomedusa about 10 cm in diameter. (C) Cunina globosa, a narcomedusa about 5 cm in diameter. (D) Benthocodon hyalinus, a trachymedusa about 3 cm in diameter. (E) Cyanea capillata, a semaeostome scyphomedusa which can attain 1 m in diameter. (F) Carybdea alata, a cubomedusa up to 15 cm high. (G) Atolla wyvillei, a coronate scyphomedusa up to 25 cm diameter. (All photographs by L. P. Madin.)
(c) 2011 Elsevier Inc. All Rights Reserved.
12
GELATINOUS ZOOPLANKTON
groups made up of feeding polyps and tentacles. In some siphonophores the stem groups break off as dispersal and sexually reproductive stages called eudoxids. The colony can be thought of as an overgrown, polymorphic juvenile stage that eventually bears the sexually reproductive adults. These are small medusoid zooids called gonophores, which produce gametes. Siphonophores range in size from a few millimeters to over 30 m in length, and occur throughout the water column. All are predators on other small zooplankton, and many genera are known to be luminescent. The colonies are fragile, and usually break up into their various units when collected in plankton nets. For this reason, much of the taxonomy is based on the morphology of the pieces, principally nectophores, and the appearance of the intact colonial stage is not always known. The Order Siphonophora is divided into three suborders and 15 families. Cystonectae This suborder includes siphonophores that possess a float but no swimming bells, so they are at the mercy of ocean currents. The Portuguese manof-war is the most familiar example. It has a float so large that the animal rests on the surface, but most cystonect species have smaller floats and are wholly submerged. Cystonects have virulent nematocysts and capture large, soft-bodied prey such as fish and squids (Figure 2A). Physonectae These siphonophores have more complex colonies, comprising a small apical float, numerous swimming bells that form a nectosome, and a stem containing several groups of gastrozooids, tentacles, bracts, etc. The stem typically contracts when the animal is swimming, and then relaxes so that the stem and tentacles extend to maximum length for fishing. This group is a major contributor to the deep scattering layer in many regions of the ocean. The largest siphonophores (the Apolemiidae, over 30 m long) are found in this group. Physonects prey mainly on small zooplankton, and many species are strong swimmers and vertical migrators. (Figure 2C). Calycophorae In this group, which contains the largest number of species, the float is absent and the nectophores are reduced, usually to two. A sequence of stem groups is budded and breaks free as eudoxids. Calycophorans are the most diverse, widely distributed, and abundant siphonophores. They catch small zooplankton and, when feeding, their tentacles form complex three-dimensional structures in the water, reminiscent of spider webs (Figure 2B).
Ctenophores
Ctenophores are exclusively marine gelatinous animals all but a few of which are holoplanktonic. Although they superficially resemble the Cnidaria, morphological and molecular studies indicate that Cnidarians and ctenophores are not closely related. Ctenophores are predators that use tentacles equipped with ‘glue’ cells or colloblasts to capture prey. The name ‘ctenophore’ is Greek for ‘comb bearer,’ referring to the comb-like plates of fused cilia that are used for propulsion. All ctenophores initially have eight meridional rows of comb plates, although in some groups these are lost or reduced during development. The vast majority of ctenophore species fall into six orders. Cydippida This group contains many species with paired tentacles that exit the body through tentacle sheaths. Species in the family Pleurobrachiidae catch prey ranging from small crustaceans to fish, while members of the Lampeidae feed mainly on large gelatinous animals like salps. The members of one species of cydippid, Haeckelia rubra, eat medusae, and retain the nematocysts of their prey (‘kleptocnidae’) for defensive use in their own tentacles. Before this behavior was known, these nematocysts were considered strong evidence for a close relationship between Cnidarians and ctenophores (Figure 2F). Platyctenida This group is primarily benthic and is distributed widely from the Arctic to the Antarctic. Members of the family Ctenoplanidae have comb rows as adults and are found in the plankton in the Indo-Pacific; all other species in the order lose their comb rows as adults, and live primarily as creeping benthic organisms. Platyctenes have functional tentacles that capture prey. Thalassocalycida This order contains a single species, Thalassocalyce inconstans, which lives in the midwater zone. It superficially resembles a medusa in overall shape, but can easily be distinguished by its eight comb rows and paired tentacles. Lobata Members of this order all have oral lobes and auricles, specialized structures that are used in feeding. Most lobates move through the water with their oral lobes widely spread to form a sort of basket. Small prey, such as crustaceans, are trapped on the mucus-covered oral lobes and tentilla, which stream over the body or extend onto the oral lobes. Ctenophores in the family Ocyropsidae lack
(c) 2011 Elsevier Inc. All Rights Reserved.
GELATINOUS ZOOPLANKTON
13
Figure 2 Siphonophores and Ctenophores. (A) Rhizophysa filiformis, a cystonect siphonophore up to 2 m long. (B) Sulculeolaria sp. a calycophoran siphonophore up to 1 m long. (C) Physophora hydrostatica, a physonect siphonophore about 10 cm high. (D) Ocyropsis maculata, a lobate ctenophore about 8 cm in diameter. (E) Beroe cucumis, a beroid ctenophore up to 25 cm long. (F) Mertensia ovum, a cydippid ctenophore about 4 cm high. (Photographs A, C, D, F by G. R. Harbison; B, E by L. P. Madin.)
(c) 2011 Elsevier Inc. All Rights Reserved.
14
GELATINOUS ZOOPLANKTON
tentacles, and capture prey by enclosing them in their muscular oral lobes (Figure 2D). Cestida These ctenophores are shaped like long, flat belts. They appear to be related to the Lobata, but lack oral lobes and auricles. There are only two genera (Cestum and Velamen) in one family (Cestidae). The comb rows extend along the aboral edge of the ribbonlike body, propelling the animal with the oral edge forward. Small prey are captured by the fine branches of the tentacles that lie over the flat sides of the body. Cestids are characteristic of oceanic, epipelagic environments. Beroida Beroids lack tentacles altogether. Their large stomodaeum occupies most of the space in the body. All beroids are predators on other ctenophores, and occasionally salps. They capture prey by engulfing them, and can bite off pieces of the prey with specialized macrocilia located immediately behind the mouth (Figure 2E).
others of which (Pseudothecosomata) have gelatinous shells and tissues. There are over 30 species of euthecosome pteropods, in two families, the Limacinidae and the Cavoliniidae. Thecosome pteropods feed by collecting particulate food on the surface of a mucous web or bubble, produced by mucus glands on the wings, and held above the neutrally buoyant and motionless animal. The mucus is periodically retrieved and ingested along with adhering particles, then replaced by a newly secreted web. Some cavoliniids have brightly colored mantle appendages that may aid in maintaining neutral buoyancy or serve as warning devices to predators. When disturbed, animals lose their neutral buoyancy and rapidly sink (Figure 3A). The Pseudothecosomata includes three families, the Peraclididae (one genus), the Cymbuliidae (three genera), and the Desmopteridae (one genus). Pseudothecosomes are larger than euthecosomes, and their mucous webs are correspondingly larger, reaching over a meter across in Gleba cordata (Figure 3C).
Heteropods
The Phylum Mollusca contains many gelatinous representatives, and the gelatinous body plan has apparently arisen independently in several groups. The Heteropoda is a superfamily of prosobranch gastropods that includes the families Atlantidae, Carinariidae, and Pterotracheidae. Heteropods are visual predators with well-developed eyes and a long proboscis containing a radula. Atlantid heteropods have thin, flattened shells into which they can completely withdraw their bodies. They feed on small crustaceans and other mollusks. The family Carinariidae includes eight species in three genera, Carinaria, Pterosoma, and Cardiapoda. These heteropods have a greatly reduced shell, enclosing only a small fraction of the body. Carinariids feed primarily on other gelatinous organisms, such as salps, doliolids and chaetognaths. In the family Pterotracheidae, with two genera — Pterotrachea (four species) and Firoloida (one species) — the shell is completely absent (Figure 3D). Pteropods
This molluskan group comprises two orders in the gastropod subclass Opisthobranchia. The foot in pteropods is modified into two wingshaped paddles responsible for swimming; their fluttering gives rise to the common name for pteropods, sea butterflies. Thecosomata This group contains the shelled pteropods, some of which (Euthocosomata) have calcareous shells and are not truly gelatinous, and
Gymnosomata Members of this order are poorly known, largely because they have no shells and contract into shapeless masses when preserved. Most species live in the deep sea, and only a few of the approximately 50 species have been observed alive. Gymnosomes appear to be highly specialized predators on particular species of thecosome pteropods. The order is divided into two suborders, the Gymnosomata and the Gymnoptera. The four families of the Gymnosomata include the Pneumodermatidae (seven genera and 22 species) with sucker-bearing arms similar to cephalopod tentacles; the Notobranchaeidae (one genus and eight species), with suckerless feeding arms called buccal cones; the Clionidae (eight genera and 16 species); and the Cliopsidae (two genera and three species) (Figure 3B). There are two families in the Gymnoptera, the Hydromylidae (one genus, one species) and the Laginiopsidae (one genus, one species). These groups are very different from each other and from other gymnosome pteropods, and some may not actually belong in the Order Gymnosomata. Cephalopods
Although many cephalopods are active, muscular swimmers, there are several gelatinous and/or transparent genera. The family Cranchiidae is composed entirely of gelatinous species, including the genera Taonius, Megalocranchia, and Teuthowenia. These relatively large, slow-moving squids probably
(c) 2011 Elsevier Inc. All Rights Reserved.
GELATINOUS ZOOPLANKTON
15
Figure 3 Mollusks and Polychaete. (A) Cuvierina columnella, a euthecosome pteropod about 1 cm high. (B) Clione limacina, a gymnosome pteropod about 2 cm high. (C) Corolla spectabilis, a pseudothecosome pteropod about 10 cm in diameter, with mucous web in background. (D) Carinaria sp. a heteropod about 10 cm long. (E) Teuthowenia megalops, a cranchiid cephalopod about 10 cm long. (F) Alciopid polychaete worm, up to 1 m long. (Photographs A, B, C, E, F by G. R. Harbison; D by L. P. Madin.)
(c) 2011 Elsevier Inc. All Rights Reserved.
16
GELATINOUS ZOOPLANKTON
capture prey through stealth rather than active pursuit. Vitreledonelliid octopods are also gelatinous (Figure 3E). Polychaete Worms
Two major groups of planktonic polychaetes are gelatinous, the Alciopidae and the Tomopteridae. Both are in the order Phyllodocida, although they are probably not closely related. Alciopids are characterized by well-developed eyes with lenses. Many have ink glands along the sides of their bodies, which may function analogously to the ink glands of cephalopods. Their habits are poorly known, but they may feed on gelatinous prey. Alciopids may attain lengths of nearly a meter. Tomopterids do not have well-developed eyes, but probably use chemoreception to locate prey. Some release luminous secretions from glands along their body when disturbed. Deep-sea tomopterids may be 25 cm long, but most shallow species are much smaller. (Figure 3F). Crustaceans
Although arthropods cannot really be considered gelatinous because of their exoskeletons, there are some examples of very transparent bodies, presumably also an adaptation for concealment. The most notable examples are species of the hyperiid amphipods Cystisoma and Phronima. Species of Cystisoma are large and transparent, and the enormous retinas of the compound eye are lightly tinted. Although the retinas of species of Phronima are darkly pigmented, the rest of the body is transparent. It is likely that the transparency of these species is a form of protective coloration, since they live on transparent gelatinous hosts. Holothurians
Although the majority of holuthurian species are rather sedentary benthic deposit feeders, there are several deep-sea genera of swimming or drifting holothurians with gelatinous bodies. Species in the genera Peniagone and Enypniastes feed on bottom deposits, but can swim up into the water column. The genus Pelagothuria appears to be wholly pelagic, with a morphology that suggests it collects and feeds on sinking particulate matter. Few pelagic holuthurians have been observed alive and little is known of their life history or behavior. Pelagic Tunicates
The subphylum Urochordata includes two classes of pelagic tunicates, the Thaliacea and the Larvacea or
Appendicularia. Thaliaceans (including the orders Pyrosomida, Doliolida, and Salpida) are relatively large animals with more or less barrel-shaped bodies. They pump a current of water through their bodies and strain phytoplankton and other small particulates from it with a filter made of mucous strands. The same current provides jet propulsion. Thaliaceans have complex life cycles with alternating generations and multiple zooid types. The class Larvacea comprises a single order of small organisms that filter food particles using an external mucous structure called a house. Both Thaliaceans and Larvaceans are widely distributed in the oceans, and are sometimes extremely abundant. Pyrosomida Pyrosomes form colonies made up of numerous small ascidian-like zooids embedded in a stiff gelatinous matrix or tunic. The colony is tubular, with a single terminal opening. Water is pumped by ciliary action through each zooid, and suspended food particles are retained on the branchial basket within the body. The excurrent water from each zooid passes into the lumen of the colony, forming a single exhalent current that provides jet propulsion. Most pyrosome colonies are a few centimeters to a meter in length, but colonies of at least one species can attain lengths of 20 m. Doliolida This order of the Thaliacea comprises six genera and 23 species of small (2–10 mm), barrel-shaped animals with circumferential muscle bands. The filter feeding mechanism is similar to that of pyrosomes, with currents generated by ciliary beating passing through a mucous net supported on the branchial basket. The life cycle involves five asexual stages and one sexual stage, several of which occur together as parts of large colonies of thousands of zooids. These colonies may attain lengths over 1 m, but are fragile and are rarely collected intact. In most genera of doliolids, the life cycle begins with a sexually produced larva, which becomes the oozooid stage. This stage feeds initially, but then begins budding off the trophozooid and phorozooid stages, thus forming the colony. During this process the oozooid loses its branchial basket and gut, and transforms into the ‘old nurse’ stage, whose function is to swim by jet propulsion and pull the attached colony along behind it. Contractions of the body muscles produce short exhalent pulses that move the colony rapidly. The trophozooids in the colony filter-feed to support themselves and the nurse. The phorozooids grow attached to the colony, but then break free to lead independent lives and produce asexually a
(c) 2011 Elsevier Inc. All Rights Reserved.
GELATINOUS ZOOPLANKTON
small group of gonozooids. These eventually break free from the phorozooid, and become the sexually reproducing stages (hermaphrodites?) that produce the larvae and begin the whole cycle again (Figure 4B). Salpida This order (with 12 genera and about 40 species) is of larger filter-feeding animals, also with
17
circumferential muscle bands. The salps alternate between two forms, an asexually budding solitary (oozooid) stage and a sexually reproducing aggregate (blastozooid) stage. The aggregate salps usually remain connected together in chains or whorls of various types. Swimming is by jet propulsion, produced by a pulsed water current generated by rhythmic contraction of body muscles.
Figure 4 Pelagic Tunicates. (A) Megalocercus huxleyi, a larvacean of about 5 mm body length, house length about 4 cm. (B) Dolioletta gegenbauri, portion of a colony showing gastrozooids and phorozooids, individuals 2–5 mm long, colonies up to 1 m. (C) Salpa maxima, solitary generation salp, up to 25 cm long. (D) Salpa maxima, chain of aggregate generation salps; orange dots are guts of salps; individuals are to 15 cm, chains up to 10 m long. (E) Pegea socia, aggregate generation salp with attached embryo of solitary generation; aggregate 7 cm, embryo about 1 cm. (F) Traustedtia multitentaculata, solitary generation salp with appendages of uncertain function, about 3 cm long. (All photographs by L. P. Madin.)
(c) 2011 Elsevier Inc. All Rights Reserved.
18
GELATINOUS ZOOPLANKTON
Food particles are strained from the water passing through the body cavity by a mucous filter, which is continuously secreted and ingested. The individual animals range in size from 5 to over 100 mm, and chains can be several meters long (Figures 4C–F). Larvacea This class (also called Appendicularia) is divided into three families (with 15 genera and 70 species) of small (1–10 mm) animals consisting of a trunk and long, flat tail. Larveaceans are also filter feeders on small particulates but are unique among tunicates in the use of an external concentrating and filtering structure called the house. The house surrounds the animal, and contains a complex set of channels and filters made of mucous fibers and sheets. Water is pumped into the house by the oscillation of the larvacean’s tail; the exhalent stream provides slow jet propulsion in some species. Particles are sieved from the flow as it passes through the internal filter; they accumulate and are aspirated at intervals into the pharynx of the larvacean via a mucous tube. The complex house is formed as a mucous secretion on the body of the larvacean, produced by specialized secretory cells. It is inflated with sea water, pumped into it by action of the tail, until it attains its full size, with all the internal structures expanding in proportion. Houses eventually become clogged with particulates and fecal pellets, and are then jettisoned. The larvacean expands a new house (there may be several house rudiments on its body, awaiting expansion) and resumes filter feeding. The abandoned houses can be an important source of marine snow and serve as food for various planktonic scavengers (Figure 4A).
Ecology of Gelatinous Zooplankton Gelatinous zooplankton are found in all of the oceans of the world, from the tropics to polar regions. They also occur at all depths, and many of the largest and most delicate species have been collected in recent years from the mesopelagic and bathypelagic parts of the ocean. The absence of turbulence in the deep sea probably allows these species to attain such large sizes, but there are also robust species that thrive in surface and coastal environments. Examples include the Portuguese man-of-war (Physalia physalis), which lives at the air–water interface and can ride out hurricanes, and the ctenophore Mnemiopsis leidyi, which lives in estuaries with strong tidal currents and turbulence. In general, gelatinous organisms have been rather neglected by zooplankton ecologists, primarily because their delicacy makes them difficult to sample
and study. Most are damaged or destroyed in conventional plankton nets, and many deep-water siphonophores and ctenophores are too delicate to be captured intact even with the most gentle of techniques. Much recent progress in understanding their biology has been based on in situ methods of study using SCUBA diving, submersibles, or remote vehicles. These methods permit observation of undisturbed behavior and collection of intact living specimens. Advances in culture techniques and laboratory measurements have improved our understanding of energetics, reproduction, and life history of some species, but most remain only partially understood. Gelatinous animals occupy every trophic niche, ranging from primary producers (symbiotic colonial radiolaria) to grazers (pteropods and pelagic tunicates) and predators (medusae, siphonophores, and ctenophores). In all these niches, the gelatinous body plan confers advantages of size and low metabolic costs. In addition to attaining large sizes with relatively little food input, gelatinous organisms such as medusa and ctenophores are able to ‘de-grow’ when deprived of food. Metabolic rates remain unchanged, and the animal simply shrinks until higher food levels allow it to resume growth. This energetic flexibility is probably important to the success of gelatinous species in the oligotrophic open ocean and deep sea. Many species of medusae and siphonophores, for example, appear able to survive at low population densities spread over very large areas. In other cases the efficiency of their feeding, growth, and reproduction allows gelatinous species to outcompete other types of zooplankton and form dense populations over large areas, which can have considerable impact on ecosystems. A dramatic recent example was the accidental introduction of the ctenophore Mnemiopsis leidyi into the Black Sea from the eastern seaboard of the Americas. In the late 1980s, this ctenophore reproduced in prodigious quantities, and the resulting predation on zooplankton and larval fishes led to the collapse of pelagic fisheries in the Black Sea. These fisheries have to some extent recovered, but seasonal blooms of this ctenophore continue to occur in the Black Sea, just as they do on the eastern shores of the Americas. Reports of Mnemiopsis in the Caspian Sea suggest that the pattern may be repeated. Many other gelatinous species form dramatic seasonal blooms, such as the medusa Chrysaora quinquecirrha in the Chesapeake Bay, the salp Thalia democratica off Florida, Georgia, and Australia, and the medusa Pelagia noctiluca in the Mediterranean. In the Southern Ocean, immense populations of the salp Salpa thompsoni alternate with those of the
(c) 2011 Elsevier Inc. All Rights Reserved.
GELATINOUS ZOOPLANKTON
Antarctic krill Euphausia superba. The formation of large aggregations through rapid reproduction appears to be a common strategy for taking advantage of favorable conditions. Dense populations are sometimes further concentrated by wind or current action, or are transported close to the coast from their normal habitats farther offshore. The combination of rapid growth and advection can cause the sudden appearance of swarms of medusae, ctenophores, or salps in coastal waters. Although these blooms may sometimes have serious or even catastrophic effects on other organisms, including fisheries or human activities, they are a natural part of the life histories of the species, and not events for which remedial action is needed, or even possible. Gelatinous zooplankton are normal components of virtually all planktonic ecosystems. They are among the most common and typical animals in the oceans, whose biology and ecological roles are now becoming better understood.
19
See also Plankton. Zooplankton Sampling with Nets and Trawls.
Further Reading Bone Q (1998) The Biology of Pelagic Tunicates. Oxford: Oxford University Press. Harbison GR and Madin MP (1982) The Ctenophora. In: Parker SB (ed.) Synopsis and Classification of Living Organisms. New York: McGraw Hill. Lalli CM and Gilmer RW (1989) Pelagic Snails. Stanford, CA: Stanford University Press. Mackie GO, Pugh PR, and Purcell JE (1987) Siphonophore biology. Advances in Marine Biology 24: 98--263. Needler-Arai M (1996) A functional biology of Scyphozoa. London: Chapman and Hall. Wrobel D and Mills CE (1998) Pacific Coast Pelagic Invertebrates: A Guide to the Common Gelatinous Animals. Monterey: Sea Challengers and Monterey Bay Aquarium.
(c) 2011 Elsevier Inc. All Rights Reserved.
GENERAL CIRCULATION MODELS G. R. Ierley, University of California San Diego, La Jolla, CA, USA
•
Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1130–1134, & 2001, Elsevier Ltd.
Introduction A general circulation model (GCM) of the ocean is nothing more than that – a numerical model that represents the movement of water in the ocean. Models, and more particularly, numerical models, play an ever-increasing role in all areas of science; in geophysics broadly, and in oceanography specifically. It was perhaps less the early advent of supercomputers than the later appearance of powerful personal workstations (tens of megaflops and megabytes) that effected not only a visible revolution in the range of possible computations but also a more subtle, less often appreciated, revolution in the very nature of the questions that scientists ask, and the answers that result. The range of length scales and timescales in oceanography is considerable. Important dynamics, such as that which creates ‘salt fingers’ and hence influences the dynamically significant profile of density versus depth, takes place on centimeter scales, while the dominant features in the average circulation cascade all the way to basin scales of several thousand kilometers. Timescales for turbulent events, like waves breaking on shore, are small fractions of seconds, while at the opposite end, scientists have reliably identified patterns in the ocean with characteristic evolution times of order several decades.1 Most simply, a ‘model’ is no more than a mathematical description of a physical system. In the case of physical oceanography, that description includes the following elements:
•
The momentum equation (F ¼ mv_ , but expressed in terms appropriate to a continuous medium), or often in its place a derivative form, the ‘vorticity equation’.
1 It is useful to distinguish extrinsic evolution times, which span geologic time, from intrinsic variation, which characterizes an isolated ocean and atmosphere, thus neglecting such secular influences as orbital variation, change in the earth’s rotation rate, variation in the solar constant, etc. Although not all causes of variability have been identified, it is possible that even documented evolution over thousands of years may reflect the latter, intrinsic, variability.
20
• • •
An equation to express the principle of mass conservation. A heat equation, which describes the advection (carrying by the fluid) and diffusion of temperature. A similar ‘advection–diffusion’ equation for salinity. An equation of state, which relates the pressure to the density.
Although it may, after the fact, sound obvious, it took scientists many years to appreciate the full role of rotation – which enters Newton’s law through a latitude dependent Coriolis force – in generating the observed large scale circulation of ocean. Elements of these equations can become quite complicated. For example, momentum in the upper ocean is imparted by a complex, not yet fully understood, process of wind–wave interaction. One has to choose whether to represent the action of the wind kinematically, which means that the spatial and temporal variability of the wind must be given beforehand, or dynamically, in which case we must solve not only the set of equations above, but a similar set that describes the simultaneous evolution of the atmosphere. Clearly the dynamical case is the more ‘realistic’ of the two, but the price is a substantially more involved computation. It is issues such as these that force one to a choice that often pits understanding and intuition on the one hand against verisimilitude and complete dynamical consistency on the other. As the common aim of most large-scale models is to generate results of maximum realism, it usual that the models are frequently corrected, or ‘steered’, on the fly by extensive use of data assimilation. It is not merely a matter of slight refinement: no large-scale model (GCM) is yet sufficiently robust that it can be used for forecasting without considerable input of such observationally derived constraints; in the oceanic component, for example, the need to force model agreement at depth by continuously ‘relaxing’ the solution to the smoothed data of the Levitus atlas (a world-wide compendium of data from many sources, smoothed and interpolated onto a regular grid). How much of this fragility is because of explicit defects (faulty ‘subgrid scale modeling’ and explicit omissions in the physics e.g., neglect of wave breaking) and how much is due to discrete numerical implementation inconsistent with plausible continuous equations remains an open question. Their limitations notwithstanding, large numerical models are one vital means by which we grapple with questions about global warming and a host of other
(c) 2011 Elsevier Inc. All Rights Reserved.
GENERAL CIRCULATION MODELS
environmental issues that affect the way that both we and future generations will live.
instead we reach a tentative conclusion of not proven wrong (yet!). While it may stimulate conjecture, and have other worthy ends, the idea of introducing artificial parameters for the express purpose of manufacturing close agreement with reality is nonetheless formally antithetical to the paradigm of testing independent, quantitative predictions. Backward compatibility. As with confinement of form, the idea of compatibility achieves a purity in mathematics that is not to be expected of the physical sciences. Each new bit of mathematics must fit perfectly into the entire edifice of results already discovered (or invented, as you will). Commonly, though not without exception, in the physical sciences, newer theories are seen to encompass the older theories as special or limiting cases. We speak, for example, of the ‘classical limit’ as a means of recovering prerelativistic or prequantum results. Indeed, it is only in light of Einstein’s theory of special relativity that we can understand the limitations of Newton’s law, which we now understand more fully as not a law, but a limiting approximation. Oceanography has families of theories, each nesting one within the next, like a series of oceanographic Matryoshka dolls, the innermost of which is often the theory of ‘quasigeostrophy,’ which dates from the 1950s.
Models in Theory and Practice Historically, computers were initially so limited that the questions posed were often, in effect, slight extensions of preexisting analytic queries. To that extent, such studies continued to conform (at least in principle) to what we may identity as four basic building blocks of most theories, which, in aiming to describe physical reality become subject to constraints. Classical modes to which physical laws are found to conform generally2 include the following.
•
•
•
Expression in quantity, extent, and duration. The language that offers itself as encompassing all distinct physical conditions and all meaningful physical relations is fundamentally that of mathematics. As others before and since, the great mathematician Eugene Wigner too had a stab at explaining what, in an eponymous essay, he termed ‘the unreasonable effectiveness of mathematics in the physical sciences’. It remains a conundrum. Confinement of form. This attains its purest expression in the Platonic view that mathematical entities are not invented, but discovered, and as such have a prior, if not physical, existence. In extension to physical theories, it corresponds to the belief that there is some true, ultimate, equation association with any natural phenomenon. One important respect in which this idea must be tempered in application to fluids is the mathematical demonstration that a variety of different microscopic laws for interaction may all yield the same generic macroscopic law applicable to behavior at large space scales or timescales. If one’s aim is solely to understand the latter, then although for a given problem we might presume that there is indeed a precise, if complicated, microscopic law, one’s effort might more profitably be spent understanding the passage to the large-scale limit. Falsifiability. Implicit in the progress of science is the idea that one makes and then tests hypotheses. But unlike in mathematics, in the physical sciences we cannot show the hypothesis is correct, only that it is wrong. A hypothesis is never vindicated,
2 We speak gently here, since to insist that those four are always either necessary or sufficient would require that we introduce a theory about theories: a metatheory. But we have no notion of how rigorously to evaluate such ideas!
21
•
The unavoidable adoption and resulting sensitivity of GCMs to ad hoc parameterizations (e.g., subgrid scale modeling) or necessity set them apart from the traditional pursuit of the scientific method. This distinction was (presciently) appreciated at least by the early 1960s and it changes, or ought to change, one’s view of such models as rigorous arbiters of precise truths. And yet while it is true that numerical experiments with GCMs are merely suggestive rather than truly predictive of future evolution of the ocean, the sheer lack of experimental data, to say nothing of the lack of a control, means that theoretical ideas are often assessed on the basis of their success in explaining strictly numerical experiments. As computers became more powerful it was natural to press for the most realistic model runs possible. And, because the growth in computing power was increasingly realized through distributed3 as well as mainframe (super), computing, the school of ‘kitchen sink’ models, which started as a specialized branch off the mainstream of oceanography – largely
3
Both virtually, through high speed and increasingly transparent networks, and physically, through desktop workstations of considerable power.
(c) 2011 Elsevier Inc. All Rights Reserved.
22
GENERAL CIRCULATION MODELS
limited in participation to those in close physical proximity to two or three central machines – became a powerful tributary in its own right: an autonomous discipline within oceanography, which naturally began to evolve its own criteria for relevance.
Limits on Numerical Models We spoke of two sources of error common to large numerical models: difficulties in numerical implementation, and poorly modeled or unrepresented physical processes. In this section we consider specific instances of each, starting with an abstract mathematical point of view. But note that if we could with the wave of hand dispense with these two issues (which one supposes are in principle tractable), the fact that we do not know the exact physical state of the ocean inevitably increases the uncertainty of the results. Moreover, even were we given that exact state at some instant in time, the intrinsic and spontaneous genesis of disorder in such a physical system must forever constrain our predictive power. There are two key mathematical features of the basic momentum (or vorticity) equation which bear comment: conservation and dissipation. The nonlinear term is fundamental to the initiation and sustenance of turbulence and by itself strictly conserves energy. (Other terms introduce explicit dissipation.) In addition, in two dimensions the term conserves ‘enstrophy’ (the square of the vorticity) and in three, both ‘circulation’ and ‘helicity’ (the dot-product of velocity and vorticity). While one might hope for all such conservation properties to be preserved in numerical implementations, some large models, often those based on curvilinear – as opposed to Cartesian – coo¨rdinates – manage only to conserve energy. Beyond the conserved quantities associated with the nonlinear term, which include the energy and, in general, the so-called ‘Noether invariants,’ there is a more subtle property associated with the exact (continuous) equation: its associated ‘multisymplectic geometry.’ Recent mathematical advances make it possible for a discrete numerical model to preserve such structure exactly, though as yet such improvements have not been incorporated into any working GCM. Is it quantitatively important that we do so when basic fluid processes, such as convection, are as yet only crudely modeled? Until the experiment is tried, no one can say. But it is pertinent to note that a similar (that is, ‘symplectic’) refinement is critical for a numerical solution of sufficient accuracy that one can decide whether planetary orbital motions are chaotic on astronomical timescales.4 Finally, strictly speaking, not only ambitious models, but even more confined ‘process’ models, rest
upon a not yet wholly secure foundation: it is still an open research question whether the basic equation of fluid mechanics (the Navier–Stokes equation) is itself ‘globally well-posed’ in three dimensions.5 (It is widely believed to be so, but belief comes cheap. A proof, however, is worth one million dollars(!) – one of seven prizes in a competition recently announced by the Clay Mathematics Institute.) Even overlooking such foundational matters on which mainly mathematicians would cavil about numerical models, the last three of the four principles above are often violated in more apparent ways in the application and development of present-day models. We illustrate this divergence from traditional norms with a few representative examples to emphasize the sometimes causal (not casual!) link between models and ‘reality.’ In delineating the borders of the known, the unknown, and the unknowable, it is important to discriminate between deduction and rationalization as competing processes for exploring and explaining those borders.
•
Although GCM simulations with a viscosity approaching that of water are at present inconceivable, at least as a thought experiment it is worth bearing in mind that those are the numerical results we would in principle compare against observation to assess a given model.6 Short of that, GCMs use various formulations that ostensibly mimic the dynamical effects of the unresolved scales of motions. In the simplest instance, this amounts to choosing a numerical viscosity several orders of magnitude larger than that of sea water. But often the value is dictated by purely heuristic numerical considerations: it is set at a threshold value, any decrease below which leads to rapid
4
On dissipation, a deep, though perhaps insufficiently appreciated, mathematical result is that the solution of a parabolic, dissipative system quickly collapses onto a finite-dimensional ‘attractor’. This is remarkable. If you think of assigning a point to every one of N molecules of water in the ocean, and tracking the velocity and position of each, the associated ‘phase space’ – just a record of that evolution – has dimension 6N. Because the momentum equation is derived on the basis that water is an infinitely divisible continuum, strictly we need to imagine that N approaches infinity. Nonetheless, even in that infinite-dimensional phase space, it remains true that the solution confines itself to only a finite, if quite large, portion. 5 Hadamard introduced the notion of ill-posedness of partial differential equations. A problem is well-posed when a solution exists, is unique, and depends continuously on the initial data. It is ill-posed when it fails to satisfy one or more of these criteria. 6 There is a curious division among physical oceanographers as to whether the large-scale flow we observe is, in the end, actually sensitive to the precise value of the viscosity of sea water. Predictably, there are two camps: yes and no.
(c) 2011 Elsevier Inc. All Rights Reserved.
GENERAL CIRCULATION MODELS
numerical blowup. It cannot be said to be satisfactory feature that a basic parameter is set not by independent dynamical considerations but for stability reasons, and those pertaining solely to the discrete form of the equations. (The solution to the continuous equation would not blow up!) At times, not only the coefficient, but the actual form of the diffusive operator is adjusted. As above, usually the motivation is intrinsically numerical, so it is not surprising that a catalogue of the various choices shows some, for example, that create artificial sources (or sinks) of vorticity in the flow. Others, subject to the given boundary conditions, do not make mathematical sense in a region where the fluid depth tends to zero (like Atlantic City). Oceanographers, unfortunately, do not have the luxury of extensive laboratory measurements from which their dissipative parameterizations can be calibrated. A program of direct observation in particular regions where dissipation is thought to be significant is just getting under way. While such measurements will help reveal deficiencies in present formulations, the largest GCMs will probably rely on a solely heuristic approach for some time to come.
•
An important, but numerically unresolved, process in the ocean is that of convection, which typically occurs at small scales in, for example, localized regions of intense surface cooling. The overall thermal structure of the ocean is sensitive to this, a means by which ‘bottom water’ is formed; a cold, relatively less salty mass that constitutes the deep Atlantic, for example. Because the horizontal resolution is too coarse to encompass the sinking motions, various schemes have been devised to mimic that effect. It has been shown that one of the most common of these leads paradoxically to unacceptable physical (and mathematical) behavior as resolution is improved; it has no verifiable correspondence to a realizable physical process. The temptation with a model that has been extensively tuned to give plausible answers for other observables is to leave well enough alone. Unhappily, a model with one or more such elements whose limits are ill-defined or nonexistent must inevitably produce end results whose errors are typically an opaque mix of effects: some physical, some numerical, some mathematical. In such circumstances, the program of falsifiability of the physical components is apt to be fatally compromised.
The point is not that one should immediately dispense with all ad hoc parametrizations; excepting those that are simply mathematically ill-posed from
23
the outset, theorists generally do not have better alternatives to suggest. But one should always bear in mind the degree to which numerical simulations are sensitive to these components, and seek independent ways in which to constrain their parameters, in isolated settings that test the limits of prediction against known measurements or, failing that, at least against fully converged, adequately resolved simulations of a local or regional character. If, within the acceptable parameter range identified, it is found that the original model no longer gives adequate large-scale predictions, then there are more basic problems to be addressed.
Summary From the numerical side, no computer improvements that can be seen on the horizon seem likely to make reasonably ambitious GCMs accessible to rigorous and extensive parametric and numerical exploration, a prerequisite to their complete understanding. From the mathematical side, it seems to be our fundamental ignorance about turbulence that most severely restricts the range of our grasp, leaving us with an often painfully narrow range of computations to which theoretical remarks can be significantly addressed. For these structural reasons, the gulf between theory and much numerical modeling will probably continue to widen for the foreseeable future, and thus there may grow to be—indeed some would say it already exists—a division akin to C.P. Snow’s ‘Two Cultures’. All the cautions about GCMs notwithstanding, they have become an integral part of the study of physical oceanography. With due regard for the novel capacities and limitations of numerical models, such scientific progress as we do make will more and more often hinge upon judicious computation.
See also Deep Convection. Double-Diffusive Convection. Forward Problem in Numerical Models. Wind Driven Circulation.
Further Reading The literature on ocean modeling is not yet productive of definitive treatises, in large measure because the field is yet young and rapidly evolving. Thus in lieu of textbooks or similar references, the reader is directed to the following series of articles. For some predictions on the perennially intriguing issue of what improvements in large-scale modeling may be
(c) 2011 Elsevier Inc. All Rights Reserved.
24
GENERAL CIRCULATION MODELS
driven by plausible increases in computing speed with massively parallel machines see Semtner A. Ocean and climate modeling Communications of the ACM 43 (43): 81–89. For a look back at the history of one of the single most influential models in physical oceanography, see A.J. Semtner’s Introduction to ‘A numerical method for the study of the circulation of the World Ocean’, which accompanies the reprinting of Kirk Bryan’s now classic 1969 article of the title indicated. This pair appears back-toback, beginning on page 149, in Journal of Computational Physics, (1997) 135 (2).
General readers may wish to consult the following succinct review, accessible to a broad audience: Semtner AJ (1995) Modeling ocean circulation Science 269 (5229): 1379–1385. Finally, for those readers desiring a more in depth appreciation of modeling issues and their implications for specific features of the large scale circulation, consult the careful review. McWilliams JC (1996) Modeling the oceanic general circulation. In: Lumley JL, Van Dyke M. Read HL (eds) Annual Review of Fluid Mechanics Vol. 28, pp. 215–248 Palo Alto, CA: Annual Reviews.
(c) 2011 Elsevier Inc. All Rights Reserved.
GEOMAGNETIC POLARITY TIMESCALE C. G. Langereis and W. Krijgsman, University of Utrecht, Utrecht, The Netherlands Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1134–1141, & 2001, Elsevier Ltd.
maintained by convective fluid motion. At the surface of the Earth, the field can conveniently be described as a dipole field, which is equivalent to having a bar magnet at the center of the Earth. Such a dipole accounts for approximately 90% of the observed field; the remaining 10% derives from
Introduction Dating and time control are essential in all geoscientific disciplines, since they allows us to date, and hence correlate, rock sequences from widely different geographical localities and from different (marine and continental) realms. Moreover, accurate time control allows to understand rates of change and thus helps in determining the underlying processes and mechanisms that explain our observations. Biostratigraphy of different faunal and floral systems has been used since the 1840s as a powerful correlation tool giving the geological age of sedimentary rocks. Radiometric dating, originally applied mostly igneous rocks, has provided numerical ages; this method has become increasingly sophisticated and can now — in favorable environments — also be used on various isotopic decay systems in sediments. We are concerned with the application of magnetostratigraphy: the recording of the ancient geomagnetic field that reveals in lava piles and sedimentary sequences, intervals with different polarity. This polarity can either be normal, that is parallel to the present-day magnetic field (north-directed), or reversed (south-directed) (Figure 1). As a rule, it appears that these successive intervals of different polarity show an irregular thickness pattern, caused by the irregular duration of the successive periods of either normal or reversed polarity of the field. This produces a ‘bar code’ in the rock record that often is distinctive. Polarity intervals have a mean duration of some 300 000 y during the last 35 My, but large variations occur, from 20 000 y to several million years. If one can construct a calibrated ‘standard’ or a so-called ‘geomagnetic polarity timescale’ (GPTS), dated by radiometric methods and/or by orbital tuning, one can match the observed pattern with this standard and hence derive the age of the sediments. Magnetostratigraphy with correlation to the GPTS has become a standard tool in ocean sciences.
The Paleomagnetic Signal The Earth’s magnetic field is generated in the liquid outer core through a dynamo process that is
N
Normal
S
Reversed Figure 1 Schematic representation of the geomagnetic field of a geocentric axial dipole (‘bar magnet’). During normal polarity of the field the average magnetic north pole is at the geographic north pole, and a compass aligns along magnetic field lines. Historically, we refer to the north pole as the pole attracting the ‘north-seeking’ needle of a compass, but physically it is a south pole. During normal polarity, the inclination is positive (downward-directed) in the Northern Hemisphere and negative (upward-directed) in the Southern Hemisphere. Conversely, during reversed polarity, the compass needle points south, and the inclination is negative in the Northern and positive in the Southern Hemisphere.
(c) 2011 Elsevier Inc. All Rights Reserved.
25
26
GEOMAGNETIC POLARITY TIMESCALE
higher-order terms: the nondipole field. At any one time, the best-fitting geocentric dipole axis does not coincide exactly with the rotational axis of the Earth, but averaged over a few thousand years we may treat the dipole as both geocentric and axial. The most distinctive property of the Earth’s magnetic field is that it can reverse its polarity: the north and south poles interchange. Paleomagnetic studies of igneous rocks provided the first reliable information on reversals. In 1906, Brunhes observed lava flows magnetized in a direction approximately antiparallel to the present geomagnetic field, and suggested that this was caused by a reversal of the field itself, rather than by a self-reversal mechanism of the rock. In 1929, Matuyama demonstrated that young Quaternary lavas were magnetized in the same direction as the present field (normal polarity), whereas older lavas were magnetized in the opposite direction. His study must be regarded as the first magnetostratigraphic investgation. Initially, it was believed that the field reversed periodically, but as more (K/Ar dating plus paleomagnetic) results of lava flows became available, it became clear that geomagnetic reversals occur randomly. It is this random character that fortuitously provides the distinctive ‘fingerprints’ and gives measured polarity sequences their correlative value. A polarity reversal typically takes several thousands of years, which on geological timescales is short and can be taken as globally synchronous. The field itself is sign invariant: the same configuration of the geodynamo can produce either a normal or reversed polarity. What causes the field to reverse is still the subject of debate, but recent hypotheses suggests that lateral changes in heat flow at the core–mantle boundary play an important role. Although polarity reversals occur at irregular times, over geological time spans the reversal frequency can change considerably. For instance, the polarity reversal frequency has increased from approximately 1 My 1 some 80 My ago to 5 My 1 in more recent times. During part of the Creataceous, no reversals occurred at all from 110 to 80 Ma: the field showed a stable normal polarity during some 30 My. Such long periods of stable polarity are called Superchrons, and the one in the Cretaceous is also recognized as the Cretaceous Normal Quiet Zone in oceanfloor magnetic anomalies. The ancient geomagnetic field can be reconstructed from its recording in rocks during the geological past. Almost every type of rock contains magnetic minerals, usually iron oxide/hydroxides or iron sulfides. During the formation of rocks, these magnetic minerals (or more accurately: their magnetic domains) statistically align with the then ambient field, and will subsequently be ‘locked in,’
preserving the direction of the field as natural remanent magnetization (NRM): the paleomagnetic signal. The type of NRM depends on the mechanism of recording the geomagnetic signal, and we distinguish three basic types: TRM, CRM, and DRM. TRM, or thermoremanent magnetization, is the magnetization acquired when a rock cools through the Curie temperature of its magnetic minerals. Curie temperatures of the most common magnetic minerals are typically in the range 350–7001C. Above this temperature, the magnetic domains align instantaneously with the ambient field. Upon cooling, they are locked: the magnetic minerals carry a remanence that usually is very stable over geological time spans. Any subsequent change of the direction of the ambient field cannot change this remanence. Typically, TRM is acquired in igneous rocks. CRM, or chemical remanent magnetization, is the magnetization acquired when a magnetic mineral grows through a critical ‘blocking diameter’ or grain size. Below this critical grain size, the magnetic domains can still align with the ambient field; above it, the field will be locked and the acquired remanence may again be stable over billions of years. CRM may be acquired under widely different circumstances, e.g., during slow cooling of intrusive rocks, during metamorphosis or (hydrothermal) fluid migration, but also during late diagenetic processes such as weathering processes through formation of new magnetic minerals. A particularly important mechanism of CRM acquisition occurs in marine sediments during early diagenesis: depending on the redox conditions, iron-bearing minerals may dissolve, and iron may become mobile. If the mobilized iron subsequently encounters oxic conditions, it may precipitate again and form new magnetic minerals that then acquire a CRM. DRM, or detrital remanent magnetization, is the magnetization acquired when magnetic grains of detrital origin already carrying TRM or CRM are deposited. The grains statistically align with the ambient field as long as they are in the water column or in the soft water-saturated topmost layer of the sediment (Figure 2). Upon compaction and dewatering, the grains are mechanically locked — somewhere in a ‘lock-in depth zone’ — and will preserve the direction of the ambient field. The Magnetic Signal in Marine Sediments
In sediments, it is often assumed that the natural remanent magnetization is due to DRM. In the past, there has been some debate on the mechanisms of DRM acquisition, and the term pDRM or postdepositional DRM has often been used, because the
(c) 2011 Elsevier Inc. All Rights Reserved.
GEOMAGNETIC POLARITY TIMESCALE
27
sedimentary environment. Depending on the role of organic matter, existing magnetic minerals may dissolve and iron may be mobilized and precipitate elsewhere. These geochemical processes may results in the (partial) removal of the original paleomagnetic signal B and in the acquisition of CRM in a particular zone (much) later than the deposition of the sediment in this zone. Clearly, such processes may severely damage the Still water fidelity of the paleomagnetic record, and may offset the position of a reversal boundary by a distance in sediment that can correspond to a time of up to tens of thousands of years. However, in ‘suitable’ sediments Sediment /water interface the damage is usually restricted, and the paleomagnetic signal of such sediments may be considered as Bioturbation reliable and near-depositional. Suitable environments DRM lock-in depth are generally those with a sufficiently high sedimentation rate, a significant detrital input, and a preCompaction dominantly oxic environment. Therefore, it is often Oxic sediment O O O O Oxygen diffusion necessary to check the origin of the NRM using rock magnetic and geochemical methods. Precipitation of iron oxides As a rule, the total NRM is composed of different and CRM formation Anoxic sediment components. Ideally, the primary NRM — that oriDissolution of iron oxides ginating from near the time of deposition — has been conserved, but often this original signal is contaminated with ‘viscous’ remanence components, Fe Fe Fe Fe Iron diffusion (Sub)oxic sediment referred to as VRM. Such a VRM may result from partial realignment of ‘soft’ magnetic domains in the Figure 2 Depositional remanent magnetization (DRM) present-day field or from low-temperature oxidation acquired in sediments involves a continuum of physical and chemical processes. Detrital magnetic minerals (black) will be of magnetic minerals. It is generally easily removed statistically aligned along the ambient geomagnetic field (B) in through magnetic ‘cleaning,’ which consists of a still water and in the soft and bioturbated water-saturated routine laboratory treatment called demagnetization; sediment just below the sediment–water interface. Upon details can be found in any standard texbook on compaction and dewatering, the grains are mechanically paleomagnetism. Despite these pitfalls in sedimentlocked, preserving the direction of the field. Early diagenetic processes such as sulfate reduction may dissolve iron-bearing ary paleomagnetic records, sediments — in contrast minerals. Upon encountering a more oxic environment, iron may to igneous rocks — have in principle the distinctive precipitate as iron oxides, which will acquire a chemical remanent advantage of providing a continuous record of the magnetisation (CRM). The thus-acquired CRM in this layer may geomagnetic field, including the history of geohave a much later age than the depositional age of this layer. magnetic polarity reversals. Paleomagnetic studies of the sediments will then reveal the pattern of geolock-in zone has a certain depth. Sediment at this magnetic reversals during deposition. depth is slightly older than that the sediment–water interface. Hence, the acquisition of NRM is always slightly delayed with respect to sediment age. For The Geomagnetic Polarity Timescale practical purposes (in magnetostratigraphy) this usually has no serious consequences. Nevertheless, the (GPTS) concept of a purely detrital remanent magnetization an Surveys over the ocean basins carried out from the oversimplification of the real world. We therefore 1950s onward found linear magnetic anomalies, prefer to use the term depositional remanent magnet- parallel to mid-ocean ridges, using magnetometers ization (DRM) to refer to a continuum of physical and towed behind research vessels. During the early chemical processes that occur during and shortly after 1960s, it was suggested, and soon confirmed, that deposition. The acquisition of a DRM thus depends these anomalies resulted from the remanent magboth on detrital input of magnetic minerals and the netization of the oceanic crust. This remanence is new formation of such minerals in the sediment acquired during the process of seafloor spreading, through early diagenetic processes (Figure 2). Early when uprising magma beneath the axis of the middiagenesis is widespread and occurs in virtually every ocean ridges cools through its Curie temperature Turbulent water
(c) 2011 Elsevier Inc. All Rights Reserved.
28
GEOMAGNETIC POLARITY TIMESCALE
(Figure 3) in the ambient geomagnetic field, thus acquiring its direction and polarity. The continuous process of rising and cooling of magma at the ridge results in magnetized crust of alternating normal and reversed polarity that produces a slight increase or decrease of the measured field — the marine magnetic anomalies. It was also found that the magnetic anomaly pattern is generally symmetric on both sides of the ridge, and, most importantly, that it provides a wonderfully continuous ‘tape recording’ of the geomagnetic reversal sequence. A major step in constructing a time series of polarity reversals was taken in 1968 by Heirztler and co-workers. They used a long profile from the southern Atlantic Ocean and, extrapolating from a known age for the lower boundary of the Gauss Chron (Figure 4), they constructed a geomagnetic polarity timescale under the assumption of constant 2 Ma
1 Ma
spreading. Subsequent revisions mostly used the anomaly profile of Heirtzler, often adding more detail from other ocean basins, and appending additional calibration points. These calibration points are derived from sections on land, which, first, must contain a clear fingerprint of magnetic reversals that can be correlated to the anomaly profile, and second, contain rocks that can be reliably dated by means of radiometric methods. The GPTS is then derived by linear interpolation of the anomaly pattern between these calibration points, again under the assumption of constant spreading rate between those points. The development of the GPTS (Figure 4) shows increasing detail and gradually improved age control. Periods of a predominant (normal or reversed) polarity are called chrons, and the four youngest ones are named after individuals: Brunhes (normal), who suggested field reversal; Matuyama (reversed),
Present
1 Ma
2 Ma
Model profile
Observed profile
Mid oceanic ridge
Sediment
Uprising magma and ‘locking in’ of magnetic polarity
Crust Lithosphere
Figure 3 Formation of marine magnetic anomalies during seafloor spreading. The oceanic crust is formed at the ridge crest, and while spreading away from the ridge it is covered by an increasing thickness of oceanic sediments. The black (white) blocks of oceanic crust represent the original normal (reversed) polarity thermoremanent magnetization (TRM) acquired upon cooling at the ridge. The black and white blocks in the drill holes represent normal and reversed polarity DRM acquired during deposition of the marine sediments. The model profile (grey) represents computed magnetic anomalies produced by the block model of TRM polarity (top); the observed profile (dark) is the observed sea-level magnetic anomaly profile due to the magnetized oceanic crust.
(c) 2011 Elsevier Inc. All Rights Reserved.
29
APTS, 2000
Cande and Kent, 1992
Berggren et al., 1985
LaBreque et al., 1977
Heirtzler et al., 1968
Cox et al., 1963
GEOMAGNETIC POLARITY TIMESCALE
Brunhes
0
Matuyama
1
Gauss
2
3
Gilbert
Age (Ma)
4
5
6
7
8 Figure 4 Development of the geomagnetic polarity timescale (GPTS) through time. The initial assumption of periodic behavior (in 1963) was soon abandoned as new data became available. The first modern GPTS based on marine magnetic anomaly patterns was established in 1968 by Heirtzler and co-workers. Subsequent revisions show improved age control and increased resolution. A major breakthrough came with the astronomical polarity timescale (APTS), in which each individual reversal is accurately dated.
who proved this; Gauss (normal), who mathematically described the field; and Gilbert (reversed), who discovered that the Earth itself is a huge magnet. Chrons may contain short intervals of opposite
polarity called subchrons, which are named after the locality where they were discovered; for example, the normal Olduvai subchron within the Matuyama reversed chron is named after Olduvai Gorge in
(c) 2011 Elsevier Inc. All Rights Reserved.
30
GEOMAGNETIC POLARITY TIMESCALE
Tanzania, and the Kaena reversed subchron within the Gauss normal chron after Kaena Point on Hawaii. Older chrons were not named but numbered, according to the anomaly numbers earlier given by Heirtzler, which has led to a confusing nomenclature of chrons and subchrons. A major step forward was taken by Cande and Kent in 1992, who thoroughly revised the magnetic anomaly template over the last 110 My. They constructed a synthetic flow line in the South Atlantic, using a set of chosen anomalies that were taken as tie points. The intervals between these tie points were designated as category I intervals. On these intervals they projected (stacks of) the bestquality profiles surveyed in this ocean basin, providing category II intervals. Since the spreading of the Atlantic is slow, they subsequently filled in the category II intervals with high-resolution profiles from fast-spreading ridges (their category III). This enabled them to include much more detail on short polarity intervals (or subchrons); see, for instance, the increase of detail around 7 Ma in Figure 4. Very short and low-intensity anomalies still have an uncertain origin. They may represent very short subchrons of the field, as has been proven for some of them (e.g., the Cobb Mt. subchron at 1.21 Ma, or the Re´union subchron at 2.13–2.15 Ma), or may just represent intensity fluctuations of the geomagnetic field causing the oceanic crust to be less (or more) strongly magnetized. Because of their uncertain or unverified nature, these were called cryptochrons. In addition, Cande and Kent developed a consistent (sub)chron nomenclature that is now used as the standard. In total, they used nine calibration points, but they made a break with tradition by using, for the first time, an astronomically dated age tie point for the youngest one: the Gauss/Matuyama boundary. The correlation of the GPTS to global biostratigraphic zonations is covered extensively by Berggren et al. (1995). The template of magnetic anomaly patterns from the ocean floor has remained central for constructing the GPTS from the late Cretaceous onward (110–0 Ma). Only recently, the younger part of the GPTS has been based on direct dating of each individual reversal through the use of orbitally tuned timescales. In their most recent version of the GPTS, Cande and Kent included the astronomical ages for all reversal boundaries for the past 5.3 My. The CK95 or Cande and Kent (1995) geomagnetic polarity timescale is at present the most widely used standard. Polarity timescales for the Mesozoic rely on some of the oldest magnetic anomaly profiles, down to the late Jurassic, and on dated magnetostratigraphies of sections on land.
The Astronomical Polarity Timescale (APTS) The latest development in constructing a GPTS comes from orbital tuning of the sediment record; for details see the article Orbitally Tuned Timescales. It differs essentially from the conventional GPTS in the sense that each reversal boundary — or any other geological boundary for that matter, e.g. biostratigraphic datum levels or stage and epoch boundaries — is dated individually. This has provided a breakthrough in dating of the geological record and has the inherent promise of increasing understanding of the climate system, since cyclostratigraphy and subsequent orbital tuning rely on decphering and understanding environmental changes driven by climate change, which in turn is orbitally forced. The fact that the age of each reversal is directly determined, rather than interpolated between calibration points, has important consequences for (changes in) spreading rates of plat pairs. Rather than having to assume constant spreading rates between calibration points, one can now accurately determine these rates, and small changes therein. Indeed, Wilson found that the use of astronomical ages resulted in very small and physically realistic spreading rate variations. As a result, the discrepancy in between plate motion rates from the global plate tectonic model (NUVEL-1) and those derived from geodesy has become much smaller. Meanwhile NUVEL-1 has been updated (to NUVEL-1A) to incorporate the new astronomical ages. Another application is the dating of Pleistocene, Pliocene, and Miocene, and older stage boundaries, many of which have been defined in the Mediterranean. The availability of a good astrochronology has effectively become a condition for the definition of a Global Boundary Stratotype Section and Point (GSSP). An example is shown in Figure 5, showing the Tortonian–Messinian GSSP that has recently been defined in the Atlantic margin basin of western Morocco. Perhaps one of the most promising areas of the application of astrochronology is in the bed-to-bed correlation of the two different realms of oceans and continent. Climate forcing may be expected to have a different expression in the different realms because of the different nature of their sedimentary environments. A recently established and refined orbital timescale for the loess sequences of northern China relies upon the correlation of detailed monsoon records to the astronomical solutions and the oceanic oxygen isotope records. An important finding was that the straightforward use of magnetostratigraphy and correlation to the GPTS cannot provide a sufficiently accurate age model for comparison with the
(c) 2011 Elsevier Inc. All Rights Reserved.
GEOMAGNETIC POLARITY TIMESCALE
31
Figure 5 Magnetostratigraphy, cyclostratigraphy, and astrochronology of the marine Oued Akrech section from the Atlantic margin of Morocco, which defines the Tortonian–Messinian Global Stratotype Section and Point (GSSP). The sedimentary cycles represent a cyclically changing environment that correlates with variations in insolation. Insolation is strongly related to climatic precession (upper left panel), which induces cyclic changes in seasonal contrast, reflected one-on-one in the sedimentary cycles. (After Hilgen FJ et al. (2000) Episodes 23(3): 172–178.)
ocean record, since the analysis of the astrochronological framework demonstrates considerable downward displacement of reversal boundaries because of delayed lock-in of the NRM. With the new chronology and its direct correlation to the oceanic record, it is now possible to analyse terrestrial paleomonsoon behavior for the past 2.6 My and compare it to climate proxies from the marine realm. This may give important information, for instance, on leads and lags of various systems in response to climate change, on phase relations with insolation, or on the relation between global ice volume and monsoonal climate.
See also Aeolian Inputs. Magnetics. Monsoons, History of. Paleoceanography, Climate Models in. Paleoceanography: Orbitally Tuned Timescales.
Further Reading Berggren WA, Kent DV, Aubry MP, and Hardenbol J (1995) Geochronology, Time Scales and Global Stratigraphic Correlations: A Unified Temporal Framework for an Historical Geology. Society of Economic Paleontologists and Mineralogist, Special Volume 54.
(c) 2011 Elsevier Inc. All Rights Reserved.
32
GEOMAGNETIC POLARITY TIMESCALE
Butler RF (1992) Paleomagnetism, Magnetic Domains to Geologic Terranes. Boston: Blackwell Scientific. Cande SC and Kent DV (1992) A new geomagnetic polarity Time-Scale for the Late Cretaceous and Cenozoic. Journal of Geophysical Research 97: 13917--13951. Cande SC and Kent DV (1995) Revised calibration of the Geomagnetic Polarity Time Scale for the Late Cretaceous and Cenozoic. Journal of Geophysical Research 100: 6093--6095. Gradstein FM, Agterberg FP, Ogg JG, et al. (1995) A Triassic, Jurassic and Cretaceous time scale. In: Berggren WA, Kent DV, Aubry M and Hardenbol J (eds) Geochronology, Time Scales and Stratigraphic Correlation. Society of Economic Paleontologists and Mineralogist, Special Volume 54: 95–128. Hilgen FJ, Krijgsman W, Langereis CG, and Lourens LJ (1997) Breakthrough made in dating of the geological record. EOS, Transactions of the AGU 78: 285--288. Heirtzler JR, Dickson GO, Herron EM, Pitman WC III, and Le Pichon X (1968) Marine magnetic anomalies,
geomagnetic field reversals, and motions of the ocean floor and continents. Journal of Geophysical Research 73: 2119--2136. Heslop D, Langereis CG, and Dekkers MJ (2000) A new astronomical time scale for the loess deposits of Northern China. Earth and Planetory Science Letters 184: 125--139. Maher BA and Thompson R (1999) Quaternary Climates, Environments and Magnetism. Cambridge: Cambridge University Press. Opdyke ND and Channell JET (1996) Magnetic Stratigraphy. San Diego: Academic Press. Shackleton NJ, Crowhurst SJ, Weedon GP, and Laskar J (1999) Astronomical calibration of Oligocene–Miocene time. Philosophical Transactions of the Royal Society of London, A 357: 1907--1929. Wilson DS (1993) Confirmation of the astronomical calibration of the magnetic polarity time scale from seafloor spreading rates. Nature 364: 788--790.
(c) 2011 Elsevier Inc. All Rights Reserved.
GEOMORPHOLOGY C. Woodroffe, University of Wollongong, Wollongong, NSW, Australia Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1142–1148, & 2001, Elsevier Ltd.
Introduction Geomorphology is the study of the form of the earth. Coastal geomorphologists study the way that the coastal zone, one of the most dynamic and changeable parts of the earth, evolves, including its profile, plan-form, and the architecture of foreshore, backshore, and nearshore rock and sediment bodies. To understand these it is necessary to examine wave processes and current action, but it may also involve drainage basins that feed to the coast, and the shallow continental shelves which modify oceanographic processes before they impinge upon the shore. Morphodynamics, study of the mutual co-adjustment of form and process, leads to development of conceptual, physical, mathematical, and simulation models, which may help explain the changes that are experienced on the coast. In order to understand coastal variability from place to place there are a number of boundary conditions which need to be considered, including geophysical and geological factors, oceanographic factors, and climatic constraints. At the broadest level plate-tectonic setting is important. Coasts on a plate margin where oceanic plate is subducted under continental crust, such as along the western coast of the Americas, are known as collision coasts. These are typically rocky coasts, parallel to the structural grain, characterized by seismic and volcanic activity and are likely to be uplifting. They contrast with trailing-edge coasts where the continental margin sits mid-plate, which are the locus of large sedimentary basins. Smaller basins are typical of marginal sea coasts, behind a tectonically active island arc. The nature of the material forming the coast is partly a reflection of these broad plate-tectonic factors. Whether the shoreline is rock or unconsolidated sediment is clearly important. Resistant igneous or metamorphic rocks are more likely to give rise to rocky coasts than are those areas composed of broad sedimentary sequences of clays or mudstones. Within sedimentary coasts sandstones may be more resistant than mudstones. Rocky coasts tend to be relatively
resistant to change, whereas coasts composed of sandy sediments are relatively easily reshaped, and muddy coasts, in low-energy environments, accrete slowly. The relative position of the sea with respect to the land has changed, and represents an important boundary condition. The sea may have flooded areas which were previously land (submergence), or formerly submarine areas may now be dry (emergence). The form of the coast may be inherited from previously subaerial landforms. Particularly distinctive are landscapes that have been shaped by glacial processes, thus fiords are glacially eroded valleys, and fields of drumlins deposited by glacial processes form a prominent feature on paraglacial coasts. The oceanographic factors that shape coasts include waves, tides, and currents. Waves occur as a result of wind transferring energy to the ocean surface. Waves vary in size depending on the strength of the wind, the duration for which it has blown and the fetch over which it acts. Wave trains move out of the area of formation and are then known as swell. The swell and wave energy received at a coast may be a complex assemblage generated by specific storms from several areas of origin. Tides represent a largewavelength wave formed as a result of gravitational attractions of the sun and moon. Tides occur as a diurnal or semidiurnal fluctuation of the sea surface that may translate into significant tidal currents particularly in narrow straits and estuaries. In addition storm-generated surges and tsunamis may cause elevated water levels with significant geomorphological consequences. The coast is shaped by these oceanographic processes working on the rocks and unconsolidated sediments of the shoreline. Climate is significant in terms of several factors. First, wind conditions lead to generation of waves and swell, and may blow sand into dunes along the backshore. Climate also influences the rate at which weathering and catchment processes operate. In addition, regional-scale climate factors such as monsoonal wind systems and the El Nin˜o Southern Oscillation phenomenon demonstrate oscillatory behavior and may reshape the coast seasonally or interannually. Gradual climate change may mean that shoreline fluctuations do not revolve around stationary boundary conditions but may exhibit gradual change themselves. In this respect the impact of perceived global climate change during recent decades as a result of human-modified environmental factors is likely to be
(c) 2011 Elsevier Inc. All Rights Reserved.
33
34
GEOMORPHOLOGY
Sea level
Sea level
(B)
(A)
Figure 1 Schematic representation of the original woodcut illustrations by Charles Darwin showing (A) the manner in which he deduced that fringing reefs around a volcanic island would develop into barrier reefs (dotted), and (B) barrier reefs would develop into an atoll, as a result of subsidence and vertical reef growth.
felt in the coastal zone. Of particular concern is anticipated sea-level rise which may have a range of geomorphological effects, as well as other significant socioeconomic impacts.
History Geomorphology has its origins in the nineteenth century with the results of exploration, and the realization that the surface of the earth had been shaped over a long time through the operation of processes that are largely in operation today (uniformitarianism). The observations by Charles Darwin during the voyage of the Beagle extended this view, particularly his remarkable deduction that fringing reefs might become barrier reefs which in turn might form atolls as a result of gradual subsidence of volcanic islands combined with vertical reef growth (Figure 1). In the first part of the twentieth century geomorphology was dominated by the ‘geographical cycle’ of erosion of William Morris Davis who anticipated landscape
denudation through a series of stages culminating in peneplanation, and subsequent rejuvenation by uplift. This highly conceptual model, across landscapes in geological time, was also applied to the coast by Davis who envisaged progressive erosion of the coast reducing shoreline irregularities with time (Figure 2). Such landscape-scale studies were extended by Douglas Johnson who emphasized the role of submergence or emergence as a result of sea-level change (Figure 3). The Davisian view was reassessed in the second half of the twentieth century with a greater emphasis on process geomorphology whereby studies focused on attempting to measure rates of process operation and morphological responses to those processes, reflecting ideas of earlier researchers such as G. K. Gilbert. The concept of the landscape as a system was examined, in which coastal landforms adjusted to equilibrium, perhaps a dynamic equilibrium, in relation to processes at work on them. Studies of
i ii iii (A) Initial
Figure 2 Schematic representation of the planform stages that W. M. Davis conceptualized through which a shoreline would progress from an initial rugged form (1) through maturity (2) to a regularized shoreline (3) that is cliffed (hatched) and infilled with sand (stippled). He envisaged this in parallel to the geographical cycle of erosion by which mountains (like the initial shoreline form, 1) would be reduced to a peneplain (like the solid ultimate shoreline, 3).
(B) Modern
Figure 3 Schematic representation of (A) the initial shoreline form envisaged by Douglas Johnson for an area in Marthas Vineyard, New England, and (B) the modern regularized, form of the shore.
(c) 2011 Elsevier Inc. All Rights Reserved.
GEOMORPHOLOGY
sediment movement and the adjustment of beach shape under different wave conditions were typical.
Scales of Study Coastal geomorphology now studies landforms and the processes that operate on them at a range of spatial and temporal scales (Figure 4). At the smallest scale geomorphology is concerned with an ‘instantaneous’ timescale where the principles of fluid dynamics apply. It should be possible to determine details of sediment entrainment, complexities of turbulent flow and processes leading to deposition of individual bedform laminae. The laws of physics apply at these scales, though they may operate stochastically. In theory, behavior of an entire embayment could be understood; in practice studies simply cannot be undertaken at that level of detail and broad extrapolations based on important empirical relationships are made. The next level of study is the ‘event’ timescale, which may cover a single event such as an individual storm or an aggregation of several lesser events over a year or more. The mechanistic relationships from instantaneous time are scaled up in a deterministic or empirical way to understand the operation of coasts at larger spatial and temporal scales. Thus, stripping of a beach during a single storm, and the more gradual reconstruction to its original state under ensuing calmer periods can be observed. Time taken for reaction to the event, and relaxation back to a more ambient state may be known from surveys of beaches, enabling definition of a ‘sweep zone’ within which the beach is regularly active.
35
At a larger scale of operation, the coastal geomorphologist is interested in the way in which coasts change over timescales that are relevant to societies and at which coastal-zone managers need to plan. This is the ‘engineering’ timescale, involving several decades. It is perhaps the most difficult timescale on which to understand coastal geomorphology and to anticipate behavior of the shoreline. The largest timescales are geological timescales. Studies over geological timescales are primarily discursive, conceptual models which recognize that boundary conditions, including climate and the rate of operation of oceanographic processes, change. It is clear that sea level itself has changed dramatically over millennia as a result of expansion and contraction of ice sheets during the Quaternary ice ages, and so the position of the coastline has changed substantially between glaciations and interglaciations.
Models of Coastal Evolution
Space
There has been considerable improvement in understanding the long-term development of coasts as a result of significant advances in paleoenvironmental reconstruction and geochronological techniques (especially radiometric dating). Incomplete records of past coastal conditions may be preserved, either as erosional morphology (notches, marine terraces, etc.) or within sedimentary sequences. Although this record is selective, reconstructions of Quaternary paleoenvironments, together with interpretation of geological sequences in older rocks, have enabled the formulation of geomorphological models based on sedimentary evidence. In the case of deltas, where there may be important hydrocarbon reserves, the complementary development of geological models GEOLOGICAL Global and study of modern deltas has led to better underReef standing of process and response at longer timescales structure ENGINEERING than those for which observations exist. Cliff Unconsolidated sediments are the key to coastal Regional retreat morphodynamics because coasts change through Delta EVENT dynamics erosion, transport, and deposition of sediment. Study of coastal systems has led to many insights, parReef-flat morphology ticularly in terms of the various pathways through Mudflat Local sedimentation which sediment may move within a coastal comBeach states partment or circulation cell. However, it is clear that INSTANTANEOUS a completely reductionist approach to coastal geomm Seconds Years Millennia morphology will not lead to understanding of all Time components and the way in which they interact. Empirical relationships and the presumed determinFigure 4 Representation of space and timescales appropriate istic nature of sediment response to forcing factors for the study of coastal geomorphology and schematic representation of some of the examples discussed in the context of remain incomplete and are all too often formulated instantaneous, event, engineering, and geological scales of on the basis of presumed uniform sediment sizes enquiry. or absence of biotic influence and are ultimately
(c) 2011 Elsevier Inc. All Rights Reserved.
36
GEOMORPHOLOGY
Bore Reflected standing wave
Breaking wave
(B)
geomorphology of each of these behaves differently. Processes operate at different rates, transitions between different states occur over different timescales and the significance of antecedent conditions varies. Rocky Coasts
Bore (A)
Breaking wave
(C)
Figure 5 The morphology of cliffs and shore platforms. Plunging cliffs (A) reflect wave energy and are generally not eroded. Shore platforms dissipate wave energy and may be either (B) ramped, or (C) subhorizontal. Change in these systems tends to occur at geological (hard rock) or engineering (soft rock) timescales.
unrealistic. More recently coastal systems have been investigated as nonlinear dynamical systems, because relationships are not linear and behavior is potentially chaotic. Nonlinear dynamical systems behave in a way that is in part dependent upon antecedent conditions. This is well illustrated by studies of beach state; beach and nearshore sediments are particularly easily reshaped by wave processes, thus a series of characteristics typical of distinct beach states can be recognized (see Figure 6). However, the beach is only partly a response to incident wave conditions, its shape being also dependent on the previous shape of the beach which was in the process of adjusting to wave conditions incident at that time. Coastal systems are inherently unpredictable. However, they may operate within a broad range of conditions with certain states being recurrent. Chaotic systems may tend towards self-organization, for instance patterns of beach cusps characteristic along low-energy beaches which reflect much of the wave energy may adopt a self-organized cuspate morphology in which swash processes, sediment sorting, and form are balanced. Models may be developed based on patterns of change inferred over geological timescales but consistent with mechanistic processes known to operate over lesser event timescales. Simulation modelling is not intended to reconstruct coastal evolution exactly, but becomes a tool for experimentation and extrapolation within which broad scenarios of change can be modeled and sensitivity to parameterization of variables examined. Coasts may be divided into rocky coasts, sandy coasts, and muddy coasts, and coral reefs and deltas and estuaries can be differentiated. The
Rocky coasts are characterized by sea cliffs, especially on tectonically active, plate-margin coasts where there are resistant rocks. Cliffed coastlines in resistant rock appear to adopt one of two forms (states): plunging cliffs where a vertical cliff extends below sea level, and shore platforms where a broad bench occurs at sea level in front of a cliff (Figure 5). Erosion of cliffs occurs where the erosional force of waves exceeds the resistance of the rocks, and where sufficient time has elapsed. Plunging cliffs occur where the rock is too resistant to be eroded. The vertical face results in a standing wave which reflects wave energy, so that there is little force to erode the cliff at water level. Waves exert a greater force if they break, or if they are already broken. This can only occur if the water depth is shallow offshore from the cliff face in which case the increased energy from the breaking waves is also able to entrain sediment from the floor (or rock fragments quarried from the foot of the cliff). This process of erosion accelerates through a positive feedback cutting a shore platform. Shore platforms thus develop in those situations where the erosive force of waves exceeds resistance of the rock. A platform widens as a result of erosion at the foot of the cliff behind the platform. The cliff oversteepens, leading to toppling and fall of detritus onto the rear of the platform, which slows further erosion of the cliff face until that talus has been removed. A series of such negative feedbacks slow the rate at which platforms widen over time, and there is often considerable uniformity of platforms up to a maximum width in any particular lithological setting. Shore platforms adopt either a gradually sloping ramped form, or a subhorizontal form often with a seaward rampart (Figure 5). There has been much discussion as to the relative roles of wave and subaerial processes in the formation of these platforms. Platforms in relatively sheltered locations appear to owe their origin to processes of water-layer leveling (physiochemical processes in pools which persist on platforms at low tide) or wetting and drying and its weakening of the rock. In other cases wave quarrying and abrasion are involved. In many cases both processes may be important. Other platforms may be polygenetic with inheritance from former stands of sea level (reflecting antecedent conditions).
(c) 2011 Elsevier Inc. All Rights Reserved.
GEOMORPHOLOGY
(A) Dissipative
(B) Intermediate
37
(C) Reflective
Plunging waves
Spilling waves
Accreted beach
Surging waves
Nearshore bar (D)
Storm
Beach volume
C
A Time
Figure 6 Beach state may be (A) dissipative, where high wave-energy impinges on low-gradient beaches; (B) intermediate (in which several states are possible); or (C) reflective, where low wave-energy surges onto steep or coarse beaches. Change between beach states occurs at event timescales. (D) A beach may react rapidly to a perturbation such as a storm changing from reflective state (C) to more dissipative state (A) but then change much more slowly through intermediate states, readjusting towards state (C) again unless further disrupted.
Coral Reefs
Sandy Coasts
A group of coastlines that are of particular interest to the geomorphologist are those formed by coral reefs. Corals are colonial animals that secrete a limestone exoskeleton that may form the matrix of a reef. Coral reefs flourish in tropical seas in high-energy settings where a significant swell reaches the shoreline. In these circumstances the reef attenuates much of the wave energy such that a relatively quiet-water environment occurs sheltered behind the reef. Not only is this reef ecosystem of enormous geomorphological significance in terms of the solid reef structure that it forms, but in addition, the carbonate reef material breaks down into calcareous gravels, sands and muds which form the sediments that further modify these coastlines. On the one hand reefs can be divided into fringing reefs, barrier reefs, and atolls, distinct morphological states that form part of an evolutionary sequence as a result of gradual subsidence of the volcanic basement upon which the reef established (see Figure 1). This powerful deduction by Darwin relating to reef structure and operating over geological timescales, has been generally supported by drilling and geochronological studies on mid-plate islands in reef-forming seas. On the other hand, the surface morphology of reefs is extremely dynamic with rapid production of skeletal sediments and their redistribution over event timescales, with the landforms of reefs responding to minor sea-level oscillations, storms, El Nin˜o, coral bleaching, and other perturbations.
Beaches form where sandy (or gravel) material is available forming a sediment wedge at the shoreline. The beach is shaped by incident wave energy, and can undergo modification particularly by formation and migration of nearshore bars which in turn modify the wave-energy spectrum. Various beach states can be recognized across a continuum from beaches which predominantly reflect wave energy, and those which dissipate wave energy across the nearshore zone. Reflective beaches are steep, waves surge up the beach and much of the energy is reflected, and the beach face may develop cusps. Dissipative beaches are much flatter, and waves spill before reaching the shoreline (Figure 6). During a storm, sand is generally eroded from the beach face and deposited in the nearshore, often forming shore-parallel or transverse bars. Waves consequently break on the bars and energy is lost, the form modifying the process in a mutual way. It may be possible to recognize a series of beach states intermediate between reflective and dissipative and any one beach may adopt one or several beach states over time (Figure 6). Beach state is clearly modified by incident wave conditions, but the rate of adjustment between states takes time, and a beach is also partly a function of antecedent beach states. Although erosion and redeposition of sand is the way in which the nearshore adjusts, there may also be long-term storage of sediment. In particular broad, flat beaches may develop dunes behind them.
(c) 2011 Elsevier Inc. All Rights Reserved.
38
GEOMORPHOLOGY
In other cases a sequence of beach ridges may develop. Geomorphological changes in state over geological timescales represented by beach-ridge plains may relate to variations in supply of sediment to the system (perhaps by rivers) as well as variations in the processes operating. Deltas and Estuaries
Where rivers bring sediment to the coast, deltas and estuaries can develop. In this case the sediment budget of the shoreline compartment or the receiving basin is augmented and there is generally a positive sediment budget. Deltas are characterized by broad wedges of sediment deposition. Delta morphology tends to reflect the processes that are dominant (Figure 7). Where wave energy is low and tides are minor, it is primarily river flows which account for sedimentation patterns. River flow may be likened to a jet, influenced by inertial forces, friction with the basin floor or buoyancy where there are density differences between outflow and the water of the receiving basin. Wave action tends to smooth the shoreline and a wave-dominated deltaic shoreline will be characterized by shore-parallel bars or ridges. Where riversupplied sediment is relatively low in volume, wave action may form a sandy barrier along the coast and rivers may supply sediment only to lagoons formed behind these barriers. Such is the case for much of the barrier-island shoreline of the eastern shores of the Americas; where barrier islands may be continually reworked landwards during relative sea-level rise. On the other hand if sea level is stable, a stable sand barrier is likely to form closing each
River
Wave
Tide
Figure 7 Deltas adopt a variety of forms, but morphology appears to reflect the relative balance of river, wave, and tide action. The broad morphology of the delta tends to be digitate where river-dominated, shore-parallel where wave-dominated, and tapering where tide-dominated.
embayment or creating a barrier estuary as in the case of coastal lagoons in Asia and Australasia. Estuaries are embayments which are likely to infill incrementally either through the deposition of riverborne sediments or through the influx of sediment from seaward by wave or tidal processes (Figure 8). The sediment from seaward may either be derived from the shelf, or from shoreline erosion. Tidal processes differ from river processes; they are bi-directional, flowing in during flood tide and out during ebb, and the flow is forced by the rising level of the sea. Small embayments are flooded and drained by a tidal prism (the volume of water between low and high tide). Longer estuaries, on the other hand, may have a series of tidal waves which progress up them. Where the tidal range is large this may flow as a tidal bore. Tide-dominated estuarine channels adopt a distinctive tapered form, with width (and depth) decreasing from the mouth upstream (Figure 8). Wave and tidal processes tend to shape deltas and estuaries to varying degrees depending on their relative operation. Many of the deltaic-estuarine processes operate over cycles of change. The hydraulic efficiency of distributaries decreases as they lengthen, until an alternative, shorter course with steeper hydraulic gradient is adopted. The abandoned distributaries of deltas often become tidally dominated with sinuous, tapering tidal creeks dominating what may be a gradually subsiding abandoned delta plain. Wave processes, in wavedominated settings, smooth and rework the abandoned delta shore, often forming barrier islands. Muddy Coasts
Muddy coasts are associated with the lowest energy environments. Mud banks may occur in high-energy, wave-exposed settings where large volumes of mud are supplied to the mouth of large rivers. Thus longshore drift north west of the Amazon, and around Bohai Bay downdrift from the Yellow River, enables mud-shoal deposition in open-water settings. Elsewhere mud flats are typical of sheltered settings, within delta interdistributary bays, around coastal lagoons, etc. These muddy environments are likely to be colonized by halophytic vegetation. Mangrove forests occur in tropical settings, whereas salt marsh occurs in higher latitudes, often extending into tropical areas also. These coastal wetlands further promote retention of fine-grained sediment. The muddy environments are areas of complex hydrodynamics and sedimentation. Sedimentation is likely to occur with a negative feedback such that as the tidal wetlands
(c) 2011 Elsevier Inc. All Rights Reserved.
GEOMORPHOLOGY
(A) Wave-dominated Tidal delta
39
(B) Tide-dominated
Wetland
Wetland
Fluvial delta
Halopytes River Send shoal River
Sand barrier
Saline mudflats
Former shoreline Mud basin
Figure 8 Estuaries are broad embayments which may adopt a wide range of morphologies. They tend to show sand-barrier accumulation where wave-dominated (A), but be prominently tapering where tide-dominated (B). Estuaries are generally sediment sinks with the rates and patterns of infill reflecting the relative dominance of river, wave, and tide processes and sediment sources.
accrete sediment and as the substrate is elevated, they are flooded less frequently and therefore sedimentation decelerates. Boundary conditions, particularly sea level, are likely to vary at rates similar to the rate of sedimentation and prograded coastal plains contain complex sedimentary records of changes in ecological and geomorphological state.
See also
Conclusion
Boyd R, Dalrymple R, and Zaitlin BA (1992) Classification of clastic coastal depositional environments. Sedimentary Geology 80: 139--150. Carter RWG and Woodroffe CD (eds.) (1994) Coastal Evolution: Late Quaternary Shoreline Morphodynamics. Cambridge: Cambridge University Press. Cowell PJ and Thom BG (1994) Morphodynamics of coastal evolution. In: Carter RWG and Woodroffe CD (eds.) Coastal Evolution: Late Quaternary Shoreline Morphodynamics, pp. 33--86. Cambridge: Cambridge University Press. Darwin C (1842) The Structure and Distribution of Coral Reefs. London: Smith Elder. Davis RA (1985) Coastal Sedimentary Environments. New York: Springer-Verlag. Johnson DW (1919) Shore Processes and Shoreline Development. New York: Prentice Hall. Trenhaile AS (1997) Coastal Dynamics and Landforms. Oxford: Clarendon Press. Wright LD and Thom BG (1977) Coastal depositional landforms: a morphodynamic approach. Progress in Physical Geography 1: 412--459.
Coastal geomorphology is the study of the evolving form of the shoreline in response to mutually adjusting processes acting upon it. It spans instantaneous and event timescales, over which beaches respond to wave energy, through engineering and geological timescales, over which deltas build seaward or switch distributaries, sea level fluctuates, and cliff morphology evolves. The morphology (state) of the coast changes in response to perturbations, particularly extreme events such as storms, but also thresholds within the system (as when a cliff oversteepens and falls, or a distributary lengthens and then switches). Human action may also represent a perturbation to the system. In each coastal setting, the influence of human modifications is being felt. Thus there are fewer coasts which are not in some way influenced by society. As anthropogenic modification of climate and sea level occurs at a global scale, the human factor increasingly needs to be given prominence in coastal geomorphology.
Beaches, Physical Processes Affecting. Coral Reefs. Rocky Shores. Salt Marshes and Mud Flats.
Further Reading
(c) 2011 Elsevier Inc. All Rights Reserved.
GEOPHYSICAL HEAT FLOW C. A. Stein, University of Illinois at Chicago, Chicago, IL, USA R. P. Von Herzen, Woods Hole Oceanographic Institution, Woods Hole, MA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1149–1157, & 2001, Elsevier Ltd.
Introduction Heat flow is a directly measurable parameter at Earth’s surface that provides important constraints on the thermal state of its interior. When combined with other data, it has proved useful for models of Earth’s origin, structure, composition, and convective motion of at least the upper mantle. Because oceanic crust is mostly thin (about 5–8 km) and deficient in heat-producing radioactive elements compared to continental crust, marine heat flow measurements are particularly illuminating for the cooling and dynamics of the oceanic lithosphere. Indeed, they have provided an important constraint in the development and refinement of the plate tectonics paradigm over the past three decades. Initially influenced by concepts of Sir Edward Bullard, the first successful marine heat flow measurements obtained shortly after the end of World War II were among the first applications of remote electronics in the deep sea. Other than a general understanding of seafloor topography, little information on the nature of the Earth beneath the oceans was then available, so theories to explain the measurements were largely unconstrained. Probably the most important initial finding was that the mean oceanic heat flux was not much different from the mean of the available continental values, contrary to conventional expectations of that era. However, the entire field of marine geophysics (especially seismic reflection and refraction, and geomagnetism) was also being transformed by instrumentation development, and the availability of ships as seagoing platforms gave rise to a ‘golden age’ of ocean exploration between about 1950 and 1970. The large-scale (>1000 km) variations of marine heat flux were shown to be consistent with the plate tectonic paradigm that explained the mean depth and age of much of the ocean floor. Heat flow is now useful for studies of seafloor tectonics and other marine geophysical problems.
40
Marine heat flow measurements now number more than 15 000, and with the exception of high latitude regions, are distributed over most of Earth’s main ocean basins (Figure 1). Values range over about two orders of magnitude (about 10–1000 mW m2), with a general systematic spatial distribution at large (41000 km) scales but also a large variability at small (o50 km) scales caused by pervasive hydrothermal circulation in mostly young (o50– 70 million years (My)) ocean crust. The spatial distribution and magnitudes of marine heat flow values may be used to deduce the geometry and intensity of such circulation.
Measurements and Techniques Instrumentation and Development
Techniques for marine measurements were developed with the realization that the temperature of the deep (42 km) ocean is mostly horizontally stratified, and relatively constant over long time scales (4100 years). In this situation, the heat flux from the Earth’s interior is reflected in a relatively uniform temperature gradient with depth beneath a flat seafloor. Heat flow is the product of the magnitude of this temperature gradient and the thermal conductivity of the material over which it is measured. Gradients are commonly measured using a vertically oriented probe with temperature sensors surmounted by a weight that is lowered from a ship to penetrate the relatively soft sediments that cover most of the seafloor. Thermal conductivity is either measured with in situ sensors during the seafloor penetration by transient heating experiments or aboard ship on sediment cores. For instrumental simplicity, the initial heat flow equipment used only two thermal sensors at either end of a 3–4 m long probe with in situ analog recording techniques on paper or film. Subsequently the number of sensors has been increased, due to a desire to measure non-linear temperature gradients and thermal conductivity that may vary with depth. The wide range of measured thermal gradients and the developments in solid-state electronics have made digital recording the present standard. ‘Pogo’ measurements, i.e., multiple penetrations on the same station, combined with nearly real-time acoustic telemetry of raw data have proven useful for investigation of small-scale (o10–20 km) marine
(c) 2011 Elsevier Inc. All Rights Reserved.
(c) 2011 Elsevier Inc. All Rights Reserved.
0° 60°E
120°E
180°
120°W
Figure 1 Locations of marine heat flow measurements (dots). Thin lines show the location of the major plate boundaries.
60°S
0°
60°N
60°W
0°
GEOPHYSICAL HEAT FLOW 41
42
GEOPHYSICAL HEAT FLOW
heat flow variability, the origins of which are still not fully understood. It was necessary for the ‘Bullard’ probe to remain undisturbed in the seafloor for most of one hour to approach thermal equilibrium with the sediments because of its relatively large diameter (about 2– 3 cm). This probe has been largely supplanted by either much smaller (about 3 mm diameter) individual probes mounted in outrigged fashion on a larger (‘Ewing’) probe to attain deep (5–6 m) penetration, or a ‘violin-bow (Lister)’ probe consisting of a small (about 1 cm) sensor string mounted parallel to, but separate from, the main strength member (Figure 2). Both are capable of in situ thermal conductivity measurements, utilizing either a constant heat source or a calibrated pulse (approximating a delta function) after the gradient measurement. The measurement time in the seafloor is 15–20 min. Depending mostly on the
desired spacing during pogo operations, the mean time between measurement is 1–2 h. Even with acoustic data telemetering, battery life can be 2–3 days, thereby allowing many measurements during a single lowering. Real-time ship location accuracy now approaches a few meters with differential Global Positioning System navigation, although the uncertainty of the probe location during pogo operations in normal ocean depths (4–5 km) is typically 200 m or more. Hence 1–2 km is a useful minimum spacing between pogo penetrations unless seafloor acoustic transponder navigation is employed for higher accuracy (about 10 m) navigation. Instrumentation is also available to determine temperatures to depths up to about 600 m below the seafloor during deep-sea drilling to establish the uniformity of heat flow to greater depths, and may be the only method for reliable measurements in shallow seafloor (o1 km) regions
Cross-section AT A_A′ Weights
Cross-section of outrigger
Instrument housing Strength member
Lister
Bullard
Ewing
Pressure vessel
Instrument housing A
A′ Weight stand
Weight stand Coupling Strength member
Pressure vessel
(A)
Tensioning device (B)
Sensor string Strength member
6 _10 m
4_ 6 m
2m
Outrigger fin with thermistor probe
Nose cone (C)
Figure 2 Diagrams of the three most commonly used marine heat flow probes. (Reproduced with permission from Louden KE and Wright JA (1989) Marine heat flow data: a new compilation of observations and brief review of its analysis. In: Wright JA and Londen KE (eds) Handbook of Seafloor Heat Flow, pp. 2–67, Boca Raton: CRC Press.)
(c) 2011 Elsevier Inc. All Rights Reserved.
GEOPHYSICAL HEAT FLOW
where ocean temperatures are more variable. Geothermal probes have also been developed for use with manned submersibles, allowing visual control of measurements in regions with large lateral gradients of heat flux or over specific geological features. A development is underway to install geothermal instrumentation on an autonomous underwater vehicle (ABE) for measurements in regions that are remote or difficult for normal ship operations. Environmental Corrections
Although most measurements on the deep seafloor with the usual instrumentation do not require corrections, they may be needed for some regions with unusual environmental parameters. As mentioned above, shallow water measurements may be subject to temporal bottom water temperature (BWT) variability, causing non-linear gradients in the seafloor. Using heat conduction theory, corrections may be applied if BWT variability is monitored for a sufficient period (months to years if possible) before the geothermal measurements. Conversely, nonlinear temperature–depth profiles can be inverted to yield BWT history assuming that the initial gradient was linear. This inverse procedure is non-unique, although closely spaced measured temperatures to a sufficient depth below the seafloor and well-determined thermal conductivity can reduce uncertainties. Continental margins and some deep-sea trenches are examples of regions where non-linear gradients have been measured, usually correlated with strong and variable deep currents that are probably focused by the seafloor topography. Topography causes lateral heat flow variability even with uniform and constant BWT, because the seafloor topography distorts isotherms that would otherwise be horizontal below a flat surface. Seafloor sedimentation reduces the heat flow measured because recently deposited sediments modify the seafloor boundary condition to which the equilibrium gradient must adjust. Corrections usually become significant when sedimentation rates exceed a few tens of meters per million years and/or the total sediment thickness exceeds 1 km. Vertical pore water flow in sediments may either enhance (for upward flow) or reduce (for downward flow) temperature gradients because the water advects some of the heat otherwise conducted upwards. The effects become significant for flow rates greater than a few centimeters per year, and nonlinear gradients may be expected if upward flow rates exceed about 10 cm year1. It is unusual for rates on normal deep seafloor to exceed the latter value because the fluid permeability of pelagic sediments is too low, but
43
coarser and sometimes rapidly deposited sediments of continental margins may support relatively rapid flows.
Range of Measured Parameters As discussed later, the plate tectonic paradigm predicts that the highest thermal gradients and heat fluxes should be found in the youngest seafloor, and the lowest values in the oldest. Although this is generally true when measurements are averaged over regions of various seafloor ages, considerable variability occurs over the youngest seafloor as a result of vigorous hydrothermal circulation. The cooling of hot and permeable upper ocean crust supports seawater convection in the crust over lateral circulation scales of at least several to a few tens of kilometers; vertical scales probably do not exceed a few kilometers. The highest gradients and heat fluxes (up to 1001C m1 and 100 W m2, respectively) are measured in sediments near seafloor vents where fluids upwell, and the lowest (o0.0051C m1 and 0.005 W m2, respectively) where fluids downwell. Mean upper basement fluid velocities may be several meters per year, and much greater in conduits that support vigorous seafloor venting. The detailed pore water flow and permeability structure for any system of seafloor hydrothermal circulation have not yet been investigated in detail and are probably very complex. After the seafloor ages to 50–70 My, hydrothermal circulation appears to decrease to a level where modulation of the surface heat flux is small to negligible (Figure 3), probably dependent on the thickness and uniformity of sediment cover. For the oldest seafloor (100–180 My), heat flux is relatively uniform at about 50 mW m2710%. This value probably reflects the secular cooling of the upper oceanic mantle. Thermal conductivity of marine sediments generally varies less than a factor of two, from about 0.7 to o1.4 Wm1 K1. The lowest values are associated with red clays or siliceous oozes, which have up to 70% or more of water by weight and the highest with carbonate oozes or coarse sediments near continental margins with a high proportion of quartz. Since water has a low thermal conductivity (0.6 W m1 K1), sediments with a large percentage of water have low thermal conductivity. Since calcium carbonate and quartz have high thermal conductivity (about 3 W m1 K1 and 4 W m1 K1, respectively), sediments with a large percentage of either of these minerals have high thermal conductivity. Conductivity variations with depth may be caused
(c) 2011 Elsevier Inc. All Rights Reserved.
44
GEOPHYSICAL HEAT FLOW
Depth (km)
3 4 5 GDH1 PSM HS
6 7 0
50
(A)
100 Age (My)
150
250
_
Heat flow (mW m 2)
GDH1 PSM
200
HS 150
Thermal Models 100
Data for Thermal Modeling
50 0
0
150
100
50 Age (My)
(B)
Fraction
1.5 1.0 0.5 0 (C)
climate (e.g., glacial) cycles. When thermal conductivity varies rapidly with depth the temperature–depth profile may not be linear. If thermal conductivity has been measured at closely spaced depth intervals, then the heat flow can be calculated using a method introduced by Bullard. This approach assumes the absence of significant heat sources or sinks and onedimensional, steady-state, conductive heat flow. Thus there is a linear relationship between temperature and the thermal resistance of the sediments (the sum from the surface to the depth of the temperature measurement of the inverse of thermal conductivity times the sediment thickness for that given thermal conductivity). The slope determined from a least-squares fit of this ‘Bullard plot’ gives the appropriate conductive heat flux.
tseal = 65 ± 10 My 0
50
100 Age (My)
150
Figure 3 Data and models for ocean depth (A) and heat flow (B) as a function of age. (Reproduced with permission from Stein and Stein, 1992.) The data are averaged in 2 My bins, and one standard deviation about the mean value for each is shown by the envelope. Shown are the plate model of Parsons and Sclater (1977) (PSM), a cooling halfspace model with the same thermal parameters (HS), and the GDH1 plate model from Stein and Stein (1992). (C) Heat flow fraction (observed heat flow/GDH1 prediction) with age averaged in 2 My bins. The discrepancy for ages o50–70 My presumably indicates the fraction of the heat transported by hydrothermal flow. The fractions for ages o50 My (closed circles), which were not used in deriving GDH1, are fit by a least squares line. The sealing age, where the line reaches the fractional value of one, is 65710 My.
by: (1) turbidity flows that sort grain sizes into layers with different proportions of pore water; (2) ice rafting in higher latitudes that may intersperse higher conductivity sediments derived from the continents and lower conductivity marine pelagic sediments; and (3) variability in the proportion of calcium carbonate in sediments near the equator deposited over
Oceanic lithosphere forms at midocean ridges, where hot magma upwells, and then cools to form plates as the material moves away from the spreading center. As the plate cools, heat flow decreases and the seafloor deepens (Figure 3). However, only shallow (less than 1 km) measurements of lithospheric temperatures are possible. Hence, the two primary data sets used to constrain models for the variation in lithospheric temperature with age are seafloor depths and heat flow. The depth, corrected for sediment load, depends on the temperature integrated over the lithospheric thickness. The heat flow is proportional to the temperature gradient. Initially seafloor depths rapidly increase, with the average increase relative to the ridge crest proportional to the square root of the crustal age. However, for ages greater than about 50–70 My, the average increase in depth is slower and the curve is said to ‘flatten.’ Mean heat flow also decreases rapidly away from the ridge crest, with values approximately proportional to the inverse of the square-root of the age, but after about 50 My this curve also ‘flattens.’ Halfspace and Plate Models
Two different mathematical models are often used to describe the thermal evolution of the oceanic lithosphere, the halfspace (or boundary layer) and plate models. For the halfspace model, the predicted lithospheric thickness increases proportionally with the square root of age. Hence depth and heat flow vary as the square root of age and the reciprocal of the square root of age, respectively. However, the halfspace model cannot explain the ‘flattening’ of the curves. Alternatively, the plate model represents the
(c) 2011 Elsevier Inc. All Rights Reserved.
GEOPHYSICAL HEAT FLOW
lithosphere as a layer with a fixed constant temperature at its base. Initially, the cooling for the plate model is the same as the halfspace model, but for older ages the influence of the lower boundary results in slower cooling, approximately predicting the observed flattening of the depths and heat flow at older ages. The plate base is assumed to represent a depth at which additional heat is supplied from the mantle below to prevent the halfspace cooling at older ages. However, the model does not directly describe how this heat is added.
45
understanding of the physics of processes affecting lithospheric thermal evolution improve, new and better reference models will be developed.
Application to Lithospheric Processes Using a reference model for the expected heat flow, regions can be examined to observe if the measured heat flow differs from that predicted and, if so, to study the causes of the discrepancy. Two primary discrepancies are assumed to reflect hydrothermal circulation and midplate swells.
Reference Models Hydrothermal Circulation
Heat flow measurements for crust of ages 0–65 My are generally lower than the predictions of all commonly used reference models (Figure 3). Because the heat flow measurements primarily reflect conductive heat flow, this difference has been attributed to hydrothermal water flow in the crust and sediments transporting some of the heat assumed in thermal models to be transferred by conduction. The missing heat transported by convection must appear somewhere, either as high conductive heat flow or as advective discharge to the sea. It is estimated that of the predicted global oceanic heat flux of 32 1012 W,
_
Heat flow (W m 1)
1.5
1.0
0.5
0 20°C
Depth (s)
Thermal models are solutions to the inverse problem of finding the temperature as a function of age that best fits the depth and heat flow. The data are used to estimate the primary model parameters (typically plate thickness, basal temperature, and thermal expansion coefficient), subject to other parameters generally specified a priori. The global average depth/heat flow variation with age, calculated for a specific model, data, and parameters is used as a reference model to represent ‘normal’ oceanic lithosphere. ‘Anomalies’, deviations from the reference model, are then investigated to see if they reflect a significant difference in thermal (or other) processes from the global average. It is important to realize that anomalies (defined as those observed minus the model-predicted values) for various reference models may differ significantly, and hence lead to different tectonic inferences. Until recently, a 125 km-thick plate model by Parsons and Sclater (denoted here PSM) was commonly used. Subsequently, as more data became available, it was noted that PSM systematically overpredicts depths and underpredicts heat flow for lithosphere older than 70–100 My, causing widespread ‘anomalies.’ A later joint inversion of the depth and heat flow data by Stein and Stein found that these ‘anomalies’ are reduced significantly by a plate model termed GDH1. GDH1 has a thinner lithosphere (95710 km), and a basal temperature of 145071001C, consistent with the PSM estimate (135072751C). GDH1 predicts heat flow in mW m2 as a function of age, t, in millions of years equal to 510 t1/2 for ages less than about 55 My, and 48 þ 96 exp( 0.0278 t) for older ages. Inversion of the same data, while prescribing a basal temperature of 13501C (a typically assumed temperature for upwelling magma at the ridge based on results from experimental petrology), also yields a thin (100 km) plate, with a somewhat higher value of lithospheric conductivity and hence makes very similar predictions. As the quality of the observed data and our
40°C 60°C 80°C
4.0 0 4.5
W
5 km
Line 88_ 30
E
Figure 4 Portion of the heat flow data from the FlankFlux survey on the Juan de Fuca Ridge sites where bare rock penetrates the sediment column have dramatically higher heat flow than the surrounding areas. Away from these outcrops, heat flow varies inversely with depth to basement rock. Expected heat flow for its age is shown with the dashed line. (Reproduced with permission from Davis EE, Chapman DS, Mottl MJ et al. (1992) FlankFlux: an experiment to study the nature of hydrothermal circulation in young oceanic crust. Canadian Journal of Earth Sciences 29: 925–952.)
(c) 2011 Elsevier Inc. All Rights Reserved.
46
GEOPHYSICAL HEAT FLOW
approximately one third occurs by hydrothermal flow. Characteristic heat flow patterns in young crust appear associated with water flow. Surveys at sites like the Galapagos Spreading Center show high heat flow associated with presumed upwelling zones at basement highs or fault scarps, and low heat flow associated with presumably down-flowing water at basement lows. Heat flow over the highs may be two to ten (or more) times greater than at nearby sites where water may be recharging the system (Figure 4). Closely spaced surveys in young crust show high scatter, presumably in part due to the local variations in sediment distribution, basement relief, and crustal permeability. The most spectacular evidence for hydrothermal circulation are vents near the ridge crest where hydrothermally altered water at temperatures up to 4001C exits the seafloor. Hot rock or magma at shallow depths provides a large heat source for the flow. The high sulfide content of the hot fluids supports unusual local biota from bacteria to tube worms. Vent areas show complex flow patterns with very high heat flow close to low heat flow. Although most vigorous vents are near spreading centers, on older crust isolated crystalline crust outcrops surrounded by a well-sedimented area can also vent measurable amounts of hydrothermal fluid (e.g., the FlankFlux survey area on the Juan de Fuca plate). The persistence of the heat flow discrepancy well away from ridges indicates that hydrothermal heat loss occurs in older crust. The ‘sealing age,’ defined as that beyond which measured heat flow approximately equals that predicted, is presumed to indicate the near cessation of heat transfer by hydrothermal circulation. The sealing age for the entire global data set is about 65710 My. Hence, although there are presumably local variations, water flow in older crust seems not, in general, to transport significant amounts of heat. However, some heat flow surveys (e.g., in the north-west Atlantic Ocean on 80 My crust and the Maderia Abyssal Plain on 90 My crust) indicate that hydrothermal circulation may continue, even if relatively little heat is lost by convection into the sea water. The reasons for the ‘sealing’ are not well understood. An earlier view was that 100– 200 m of sediment would be sufficiently impermeable to ‘seal’ off the crust from the sea, so that heat flow at these sites would yield a ‘reliable’ value, i.e., that predicted by conduction-only thermal models. However, many such sites have lower than expected values. A simple way to reconcile these observations with the present understanding of seafloor hydrology to assume that if water cannot flow vertically through thick sediments at a particular site, it may
flow laterally to a fault or basement outcrop, and then be manifested as either high conductive heat flow or hot water exiting to the sea, such as observed for the FlankFlux area. It has been suggested that the porosity and permeability of the crust decrease with increasing age, thus significantly reducing the water flow. However, recent studies of permeability and seismic velocity suggest that most of the rapid change occurs within the first 5–15 My and that older crust may still retain some relatively permeable pathways. Hydrothermal circulation has profound implications for the chemistry of the oceanic crust and sea water, because sea water reacts with the crust, giving rise to hydrothermal fluid of significantly different composition. The primary geochemical effects are thought to result from the high-temperature water flow observed at ridge axes. Circulation of sea water through hot rock removes magnesium and sulfate from sea water and enriches calcium, potassium, silica, iron, manganese, and other elements within the hydrothermal solution. Estimating the volume of water flowing through the crust depends on the heat capacity of the water, the heat lost by convection, and the assumed temperature of the water, the latter having the largest uncertainties. Although the heat flow anomaly is greatest in young lithosphere, only about 30% of the hydrothermal heat loss and 7% of the water flow occurs in crust younger than 1 My. This effect should be significant for ocean chemistry because a water volume equivalent to that of the total ocean is estimated to circulate through the oceanic crust about once every 0.5–5 My. Hot Spots
Oceanic midplate swells are identified by seafloor depths shallower than expected for their lithospheric age. Thus, models of the processes giving rise to these regions rely on assessments of how their heat flow and other properties differ from unperturbed lithosphere. The origin of these swells is generally thought to be related to upwelling mantle plumes (hot spots) that result in uplift and volcanism. Two basic types of models have been proposed for their origin. In one, the swell is a thermal effect due to the hot spot thinning and heating the lithosphere at depth. In the second, the uplift is primarily due to the dynamic effects of the upwelling plume, which may largely reflect thermal buoyancy forces within the upwelling mantle. The thermal models predict significantly larger heat flow anomalies than do the dynamic models. Detailed heat flow measurements have been made for Hawaii, Cape Verde, Reunion, Bermuda, and Crozet swells. Relative to the PSM reference curve, large heat flow anomalies are suggested.
(c) 2011 Elsevier Inc. All Rights Reserved.
GEOPHYSICAL HEAT FLOW
47
0 1
Mid way 2
5 4
3 4
25
H 7
5 G
4
6
8 20
2000
5
Depth (m)
30°N
E DB
F
5
5
4
4000
8
7
6
5
4
3
2
1
Hawaii
C
GDH1 PSM
A 6000 175
165
170
160
155˚W
200
8
7
400 600 800 1000 1200 Distance (km)
G
F
E C B A
_
H
Heat flow (W m 2)
70
_
Heat flow (mW m 2 )
70
0
60 GDH1 50
PSM
40
4 6
60
5
1 3
2 GDH1
50 PSM 40
30
20 10 Time since hot spot passage (My)
0
0
200
400 600 800 1000 1200 Distance (km)
Figure 5 Heat flow data for transects along (lower left) and across (lower right) the Hawaiian Swell at the locations shown (upper left). The heat flow, though anomalously high with respect to the Parsons and Sclater (PSM) model, is at most slightly above that expected for GDH1. The predicted heat flow values are for 100 My (lower right) and 95–110 My (lower left). Figure modified from Von Herzen et al. (1982, 1989). (Reproduced with permission from Stein CA and Stein S (1993) Constraints of Pacific midplate swells from global depth-age and heat flow-age models. The Mesozoic Pacific, American Geophysical Union Monograph 77: 53–76.)
However, relative to both measured heat flow from off-swell lithosphere and the GDH1 reference curve, smaller heat flow anomalies (less than 5–8 mW m2) are deduced, supporting the idea of a primary dynamic mechanism for hotspots (Figure 5).
Application to Marine Margin Studies Subduction Zones
Many extensive marine heat flow surveys have been done near the Japanese subduction zones. On average, observed heat flow from the trench axis to the forearc area is lower than that characteristic of the crust seaward of the trench. However, heat flow is higher and more variable over the volcanic arc and back arc region compared to the area seaward of the trench. Recent surveys for subduction zones, including Barbados, Nankai, and Cascadia, show that heat flow is highly variable. Within accretionary prisms, high values are often associated with upward advection of pore fluids, typically found along faults and the bottom decollement. Although they have not been
investigated extensively, the active nonaccretionary forearcs seem to have the lowest mean heat fluxes (20– 30 mW m2). Relatively low heat flow values over the subducting lithosphere are the thermal consequence of the subduction of one plate beneath another. Overall, heat flow depends on the age, rate, geometry, and thermal structure of the subducting plate and the sediment thickness and deformation history of the region. Heat flow in marginal basins behind the volcanic arcs is often high. Many back arc basins with high heat flow appear to have formed by back-arc spreading processes similar to those at midocean ridges, within the last 50 My. Passive Continental Margins
Passive continental margins form after continental crust is rifted and seafloor spreading occurs. Because rifting heats the lithosphere, heat flow data can be used to constrain models of rifting. Subsequent to rifting, margins slowly subside as cooling occurs. Simple models for this process suggest that subsequent to rifting the additional heat will almost completely dissipate within 100 My. Although the
(c) 2011 Elsevier Inc. All Rights Reserved.
48
GEOPHYSICAL HEAT FLOW
amount of radioactive heating from continental crust may introduce some variability, older passive margins typically have heat flows similar to old oceanic crust, whereas young margins, such as in the Red Sea, have high heat flow.
Gas Hydrates
Gas hydrates, solid composites of biogenically derived gasses (mainly methane) combined with water ice, may be present in marine sediments, especially at continental margins, under the correct pressure and temperature conditions. Gas hydrates are detected by direct sampling, and inferred from seismic reflection data when a strong bottom-simulating reflector (BSR) is produced by the seismic velocity contrast between the gas hydrate and the sediment below. Given a depth (and thus pressure) of a bottom-simulating reflector, its temperature can be predicted from known relationships and a heat flow calculated. Alternatively, heat flow data can be used to estimate the bottom-simulating reflector temperature. The volume of hydrocarbons contained in marine gas hydrates is large. Changes in eustatic sea level (hence pressure) or bottom seawater temperatures could result in their release and thus increase greenhouse gasses and affect global climate.
Further Reading Humphris S, Mullineaux L, Zierenberg R, and Thomson R (eds.) (1995) Seafloor hydrothermal systems, physical, chemical, biological, and geological interactions, Geophysical Monographs, 91, 425--445. Hyndman RD, Langseth MG, and Von Herzen RP (1987) Deep Sea Drilling Project geothermal measurements: a review. Review of Geophysics 25: 1563--1582. Langseth MG, Jr and Von Herzen RP (1971) Heat Sow through the Soor of the world oceans. In: Maxwell AE (ed.) The Sea, vol. IV, part 1, pp. 299--352. New York: Wiley-Interscience. Lowell RP, Rona PA, and Von Herzen RP (1995) Seafloor hydrothermal systems. Journal of Geophysical Research 100: 327--352. Parsons B and Sclater JG (1977) An analysis of the variation of ocean floor bathymetry and heat flow with age. Journal of Geophysical Research 82: 803--827. Pollack HN, Hurter SJ, et al. (1993) Heat flow from the earth’s interior: analysis of the global data set. Review of Geophysics 31: 267--280. Stein CA and Stein S (1992) A model for the global variation in oceanic depth and heat flow with lithospheric age. Nature 359: 123--129. Von Herzen RP (1987) Measurement of oceanic heat flow. In: Sammis C and Henyey T (eds.) Methods of Experimental Physics – Geophysics, vol. 24, Part B, pp. 227--263. London: Academic Press. Wright JA and Louden KE (eds.) (1989) CRC Handbook of Seafloor Heat Flow. Boca Raton: CRC Press.
(c) 2011 Elsevier Inc. All Rights Reserved.
GLACIAL CRUSTAL REBOUND, SEA LEVELS, AND SHORELINES K. Lambeck, Australian National University, Canberra, ACT, Australia Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1157–1167, & 2001, Elsevier Ltd.
Introduction Geological, geomorphological, and instrumental records point to a complex and changing relation between land and sea surfaces. Elevated coral reefs or wave-cut rock platforms indicate that in some localities sea levels have been higher in the past, while observations elsewhere of submerged forests or flooded sites of human occupation attest to levels having been lower. Such observations are indicators of the relative change in the land and sea levels: raised shorelines are indicative of land having been uplifted or of the ocean volume having decreased, while submerged shorelines are a consequence of land subsidence or of an increase in ocean volume. A major scientific goal of sea-level studies is to separate out these two effects. A number of factors contribute to the instability of the land surfaces, including the tectonic upheavals of the crust emanating from the Earth’s interior and the planet’s inability to support large surface loads of ice or sediments without undergoing deformation. Factors contributing to the ocean volume changes include the removal or addition of water to the oceans as ice sheets wax and wane, as well as addition of water into the oceans from the Earth’s interior through volcanic activity. These various processes operate over a range of timescales and sea level fluctuations can be expected to fluctuate over geological time and are recorded as doing so. The study of such fluctuations is more than a scientific curiosity because its outcome impacts on a number of areas of research. Modern sea level change, for example, must be seen against this background of geologically–climatologically driven change before contributions arising from the actions of man can be securely evaluated. In geophysics, one outcome of the sea level analyses is an estimate of the viscosity of the Earth, a physical property that is essential in any quantification of internal convection and thermal processes. Glaciological modeling of the behavior of large ice sheets during the last cold
period is critically dependent on independent constraints on ice volume, and this can be extracted from the sea level information. Finally, as sea level rises and falls, so the shorelines advance and retreat. As major sea level changes have occurred during critical periods of human development, reconstructions of coastal regions are an important part in assessing the impact of changing sea levels on human movements and settlement.
Tectonics and Sea Level Change Major causes of land movements are the tectonic processes that have shaped the planet’s surface over geological time. Convection within the high-temperature viscous interior of the Earth results in stresses being generated in the upper, cold, and relatively rigid zone known as the lithosphere, a layer some 50–100 km thick that includes the crust. This convection drives plate tectonics — the movement of large parts of the lithosphere over the Earth’s surface — mountain building, volcanism, and earthquakes, all with concomitant vertical displacements of the crust and hence relative sea level changes. The geological record indicates that these processes have been occurring throughout much of the planet’s history. In the Andes, for example, Charles Darwin identified fossil seashells and petrified pine trees trapped in marine sediments at 4000 m elevation. In Papua New Guinea, 120 000-year-old coral reefs occur at elevations of up to 400 m above present sea level (Figure 1). One of the consequences of the global tectonic events is that the ocean basins are being continually reshaped as mid-ocean ridges form or as ocean floor collides with continents. The associated sea level changes are global but their timescale is long, of the order 107 to 108 years, and the rates are small, less than 0.01 mm per year. Figure 2 illustrates the global sea level curve inferred for the past 600 million years from sediment records on continental margins. The long-term trends of rising and falling sea levels on timescales of 50–100 million years are attributed to these major changes in the ocean basin configurations. Superimposed on this are smaller-amplitude and shorter-period oscillations that reflect more regional processes such as large-scale volcanism in an ocean environment or the collision of continents. More locally, land is pushed up or down episodically
(c) 2011 Elsevier Inc. All Rights Reserved.
49
50
GLACIAL CRUSTAL REBOUND, SEA LEVELS, AND SHORELINES
and, in any discussion of global changes in sea level, information from such localities is best set aside in favor of observations from more stable environments.
Glacial Cycles and Sea Level Change
600 400 200 0 590
500
440 410
360 290 250 210 Time (million years BP)
140
65
35
5
Sea level above present level (m)
Pliocene_Pleistocene
Miocene
Paleocene & Eocene
Oligocene
Cretaceous
Permian
Carboniferous
Devonian
Silurian
Cambrian
Ordovician
in response to the deeper processes. The associated vertical crustal displacements are rapid, resulting in sea level rises or falls that may attain a few meters in amplitude, but which are followed by much longer periods of inactivity or even a relaxation of the original displacements. The raised reefs illustrated in Figure 1, for example, are the result of a large number of episodic uplift events each of typically a meter amplitude. Such displacements are mostly local phenomena, the Papua New Guinea example extending only for some 100–150 km of coastline. The episodic but local tectonic causes of the changing position between land and sea can usually be identified as such because of the associated seismic activity and other tell-tale geological signatures. The development of geophysical models to describe these local vertical movements is still in a state of infancy
Jurassic
Figure 1 Raised coral reefs from the Huon Peninsula, Papua New Guinea. In this section the highest reef indicated (point 1) is about 340 m above sea level and is dated at about 125 000 years old. Elsewhere this reef attains more than 400 m elevation. The top of the present sea cliffs (point 2) is about 7000 years old and lies at about 20 m above sea level. The intermediate reef tops formed at times when the rate of tectonic uplift was about equal to the rate of sea level rise, so that prolonged periods of reef growth were possible. Photograph by Y. Ota.
Triassic
More important for understanding sea level change on human timescales than the tectonic contributions – important in terms of rates of change and in terms of their globality – is the change in ocean volume driven by cyclic global changes in climate from glacial to interglacial conditions. In Quaternary time, about the last two million years, glacial and interglacial conditions have followed each other on timescales of the order of 104–105 years. During interglacials, climate conditions were similar to those of today and sea levels were within a few meters of their present day position. During the major glacials, such as 20 000 years ago, large ice sheets formed in the northern hemisphere and the Antarctic ice sheet expanded, taking enough water out of the oceans to lower sea levels by between 100 and 150 m. Figure 3 illustrates the changes in global sea level over the last 130 000 years, from the last interglacial, the last time that conditions were similar to those of today, through the Last Glacial Maximum and to the present. At the end of the last interglacial, at 120 000 years ago, climate began to fluctuate; increasingly colder conditions were reached, ice sheets over North America and Europe became more or less permanent features of the landscape, sea levels reached progressively lower values, and large parts of today’s coastal shelves were exposed. Soon after the culmination of maximum glaciation, the ice sheets again disappeared, within a period of about
_200
Figure 2 Global sea level variations through the last 600 million years estimated from seismic stratigraphic studies of sediments deposited at continental margins. Redrawn with permission from Hallam A (1984) Pre-Quaternary sea-level changes. Annual Review of Earth and Planetary Science 12: 205–243.
(c) 2011 Elsevier Inc. All Rights Reserved.
GLACIAL CRUSTAL REBOUND, SEA LEVELS, AND SHORELINES
Relative sea level (m)
20 0 _20 _40 _60 _80 _100 _120
0
20
60 80 100 40 Time (thousand years BP)
120
Figure 3 Global sea level variations (relative to present) since the time of the last interglacial 120 000–130 000 years ago when climate and environmental conditions were last similar to those of the last few thousand years. Redrawn with permission from Chapell J et al. (1996) Reconciliation of Late Quaternary sea level changes derived from coral terraces at Huon Peninsula with deep sea oxygen isotope records. Earth and Planetary Science Letters 141b: 227–236.
10 000 years, and climate returned to interglacial conditions. The global changes illustrated in Figure 3 are only one part of the sea level signal because the actual ice– water mass exchange does not give rise to a spatially uniform response. Under a growing ice sheet, the Earth is stressed; the load stresses are transmitted through the lithosphere to the viscous underlying mantle, which begins to flow away from the stressed area and the crust subsides beneath the ice load. When the ice melts, the crust rebounds. Also, the meltwater added to the oceans loads the seafloor and the additional stresses are transmitted to the mantle, where they tend to dissipate by driving flow to unstressed regions below the continents. Hence the seafloor subsides while the interiors of the continents rise, causing a tilting of the continental margins. The combined adjustments to the changing ice and water loads are called the glacio- and hydro-isostatic effects and together they result in a complex pattern of spatial sea level change each time ice sheets wax and wane.
Observations of Sea Level Change Since the Last Glacial Maximum Evidence for the positions of past shorelines occurs in many forms. Submerged freshwater peats and tree stumps, tidal-dwelling mollusks, and archaeological sites would all point to a rise in relative sea level since the time of growth, deposition, or construction. Raised coral reefs, such as in Figure 1, whale bones cemented in beach deposits, wave-cut rock platforms
51
and notches, or peats formed from saline-loving plants would all be indicative of a falling sea level since the time of formation. To obtain useful sea level measurements from these data requires several steps: an understanding of the relationship between the feature’s elevation and mean sea level at the time of growth or deposition, a measurement of the height or depth with respect to present sea level, and a measurement of the age. All aspects of the observation present their own peculiar problems but over recent years a substantial body of observational evidence has been built up for sea level change since the last glaciation that ended at about 20 000 years ago. Some of this evidence is illustrated in Figure 4, which also indicates the very significant spatial variability that may occur even when, as is the case here, the evidence is from sites that are believed to be tectonically stable. In areas of former major glaciation, raised shorelines occur with ages that are progressively greater with increasing elevation and with the oldest shorelines corresponding to the time when the region first became ice-free. Two examples, from northern Sweden and Hudson Bay in Canada, respectively, are illustrated in Figure 4A. In both cases sea level has been falling from the time the area became ice-free and open to the sea. This occurred at about 9000 years ago in the former case and about 7000 years later in the second case. Sea level curves from localities just within the margins of the former ice sheet are illustrated in Figure 4B, from western Norway and Scotland, respectively. Here, immediately after the area became ice-free, the sea level fell but a time was reached when it again rose. A local maximum was reached at about 6000 years ago, after which the fall continued up until the present. Farther away from the former ice margins the observed sea level pattern changes dramatically, as is illustrated in Figure 4C. Here the sea level initially rose rapidly but the rate decreased for the last 7000 years up to the present. The two examples illustrated, from southern England and the Atlantic coast of the United States, are representative of tectonically stable localities that lie within a few thousand kilometers from the former centers of glaciation. Much farther away again from the former ice sheets the sea level signal undergoes a further small but significant change in that the present level was first reached at about 6000 years ago and then exceeded by a small amount before returning to its present value. The two examples illustrated in Figure 4D are from nearby localities in northern Australia. At both sites present sea level was reached about 6000 years ago after a prolonged period of rapid rise with resulting highstands that are small in amplitude but geographically variable.
(c) 2011 Elsevier Inc. All Rights Reserved.
52
GLACIAL CRUSTAL REBOUND, SEA LEVELS, AND SHORELINES
Shoreline elevation (m)
250 140
Angermanland, northern Sweden
200
Ottawa Island, Hudson Bay, Canada
120 100
150 80 100
60 40
50 20 0 10
(A)
8
6
4
2
0
Shoreline elevation (m)
50 Andoya, northern Norway
40
7
5
6
4
0
30 20 10 10
0
15
5
10
0
0 15
5
10
0
10
0
0
_ 10 _ 10 _ 20 _ 30 _ 40 10
Shoreline elevation (m)
1
20
_ 20
Dungeness, Kent,
8
6
4
2
Cape Charles, Virginia, USA
0
_ 30 12
10
8
6
4
2
0
3
0
Karumba, Gulf of Carpentaria, North Queensland
_ 10 2 _ 20
2
_ 30
0
1 −1 6 5 4 3 2 1 0 North Queensland, Australia
_ 40 _ 50 12
(D)
2
30
southern England (C)
3
Firth of Forth, Scotland
40
10
Shoreline elevation (m)
8
50
_ 10 20
(B)
0
10
8
6
1
0 0 6 5 4 2 4 Time (1000 radiocarbon years BP)
3
2
1
0
˚ ngermanland, Figure 4 Characteristic sea level curves observed in different localities (note the different scales used). (A) From A Gulf of Bothnia, Sweden, and from Hudson Bay, Canada. Both sites lie close to centers of former ice sheet, over northern Europe and North America respectively. (B) From the Atlantic coast of Norway and the west coast of Scotland. These sites are close to former icesheet margins at the time of the last glaciation. (C) From southern England and Virginia, USA. These sites lie outside the areas of former glaciation and at a distance from the margins where the crust is subsiding in response to the melting of the nearby icesheet. (D) Two sites in northern Australia, one for the Coral Sea coast of Northern Queensland and the second from the Gulf of Carpentaria some 200 km away but on the other side of the Cape York Peninsula. At both localities sea level rose until about 6000 years ago before peaking above present level.
(c) 2011 Elsevier Inc. All Rights Reserved.
GLACIAL CRUSTAL REBOUND, SEA LEVELS, AND SHORELINES
The examples illustrated in Figure 4 indicate the rich spectrum of variation in sea level that occurred when the last large ice sheets melted. Earlier glacial cycles resulted in a similar spatial variability, but much of the record before the Last Glacial Maximum has been overwritten by the effects of the last deglaciation.
Glacio-Hydro-Isostatic Models If, during the decay of the ice sheets, the meltwater volume was distributed uniformly over the oceans, then the sea level change at time t would be Dze ðtÞ ¼ ðchange in ice volume=ocean surface areaÞ ri =rw
½1a
where ri and rw are the densities of ice and water, respectively. (See end of article – symbols used.) This term is the ice-volume equivalent sea level and it provides a measure of the change in ice volume through time. Because the ocean area changes with time a more precise definition is Dze ðtÞ ¼
ð dVi ri ri dt dt r0 A0 ðtÞ
½1b
t
where A0 is the area of the ocean surface, excluding areas covered by grounded ice. The sea level curve illustrated in Figure 3 is essentially this function. However, it represents only a zero-order approximation of the actual sea level change because of the changing gravitational field and deformation of the Earth. In the absence of winds or ocean currents, the ocean surface is of constant gravitational potential. A planet of a defined mass distribution has a family of such surfaces outside it, one of which – the geoid – corresponds to mean sea level. If the mass distribution on the surface (the ice and water) or in the interior (the load-forced mass redistribution in the mantle) changes, so will the gravity field, the geoid, and the sea level change. The ice sheet, for example, represents a large mass that exerts a gravitational pull on the ocean and, in the absence of other factors, sea level will rise in the vicinity of the ice sheet and fall farther away. At the same time, the Earth deforms under the changing load, with two consequences: the land surface with respect to which the level of the sea is measured is time-dependent, as is the gravitational attraction of the solid Earth and hence the geoid. The calculation of the change of sea level resulting from the growth or decay of ice sheets therefore
53
involves three steps: the calculation of the amount of water entering into the ocean and the distribution of this meltwater over the globe; the calculation of the deformation of the Earth’s surface; and the calculation of the change in the shape of the gravitational equipotential surfaces. In the absence of vertical tectonic motions of the Earth’s surface, the relative sea level change Dzðj; tÞ at a site j and time t can be written schematically as Dzðj; tÞ ¼ Dze ðtÞ þ DzI ðj; tÞ
½2
where DzI(j, t) is the combined perturbation from the uniform sea level rise term [1]. This is referred to as the isostatic contribution to relative sea level change. In a first approximation, the Earth’s response to a global force is that of an elastic outer spherical layer (the lithosphere) overlying a viscous or viscoelastic mantle that itself contains a fluid core. When subjected to an external force (e.g., gravity) or a surface load (e.g., an ice cap), the planet experiences an instantaneous elastic deformation followed by a timedependent or viscous response with a characteristic timescale(s) that is a function of the viscosity. Such behavior of the Earth is well documented by other geophysical observations: the gravitational attraction of the Sun and Moon raises tides in the solid Earth; ocean tides load the seafloor with a time-dependent water load to which the Earth’s surface responds by further deformation; atmospheric pressure fluctuations over the continents induce deformations in the solid Earth. The displacements, measured with precision scientific instruments, have both an elastic and a viscous component, with the latter becoming increasingly important as the duration of the load or force increases. These loads are much smaller than the ice and water loads associated with the major deglaciation, the half-daily ocean tide amplitudes being only 1% of the glacial sea level change, and they indicate that the Earth will respond to even small changes in the ice sheets and to small additions of meltwater into the oceans. The theory underpinning the formulation of planetary deformation by external forces or surface loads is well developed and has been tested against a range of different geophysical and geological observations. Essentially, the theory is one of formulating the response of the planet to a point load and then integrating this point-load solution over the load through time, calculating at each epoch the surface deformation and the shape of the equipotential surfaces, making sure that the meltwater is appropriately distributed into the oceans and that the total ice–water mass is preserved. Physical inputs into the
(c) 2011 Elsevier Inc. All Rights Reserved.
54
GLACIAL CRUSTAL REBOUND, SEA LEVELS, AND SHORELINES
formulation are the time–space history of the ice sheets, a description of the ocean basins from which the water is extracted or into which it is added, and a description of the rheology, or response parameters, of the Earth. For the last requirement, the elastic properties of the Earth, as well as the density distribution with depth, have been determined from seismological and geodetic studies. Less well determined are the viscous properties of the mantle, and the usual procedure is to adopt a simple parametrization of the viscosity structure and to estimate the relevant parameters from analyses of the sea level change itself. The formulation is conveniently separated into two parts for schematic reasons: the glacio-isostatic effect representing the crustal and geoid displacements due to the ice load, and the hydro-isostatic effect due to the water load. Thus ½3 Dzðj; tÞ ¼ Dze ðtÞ þ Dzi ðj; tÞ þ Dzw ðj; tÞ where Dzi(j,t) is the glacio-isostatic and Dzw(j,t) the hydro-isostatic contribution to sea level change. (In reality the two are coupled and this is included in the formulation). If Dzi(j,t) is evaluated everywhere, the past water depths and land elevations H(j,t) measured with respect to coeval sea level are given by ½4 H ðj; tÞH0 ðjÞ Dzðj; tÞ where HðfÞ is the present water depth or land elevation at location j. A frequently encountered concept is eustatic sea level, which is the globally averaged sea level at any time t. Because of the deformation of the seafloor during and after the deglaciation, the isostatic term Dzðj; tÞ is not zero when averaged over the ocean at any time t, so that the eustatic sea level change is Dzeus ðtÞ ¼ Dze ðtÞ þ hDzI fj; tgi0 where the second term on the right-hand side denotes the spatially averaged isostatic term. Note that Dze ðtÞ relates directly to the ice volume, and not Dzeus ðtÞ.
The Anatomy of the Sea Level Function The relative importance of the two isostatic terms in eqn. [3] determines the spatial variability in the sea level signal. Consider an ice sheet of radius that is much larger than the lithospheric thickness: the limiting crustal deflection beneath the center of the load is Iri =Irm where I is the maximum ice thickness and Irm is the upper mantle density. This is the local isostatic approximation and it provides a reasonable approximation of the crustal deflection if the loading time is long compared with the relaxation time of the mantle. Thus for a 3 km thick ice sheet the maximum
deflection of the crust can reach 1 km, compared with a typical ice volume for a large ice sheet that raises sea level globally by 50–100 m. Near the centers of the formerly glaciated regions it is the crustal rebound that dominates and sea level falls with respect to the land. This is indeed observed, as illustrated in Figure 4A, and the sea level curve here consists of essentially the sum of two terms, the major glacio-isostatic term and the minor ice-volume equivalent sea level term Dze ðtÞ (Figure 5A). Of note is that the rebound continues long after all ice has disappeared, and this is evidence that the mantle response includes a memory of the earlier load. It is the decay time of this part of the curve that determines the mantle viscosity. As the ice margin is approached, the local ice thickness becomes less and the crustal rebound is reduced and at some stage is equal, but of opposite sign, to Dze(t). Hence sea level is constant for a period (Figure 5B) before rising again when the rebound becomes the minor term. After global melting has ceased, the dominant signal is the late stage of the crustal rebound and levels fall up to the present. Thus the oscillating sea level curves observed in areas such as Norway and Scotland (Figure 4B) are essentially the sum of two effects of similar magnitude but opposite sign. The early part of the observation contains information on earth rheology as well as on the local ice thickness and the globally averaged rate of addition of meltwater into the oceans. Furthermore, the secondary maximum is indicative of the timing of the end of global glaciation and the latter part of the record is indicative mainly of the mantle response. At the sites beyond the ice margins it is the meltwater term Dze(t) that is important, but it is not the sole factor. When the ice sheet builds up, mantle material flows away from the stressed mantle region and, because flow is confined, the crust around the periphery of the ice sheet is uplifted. When the ice sheet melts, subsidence of the crust occurs in a broad zone peripheral to the original ice sheet and at these locations the isostatic effect is one of an apparent subsidence of the crust. This is illustrated in Figure 5C. Thus, when the ocean volumes have stabilized, sea level continues to rise, further indicating that the planet responds viscously to the changing surface loads. The early part of the observational record (e.g., Figure 4C) is mostly indicative of the rate at which meltwater is added into the ocean, whereas the latter part is more indicative of mantle viscosity. In all of the examples considered so far it is the glacial-rebound that dominates the total isostatic adjustment, the hydro-isostatic term being present but comparatively small. Consider an addition of water that raises sea level by an amount D. The local
(c) 2011 Elsevier Inc. All Rights Reserved.
GLACIAL CRUSTAL REBOUND, SEA LEVELS, AND SHORELINES
55
300
300
Relative sea level change (m)
Prediction
ice
200
100
100
0
0 ESL
_ 100 (A)
200
20
Angerman River, Gulf of Bothnia
15
10
5
_ 100
0
20
Relative sea level change (m)
150
10
5
0
15
10
5
0
150
ice
100
100
50
50
0
0 _ 50
_ 50 ESL
_ 100 _ 150 20
15
200
200
(B)
Observation
Upper Firth of Forth, Scotland
15
5
10
_ 100 _ 150
0
20
20
20 water
0
0 Relative sea level change (m)
_ 20
ice
_ 20
_ 40
_ 40
_ 60
_ 60 _ 80
ESL
_ 80
_ 100
_ 100 _ 120
Dungeness, south of England
_ 140 (C)
15
20
Relative sea level change (m)
10
(D)
0 _ 10
10
5
20
10
15
5
0
0 _ 10
ice ESL
Karumba, Gulf of Carpentaria
_ 30 8
_ 140
10
water
_ 20 _ 40 10
0
_ 120
6
4
2
0
_ 20 _ 30 _ 40 10
8
6
4
2
0
Time (1000 years BP)
Figure 5 Schematic representation of the sea level curves in terms of the ice-volume equivalent function Dze ðtÞ (denoted by ESL) and the glacio- and hydro-isostatic contributions Dzi ðf; tÞ, Dzw ðf; tÞ (denoted by ice and water, respectively). The panels on the left indicate the predicted individual components and the panels on the right indicate the total predicted change compared with the observed values (data points). (A) For the sites at the ice center where Dzi ðf; tÞbDze ðtÞbDzw ðf; tÞ. (B) For sites near but within the ice margin where jDzi ðf; tÞjBjDze ðtÞjbjDzw ðf; tÞj, but the first two terms are of opposite sign. (C) For sites beyond the ice margin where jDze ðtÞj > jDzi ðf; tÞjbjDzw ðf; tÞj. (D) For sites at continental margins far from the former ice sheets where jDzw ðf; tÞj > jDzi ðtÞj. Adapted with permission from Lambeck K and Johnston P (1988). The viscosity of the mantle: evidence from analyses of glacial rebound phenomena. In: Jackson ISN (ed.) The Earth’s Mantle – Composition, Structure and Evolution, pp. 461–502. Cambridge: Cambridge University Press.
(c) 2011 Elsevier Inc. All Rights Reserved.
GLACIAL CRUSTAL REBOUND, SEA LEVELS, AND SHORELINES
Sea Level Change and Shoreline Migration: Some Examples Figure 6 illustrates the ice-volume equivalent sea level function Dze ðtÞ since the time of the last maximum glaciation. This curve is based on sea level
indicators from a number of localities, all far from the former ice sheets, with corrections applied for the isostatic effects. The right-hand axis indicates the corresponding change in ice volume (from the relation [1b]). Much of this ice came from the ice sheets over North America and northern Europe but a not insubstantial part also originated from Antarctica. Much of the melting of these ice sheets occurred within 10 000 years, at times the rise in sea level exceeding 30 mm per year, and by 6000 years ago most of the deglaciation was completed. With this sea level function, individual ice sheet models, the formulation for the isostatic factors, and a knowledge of the topography and bathymetry of the world, it becomes possible to reconstruct the paleo shorelines using the relation [4]. Scandinavia is a well-studied area for glacial rebound and sea level change since the time the ice retreat began about 18 000 years ago. The observational evidence is quite plentiful and a good record of ice margin retreat exists. Figure 7 illustrates examples for two epochs. The first (Figure 7A), at 16 000 years ago, corresponds to a time after onset of deglaciation. A large ice sheet existed over
20
Time (radiocarbon, 1000 years BP) 15 10 5
0
0
0
_20
7.7
_40 _60
23.1
_80 _100
38.6
Ice volume (106 km3 )
isostatic response to this load is Drw =rm , where rw is the density of ocean water. This gives an upper limit to the amount of subsidence of the sea floor of about 30 m for a 100 m sea level rise. This is for the middle of large ocean basins and at the margins the response is about half as great. Thus the hydro-isostatic effect is significant and is the dominant perturbing term at margins far from the former ice sheets. This occurs at the Australian margin, for example, where the sea level signal is essentially determined by Dze ðtÞ and the water-load response (Figure 4D). Up to the end of melting, sea level is dominated by Dze ðtÞ but thereafter it is determined largely by the water-load term Dzw ðj; tÞ, such that small highstands develop at 6000 years. The amplitudes of these highstands turn out to be strongly dependent on the geometry of the waterload distribution around the site: for narrow gulfs, for example, their amplitude increases with distance from the coast and from the water load at rates that are particularly sensitive to the mantle viscosity. While the examples in Figure 5 explain the general characteristics of the global spatial variability of the sea level signal, they also indicate how observations of such variability are used to estimate the physical quantities that enter into the schematic model (3). Thus, observations near the center of the ice sheet partially constrain the mantle viscosity and central ice thickness. Observations from the ice sheet margin partially constrain both the viscosity and local ice thickness and establish the time of termination of global melting. Observations far from the ice margins determine the total volumes of ice that melted into the oceans as well as providing further constraints on the mantle response. By selecting data from different localities and time intervals, it is possible to estimate the various parameters that underpin the sea level eqn. [3] and to use these models to predict sea level and shoreline change for unobserved areas. Analyses of sea-level change from different regions of the world lead to estimates for the lithospheric thickness of between 60–100 km, average upper mantle viscosity (from the base of the lithosphere to a depth of 670 km) of about (1–5)1020 Pa s and an average lower mantle viscosity of about (1–5)1022 Pa s. Some evidence exists that these parameters vary spatially; lithosphere thickness and upper mantle viscosity being lower beneath oceans than beneath continents.
Ice-volume equivalent sea level (m)
56
_120 _140 25
54.0 5 20 15 10 Time (calibrated, 1000 years BP)
0
Figure 6 The ice-volume equivalent function Dze ðtÞ and ice volumes since the time of the last glacial maximum inferred from corals from Barbados, from sediment facies from north-western Australia, and from other sources for the last 7000 years. The actual sea level function lies at the upper limit defined by these observations (continuous line). The upper time scale corresponds to the radiocarbon timescale and the lower one is calibrated to calendar years. Adapted with permission from Fleming K et al. (1998) Refining the eustatic sea-level curve since the LGM using the far- and intermediate-field sites. Earth and Planetary Science Letters 163: 327–342; and Yokoyama Y et al. (2000) Timing of the Last Glacial Maximum from observed sea-level minimum. Nature 406: 713–716.
(c) 2011 Elsevier Inc. All Rights Reserved.
GLACIAL CRUSTAL REBOUND, SEA LEVELS, AND SHORELINES
57
Figure 8 Shoreline reconstructions for South East Asia and northern Australia at the time of the last glacial maximum, (A) at about 18 000 years ago and (B) at 12 000 years ago. The water depth contours are the same as in Figure 7. The water depths in the inland lakes at 18 000 years ago are contoured at 25 m intervals relative to their maximum levels that can be attained without overflowing.
Figure 7 Shoreline reconstructions for Europe (A) at 16 000 and (B) at 10 500 years ago. The contours are at 400 m intervals for the ice thickness (white) and 25 m for the water depths less than 100 m. The orange, yellow, and red contours are the predicted lines of equal sea level change from the specified epoch to the present and indicate where shorelines of these epochs could be expected if conditions permitted their formation and preservation. The zero contour is in yellow; orange contours, at intervals of 100 m, are above present, and red contours, at 50 m intervals, are below present. At 10 500 years ago the Baltic is isolated from the Atlantic and its level lies about 25 m above that of the latter.
Scandinavia with a smaller one over the British Isles. Globally sea level was about 110 m lower than now and large parts of the present shallow seas were exposed, for example, the North Sea and the English Channel but also the coastal shelf farther south such as the northern Adriatic Sea. The red and orange contours indicate the sea level change between this period and the present. Beneath the ice these rebound contours are positive, indicating that if shorelines could form here they would be above sea level today. Immediately beyond the ice margin, a broad but shallow bulge develops in the topography, which will subside as the ice sheet retreats. At the second epoch selected (Figure 7B) 10 500 years ago, the ice has retreated and reached a temporary halt as the climate briefly returned to colder conditions; the Younger Dryas time of Europe. Much of the Baltic was then ice-free and a freshwater lake developed at some
(c) 2011 Elsevier Inc. All Rights Reserved.
58
GLACIAL CRUSTAL REBOUND, SEA LEVELS, AND SHORELINES
25–30 m above coeval sea level. The flooding of the North Sea had begun in earnest. By 9000 years ago, most of the ice was gone and shorelines began to approach their present configuration. The sea level change around the Australian margin has also been examined in some detail and it has been possible to make detailed reconstructions of the shoreline evolution there. Figure 8 illustrates the reconstructions for northern Australia, the Indonesian islands, and the Malay–IndoChina peninsula. At the time of the Last Glacial Maximum much of the shallow shelves were exposed and deeper depressions within them, such as in the Gulf of Carpentaria or the Gulf of Thailand, would have been isolated from the open sea. Sediments in these depressions will sometimes retain signatures of these pre-marine conditions and such data provide important constraints on the models of sea level change. Part of the information illustrated in Figure 6, for example, comes from the shallow depression on the Northwest Shelf of Australia. By 12 000 years ago the sea has begun its encroachment of the shelves and the inland lakes were replaced by shallow seas.
Symbols used Ice-volume equivalent sea level. Uniform change in sea level produced by an ice volume that is distributed uniformly over the ocean surface. Perturbation in sea level due to glacioiDzI ðtÞ sostatic Dzi(t) and hydro-isostatic Dzw(t) effects. Dzeus ðtÞ Eustatic sea level change. The globally averaged sea level at time t. Dze ðtÞ
See also Beaches, Physical Processes Affecting. Coral Reefs. Fiordic Ecosystems. Geomorphology. Lagoons. Mangroves. Past Climate from Corals. Rocky Shores. Salt Marshes and Mud Flats. Salt Marsh Vegetation. Sandy Beaches, Biology of. Sea Level Change.
Further Reading Lambeck K (1988) Geophysical Geodesy: The Slow Deformations of the Earth. Oxford: Oxford University Press. Lambeck K and Johnston P (1999) The viscosity of the mantle: evidence from analysis of glacial-rebound phenomena. In: Jackson ISN (ed.) The Earth’s Mantle – Composition, Structure and Evolution, pp. 461--502. Cambridge: Cambridge University Press. Lambeck K, Smither C, and Johnston P (1998) Sea-level change, glacial rebound and mantle viscosity for northern Europe. Geophysical Journal International 134: 102--144. Peltier WR (1998) Postglacial variations in the level of the sea: implications for climate dynamics and solid-earth geophysics. Reviews in Geophysics 36: 603--689. Pirazzoli PA (1991) World Atlas of Holocene Sea-Level Changes. Amsterdam: Elsevier. van de Plassche O (ed.) (1986) Sea-Level Research: A Manual for the Collection and Evaluation of Data. Norwich: Geo Books. Sabadini R, Lambeck K, and Boschi E (1991) Glacial Isostasy, Sea Level and Mantle Rheology. Dordrecht: Kluwer.
(c) 2011 Elsevier Inc. All Rights Reserved.
GLIDERS C. C. Eriksen, University of Washington, Seattle, WA, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction Underwater gliders are a recently developed class of autonomous underwater vehicle (AUV) driven by buoyancy changes to fly along saw-tooth trajectories through the ocean. Gliders quickly have become the AUVs with the highest endurance and longest range. They are able to sample the ocean interior at comparatively low cost because they can operate independently of ships, some for the better part of a year under global remote control. Sampling the ocean on space and timescales as fine as those dominating its variability and over long ranges and durations has been a challenge throughout the history of ocean science. Traditionally, most knowledge of the ocean interior has been collected by ships. Starting in the late twentieth century, moored observations have complemented those from ship-based surveys, and more recently satellite remote sensing has provided images of sea surface conditions globally on ever-finer space and timescales. Were these techniques inexpensive, there would be no particular need for autonomous data collection by moving platforms. Unfortunately, ships stay at sea only for a month or two and rarely are directed to sample the same region for multiple cruises. Moorings sampling the open ocean number in the hundreds globally, typically in the dozens in the deep sea. This means variability of the ocean interior is generally undersampled. Undersampling is arguably the principal impediment to description and understanding of the ocean. Gliders have been developed to address the shortcomings of conventional observing means so that such dominant sources of variability as mesoscale eddies, fronts, and boundary currents can be resolved simultaneously in space and time and over sufficiently long periods and wide domains to allow them to be understood. They are able to make progress in solving the sampling problem simply by the numbers their economy affords. The salient characteristic of gliders is that they cost roughly the equivalent of a few days of research vessel operation to build and can be operated for a few months for the cost of a day of ship time.
The tasks gliders do best are those for which ships are least suited: intensive, regular, and sustained observations of oceanic properties that are readily measured by electronic means. Observations that require large or power-hungry instruments, physical collection of samples, or specialized labor are unsuitable for gliders and continue the need for ship support. Gliders excel at measuring standard properties typically collected by a conductivity–temperature–depth (CTD) package. Among their many advantages is their ability to function well through the most severe seas the ocean has to offer, day and night, around the globe. They are a means to erasing the fair-weather bias of ship-based observations. One may sample the ocean at deliberately chosen locations and times with gliders, examine the data very shortly after it is collected, and alter the sampling plan as often as a glider reaches the sea surface and communicates. The density of observations can readily be scaled to the phenomena of interest by adjusting the number and distribution of platforms to address local, regional, or global variability on diurnal, fortnightly, or seasonal scales, for example. Of course, gliders are not without limitation: long range and high endurance are achieved at the cost of traveling slowly through the ocean. When they cease to communicate, they are lost.
A Short History John Swallow began the era of exploring the ocean interior autonomously in the 1950s with acoustically tracked floats ballasted to follow water motions at depth, famously deployed in the various North Atlantic locales. In following decades, acoustic floattracking technology extended to ocean basin scale. Russ Davis and Doug Webb turned to satellite tracking to follow the motion of Autonomous Lagrangian Circulation Explorer (ALACE) floats that periodically adjusted their buoyancy to reach the sea surface for communication, and then returned to a ‘parking’ depth to drift. ALACE floats were the precursor to Argo floats, platforms used in an ambitious international program to seed the ocean with 3000 profiling floats. While Argo floats continue to track horizontal water movements, the focus of the Argo program is largely on CTD profiles collected periodically to describe large-scale features of ocean circulation. Gliders are the descendants of profiling floats. They can be thought of as profiling floats with wings.
(c) 2011 Elsevier Inc. All Rights Reserved.
59
60
GLIDERS
Their essential difference is that their sampling location is controlled, subject to a limited ability to navigate through currents. Floats randomly sample where currents carry them. While gliders do not offer a Lagrangian view of the ocean, they are commanded to describe oceanic conditions at deliberate times and locations, adding to the Eulerian description of the ocean. Henry Stommel recognized the potential of underwater gliders (the idea was shared with him by his friend, neighbor, and colleague Doug Webb), and published a science fiction article describing them in 1989. Written as the retrospective of a scientist working at a glider network control station in 2021, the article has proved remarkably prescient in describing the gliders that were developed over the next decade and a half. Stommel envisioned a fleet of about 1000 gliders plying the oceans, continuously collecting profiles and communicating them ashore via satellite, while being controlled to variously occupy hydrographic sections or chase features within the ocean. The fundamental vision was of a distributed network of relatively inexpensive platforms. His optimism on the pace of development can understandably be blamed on his enthusiasm, for he predicted the maiden voyage at sea of a glider to have taken place in 1994 and last 198 days. In reality, the first ocean deployment longer than a week took place in 1999 and the longest mission as of this writing lasted 217 days, accomplished in 2005. The Autonomous Oceanographic Sampling Network program set up by the US Office of Naval Research (ONR) in 1995 provided 5 years of support that resulted in gliders becoming a reality. The three operational gliders, Slocum (Webb Research Corp.), Spray (Scripps Institution of Oceanography and Woods Hole Oceanographic Institution), and Seaglider (University of Washington), were developed under this program and have subsequently been supported both by ONR and the US National Science Foundation. Four key technical elements enabled the successful development of underwater gliders: small reliable buoyancy engines, low-power computer microprocessors, Global Positioning System (GPS) navigation, and low-power duplex satellite communication. Seaglider, Slocum, and Spray were designed to fulfill similar missions, hence understandably share many elements (Figure 1). All three are of similar size, roughly 50 kg in mass, in order that both manufacturing and operating costs would be relatively small. The vehicles can be launched and recovered from small vessels by two people without power-assisted equipment. They are designed to carry out missions from a few weeks (Slocum) to
several months (Seaglider and Spray) duration while traveling at B0.5 kt (B0.25 m s 1). All are batterypowered, navigate by dead-reckoning underwater between GPS navigational fixes obtained at the sea surface, and send data and receive commands via a constellation of low Earth-orbit satellites on longrange missions. Derivatives of these three gliders are currently under development, including a version powered in part by extracting thermal energy from ocean stratification, one that operates under ice, and one capable of making open ocean full-depth dives. In addition, parallel development of a vehicle considerably larger, faster, and capable of flying along a shallower glider path is underway.
Design Considerations The three operational gliders were designed to address the sampling deficiencies of infrequent shipbased surveys and widely separated moorings in describing oceanic internal structure. Mesoscale eddies with characteristic space and timescales of O(100 km) and O(1 month) provide the principal noise to ocean general circulation and climate variability. Fronts, eddies, internal waves, diurnal cycles, and tidal fluctuations contribute to variability on smaller scales in problems of regional focus. Measurement platforms capable of resolving these features require sampling on scales O(10 km) and O(10 h) while the need to observe many realizations of random processes calls for sampling to be carried out for several weeks or months over tens or hundreds of kilometers. In order to observe at much lower cost than with ships or moorings, the three operational gliders were designed with endurance and range capability to meet these demands. Gliding is the conversion of a buoyancy force into a body’s forward as well as vertical motion. While aeronautical gliders are heavier than the air around them, hence always fall through it, underwater gliders are designed to alternately be heavier and lighter than the surrounding ocean so that they may glide both when diving and climbing. On average, underwater gliders displace an amount of water equal to their mass. They accomplish gliding by adjusting their volume displacement downward to dive and upward to climb. Gliders pitch themselves downward and are supported by wings to descend along a slanting path (conversely to ascend). Wings are glider propellers. They convert the vertical force of buoyancy into forward motion. The desire to obtain vertical as well as horizontal structure of the ocean is conveniently and efficiently
(c) 2011 Elsevier Inc. All Rights Reserved.
Figure 1 The operational gliders (clockwise from left): Seaglider pitched down at the sea surface with its trailing antenna raised (photo by Keith Van Thiel, University of Washington), top view of Spray at the sea surface (photo by Robert Todd, Scripps Institution of Oceanography), and rear view of Slocum climbing (photo by David Fratantoni, Woods Hole Oceanographic Institution).
GLIDERS
(c) 2011 Elsevier Inc. All Rights Reserved.
61
62
GLIDERS
provided by gliding, provided glide slopes are steep compared to slopes of oceanic property surfaces. The operational gliders variously operate along glide slopes in the range 0.25 to unity (e.g., B14–451), similar to the glide slope of a NASA space shuttle. By contrast, high-performance sail planes attain slopes as gentle as 0.02, not far from the higher slopes of property surfaces within the ocean. Sinking through the ocean is gratis, while climbing through it requires energy input. The force of ambient pressure can be used to partially collapse vehicle volume, raise its mass density over that of seawater surrounding it, and make it sink. Diving requires only the energy to open a valve to allow a swim bladder to drain into a reservoir within the hull. To stop diving and begin ascent, displacement volume must be increased, requiring work against pressure to expand a swim bladder, for example. Work is required not only to provide sufficient potential energy to bring a glider to the sea surface, but also to provide kinetic energy to fly through the ocean. Once made buoyant, gliding eventually brings a vehicle to the sea surface where it can raise an antenna to navigate and communicate. Kinetic energy provided through buoyancy forcing is ultimately dissipated by hydrodynamic drag in gliding. Since drag is generally quadratic with speed, halving speed quadruples endurance and doubles range. Underwater gliders achieve endurance and range 2 orders of magnitude higher than conventional propeller-driven AUVs by virtue of employing speeds an order of magnitude smaller. While a slow propeller-driven AUV might attain considerable endurance and range traveling horizontally, it too must work against the ocean’s potential energy gradient (gravity) to reach the sea surface to navigate or communicate electromagnetically. Slocum, Spray, and Seaglider were each designed to operate at speeds barely faster than typical ocean currents. Since currents are typically surface-intensified due to baroclinicity, deeper dives tend to reduce average current encountered by a glider, hence enhance the ability to navigate effectively. Gliders can navigate more effectively in regions where currents reverse with depth, despite being swift, than in regions of strong barotropic flow. For example, glider navigation in a swift equatorial current system or within a strait with strongly baroclinic flow is expected to be more effective than in weakly stratified coastal jets over the continental shelf. Gliders have operated across the Kuroshio and Gulf Stream by gliding with an upstream component to ferry across them, much as a canoeist paddles both upstream and across stream in order to cross a swift river. The finite spatial extent of ocean currents can be exploited to
make up for downstream drift within them by gliding upstream in weak or opposing flows on their flanks. Navigating the swirling currents of a strong eddy is also possible if glider position within an eddy is recognized. The operational gliders all carry a fixed energy supply in the form of primary batteries. Packs consisting of lithium cells provide the highest energy density, but are classified as hazardous material. Gliders using these are subject to various shipping restrictions, particularly as air freight. Nevertheless, lithium battery packs enable glider missions of half a year or more. Up to about one quarter of the total mass of these gliders can be devoted to carrying batteries, so packs of the equivalent of B100 lithium D-cells can be used, providing roughly 10 MJ of energy, enough for B1000 dive cycles using B10 kJ per cycle. The total average power consumption may be as little as 0.5 W, resulting in mission endurance of several months. Recent improvements suggest that missions longer than a year are feasible for some battery-powered gliders. A buoyancy engine powered by ocean thermal stratification, as envisioned by Stommel and Webb, is under development for Slocum, but will be restricted to regions where sufficient ocean temperature stratification is found. In principle, longer endurance and range can be attained by larger vehicles, since drag is determined by area while energy storage scales with volume. Instrument payload also scales with volume, and also expense. Practically, AUVs larger than B60 kg require power assistance for handling at sea. Both the significant economy of manual handling by a pair of operators and the smaller unit manufacture and service costs associated with small platforms are consistent with the vision expressed by Stommel’s 1989 article and also that guiding the Argo project: that of a network of a large number of relatively small, inexpensive platforms. By keeping gliders small, networks of them are tolerant to occasional vehicle loss.
Glider Dive Cycles Gliders are fully autonomous from the time they leave the sea surface until they return. At the surface, they use a GPS fix to compute a desired heading to a target (or use a preset or communicated heading), then reduce volume displacement to start diving. Retracting a piston accomplishes buoyancy reduction in the 200-m-depth rated Slocum, while in Spray and Seaglider an external bladder is deflated to reduce buoyancy by opening a valve, since subbarometric pressure is maintained inside the pressure
(c) 2011 Elsevier Inc. All Rights Reserved.
GLIDERS
hull. Glider pitch attitude is adjusted by redistributing mass within the vehicles, both by changing buoyancy and moving fixed masses such as battery packs. Gliders turn to attain the desired heading through the use of an active rudder or by redistributing mass to roll the wings. Typical turn radii of the operational gliders are tens of meters. Under water, navigation is by dead reckoning: simply following the chosen compass course. While course through the water varies due to environmental effects such as water-column shear of horizontal current as well as to vehicle control, the overall horizontal displacement due to gliding at various headings and speeds can be measured or inferred. The difference between displacement between surface GPS fixes and gliding displacement gives an estimate of depth-averaged current. The depthaveraged current for each dive cycle can be used to guide the choice of vehicle heading to approach a target through a prediction scheme, either on board or communicated from afar.
Hydrodynamics The existing operational underwater gliders attain effectively steady flight in some tens of seconds due to relatively slow acceleration relative to hydrodynamic drag. The balance of buoyancy by lift and drag forces determines the glide slope. Glider design requires sufficiently small drag that the ratio of drag to lift, equal in steady flight to glide slope, can provide the desired performance. The wings of Seaglider, Slocum, and Spray are somewhat small relative to body vehicle size in comparison to those of typical aeronautical sail planes. Both considerations of how steeply to dive through the ocean’s structure and ease of handling on launch and recovery guided wing sizing. Vehicle drag depends on both shape and surface texture, form, and skin drag, respectively. A significant component of drag is that induced by lift, typically parametrized as proportional to the square of attack angle. Considerations of drag led Spray to approximate a slender ellipsoidal form with a long cylinder capped by ellipsoidal ends and Seaglider to use a long smooth forebody (to maintain laminar boundary layer flow) and short after-section (where turbulence is tolerated). Tow tank, wind tunnel, and open-ocean experience have demonstrated that small appendages can account for disproportionately large contributions to total vehicle drag. A glider with a given set of lift, drag, and induced drag coefficients can be flown over a range of speed and glide slope, the latter limited by stall at the
63
maximum lift to drag ratio. As an example, buoyancy, power, and attack angle dependence on vertical and horizontal speed are shown in Figure 2 for a typical Seaglider. Contours of power (solid curves) reflect the nearly quadratic nature of the drag law for this vehicle: operating at speeds of B0.2 m s 1 costs B0.1 W, while operating at B0.4 m s 1 costs B0.4 W (of delivered power – the buoyancy engines are at most about 50% efficient). The maximum horizontal speed component attainable for a given power use is at a glide slope only slightly steeper than the stall glide slope, as can be seen from the departure of the power curve shape from near-circular about the origin to steeply curved to form a cone of exclusion at the stall slope. The extreme marks the most efficient glide slope for range, B0.28 for speeds of B0.25 m s 1 to B0.20 for speeds of B1 m s 1 for the example given in Figure 2. A given buoyancy force leads to a wide range of speed, depending on glide slope, as illustrated by the dashed curves in Figure 2. For example, a buoyancy of 150 g can provide a horizontal speed of 0.22 m s 1 near the glide slope of maximum efficiency, or 0.33 m s 1 at a glide slope of 0.8 (using about 4 times as much power). The maximum total speed is highest for a given buoyancy if a glider were to be able to rise perfectly vertically (requiring no wing lift), but the maximum horizontal speed is attained at slopes of 0.78–0.84 (glide angles of 30–401 above horizontal) for the particular combination of lift, drag, and induced drag coefficients in the Seaglider example. The buoyancy range plotted in Figure 2 exceeds what the Seaglider buoyancy engine is capable of providing in both positive and negative buoyancy (the end-to-end range is B850 g), but diving through a light near surface layer may produce instantaneously buoyancy as high as is plotted in the figure. Attack angles for gliders (the difference between pitch and glide angle, where the latter is always steeper) range from less than 11 for steep dive slopes to several degrees for glide slopes near stall. For an underwater flying wing, for example, glider pitch can be in the opposite sense to glide angle to produce large lift. For the Seaglider example, typical attack angles are a few degrees for the most efficient glide slope. Attack angle affects flow around the hull, amplifying it as a function of position.
Instrumentation Size, power, hydrodynamic smoothness, and mass invariance are the principal constraints on instrumentation carried on gliders. For the operational
(c) 2011 Elsevier Inc. All Rights Reserved.
64
GLIDERS
Seaglider performance model 1 0
6
750
7 700
0
650
55
0
3
300
0.5
−
45
0.9
5
4
0
0 50
3
600
40
350
2250
0.8 4
2
600 650 700 750 80 0
−1
450
2
500 550
−1
3
0.
8
−1.5
−2
40 350 0
0.3
1 9 0.
7
0.
−1
250
0.3
150
0.4
4
0.6
0.5 100 0.4
−0 .5
0
0.5
0
1 0.9
0.8 0.7
20
Vertical speed (m s−1)
150
.5
−0
400
0
5
25
300
0.6
200
0 35
0.7
3
5
0.
0.2 −2 50 −2.5 0 1.5 0 00.0.1 −3 9 00500 75800 .08 0 0 6 −2.5 5055 3670 0.2 0 2 0 . 0 0 0.05.06 7 −4 −4.5 3.5 45 0 0.0.0341 −5.5 4 3 5 5 4. −5 0.02 −2. 3.5 −78.5 5.5 −6 −6.57.5 10 0.1 0.0 9.5 4 9 8 0.9 1 700 600 650 0 3 550 0500 1 0.2 10 7 0.07 0.02 0.06 0.03 0.01 0.04 0.6 9 .5 45 0.05 98387 5 6.5 9.516350 5−5.5 00.1 .3 −4 4.5 6.5 .0 72.0 80 0.8 3.5 0.7 34.5 15 8.5 50 510 0.4 0.5 8310700 400 .5250 02 8.5 8 7.5 0.00 2.5 5.5550 6 0 00010 50 0535 09.5 10 8 5 9 .5 1 12. 00 47 24150 .2 1 0. 00.4 800 5 0.6 750 63 7 5 5 .0 0.3 450 600 0 5 300 .7 .57 9.5 93049 200 03.5 1.0 6 .0 .2858 0.08 0.09 0.1 002 100 0087.5 7 0.5 0.0 0 0. .5 0 0 0 7 0 5 7 3 .5 0 6 .5 6 0 38 7 6 4 0 8 6 5 .5 8 5 9 350 5 8. 5 0 7 1.5 0 0 50 53 .5 0 10 0.2 5 00 4 0 0 2 5 5200 10 4 01 1 .5 23.5 63 07 1 520.0 5.5 95.0 204 .0 0.0 05 0. 730 .0 40.1 0 80.0 010 0 .5 100 .59785.5 35.56.5 3994.5 79. 36 87 610 1.5 35683 8 52810 5.50.505 4200 0. 0 Buoyancy (g) 0 0.1 0.2 0.3 0.4 0.5 0.6 0.7 0.8 0.9 1 Horizontal speed (m s−1) Power (W) Attack angle (°) a = 0.002 758 9 b = 0.009 257 5 c = 9.228 2 e −06 Λ0.25 = 211.863 4 0
30
20
0
2
100
0.6
4
0.
4
−0. 5 0
0.3
Figure 2 Model Seaglider performance based on steady flight with lift, drag, and induced drag coefficients a, b, and c where lift and drag forces follow L ¼ ql 2aa and D ¼ ql 2(bq 1/4 þ c a2), where l ¼ 1.8 m is the Seaglider body length, a is attack angle, and dynamic pressure is given by q ¼ r(u2 þ w2)/2 for horizontal and vertical speed components u and w for water density r. Solid, dashed, and dotted curves contour power, buoyancy, and attack angle, respectively, as functions of horizontal and vertical speed. Contours bunch at the stall limit. Glide slope is the ratio w/u and attack angles are labeled as the difference between vehicle pitch and glide angle, referenced to the horizon (e.g., negative attack for upward glide).
gliders, sensors and their packaging are practically limited to perhaps B500-cc volume and power consumption a fraction of that used for glider propulsion. On long missions, as much as 85% of glider battery energy might be devoted to running the buoyancy engine, limiting the average power consumption available to instrumentation and operating the onboard microprocessor to less than B0.1 W. Average power of this level is achieved by lowering the duty cycle of typical instruments considerably; that is, most of the time sensors are turned off and only briefly are energized for sampling. The low-power sleep current of the microprocessor is a significant constraint on mission endurance. Oceanic fields that are readily measured electronically are well suited to gliders. Temperature, salinity, and pressure are readily measured electronically at little energy cost. In addition to scientific
interest, together with measurements of vehicle pitch, volume, and heading, these can be used to estimate glider speed through the water. Other instrumentation routinely installed on gliders include dissolved oxygen sensors, bio-optical sensors for chlorophyll fluorescence, optical backscatter, and transmissivity, and acoustic sensors for both active and passive measurements. Specially adapted acoustic Doppler current profilers have also been carried on gliders. All of these instruments involve tradeoffs between capability, sampling rate, and mission endurance.
Missions The track of a Seaglider in the Labrador Sea (Figure 3) illustrates the ability of a glider to make transects in regions where depth-averaged flow is
(c) 2011 Elsevier Inc. All Rights Reserved.
GLIDERS
65° W
60° W
55° W
50° W
65
45° W
70
0 60
GPS fix and iridium call
40
05 0
0
400
0
0 00 60 5
20 cm s−1 current 0
300
200
upper 1-km average
10
0
20
0
20
100
0
70
300
800
0
300
90
20
65° N
100
Nuuk
1000
0
Greenland 400
300
200
600
100
65° N
0 40
00 04 0 6080
2000 00
00 5070
5
30
0
0
90 100
30
00
00
10
0
30
600
0
20 00
300
60° N
20
200
60° N 00
000 000 91 60 00080 450
30
700
0
0
200
100
00
30
40
0
0
20
0 30
50
0
20
55° N
55° N 30 0
0
65° W
60° W
55° W
50° W
45° W
Figure 3 Track of a Seaglider in the Labrador Sea from 24 Sep. 2004 to 29 Apr. 2005, the longest AUV mission to date. This vehicle was launched from R/V Knorr in Davis Strait and recovered more than 7 months later offshore Nuuk, Greenland. Pink symbols mark the location of GPS fixes and Iridium calls between dive cycles. Blue arrows indicate depth-averaged currents (to the shallower of bottom depth and 1 km), where the scale length is typical glider speed through the water. Depth contour intervals are 100–1000 m, 500 m thereafter with heavy contours every 1 km.
(c) 2011 Elsevier Inc. All Rights Reserved.
66
GLIDERS
sufficiently weak, but be hampered in its ability to navigate within strong currents. Launched from a research vessel offshore west Greenland, this vehicle made two transects across Davis Strait before being sent south to Hamilton Bank offshore southeast Labrador. It was swept across part of the continental shelf region before escaping the Labrador Current to continue northeastward on its return to Greenland near Kap Desolation. On its way north toward recovery, it was caught in an anticyclonic eddy that drifted offshore. This glider was directed mainly northward as it cycled around the eddy core four times before it finally escaped by being guided radially outward across the eddy. It continued northeastward to Fyllas Bank, offshore Nuuk, Greenland, where it was recovered from a fishing vessel, having completed the longest AUV mission to date as of this writing, lasting 217 days, traveling 3750 km through the water, and making 663 dives, the majority to 1000-m depth. Several different types of glider missions have been undertaken, of which the example in Figure 3 is just one: a solitary glider survey. In other applications, gliders have been used to make repeat surveys along the same track, controlled to maintain position (a ‘virtual mooring’), used in numbers to survey a region intensively, used to interact with one another, to communicate with moored sensors, act as a courier for data recorded by these devices, and to control them.
See also Drifters and Floats. Underwater Vehicles.
Platforms:
Autonomous
Further Reading Davis RE, Eriksen CC, and Jones CP (2003) Autonomous buoyancy-driven underwater gliders. In: Griffiths G (ed.) Technology and Applications of Autonomous Underwater Vehicles, pp. 37--58. New York: Taylor and Francis. Rudnick DL, Davis RE, Eriksen CC, Fratantoni DM, and Perry MJ (2004) Underwater gliders for ocean research. Journal of the Marine Technology Society 38: 73--84. Stommel H (1989) The Slocum mission. Oceanography 2: 22--25.
Relevant Websites http://iop.apl.washington.edu – Operations Summary: Custom View, Seaglider Status, Integrative Observational Platforms, Ocean Physics Department, Applied Physics Laboratory, University of Washington. http://spray.ucsd.edu – SIO IDG Spray Home. http://www.webbresearch.com – Slocum Glider, Webb Research Corporation. http://seaglider.washington.edu – The Seaglider Fabrication Center of the University of Washington.
(c) 2011 Elsevier Inc. All Rights Reserved.
GLOBAL MARINE POLLUTION A. D. Mclntyre, University of Aberdeen, Aberdeen, UK & 2009 Elsevier Ltd. All rights reserved.
This aritcle is designed to provide an entry to the historical and global context of issues related to pollution of the oceans. Until recently, the size and mobility of the oceans encouraged the view that they could not be significantly affected by human activities. Freshwater lakes and rivers had been degraded for centuries by effluents, particularly sewage, but, although from the 1920s coastal oil pollution from shipping discharges was widespread, it was felt that in general the open sea could safely dilute and disperse anything added to it. Erosion of this view began in the 1950s, when fallout from the testing of nuclear weapons in the atmosphere resulted in enhanced levels of artificial radionuclides throughout the world’s oceans. At about the same time, the effluent from a factory at Minamata in Japan caused illness and deaths in the human population from consumption of mercurycontaminated fish, focusing global attention on the potential dangers of toxic metals. In the early 1960s, buildup in the marine environment of residues from synthetic organic pesticides poisoned top predators such as fish-eating birds, and in 1967 the first wreck of a supertanker, the Torrey Canyon, highlighted the threat of oil from shipping accidents, as distinct from operational discharges. It might therefore be said that the decades of the 1950s and 1960s saw the beginnings of marine pollution as a serious concern, and one that demanded widespread control. It attracted the efforts of national and international agencies, not least those of the United Nations. The fear of effects of radioactivity focused early attention, and initiated the establishment of the International Commission on Radiological Protection (ICRP), which produced a set of radiation protection standards, applicable not just to fallout from weapons testing but also to the increasingly more relevant issues of operational discharges from nuclear reactors and reprocessing plants, from disposal of low-level radioactive material from a variety of sources including research and medicine, and from accidents in industrial installations and nuclear-powered ships. Following Minamata, other metals, in particular cadmium and lead, joined mercury on the list of concerns. Since this was seen, like radioactive wastes, as a public health problem, immediate action was taken. Metals
in seafoods were monitored and import regulations were put in place. As a result, metal toxicity in seafoods is no longer a major issue, and since most marine organisms are resilient to metals, this form of pollution affects ecosystems only when metals are in very high concentrations, such as where mine tailings reach the sea. Synthetic organics, either as pesticides and antifoulants (notably tributyl tin (TBT)) or as industrial chemicals, are present in seawater, biota, and sediments, and affect the whole spectrum of marine life, from primary producers to mammals and birds. The more persistent and toxic compounds are now banned or restricted, but since many are resistant to degradation and tend to attach to particles, the seabed sediment acts as a sink, from which they may be recirculated into the water column. Other synthetic compounds include plastics, and the increasing use of these has brought new problems to wildlife and amenities. Oil contaminates the marine environment mainly from shipping and offshore oil production activities. Major incidents can release large quantities of oil over short periods, causing immense local damage; but in the longer term, more oil reaches the sea via operational discharges from ships. These and other threats to the ocean are controlled by the International Convention for the Prevention of Pollution from Ships, administered by the International Maritime Organisation of the UN. Pollution is generated by human activities, and the most ubiquitous item is sewage, which is derived from a variety of sources: as a direct discharge; as a component of urban wastewater; or as sludge to be disposed of after treatment. Sewage in coastal waters is primarily a public health problem, exposing recreational users to pathogens from the local population. The dangers are widely recognized, and many countries have introduced protective legislation. At the global level, the London Dumping Convention controls, among other things, the disposal of sewage sludge. As well as introducing pathogens, sewage also contributes carbon and nutrients to the sea, adding to the substantial quantities of these substances reaching the marine environment from agricultural runoff and industrial effluents. The resulting eutrophication is causing major ecosystem impacts around the world, resulting in excessive, and sometimes harmful, algal blooms. The need for a global approach to ocean affairs was formally brought to the attention of the United
(c) 2011 Elsevier Inc. All Rights Reserved.
67
68
GLOBAL MARINE POLLUTION
Nations in 1967, and over the next 15 years, while sectoral treaties and agreements were being introduced, negotiations for a comprehensive regime led in 1982 to the adoption of the UN Convention on the Law of the Sea (UNCLOS). This provided a framework for the protection and management of marine resources. In the ensuing years, concepts such as sustainable development and the precautionary approach came to the fore and were endorsed in the Declaration of the Rio Summit, while proposals for Integrated Coastal Zone Management (ICZM) are being widely explored nationally and advanced through the Intergovernmental Oceanographic Commission. Early ideas of pollution were focused on chemical inputs, but following the Group of Experts on the Scientific Aspects of Marine Pollution (GESAMP) definition pollution is now seen in a much wider context, encompassing any human activity that damages habitats and amenities and interferes with legitimate users of the sea. Thus, manipulation of terrestrial hydrological cycles, and other hinterland activities including alterations in agriculture, or afforestation, can profoundly influence estuarine regimes, and are seen in the context of pollution. In particular, it is now recognized that excessive fishing can do more widespread damage to marine ecosystems than most chemical pollution, and the need for ecosystem-based fisheries management is widely recognized. Over the years, assessments of pollution effects have altered the priority of concerns, which today are very different from those of the 1950s. Thanks to the rigorous control of radioactivity and metals, these are not now major worries. Also, decades of experience with oil spills have shown that, after the initial damage, oil degrades and the resilience of natural communities leads to their recovery. Today, while the effects of sewage, eutrophication, and harmful algal blooms top the list of pollution concerns, along with the physical destruction of habitats by coastal construction, another item has been added: aquatic invasive species. The particular focus is on the transfer of harmful organisms in ships’ ballast water and sediments, which can cause disruption of fisheries, fouling of coastal industry, and reduction of human amenity. In conclusion, most of the impacts referred to above are on the shallow waters and the shelf, associated with continental inputs and activities. The open ocean, although subject to contamination from the atmosphere and from vessels in shipping lanes, is relatively less polluted United Nations Environment Programme (UNEP).
See also Anthropogenic Trace Elements in the Ocean. Antifouling Materials. Atmospheric Input of Pollutants. Chlorinated Hydrocarbons. Eutrophication. Exotic Species, Introduction of. International Organizations. Marine Policy Overview. Metal Pollution. Oil Pollution. Phytoplankton Blooms. Pollution: Approaches to Pollution Control. Pollution: Effects on Marine Communities. Pollution, Solids. Radioactive Wastes. Thermal Discharges and Pollution. Viral and Bacterial Contamination of Beaches.
Further Reading Brackley P (ed.) (1990) World Guide to Environmental Issues and Organisations. Harlow: Longman. Brune D, Chapman DV, and Gwynne DW (1997) Eutrophication. In: The Global Environment, ch. 30. Weinheim: VCH. Coe JM and Rogers DB (1997) Marine Debris. Berlin: Springer. de Mora SJ (ed.) (1996) Tributyltin Case Study of an Environmental Contaminant. Cambridge, UK: Cambridge University Press. GESAMP (1982) Reports and Studies No. 15: The Review of the Health of the Oceans. Paris: UNESCO. Grubb M, Koch M, and Thomson K (1993) The ‘Earth Summit’ Agreements: A Guide and Assessment. London: Earth Scan Publications. HMSO (1981) Eighth Report of the Royal Commission on Environmental Pollution, Oil Pollution of the Sea, Cmnd 8358. London: HMSO. Hollingworth C (2000) Ecosystem effects of fishing. ICES Journal of Marine Science 57(3): 465--465(1). International Maritime Organization (1991) IMO MARPOL 73/78 Consolidated Edition MARPOL 73/78. London: International Maritime Organization. International Oceanographic Commission (1998) Annual Report of the International Oceanographic Commission. Paris: UNESCO. Kutsuna M (ed.) (1986) Minamata Disease. Kunamoto: Kunammoyo University. Matheickal J and Raaymakers S (eds.) (2004) Second International Ballast Water Treatment R & D Symposium, 21–23 July 2003: Proceedings. GloBallast Monograph Series No. 15. London: International Maritime Organization. National Research Council (1985) Oil in the Sea. Washington, DC: National Academies Press. Park PK, Kester DR, and Duedall IW (eds.) (1983) Radioactive Wastes and the Ocean. New York: Wiley. Pravdic V (1981) GESAMP the First Dozen Years. Nairobi: UNEP. Pritchard SZ (1987) Oil Pollution Control. Wolfeboro, NH: Croom Helm. Sinclair M and Valdimarsson G (eds.) (2003) Responsible Fisheries in the Marine Ecosystem. Oxford, UK: CABI.
(c) 2011 Elsevier Inc. All Rights Reserved.
GLOBAL MARINE POLLUTION
Tolba MK, El-Kholy OA, and El-Hinnawi E (1992) The World Environment 1972–1992: Two Decades of Challenge. London: Chapman and Hall. UNEP (1990) The State of the Marine Environment. UNEP Regional Seas Reports and Studies No. 115. Nairobi: UNEP.
69
UNO (1983) The Law of the Sea: Official Text of UNCLOS. New York: United Nations. World Commission on Environment and Development (1987) Our Common Future. Oxford, UK: Oxford University Press.
(c) 2011 Elsevier Inc. All Rights Reserved.
GRABS FOR SHELF BENTHIC SAMPLING P. F. Kingston, Heriot-Watt University, Edinburgh, UK & 2009 Elsevier Ltd. All rights reserved.
Introduction The sedimentary environment is theoretically one of the easiest to sample quantitatively and one of the most convenient ways to secure such samples is by means of grabs. Grab samplers are used for both faunal samples, when the grab contents are retained in their entirety and then sieved to remove the biota from the sediment, and for chemical/physical samples when a subsample is usually taken from the surface of the sediment obtained. In both cases, the sampling program is reliant on the grab sampler taking consistent and relatively undisturbed sediment samples.
Conventional Grab Samplers The forerunner of the grab samplers used today is the Petersen grab, designed by C.G.J. Petersen to conduct benthic faunal investigations in Danish fiords in the early part of the twentieth century. It consisted of two quadrant buckets that were held in an open
Figure 1 Petersen grab.
70
position and lowered to the seabed (Figure 1). On the bottom, the relaxing of the tension on the lowering warp released the buckets and subsequent hauling caused them to close before they left the bottom. The instrument is still used today but is seriously limited in its range of usefulness, working efficiently only in very soft mud. Petersen’s grab formed the basis for the design of many that came after. One enduring example is the van Veen grab, a sampler that is in common use today (Figure 2). The main improvement over Petersen’s design is the provision of long arms attached to the buckets to provide additional leverage to the closing action. The arms also provided a means by which the complex closing mechanism of the Petersen grab could be simplified with the hauling warp being attached to chains on the ends of the arms. The mechanical advantage of the long arms can be improved further by using an endless warp rig; this has the added advantage of helping to prevent the grab being jerked off the bottom if the ship rolls as the grab is closing. The van Veen grab was designed in 1933 and is still widely used in benthic infaunal studies owing to its simple design, robustness, and digging efficiency. The van Veen grab
Figure 2 Van Veen grab.
(c) 2011 Elsevier Inc. All Rights Reserved.
GRABS FOR SHELF BENTHIC SAMPLING
71
Figure 3 Diagram of a Hunter grab. Reproduced from Hunter B and Simpson AE (1976) A benthic grab designed for easy operation and durability. Journal of the Marine Biological Association 56: 951–957.
typically covers a surface area of 0.1 m2, although instruments of twice this size are sometimes used. A more recent design of frameless grab is the Hunter grab (Figure 3). This is of a more compact design than the van Veen. The jaws are closed by levers attached to the buckets in a parallelogram configuration, giving the mechanism a good overall mechanical advantage. The closing action requires no chains or pulleys and the instrument can be operated by one person. Its disadvantage is that the bucket design does not encourage good initial penetration of the sediment, which is important in hardpacked sediments. A disadvantage of the grab samplers discussed so far is that there is little latitude for horizontal movement of the ship while the sample is being secured: the smallest amount of drift and the sampler is likely to be pulled over. The Smith–McIntyre grab was designed to reduce this problem by mounting the grab buckets in a stabilizing frame (Figure 4). Initial penetration of the leading edge of the buckets is assisted by the use of powerful springs and the buckets are closed by cables pulling on attached short arms in a similar way to that on the van Veen grab. The driving springs are released by two trigger plates, one on either side of the supporting frame to ensure that the sampler is resting flat on the seabed before the sample is taken. In firm sand the Smith–McIntyre grab penetrates to about the same depth of sediment as the van Veen grab. Its main disadvantage is the need to cock the spring mechanism on deck before deploying the sampler, a process that can be quite hazardous in rough weather. The Day grab is a simplified form of the Smith– McIntyre instrument in which the trigger and closing mechanism remains the same, but without spring assistance for initial penetration of the buckets
Figure 4 Smith–McIntyre grab.
Figure 5 Day grab.
(Figure 5). The Day grab is widely used, particularly for monitoring work, despite its poor performance in hard-packed sandy sediments. Most of the grabs thus far discussed have been designed to take samples with a surface area of 0.1 or 0.2 m2. The Baird grab, however, takes samples of 0.5 m2 by means of two inclined digging plates that are pulled together by tension on the warp (Figure 6). The grab is useful where a relatively large surface area needs to be covered, but has the disadvantage of taking a shallow bite and having the surface of the sample exposed while it is being hauled in.
(c) 2011 Elsevier Inc. All Rights Reserved.
72
GRABS FOR SHELF BENTHIC SAMPLING
Warp Activation All the grabs described above use the warp acting against the weight of the sampler to close the jaws. However, direct contact with the vessel on the surface during the closure of the grab mechanism poses several problems.
Warp Heave
As tension is taken up by the warp to close the jaws, there is a tendency for the grab to be pulled up off the bottom, resulting in a shallower bite than might be expected from the geometry of the sampler. This tendency is related to the total weight of the sampler and the speed of hauling and is exacerbated by firm sediments. For example, the theoretical maximum depth of bite of a 120-kg Day grab is 13 cm (based
on direct measurements of the sampler); however, in medium sand, the digging performance is reduced to a maximum depth of only 8 cm (Figure 7(e)). The influence of warp action on the digging efficiency of a grab sampler can also depend on the way in which the sampler is rigged. This is particularly true of the van Veen grab. Figure 7(b) shows the bite profile of the chain-rigged sampler in which the end of each arm is directly connected to the warp by a chain. The vertical sides of the profile represent the initial penetration of the grab while the central rise shows the upward movement of the grab as the jaws close. Figure 7(c) shows the bite profile of a van Veen of similar size and weight (30 kg) rigged with an endless warp in which the arms are closed by a loop of wire passing through a block at the end of each arm (as in Figure 2). The vertical profile of the initial penetration is again apparent; however, in this case, the overall depth of the sampler in the sediment is maintained as the jaws close. The endless warp rig increases the mechanical advantage of the pull of the warp while decreasing the speed at which the jaws are closed. The result is that the sampler is ‘insulated’ from surface conditions to a greater extent than when chain-rigged, giving a better digging efficiency. Grab ‘Bounce’
In calm sea conditions it is relatively easy to control the rate of warp heave and obtain at least some consistency in the volume of sediment secured. However, such conditions are seldom experienced in the open sea where it is more usual to encounter wave action. Few ships used in offshore benthic
Figure 6 Baird grab.
(a)
(b)
(c)
2 4 6 8 10
2 4 6 8 10
2 4 6 8 10
Petersen grab (30 kg)
Chain-rigged van Veen grab (30 kg)
Endless-warp-rigged van Veen grab (short-armed, 30 kg)
(d)
(e)
(f)
2 4 6 8 10 12 14
2 4 6 8 10
2 4 6 8 10
Endless-warp-rigged van Veen grab (long-armed, 70 kg)
Day grab (120 kg)
Smith−mcintyre grab (120 kg)
Figure 7 Digging profiles of a range of commonly used benthic grab samplers obtained in a test tank using a fine sand substratum. (a) Peterson grab (30 kg); (b) chain-rigged van Veen grab (30 kg); (c) endless-warp-rigged van Veen grab (short-armed, 30 kg); (d) endless-warp-rigged van Veen grab (long-armed, 70 kg); (e) Day grab (120 kg); (f) Smith–McIntyre grab (120 kg).
(c) 2011 Elsevier Inc. All Rights Reserved.
GRABS FOR SHELF BENTHIC SAMPLING
studies are fitted with winches with heave compensators so that the effect of ship’s roll is to introduce an erratic motion to the warp. This may result in the grab ‘bouncing’ off the bottom where the ship rises just as bottom contact is made, or in the grab being snatched off the bottom where the ship rises just as hauling commences. In the former instance, it is unlikely that any sediment is secured; in the latter, the amount of material and its integrity as a sample will vary considerably, depending on the exact circumstances of its retrieval. The intensity of this effect will depend on the severity of the weather conditions. Figure 8 shows the relationship between wind speed and grab failure rate, which is over 60% of hauls at wind force 8. What is of more concern to the scientist attempting to obtain quantitative samples is the dramatic increase in variability with increase in wind speed with a coefficient of variation between 20 and 30 at force 7. The high cost of ship-time places considerable pressure on operators to work in as severe weather conditions as possible and it is not unusual for sampling to continue in wind force 7 conditions with all its disadvantages. Drift
For a warp-activated grab sampler to operate efficiently it should be hauled with the warp positioned vertically above. Where there is a strong wind or current, these conditions may be difficult to achieve. The result is that the grab samplers are pulled on to their sides. This is a particular problem with samplers,
73
such as the van Veen grab, that do not have stabilizing frames. Diver observations have shown, however, that at least in shallow water, where the drift effect is at its greatest on the bottom, even the framed heavily weighted Day and Smith–McIntyre grabs can be toppled. Initial Penetration
It is clear that the weight of the sampler is an important element in determining the volume of the sample secured. Much of the improved digging efficiency of the van Veen grab shown in Figure 7(d) can be attributed to the addition of an extra 40 kg of weight which increased the initial penetration of the sampler on contact with the sediment surface. Initial penetration is one of the most important factors in the sequence of events in grab operation, determining the final volume of sediment secured. Figure 9 shows the relationship between initial penetration and final sample volume obtained for a van Veen grab. Over 70% of the final volume is determined by the initial penetration. Subsequent digging of the sampler is hampered, as already shown, by the pull of the warp. For most benthic faunal studies it is important for the sampler to penetrate at least 5 cm into the sediment (for a 0.1 m2 surface area sample this gives 5 l of sample). In terms of number of species and individuals, over 90% of benthic macrofauna are found in the top 4–5 cm of sediment. Figure 10 shows how the number of individuals relates to average sample 20 18
70
16 Initial penetration depth (cm)
Number of failures due to motion as the % of number of attempts
80
60 50 40 30 20
14 12 10 8 6 4
10
2
0
0 0
1
2
3
4
5
6
7
8
0
2
4
Sea state (Beaufort scale) Figure 8 Relationship between wind speed and grab failure rate.
6
8
10 12 14 16 18 20 22 24
Sample volume (l) Figure 9 Relationship between initial penetration of a van Veen grab sampler and volume of sediment secured.
(c) 2011 Elsevier Inc. All Rights Reserved.
74
GRABS FOR SHELF BENTHIC SAMPLING
1500
Number of individuals per 0.5 m2
1400 1300 1200 1100 1000 900 800 700 600
Figure 11 Shipek grab.
500 4.5
5
5.5
6
6.5
7
7.5
Average volume of grab sample (l) Figure 10 Relationship between number of benthic fauna individuals captured and sample volume for a boreal offshore sand substratum.
volume for 18 stations in the southern North Sea (each liter recorded represents 1 cm of penetration). Although there is a considerable variation in the numbers of individuals between stations, there is no significant trend linking increased abundance with increased sample volume penetration. No sample volumes of less than 4.5 l were taken, indicating that at that level of penetration most of the fauna were being captured. Samplers in which the jaws are held rigidly in a frame have no initial penetration if the edges of the jaw buckets, when held in the open position, are on a level with the base of the frame. The lack of any initial penetration in such instruments has the added disadvantage in benthic fauna work of undersampling at the edges of the bite profile (see Figures 7(e) and 7(f)), although the addition of weight will usually increase the sample volume obtained. Pressure Wave Effect
The descent of the grab necessarily creates a bow wave. Under field conditions, it is usually impracticable to lower the grab at a rate that will eliminate a preceding bow wave, even if the sea were flat calm. There have been several investigations of the effects of ‘downwash’ both theoretical, using artificially placed surface objects, and in situ. The effects of downwash can be reduced by replacing the upper surface of the buckets with an open mesh. Although there is still a considerable effect on the surface flock layer (rendering the samples of dubious
value for chemical contamination studies), the effect on the numbers of benthic fauna is generally very small. Self-Activated Bottom Samplers
There can be little doubt that one of the most important factors responsible for sampler failure or sample variability in heavy seas is the reliance of most presently used instruments on warp-activated closure. The most immediate and obvious answer to this problem is to make the closing action independent of the warp by incorporating a selfpowering mechanism. Spring-Powered Samplers
One solution to the problem is to use a spring to actuate the sampler buckets. Such instruments are in existence, possibly the most widely used being the Shipek grab, a small sampler (0.04 m2) consisting of a spring-loaded scoop (Figure 11). This instrument is widely used where small superficial sediment samples are required for physical or chemical analysis. The use of a pretensioned spring unfortunately sets practical limits on the size of the sampler, since to cock a spring in order to operate a sampler capable of taking a 0.1 m2 sample would require a force that would be impracticable to apply routinely on deck. In addition, in rough weather conditions, a loaded sampler of this size would be very hazardous to deploy. Compressed-Air-Powered Samplers
Another approach has been to use compressed air power. In the 1960s, Flury fitted a compressed air ram to a modified Petersen grab with success. However, the restricted depth range of the instrument and the inconvenience of having to recharge the
(c) 2011 Elsevier Inc. All Rights Reserved.
GRABS FOR SHELF BENTHIC SAMPLING
air reservoir for each haul limited its potential for routine offshore work. Piston
Hydraulically Powered Samplers
Hydraulically powered grabs are commonly used for large-scale sediment shifting operations such as seabed dredging. The Bedford Institute of Oceanography, Nova Scotia, successfully scaled down this technology to that of a practical benthic sampler. Their instrument is relatively large, standing 2.5 m high and weighing some 1136 kg. It covers a surface area of 0.5 m2 and samples to a maximum sediment depth of 25 cm. At full penetration, the sediment volume taken is about 100 l. The buckets are driven closed by hydraulic rams powered from the surface. The grab is also fitted with an underwater television camera which allows the operator to visually select the precise sampling area on the seabed, close and open the bucket remotely, and verify that the bucket closed properly prior to recovery. The top of the buckets remain open during descent to minimize the effect of downwash and close on retrieval to reduce washout of the sample on ascent. The current operating depth of the instrument is 500 m. The instrument has been successfully used on several major offshore studies, but does require the use of a substantial vessel for its deployment. Hydrostatically Powered Samplers
Hydrostatically powered samplers use the potential energy of the difference in hydrostatic pressure at the sea surface and the seabed. The idea of using this power source is not new. In the early part of the twentieth century, a ‘hydraulic engine’ was in use by marine geologists that harnessed hydrostatic pressure to drive a rock drill. Hydrostatic power has also been used to drive corers largely for geological studies. However, these instruments were principally concerned with deep sediment corers and were not designed to collect macrofauna or material at the sediment–water interface. A more recent development has been that of a grab built by Heriot-Watt University, Edinburgh. The sampler uses water pressure difference to operate a hydraulic ram that is activated when the grab reaches the seabed. Figure 12 shows the general layout of the instrument. Water enters the upper chamber of the cylinder when the sampler is on the seabed, forcing down a piston that is connected to a system of levers that close the jaws. The actuating valve is held shut by the weight of the sampler and there is a delay mechanism to prevent premature closure of the jaws resulting from ‘bounce’. Back on the ship, the jaws are held shut by an overcenter locking mechanism
75
Water inlet valve with release delay mechanism held shut by weight or sampler
Ram operated by hydrostatic pressure Connecting rod
Figure 12 Diagram of a grab sampler using hydrostatic pressure to close the jaws.
and, on release, are drawn open by reversal of the piston motion from air pressure built up on the underside of the piston during its initial power stroke. Since the powering of the grab jaws is independent of the warp, the sampler may be used successfully in a much wider range of surface weather conditions than conventional grabs.
Alternatives to Grab Samplers Ideally a benthic sediment sample for faunal studies should be straight-sided to the maximum depth of its excavation and should retain the original stratification of the sediment. Grab samplers by the very nature of their action will never achieve this end. Suction Samplers
One answer to this problem is to employ some sort of corer designed to take samples of sufficient surface area to satisfy the present approaches to benthic studies. The Knudsen sampler is such a device and is theoretically capable of taking the perfect benthic sample. It uses a suction technique to drive a core tube of 0.1 m2 cross-sectional area 30 cm into the sediment. Water is pumped out of the core tube on the seabed by a pump that is powered by unwinding a cable from a drum. The sample is retrieved by pulling the core out sideways using a wishbone arrangement and returning it to the surface bottomside up (Figures 13 and 14). Under ideal conditions, the device will take a straight-sided sample to a depth
(c) 2011 Elsevier Inc. All Rights Reserved.
76
GRABS FOR SHELF BENTHIC SAMPLING
Figure 13 Knudsen sampler in descent position.
of 30 cm. However, conditions have to be flat calm in order to allow time for the pump to operate on the seabed and evacuate the water from the core. This limits the use of the Knudsen sampler and it is generally not suitable for sampling in unsheltered conditions offshore. Mounting the sampler in a stabilizing frame can improve its success rate and it is used regularly for inshore monitoring work where it is necessary to capture deep burrowing species. Spade Box Samplers
Another approach to the problem is to drive an open-ended box into the sediment, using the weight of the sampler, and arrange for a shutter to close off the bottom end. The most widespread design of such an instrument is that of the spade box sampler, first described by Reineck in the 1950s and later subjected to various modifications. The sampler consists of a removable steel box open at both ends and driven into the sediment by its own weight. The lower end of the box is closed by a shutter supported on an arm pivoted in such a way as to cause it to slide through the sediment and across the mouth of the box (Figure 15). As with the grab samplers previously described, the shutter is driven by the act of hauling on the warp with all the attendant disadvantages. Nevertheless, box corers are very successful and are used widely for obtaining relatively undisturbed samples of up to 0.25 m2 surface area (Figure 16). One big advantage of the box sampler is that the box can usually be removed with the sample and its overlying water left intact, allowing detailed studies of the sediment surface. Furthermore, it is possible to subsample using small-diameter corers for studies of chemical and physical characteristics. Despite their potential of securing the ‘ideal’ sediment sample, box corers are rarely used for routine benthic monitoring work. This is largely because of their size (a box corer capable of taking a 0.1 m2
Figure 14 Knudsen sampler in ascent position.
Lead weights
Box closure (‘spade’) Box holder Box
Figure 15 Diagram of a Reineck spade box sampler.
(c) 2011 Elsevier Inc. All Rights Reserved.
GRABS FOR SHELF BENTHIC SAMPLING
77
Cable drum
Scoops Figure 17 Diagram of a Holme scoop. Reproduced from Holme NA (1953) The biomass of the bottom fauna in the English Channel off Plymouth. Journal of the Marine Biological Association 32: 1–49.
macrobenthos sampling. At present, there is no instrument that fulfils the requirements for a quick turnaround precision multiple corer for offshore sampling.
Sampling Difficult Sediments
Figure 16 A 0.25 m2 spade box sampler.
sample weighs over 750 kg and stands 2 m high) and the difficulty in deployment and recovery in heavy seas. Precision Corers
For chemical monitoring, it is important that the sediment–water interface is maintained intact, for it is the surface flock layer that will contain the most recently deposited material. Unfortunately, such undisturbed samples are rarely obtained using grab samplers or box corers. Precision corers are capable of securing undisturbed surface sediment cores; however, they are unsuitable for routine offshore work because of the time taken to secure a sample on the seabed and dependence on warp-activated closure. Additionally, the cross-sectional area of the core (0.002–0.004 m2) would necessitate the taking of large numbers of replicate samples in order to capture sufficient numbers of benthic macrofauna to be useful. This would be impracticable given the time taken to take a single sample. Large multiple precision corers have been constructed; these are usually too large and difficult to deploy for routine
Most of the samplers so far discussed operate reasonably well in mud or sand substrata. Few operate satisfactorily in gravel or stony mixed ground either because the bottom is too hard for the sampler to penetrate the substratum or because of the increased likelihood of a stone holding the jaws open when they are drawn together. To get around this problem various types of scoops have been devised. The Holme grab has a double scoop action with two buckets rotating in opposite directions to minimize any lateral movement during digging. The scoops are closed by means of a cable and pulley arrangement (Figure 17) and simultaneously take two samples of 0.05 m2 surface area. The Hamon grab, which has proved to be very effective in coarse, loose sediments, takes a single rectangular scoop of the substratum covering a surface area of about 0.29 m2. The scoop is forced into the sediment by a long lever driven by pulleys that are powered by the pull of the warp (Figure 18). Although the samples may not always be as consistent as those from a more conventional grab sampler, the Hamon grab has found widespread use where regular sampling on rough ground is impossible by any other means.
(c) 2011 Elsevier Inc. All Rights Reserved.
78
GRABS FOR SHELF BENTHIC SAMPLING
A generic disadvantage of most samplers presently in use is that they rely on slackening of the warp to trigger the action and the heave of the warp to drive the closing mechanism. In calm conditions, this presents no great problem, but with increasing sea state the vertical movement of the warp decreases reliability dramatically until the variability and finally failure rate of hauls make further sampling effort fruitless.
Specific Problems, Requirements, and Future Developments Scoop
Release Stop plate hook Figure 18 Diagram of a Hamon scoop.
Present State of Technology It is perhaps surprising that given the high state of technology of survey vessels, position fixing, and analytical equipment, the most commonly used samplers are relatively primitive (being designed some 40 or more years ago). Yet the quality of the sample is of fundamental importance to any research or monitoring work. Currently the most popular instruments are grab samplers, probably because of their wide operational weather window and apparent reliability. Samplers such as the Day, van Veen, and Smith–McIntyre grab samplers are still routinely used for sampling sediment for chemical and biological analysis despite their well-documented shortcomings. As discussed earlier, the most important of these are the substantial downwash that precedes the sampler as it descends and the disturbance of the trapped sediment layers by the closing action of the jaws. Both chemical and meiofaunal studies are particularly vulnerable to these. Although box corers go some way to reducing disturbance of the sediment strata, the all-important surface flocculent layer is invariably washed away. A big disadvantage of the box corer for routine offshore work is that it is sensitive to weather conditions; in addition, the larger instruments do not perform well on sand substrata.
Chemical Studies
Studies involving sediment chemistry require precision sampling if undisturbed samples at the sediment/water interface are to be obtained. This can be critical, particularly when the results of recent sedimentation are of interest. The impracticality of using existing hydraulically damped corers such as the Craib corer for offshore work has led to the widespread use of less ‘weather sensitive’ devices such as spade box corers and grabs for routine monitoring purposes. However, studies carried out have shown that these samplers produce a considerable ‘downwash effect’, blowing the surface flocculent layer away before the sample is secured. This can have serious consequences if any meaningful estimation of the surface chemistry of the sediment is desired. Repetitive and accurate sampling is also a prerequisite for determining spatial and temporal change in sediment chemistry.
Meiofauna Studies Meiofauna has increasingly been shown to have potential as an important tool in benthic monitoring. One of the major factors limiting its wider adoption is the lack of a suitable sampler. Although instruments such as the Craib corer and its multicorer derivatives are capable of sampling the critically important superficial sediment layer, these designs provide a poor level of success on harder sediments and in anything but near-perfect weather conditions. They also have a slow turnaround time. Box corers are widely used as an alternative; however, they are known to be unreliable in their sampling of meiofauna. As with the macrobenthos, meiobenthic patchiness results in low levels of precision of abundance estimates unless large numbers of samples are taken.
(c) 2011 Elsevier Inc. All Rights Reserved.
GRABS FOR SHELF BENTHIC SAMPLING
Macrobenthic Studies
The measurement and prediction of spatial and temporal variation in natural populations are of great importance to population biologists, both for fundamental research into population dynamics and productivity and in the characterization of benthic communities for determining change induced by environmental impact. Though always an important consideration, cost-effectiveness of sampling and sample processing is not so crucial in fundamental research, since time and funding may be tailored to fit objectives. This is rarely the case in routine monitoring work where often resolution and timescales have to fit the resources available. Benthic fauna are contagiously distributed and to sample such communities with a precision that will enable distinction between temporal variation and incipient change resulting from pollution effects, it is generally accepted that five replicate 0.1 m2 hauls from each station are necessary (giving a precision at which the standard error is no more than 20% of the mean). This frequency of sampling requires approximately 10–15 man-days of sediment faunal analysis per sample station. While this may be acceptable in community structure studies in which time and manpower (and thus cost) are not a primary consideration, this high cost of analyzing samples is of importance in routine monitoring studies and has led monitoring agencies and offshore operators to reduce sampling frequency on cost grounds to as few as two replicates per station. This reduces the precision with which faunal abundance can be estimated to a level at which only gross change can be demonstrated. However, sampling to an acceptable precision may be achieved from an area equivalent to 0.1 m2 if a smaller sampling unit is used. For example, 50 5-cm core samples (with a similar total surface area) have been shown to give a similar precision to that of 5 to 12 0.1-m2 grab samples. Thus a similar degree of precision may be obtained for around one-fifth to one-twelfth the analytical costs using a conventional approach. The problem is to be able to secure the 50 core samples per site that would be needed for the macrofaunal monitoring in a timescale that would be realistic offshore. At present, there is no instrument capable of supporting such a sampling demand and operating in the range of sediment types that wide-
79
scale monitoring studies demand. Clearly a single core sampler would be impracticable, and one must look to the future development of a multiple corer that is capable of flexibility in its operation, which allows a quick on-deck turnaround between hauls.
See also Benthic Boundary Organisms Overview.
Layer
Effects.
Benthic
Further Reading Ankar S (1977) Digging profile and penetration of the van Veen grab in different sediment types. Contributions from the Asko¨ Laboratory, University of Stockholm, Sweden 16: 1--12. Beukema JJ (1974) The efficiency of the van Veen grab compared with the Reineck box sampler. Journal du Conseil Permanent International pour l’Exploration de la Mer 35: 319--327. Eleftheriou A and McIntyre AD (eds.) (2005) Methods for the Study of the Marine Benthos. Oxford, UK: Blackwell. Flury JA (1967) Modified Petersen grab. Journal of the Fisheries Research Board of Canada 20: 1549--1550. Holme NA (1953) The biomass of the bottom fauna in the English Channel off Plymouth. Journal of the Marine Biological Association 32: 1--49. Hunter B and Simpson AE (1976) A benthic grab designed for easy operation and durability. Journal of the Marine Biological Association 56: 951--957. Riddle MJ (1988) Bite profiles of some benthic grab samplers. Estuarine, Coastal and Shelf Science 29(3): 285--292. Thorsen G (1957) Sampling the benthos. In: Hedgepeth JW (ed.) Treatise on Marine Ecology and Paleoecology, Vol. 1: Ecology. Washington, DC: The Geological Society of America.
(c) 2011 Elsevier Inc. All Rights Reserved.
GRAVITY M. McNutt, MBARI, Moss Landing, CA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1177–1185, & 2001, Elsevier Ltd.
Introduction The gravity field varies over the oceans on account of lateral variations in density beneath the ocean surface. The most prominent anomalies arise from undulations on density interfaces, such as occur at the water–rock interface at the seafloor or at the crust– mantle interface, also known as the Moho discontinuity. Because marine gravity is relatively easy to measure, it serves as a remote sensing tool for exploring the earth beneath the oceans. The interpretation of marine gravity anomalies in terms of the Earth’s structure is highly nonunique, however, and thus requires simultaneous consideration of other geophysically observed quantities. The most useful auxiliary measurements include depth of the ocean from echo sounders, the shape of buried reflectors from marine seismic reflection data, and/or the density of ocean rocks as determined from dredge samples or inferred from seismic velocities. Depending on the spatial wavelength of the observed variation in the gravity field, marine gravity observations are applied to the solution of a number of important problems in earth structure and dynamics. At the very longest wavelengths of 1000 to 10 000 km, the marine gravity field is usually combined with anomalies over land to infer the dynamics of the entire planet. At medium wavelengths of several tens to hundreds of kilometers, the gravity field contains important information on the thermal and mechanical properties of the lithospheric plates and on the thickness of their sedimentary cover. At even shorter wavelengths, the field reflects local irregularities in density, such as produced by seafloor bathymetric features, magma chambers, and buried ore bodies. On account of the large number of potential contributions to the marine gravity field, modern methods of analysis include spectral filtering to remove signals outside of the waveband of interest and interpretation within the context of models that obey the laws of phyics.
Units Gravity is an acceleration. The acceleration of gravity on the earth’s surface is about 9.81 m s2.
80
Gravity anomalies (observed gravitational acceleration minus an expected value) are typically much smaller, about 0.5% of the total field. The SI-compatible unit for gravity anomaly is the gravity unit (gu): 1 gu ¼ 106 m s2. However, the older c.g.s unit for gravity anomaly, the milligal (mGal), is still in very wide use: 1 mGal ¼ 10 gu. Typical small-scale variations in gravity over the ocean range from a few tens to a few hundreds of gravity units. Lateral variations in gravitational acceleration (gravity gradients) are measured in Eo¨tvo¨s units (E): 1 E ¼ 109 s2. Another quantity useful in gravity interpretation is the density of earth materials, measured in kg m3. In the marine realm, relevant densities range from about 1000 kg m3 for water to more than 3300 kg m3 for mantle rocks. A close relative of the marine gravity field is the marine geoid. The geoid, measured in units of height, is the elevation of the sea level equipotential surface. Geoid anomalies are measured in meters and are the departure of the true equipotential surface from that predicted for an idealized spheroidal Earth whose density structure varies only with radius. Geoid anomalies range from 0 to more than 7100 m. The direction of the force of gravity is everywhere perpendicular to the geoid surface, and the magnitude of the gravitational attraction is the vertical derivative of the geopotential U (eqn [1]). g¼
@U @z
½1
Geoid height N is related to the same equipotential U via Brun’s formula eqn ([2]). N¼
U g0
½2
in which g0 is the acceleration of gravity on the spheroid. For a ship sailing on the sea surface (the equipotential), it is easier to measure gravity. From a satellite in free-fall orbit high over the Earth’s surface, radar altimeters can measure with centimeter precision variations in the elevation of sea level, an excellent approximation to the true geoid that would follow the surface of a motionless ocean. Regardless of whether geoid or gravity is the quantity measured directly, simple formulas in the wavenumber domain allow gravity to be computed from geoid and vice versa. Given the same equipotential, the gravity representation emphasizes the power in
(c) 2011 Elsevier Inc. All Rights Reserved.
GRAVITY
the high-frequency (short-wavelength) part of the spectrum, whereas the geoid representation emphasizes the longer wavelengths. Therefore, for investigations of high-frequency phenomena, gravity is generally the quantity interpreted even if geoid is what was measured. The opposite is true for longwavelength phenomena.
Measurement of Marine Gravity Marine gravity measurements can be and have been acquired with several different sorts of sensors and from a variety of platforms, including ships, submarines, airplanes, and satellites. The ideal combination of sensor system and platform depends upon the needed accuracy, spatial coverage, and available time and funds. Gravimeters
The design for most marine gravimeters is borrowed from their terrestrial counterparts and are either absolute or relative in their measurements. Absolute gravimeters measure the full acceleration of gravity g at the survey site along the direction of the local vertical. Modern marine absolute gravimeters measure precisely the vertical position z of a falling mass (e.g., a corner cube reflector) as a function of time t in a vacuum cylinder using laser interferometry and an atomic clock. The acceleration is then calculated as the second derivative of the position of the falling mass as a function of time (eqn [3]). g¼
d2 zðtÞ dt2
½3
Absolute gravimeters tend to be larger, more difficult to deploy, costlier to build, and more expensive to run than relative gravimeters, and thus are only used when relative gravimeters are inadequate for the problem being addressed. Most gravity measurements at sea are relative measurements, Dg: the instrument measures the difference between gravity at the study site and at another site where absolute gravity is known (e.g., the dock where the expedition originates). Modern relative gravimeters are based on Hooke’s law for the force F required to extend a spring a distance x (eqn [4]), where ks is the spring constant, calculated by extending the spring under a known force. F ¼ ks x ¼ mg
81
at other locations can be calculated by observing how much more or less that same spring is stretched at other locations by the same mass. Although such systems are relatively inexpensive to build and easy to deploy, they suffer from drift: in effect, the spring constant changes with time because no physical spring is perfectly elastic. To first order, the drift can be corrected by returning to the same or another base station with the same instrument, and assuming that the drift was linear with time in between. The accuracy of this linear drift assumption improves with more frequent visits to the base station, but this is usually impractical for marine surveys. Through clever design, the latest generation of marine gravimeters has greatly reduced the drift problem as compared with earlier instruments. The measurement of the gravity gradient tensor was widely used early in the twentieth century for oil exploration, but fell into disfavor in the 1930s as scalar gravimeters became more reliable and easy to use. Gravity gradiometry at sea is currently making a comeback as the result of declassification of military gradiometer technology developed for use in submarines during the Cold War. Gravity gradiometers measure the three-dimensional gradient in the gravity vector using six pairs of aligned gravimeters, with accuracies reaching better than 1 Eo¨tvo¨s. In comparison with measurements of gravity, the gravity gradient has more sensitivity to variations at short wavelenghts (B5 km or less), making it useful for delimiting shallow structures buried beneath the seafloor. Geoid anomalies can be directly measured from orbiting satellites carrying radar altimeters. The altimeters measure the travel time of a radar pulse from the satellite to the ocean surface, from which it is reflected and bounced back to the satellite. Tracking stations on Earth solve for the position of the satellite with respect to the center of the Earth. These two types of information are then combined to calculate the height of the sea level equipotential surface above the center of the Earth. Because the solid land surface does not follow an equipotential, altimeters cannot be used to constrain the terrestrial geoid. Furthermore, it is difficult to extract geoid from ocean areas covered by sea ice. However, in the near future, laser altimeters deployed from satellites hold the promise of extracting geoid information even over ocean surfaces marred with sea ice, on account of their enhanced resolution.
½4
If a mass m is suspended from this spring at a site where gravity g is known (e.g., by deploying an absolute gravimeter at that base station), then gravity
Platforms
Marine gravity data can be acquired either from moving or from stationary platforms. Because the
(c) 2011 Elsevier Inc. All Rights Reserved.
82
GRAVITY
gravity field from variations in the depth of the sea floor is such a large component of the observed signal, most marine gravity surveys have relied on ships or submarines that enable the simultaneous acquisition of depth observations. However, airborne gravity measurements have been acquired successfully over ice-covered areas of the polar oceans, and orbiting satellites have measured the marine geoid from space. A major challenge in acquiring gravity data from a floating platform at the sea surface is in separating the acceleration of the platform in the dynamic ocean from the acceleration of gravity. This problem is overcome by mounting marine gravimeters deployed from ships on inertially stabilized tables. These tables employ gyroscopes to maintain a constant attitude despite the pitching and rolling of the ship beneath the table. The nongravitational acceleration is somewhat mitigated by mounting marine gravimeters deep in the hold and as close to the ship’s center of motion as possible. Special damping mechanisms also prevent the spring in the gravimeter from responding to extremely high-frequency changes in the force on the suspended mass. Instruments deployed in submersibles resting on the bottom of the ocean or in instrument packages lowered to the bottom of the ocean do not suffer from the dynamic accelerations of the moving ocean surface, but bottom currents can also be an important source of noise in submarine gravimetry. Installing instruments in boreholes is the most effective way to counter this problem, but it is also an expensive solution.
Reduction of Marine Data A number of standard corrections must be applied to the raw gravity data (either g or Dg) prior to interpretation. In addition to any drift correction, as mentioned above for relative gravity measurements, a latitude correction is immediately applied to account for the large change in gravity between the poles and the Equator caused by Earth’s rotation. Near the Equator, the centrifugal acceleration from the Earth’s spin is large, and gravity is about 50 000 gu less, on average, than at the poles. Because this effect is 5000 times larger than typical regional gravity signals of interest, it must be removed from the data using a standard formula for the variation of gravity g0 on a spheroid of revolution best fitting the shape of the Earth (eqn [5]; Y ¼ latitude). g0 ðYÞ ¼ 9:7803185 1 þ 5:278895 103 sin2 Y þ2:3462 105 sin4 Y ms2 ½5
A second correction that must be made if the gravity is measured from a moving vehicle, such as a ship or airplane, accounts for the effect on gravity of the motion of the vehicle with respect to the Earth’s spin. A ship steaming to the east is, in effect, rotating faster than the Earth. The centrifugal effect of this increased rate of rotation causes gravity to be less than it would be if the ship were stationary. The opposite effect occurs for a ship steaming to the west. This term, called the Eo¨tvo¨s correction, is largest near the Equator and involves only the east–west component vEW of the ship’s velocity vector (eqn [6]), in which o is the angular velocity of the Earth’s rotation. gEOT ¼ 2ovEW cos Y
½6
The free air gravity correction, which accounts for the elevation of the measurement above the Earth’s sea level equipotential surface, is obviously not needed if the measurement is made on the sea surface. The free air correction gFA is required if the measurement is made from a submersible or an airplane: eqn [7], where h is elevation above sea level in meters. gFA ¼ 3:1h
gu
½7
This correction is added to the observation if the sensor is deployed above the Earth’s surface, and subtracted for stations below sea level. For land surveys, the Bouguer correction accounts for the extra mass of the topography between the observation and sea level. For its marine equivalent, it adds in the extra gravitational attraction that would be present if rock rather than water existed between sea level and the bottom of the ocean. Except in areas of rugged bathymetry, the Bouguer correction gB is calculated using the slab formula (eqn [8]). gB ¼ 2pDrGz
½8
Here Dr is the density difference between oceanic crust and sea water, G is Newton’s constant, and z is the depth of the sea floor. This correction is seldom used because it produces very large positive gravity anomalies. Furthermore, there are more accurate corrections for the effect of bathymetry that do not make the unrealistic assumption that the expected state for the oceans should be that the entire depth is filled with crustal rocks displacing the water. The Bouguer correction is necessary, however, when gravity measurements are made from a submarine, in order to combine those data with more conventional observations from the sea surface. In this case, the Bouguer correction is applied twice: once to remove
(c) 2011 Elsevier Inc. All Rights Reserved.
GRAVITY
the upward attraction of the layer of water above the submarine, and once more to add in that layer’s gravitation field below the sensor. Satellite measurements of sea surface height go through a different processing sequence to recover marine geoid anomalies. The most important step is in calculating precise orbits. Information from tracking stations is supplemented with a ‘crossover analysis’ that removes long-wavelength bias in orbit elevation by forcing the height valves to agree wherever orbits cross. Corrections are then made for known physical oceanographic effects such as tides, and wave action is averaged out. The height of the sea level geoid above the Earth’s center, assuming the standard spheroid, is subtracted from the data to create geoid anomalies.
History A principal impediment to the acquisition of useful gravity observations at sea was the difficulty in separating the desired acceleration of gravity from the acceleration of the platform floating on the surface of the moving ocean. For this reason, the first successful gravity measurements to be acquired at sea were taken from a submarine by the Dutch pioneer, Vening Meinesz, in 1923. He used a pendulum gravimeter, which was the state of the art for measuring absolute gravity at that time. By accurately timing the period, T, of the swinging pendulum, the acceleration of gravity, g, can be recovered according to eqn [9], in which l is the length of the pendulum arm. sffiffiffi l T ¼ 2p g
½9
By 1959, five thousand gravity measurements had been acquired from submarines globally. These measurements were instrumental in revealing the large gravity anomalies associated with the great trenches along the western margin of the Pacific. However, these gravity observations were very timeconsuming to acquire because of the long integration times needed to achieve a high-precision estimate of the pendulum’s period, and could not be adapted for use on a surface ship. Gravity measurements at sea became routine and reliable in the late 1950s with the development of gyroscopically stabilized platforms and heavily damped mass-and-spring systems constrained to move only vertically. The new platforms compensated for the pitch and roll of the ship such that simple mass-and-spring gravimeters could collect
83
time series of variations in gravity over the oceans from vessels under way. Without any need to stop the ship on station, a time series of gravity measurements could be obtained at only small incremental cost to ship operations. With the advent of the new instrumentation, the catalogue of marine gravity values has grown in the past 40 years to more than 2.5 million measurements. A new era of precision in marine gravity began with the advent of the Global Positioning System (GPS) in the late 1980s. Prior to this time, the largest source of uncertainty in marine gravity lay in the Eo¨tvo¨s correction. Older navigation systems (dead reckoning, celestial, and even the TRANSIT satellite system) were too imprecise in the absolute position of the ship and too infrequently available to allow accurate velocity estimation from minute to minute, especially if the ship was maneuvering. Typically, gravity data had to be discarded for an hour or so near the time of any change in course. The high positioning accuracy and frequency of GPS fixes now allows such precise calculation of the Eo¨tvo¨s correction that it is no longer the limiting factor in the accuracy of marine gravity data. A breakthough in determining the global marine gravity field was achieved with the launching of the GEOS-3 (1975–1977) and Seasat (1978) satellites, which carried radar altimeters. Altimeters were deployed for the purpose of measuring dynamic sea surface elevation associated with physical oceanographic effects. The Seasat satellite carried a new, high-precision altimeter that characterized the variations in sea surface elevation with unprecedented detail. The satellite failed prematurely, but not before it returned a wealth of data on the marine gravity field from its observations of the marine geoid. The geoid variations at mid- and short-wavelength were so large that the dynamic oceanographic effects motivating the mission could be considered a much smaller noise term. The success of the Seasat mission led to the launch of Geosat, which measured the geoid at even higher precision and resolution. Unfortunately, most of that data remained classified by the US military until the results from a similar European mission were about to be released into the public domain. The declassification of the Geosat data in 1995 fueled a major revolution in our understanding of the deep seafloor (Figure 1). The latest developments in marine gravity stem from the desire to detect the shortest spatial wavelengths of gravity variations by taking gravimeters to the bottom of the ocean. Gravity is one example of a potential field, and as such the amplitude, A, of the signal of interest decays with distance, z, between source and detector as in eqn [10], where k is the
(c) 2011 Elsevier Inc. All Rights Reserved.
84
GRAVITY
Figure 1 Gravity field over the North Pacific. This view is based on satellite altimetry data from the Geosat and other missions. (Data from Sandwell and Smith (1997).)
modulus of the spatial wavenumber, the reciprocal of the spatial wavelength. ABe2pkz
½10
For sensors located on a ship at the sea surface in average ocean depths of 4.5 km, it is extremely difficult to detect short-wavelength variations in gravity of a few kilometers or less. Even lowering the gravimeter to the cruising depth of most submarines (a few hundred meters) does little to overcome the upward attenuation of the signal from localized sources on and beneath the seafloor. The solution to this problem recently has been to take gravimeters to the bottom of the ocean, either in a deep-diving submersible such as Alvin, or as an instrument package lowered on a cable. Most gravity measurements at sea are relative measurements. However, recent advances in instrumentation now allow absolute gravimeters to be deployed on the bottom of the ocean, avoiding the problem of instrument drift that adds error to relative gravity measurements. However, noise associated with the short baselines required for operation in the deep sea remains problematic.
Interpretation of Marine Gravity Short-Wavelength Anomalies
The shortest-wavelength gravity anomalies over the oceans (less than a few tens of kilometers) are the
least ambiguous to interpret since they invariably are of shallow origin. The upward continuation factor guarantees that any spatially localized anomalies with deep sources will be undetectable at the ocean surface. Near-bottom gravity measurements are able to improve somewhat the detection of concentrated density anomalies buried at deeper levels, but most are assumed to lie within the oceanic crust. One of the most useful applications of shortwavelength gravity anomalies has been to predict ocean bathymetry (Figure 2). Radar altimeters deployed on the Seasat and Geosat missions measured with centimeter accuracy the height of the underlying sea surface, an excellent approximation to the marine geoid, over all ice-free marine regions. The accuracy and spatial coverage was far better than had been provided from more than a century of marine surveys from ships. At short wavelengths, undulations of the rock–water interface are the largest contribution to the short-wavelength portion of the geoid spectrum, which opened up the possibility of predicting ocean depth from the excellent geoid data. For example, an undersea volcano, or seamount, represents a mass excess over the water it displaces. The extra mass locally raises the equipotential surface, such that positive geoid anomalies are seen over volcanoes and ridges while geoid lows are seen over narrow deeps and trenches. The prediction of bathymetry from marine geoid or gravity data is tricky: the highest frequencies in the bathymetry cannot be estimated because of the upward attenuation
(c) 2011 Elsevier Inc. All Rights Reserved.
GRAVITY
85
(A)
(B)
Figure 2 Example of bathymetric prediction from gravity anomalies in a largely unexplored region of the South Pacific. (A) The best available bathymetry from sparse echo soundings available in the early 1990s. (B) A diagram shows a dramatic improvement in definition of the bathymetry when satellite gravity observations are used to constrain the short-wavelength component of the bathymetry. (Adapted from McNutt and Bonneville (1996).)
problem, and the longer wavelengths are canceled out in the geoid by their isostatic compensation (see following section). These longer wavelengths in the bathymetry must be introduced into the solution using traditional echo soundings from sparse ship tracks. Nevertheless, the best map we currently have of the depth of the global ocean is courtesy of satellite altimetry. Mid-Wavelength Anomalies
The mid-wavelength part of the gravity spectrum (tens to hundreds of kilometers) is dominated by the effects of isostatic compensation. Isostasy is the process by which the Earth supports variations in topography or bathymetry in order to bring about a condition of hydrostatic equilibrium at depth. The definition of isostasy can be extended to include both static and dynamic compensation mechanisms, but at these wavelengths the static mechanisms are most important. There are a number of different types of isostatic compensation at work in the oceans, and the details of the gravity field can be used to distinguish them and to estimate the thermomechanical behavior of oceanic plates. One of the simplest mechanisms for isostatic compensation is Airy isostasy: the oceanic crust is thickened beneath areas of shallow bathymetry. The thick crustal roots displace denser mantle material, such that the elevated features float on the mantle much like icebergs float in the ocean. Of the various
methods of isostatic compensation, this mechanism predicts the smallest gravity anomalies over a given feature. From analysis of marine gravity, we now know that this sort of compensation mechanism is only found where the oceanic crust is extremely weak, such as on very young lithosphere near a midocean ridge. For example, large plateaus formed when hotspots intersect midocean ridges are largely supported by Airy-type isostasy. Elsewhere the oceanic lithosphere is strong enough to exhibit some lateral strength in supporting superimposed volcanoes and other surface loads. An extremely common form for support of bathymetric features in the oceans is elastic flexure. Oceanic lithosphere has sufficient strength to bend elastically, thus distributing the weight of a topographic feature over an area broader than that of the feature itself (Figure 3). Analysis of marine gravity has been instrumental in establishing that the elastic strength of the oceanic lithosphere increases with increasing age. Young lithosphere near the midocean ridge is quite weak, in some cases hardly distinguished from Airytype isostasy. The oldest oceanic lithosphere displays an effective thickness equivalent to that of a perfectly elastic plate 40 km thick. The fact that this thickness is less than that of the commonly accepted value for the thickness of the mechanical plate that drifts over the asthenosphere indicates that the base of the oceanic lithosphere is not capable of sustaining large deviatoric stresses (of the order of 100 MPa or more) over million-year timescales.
(c) 2011 Elsevier Inc. All Rights Reserved.
86
GRAVITY
Net gravity anomaly Contribution from volcano
Contribution from flexure of plate Volcano
Light crust
Elastic plate Dense mantle
Figure 3 Cartoon showing how the seafloor is warped as an elastic plate under the weight of a small volcano. The gravity anomaly that would be detected by a ship sailing along the sea surface over this feature is the net difference between the positive gravity perturbation from the extra mass of the volcano and the negative gravity perturbation produced when the elastically flexed light crust replaces denser mantle. (Adapted from McNutt and Bonneville (1996).)
Another very important method of isostatic compensation in the ocean is Pratt isostasy. This method of support supposes that the height of a vertical column of bathymetry is inversely proportional to its density. Low-density columns can be higher because they are lighter, whereas heavy columns must be short in order to produce the same integrated mass at some assumed depth of compensation. In the oceans, variations in the temperature of the lithosphere produce elevation changes in the manner of Pratt isostasy. For example, ridges stand 4 km above the deep ocean basins because the underlying lithosphere is hotter when the plate is young. The bathymetric swells around young hotspot volcanoes may also be supported by Pratt-type isostasy, although some combination of crustal thickening and dynamic isostasy may be operating as well. Again, gravity and geoid anomalies have been principal constraints in arguing for the mechanism of support for bathymetric swells. Long-Wavelength Anomalies
At wavelengths from 1000 to several thousand kilometers, gravity anomalies are usually derived from satellite observations and interpreted using equations appropriate for a spherical earth. Geoid is interpreted more commonly than gravity directly, as it emphasizes the longer wavelengths in the geopotential field. Isostatic compensation for smaller-scale bathymetric features, such as seamounts, can usually be ignored in that the gravity anomaly from bathymetry is canceled out by that from its compensation when spatially averaged over longer wavelengths.
The principal signal at these wavelengths arises from the subduction of lithospheric slabs and other sorts of convective overturn within the mantle. Three sorts of gravitational contributions must be considered: (1) the direct effect of mass anomalies within the mantle, either buoyant risers or dense sinkers which drive convection; (2) the warping of the surface caused by viscous coupling of the risers or sinkers to the earth’s surface; and (3) the warping of any deeper density discontinuities (such as the core– mantle boundary) also caused by viscous coupling. In the 1980s, estimates of the locations and densities of mass anomalies in the mantle responsible for the first contribution above began to become available courtesy of seismic tomography. Travel times of earthquake waves constrained the locations of seismically fast and slow regions in the mantle. By assuming that the seismic velocity variations were caused by temperature differences between hot, rising material and cold, sinking material, it was possible to convert velocity to density using standard relations. Knowing the locations of the mass anomalies driving convection inside the Earth led to a breakthrough in understanding the long-wavelength gravity and geoid fields. The amount of deformation on density interfaces above and below the mass anomalies inferred from tomography (contributions (2) and (3) above) depends upon the viscosity structure of Earth’s mantle. Coupling is more efficient with a more viscous mantle, whereas a weaker mantle is able to soften the transmission of the viscous stresses from the risers and sinkers. Therefore, one of the principal uses of marine gravity anomalies at long wavelengths has been to calibrate the viscosity structure of the oceanic upper mantle. This interpretation must be constrained by estimates of the dynamic surface topography over the oceans, which is actually easier to estimate than over the continents because of the relatively uniform thickness of oceanic crust. A fairly common result from this sort of analysis is that the oceanic upper mantle must be relatively inviscid. The geoid shows that there are large mass anomalies within the mantle driving convection that are poorly coupled to variations in the depth of the seafloor. If the upper mantle were more viscous, there should be a stronger positive correlation between marine geoid and depth of the seafloor at long wavelengths.
See also Manned Submersibles, Deep Water. Satellite Altimetry. Satellite Oceanography, History and Introductory Concepts.
(c) 2011 Elsevier Inc. All Rights Reserved.
GRAVITY
Further Reading Bell RE, Anderson R, and Pratson L (1997) Gravity gradiometry resurfaces. The Leading Edge 16: 55--59. Garland GD (1965) The Earth’s Shape and Gravity. New York: Pergamon Press. McNutt MK and Bonneville A (1996) Mapping the seafloor from space. Endeavour 20: 157--161. McNutt MK and Bonneville A (2000) Chemical origin for the Marquesas swell. Geochemistry, Geophysics and Geosystems 1. Sandwell DT and Smith WHF (1997) Marine gravity for Geosat and ERS-1 altimetry. Journal of Geophysical Research 102: 10039--10054. Smith WHF and Sandwell DT (1997) Global seafloor topography from satellite altimetry and ship depth soundings. Science 277: 1956--1962.
87
Turcotte DL and Schubert G (1982) Geodynamics: Applications of Continuum Physics to Geological Problems. New York: Wiley. Watts AB, Bodine JH, and Ribe NM (1980) Observations of flexure and the geological evolution of the Pacific Ocean basin. Nature 283: 532--537. Wessel P and Watts AB (1988) On the accuracy of marine gravity measurements. Journal of Geophysical Research 93: 393--413. Zumberg MA, Hildebrand JA, Stevenson JM, et al. (1991) Submarine measurement of the Newtonian gravitational constant. Physical Review Letters 67: 3051--3054. Zumberg MA, Ridgeway JR, and Hildebrand JA (1997) A towed marine gravity meter for near-bottom surveys. Geophysics 62: 1386--1393.
(c) 2011 Elsevier Inc. All Rights Reserved.
GROUNDWATER FLOW TO THE COASTAL OCEAN A. E. Mulligan and M. A. Charette, Woods Hole Oceanographic Institution, Woods Hole, MA, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction Water flowing through the terrestrial landscape ultimately delivers fresh water and dissolved solutes to the coastal ocean. Because surface water inputs (e.g., rivers and streams) are easily seen and are typically large point sources to the coast, they have been well studied and their contributions to ocean geochemical budgets are fairly well known. Similarly, the hydrodynamics and geochemical importance of surface estuaries are well known. Only recently has significant attention turned toward the role of groundwater inputs to the ocean. Historically, such inputs were considered insignificant because groundwater flow is so much slower than riverine flow. Recent work however has shown that groundwater flow through coastal sediments and subsequent discharge to the coastal ocean can have a significant impact on geochemical cycling and it is therefore a process that must be better understood. Groundwater discharge into the coastal ocean generally occurs as a slow diffuse flow but can be found as large point sources in certain terrain, such as karst. In addition to typically low flow rates, groundwater discharge is temporally and spatially variable, complicating efforts to characterize site-specific flow regimes. Nonetheless, the importance of submarine groundwater discharge (SGD) as a source of dissolved solids to coastal waters has become increasingly recognized, with recent studies suggesting that SGDderived chemical loading may rival surface water inputs in many coastal areas. So while the volume of water discharged as SGD may be small relative to surface discharge, the input of dissolved solids from SGD can surpass that of surface water inputs. For example, SGD often represents a major source of nutrients in estuaries and embayments. Excess nitrogen loading can result in eutrophication and its associated secondary effects including decreased oxygen content, fish kills, and shifts in the dominant flora. First, let us define ‘groundwater’ in a coastal context. We use the term to refer to any water that resides in the pore spaces of sediments at the land–ocean boundary. Hence, such water can be fresh terrestrially derived water that originates as
88
rainwater infiltrating through the subsurface or it can represent saline oceanic water that flows through the sediments (Figure 1). Therefore, groundwater discharging to coastal waters can have salinity that spans a large range, being some mixture of the two end members. We therefore use the terms fresh SGD and saline SGD to distinguish these sources of fluid and brackish SGD to mean a mixture of the fresh and saline end members.
Basics of Groundwater Flow Groundwater flow in the subsurface is driven by differences in energy – water flows from high energy areas to low energy. The mechanical energy of a unit volume of water is determined by the sum of gravitational potential energy, pressure energy, and kinetic energy: Energy per unit volume ¼ rgz þ P þ
rV 2 2
½1
where r is fluid density, g is gravitational acceleration, z is elevation of the measuring point relative to a datum, P is fluid pressure at the measurement point, and V is fluid velocity. Because groundwater flows very slowly (on the order of 1 m day 1 or less), its kinetic energy is very small relative to its gravitational potential and pressure energies and the kinetic energy term is therefore ignored. By removing the kinetic energy term and rearranging eqn [1] to express energy in terms of mechanical energy per unit weight, the concept of hydraulic head is developed: Energy per unit weight ¼ hydraulic head ¼ z þ
P rg ½2
Groundwater therefore flows from regions of high hydraulic head to areas of low hydraulic head. Because groundwater flows through a porous media, the rate of flow depends on soil properties such as the degree to which pore spaces are interconnected. The property of interest in groundwater flow is the permeability, k, which is a measure of the ease with which a fluid flows through the soil matrix. Groundwater flow rate can then be calculated using Darcy’s law, which says that the flow rate is linearly proportional to the hydraulic gradient:
(c) 2011 Elsevier Inc. All Rights Reserved.
q¼
rgk ðrhÞ m
½3
GROUNDWATER FLOW TO THE COASTAL OCEAN
89
Precipitation Evaporation
Overland flow Infiltratio
n
Septic system
Ocean
Seawater circulation
ate
r
Water table
tw
Groundwater
l sa r− ce e t a fa shw Inter Fre
Figure 1 Simplified water cycle at the coastal margin. Modified from Heath R (1998) Basic ground-water hydrology. US Geological Survey Water Supply Paper 2220. Washington, DC: USGS.
where q is the Darcy flux, or flow rate per unit surface area, and m is fluid viscosity. A more general expression of Darcy’s law is: k q ¼ ðrP þ rgrzÞ m
Runup and waves
z
½4
In inland aquifers, the density of groundwater is constant and eqn [4] is reduced to the simpler form of Darcy’s law (eqn [3]). In coastal aquifers, however, the presence of saline water along the coast means that the assumption of constant density is not valid and so the more inclusive form of Darcy’s law, eqn [4], is required.
Groundwater Flow at the Coast Several forces drive groundwater flow through coastal aquifers, leading to a complex flow regime with significant variability in space and time (Figure 2). The primary driving force of fresh SGD is the hydraulic gradient from the upland region of a watershed to the surface water discharge location at the coast. Freshwater flux is also influenced by several other forces at the coastal boundary that also drive seawater through the sediments. For example, seawater circulates through a coastal aquifer under the force of gravity, from oceanic forces such as
Fresh Tidal pumping 40z
Saline
Dispersive circulation
Figure 2 Simplification of an unconfined coastal groundwater system. Water flow is driven by the inland hydraulic gradient, tides, beach runup and waves, and dispersive circulation. Other forcing mechanisms can drive fluid through coastal sediments, including seasonal changes in recharge to the inland groundwater system and tidal differences across islands.
waves and tides, as a result of dispersive circulation along the freshwater–saltwater boundary within the aquifer, and from changes in upland recharge. Several other forcing mechanisms exist, but they are generally only present in specific settings. For example, tidal height differences across many islands can drive flow through the subsurface. All of these forcing mechanisms affect the rate of fluid flow for
(c) 2011 Elsevier Inc. All Rights Reserved.
90
GROUNDWATER FLOW TO THE COASTAL OCEAN
both fresh and saline groundwater and are ultimately important in controlling the submarine discharge of both fluids. The analysis of coastal groundwater flow must account for the presence of both fresh- and saline water components. When appropriate, such as in regional-scale analysis or for coarse estimation purposes, an assumption can be made that there is a sharp transition, or interface, between the fresh and saline groundwater. While this assumption is not strictly true, it is often appropriate, and invoking it results in simplifying the analysis. For example, we can estimate the position of the freshwater–saltwater interface by assuming a sharp interface, no flow within the saltwater region, and only horizontal flow within the fresh groundwater. Invoking these assumptions means that the pressures at adjacent points along the interface on both the freshwater and saltwater sides are equal. Equating these pressures and rearranging, the depth to the interface can be calculated as follows:
Interface depth ¼
r1 z ¼ 40z r2 r1
½5
where r1 is the density of fresh water (1000 kg m 3) and r2 is the density of seawater (1025 kg m 3). This equation states that the depth of the interface is 40 times the elevation of the water table relative to mean sea level. While eqn [5] is only an approximation of the interface location, it is very helpful in thinking about freshwater and saltwater movement in response to changes in fresh groundwater levels. As recharge to an aquifer increases, water levels increase and the interface is pushed downward. This is also equivalent to pushing the interface seaward and the net effect is to force saltwater out of the subsurface and to replace it with fresh water. The opposite flows occur during times of little to no recharge or if groundwater pumping becomes excessive. While the sharp-interface approach is useful for conceptualizing flow at the coast, particularly in large-scale problems, the reality is more complex. Not only does the saline groundwater flow but also a zone of intermediate salinity extends between the fresh and saline end members, establishing what many refer to as a ‘subterranean estuary’. Like their surface water counterparts, these zones are hotbeds of chemical reactions. Because the water in the interface zone ultimately discharges into coastal waters, the flow and chemical dynamics within the zone are critically important to understand. Research into these issues has only just begun.
Detecting and Quantifying Submarine Groundwater Discharge As a first step in quantifying chemical loads to coastal waters, the amount of water flowing out of the subsurface must be determined. This is particularly challenging because groundwater flow is spatially and temporally variable. A number of qualitative and quantitative techniques have been developed to sample SGD, with each method sampling a particular spatial and temporal scale. Because of limitations with each sampling method, several techniques should be used at any particular site. Physical Approaches
Infrared thermography Infrared imaging has been used to identify the location and spatial variability of SGD by exploiting the temperature difference between surface water and groundwater at certain times of the year. While this technique is quite useful for identifying spatial discharge patterns, it has not yet been applied to estimating flow rates. An example of thermal infrared imagery is shown in Figure 3, an image of the head of Waquoit Bay, a small semi-enclosed estuary on Cape Cod, Massachusetts, USA. In late summer, the groundwater temperature is approximately 13 1C, whereas surface water is about 7–10 1C warmer. Locations of SGD can be seen in the infrared image as locations along the beach face with cooler temperature than the surrounding surface water. The image clearly shows spatial variability in SGD along the beach face, information that is extremely valuable in designing an appropriate field sampling campaign. Hydrologic approaches There are two hydrologic approaches to estimating SGD: the mass balance method and Darcy’s law calculation. Both methods are typically applied to estimating fresh groundwater discharge, although Darcy’s law can also be used to estimate saline flow into and out of the seafloor. To apply Darcy’s law, one must measure the soil permeability and hydraulic head at several locations (at least two) at the field site. Data must also be gathered to determine the cross-sectional flow area. The field data are then used with Darcy’s law to calculate a groundwater flow rate into the coastal ocean. The main disadvantages of this approach include the fact that permeability is highly heterogeneous, often ranging over several orders of magnitude, and so an ‘average’ value to use with Darcy’s law is seldom, if ever, well known. Furthermore, hydraulic head measurements require invasive, typically expensive, well installations. Finally, hydraulic head
(c) 2011 Elsevier Inc. All Rights Reserved.
GROUNDWATER FLOW TO THE COASTAL OCEAN
91
Temp. (°C) 2.3 2.4−7.4 7.5−8 8.1−9 9.1−9.5 9.6−10 top o con graph tou ic r
10.1−10.5 10.6−11 11.1−11.5
3-m
11.6−12 12.1−12.5 12.6−13 13.1−13.5 13.6−14 14.1−14.5
15
12
18
9
9
11 11
9
11
18
14.6−15 14
15.1−15.5 15.6−16
25
Waquoit Bay
16.1−16.5
Beach/water line
16.6−17 17.1−17.5 0
50
100
200 m
17.6−18 18.1−18.5 18.6−19 19.1−19.5 19.6−20 20.1−20.5
Figure 3 Thermal infrared image of the head of Waquoit Bay, Massachusetts. The beach/water line is shown as a red curve. Light grays imply higher temperatures and darker shades show lower temperatures. Lower temperatures within the bay indicate regions influenced by SGD. The bars show average groundwater seepage rates as measured by manual seep meters from high tide to low tide. Numbers below the bars are average rates in cm d 1. Modified from Mulligan AE and Charette MA (2006) Intercomparison of submarine ground water discharge estimates from a sandy unconfined aquifer. Journal of Hydrology 327: 411–425.
is a point measurement and capturing the spatial variability therefore requires installing many wells. The primary advantage of this approach is that it is well established and easy to implement: head measurements are easy to collect once wells are installed and the flux calculations are simple. The mass balance approach to estimating SGD requires ascertaining all inputs and outputs of water flow, except SGD, through the groundwatershed. Assuming a steady-state condition over a specified time frame, the groundwater discharge rate is calculated as the difference between all inputs and all outputs. Implementing this approach can be quite simple or can result in complex field campaigns, but the quality of the data obviously affect the level of uncertainty. Even with extensive field sampling, water budgets are seldom known with certainty and
so should be used with that in mind. Furthermore, if the spatial and temporal variability of SGD is needed for a particular study, the mass balance approach is not appropriate. Direct measurements: Seepage meters SGD can be measured directly with seepage meters. Manual seepage meters (Figure 4) are constructed using the tops of 55 gallon drums, where one end is open and placed into the sediment. The top of the seepage meter has a valve through which water can flow; a plastic bag prefilled with a known volume of water is attached to the valve, so that inflow to or outflow from the sediments can be determined. After a set length of time, the bag is removed and the volume of water in the bag is measured. The change in water volume over the sampling period is
(c) 2011 Elsevier Inc. All Rights Reserved.
92
GROUNDWATER FLOW TO THE COASTAL OCEAN
then used to determine the average flow rate of fluid across the water–sediment boundary over the length of the sampling period. These meters are very simple to operate, but they are manually intensive and are sensitive to wave disturbance and currents. Furthermore, they only sample a small flow area and so many meters are needed to characterize the spatial variability seen at most sites (Figure 3). Recently, several other technologies have been applied toward developing automated seep meters. Technologies include the heat-pulse method, continuous
Water surface
Chemical Tracers
Collection bag
Connectors
heat, ultrasonic, and dye dilution. These meters can be left in place for days and often weeks and will measure seepage without the manual intervention needed using traditional seepage meters. The trade-off with these meters is that they are expensive and therefore only a limited number are typically employed at any given time. An example of seepage measurements made using the dye dilution meter is shown in Figure 5. Note that the seepage rates are inversely proportional to tidal height. As the tide rises, the hydraulic gradient from land to sea is reduced, seepage slows, and the flow reverses, indicating that seawater is flowing into the aquifer at this location during high tide. Conversely, at low tide, the hydraulic gradient is at its steepest and SGD increases.
Seepage meter
Fluid flow Sediment
Figure 4 Graphic of a manual seepage meter deployed at the sediment–water interface. These meters can be constructed using the top of a 55 gallon drum. The collection bag serves as a fluid reservoir so that both inflow to and outflow from the sediments can be measured. Modified from Lee DR (1977) A device for measuring seepage in lakes and estuaries. Limnology and Oceanography 22: 140–147.
The chemical tracer approach to quantifying SGD has an advantage over seepage meters in that it provides an integrated flux over a wide range of spatial scales from estuaries to continental shelves. The principle of using a chemical tracer is simple: find an element or isotope that is highly enriched (or depleted) in groundwater relative to other sources of water, like rivers or rainfall, to the system under study. If SGD is occurring, then the flux of this element via groundwater will lead to enrichment in the coastal zone that is well above background levels in the open ocean (Figure 6). A simple mass balance/box model for the system under study can be performed, where all sources of the tracer other than groundwater are subtracted from the total inventory of the chemical. The residual inventory, or ‘excess’, is then
75
Flow (cm d−1) Tide (cm)
50 Submarine groundwater discharge
25
0 −25 −50 6 Sep.
Flow Tide 7 Sep.
Seawater flow into sediments
8 Sep.
Date 2002 Figure 5 Seepage rates at Waquoit Bay, Massachusetts, USA, measured using a dye dilution automated seepage meter. Modified from Sholkovitz ER, Herbold C, and Charette MA (2003) An automated dye-dilution based seepage meter for the time-series measurement of submarine groundwater discharge. Limnology and Oceanography: Methods 1: 17–29.
(c) 2011 Elsevier Inc. All Rights Reserved.
GROUNDWATER FLOW TO THE COASTAL OCEAN
34 Myrtle 20 Beach 23 20 17 15 Inner 13 shelf 12 10 14 15
(dpm 100 l−1) Jul. 1994
226Ra
Winyah Bay
17
25 20 15 Inner 13 shelf 11 10 14 13
Sh
am
re
t lf S
Gu
Open ocean = 8.0 ± 0.5
am
tre
Sh
12 11 10
31 81
ak
re
b elf
8.9
el
32
11 8.9
ak
20 18
15 15 11 9.3 8.7 10 10
re
26
17
Cape Fear
10
17 19 16 13 20 19 10 25 8.2 Charleston 28 24 21 7.8 10 24 20 8.7 20 11 14 11 25 15 10 10 9.7 19 11 11 11 10 9.4 10 11
fb
Sav an Riv nah er
Latitude (°N)
33
25
93
S ulf
0
20
40
G
80
79
78
60 km
80
100
77
Longitude (° W) Figure 6 The distribution of 226Ra offshore South Carolina reveals high activities on the inner shelf that decrease offshore. Moore (1996) used the excess 226Ra to estimate regional SGD fluxes.
divided by the concentration of the tracer in the discharging groundwater to calculate the groundwater flow rate. Naturally occurring radionuclides such as radium isotopes and radon-222 have gained popularity as tracers of SGD due to their enrichment in groundwater relative to other sources and their built-in radioactive ‘clocks’. The enrichment of these tracers is owed to the fact that the water–sediment ratio in aquifers is usually quite small and that aquifer sediments (and sediments in general) are enriched in many U and Th series isotopes; while many of these isotopes are particle reactive and remain bound to the sediments, some like Ra can easily partition into the aqueous phase. Radon-222 (t1/2 ¼ 3.82 days) is the daughter product of 226Ra (t1/2 ¼ 1600 years) and a noble gas; therefore, it is even more enriched in groundwater than radium.
A key issue when comparing different techniques for measuring SGD is the need to define the fluid composition that each method is measuring (i.e., fresh, saline, or brackish SGD). For example, whereas hydrogeological techniques are estimates of fresh SGD, the radium and radon methods include a component of recirculated seawater. Therefore, it is often not possible to directly compare the utility of these techniques. Instead, they should be regarded as complementary. One of the seminal studies that showed how SGD can impact chemical budgets on the scale of an entire coastline was conducted off North and South Carolina, USA. Literature estimates of water residence time, riverine discharge/suspended sediment load, and the activity of desorbable 226Ra on riverine particles were used to determined that only B50% of the 226Ra inventory on the inner continental shelf
(c) 2011 Elsevier Inc. All Rights Reserved.
GROUNDWATER FLOW TO THE COASTAL OCEAN
222Rn
140
7
120
6
100
5
80
4
60
3
40
2
20
1 18 May
19 May
20 May
40
30
0 21 May
20
10
Seep meter (cm d−1)
8
Water level (cm)
(dpm l−1)
94
0
Date 2002 Tide Seep meter Radon Figure 7 Plot comparing variations in SGD at Shelter Island, NY, between an automated seepage meter and the continuous radon method.
off North and South Carolina could be explained by surface inputs (Figure 6). The remaining inventory enters the system via SGD. Using an estimate of groundwater 226Ra, it is calculated that the groundwater flux to this region of the coastline is on the order of 40% of the river water flux. The approach for quantifying SGD using 222Rn is similar to that for radium (226Ra), except for a few key differences: (1) 222Rn loss to the atmosphere must be accounted for in many situations, (2) there is no significant source from particles in rivers, and (3) decay must be accounted for owing to its relatively short half-life. The first example of 222Rn use to quantify SGD to the coastal zone occurred in a study in the northeastern Gulf of Mexico. The strong pycnocline that develops in the summer time means that fluid that flows from the sediments into the bottom boundary layer is isolated from the atmosphere and therefore no correction for the air–sea loss of 222Rn is needed. Using this approach, diffuse SGD in a 620 km2 area of the inner shelf was estimated to be equivalent to B20 first-magnitude springs (a firstmagnitude spring has a flow equal to or greater than 245 000 m3 water per day, which is equivalent to B60% of the daily water supply for the entire state of Rhode Island, USA). Since 2000, a number of SGD estimation technique intercomparison experiments have been conducted through a project sponsored by the Scientific Committee on Oceanic Research (SCOR) and the Land– Ocean Interaction in the Coastal Zone (LOICZ)
Project. During these intercomparisons, several methods (chemical tracers, different types of seep meters, hydrogeologic approaches, etc.) were run side by side to evaluate their relative strengths and weaknesses. Figure 7 displays a comparison from the Shelter Island, NY, experiment of calculated radon fluxes (based on measurements from a continuous radon monitor), with seepage rates measured directly via the dye dilution seepage meter. During the period (17–20 May) when both devices were operating at the same time, there is a clear and reproducible pattern of higher radon and water fluxes during the low tides. There is also a suggestion that the seepage spikes slightly led the radon fluxes, which is consistent with the notion that the groundwater seepage is the source of the radon (the radon monitor was located offshore of the seepage meter). The excellent agreement in patterns and overlapping calculated advection rates (seepage meter: 2–37 cm day 1, average ¼ 127 8 cm day 1; radon model: 0–34 cm day 1, average ¼ 1177 cm day 1) by these two completely independent assessment tools is reassuring.
Geochemistry of the Subterranean Estuary The magnitude of chemical fluxes carried by SGD is influenced by biogeochemical processes occurring in the subterranean estuary, defined as the mixing zone between groundwater and seawater in a coastal
(c) 2011 Elsevier Inc. All Rights Reserved.
GROUNDWATER FLOW TO THE COASTAL OCEAN
95
(a)
Fe (ppm) 4500
0
0
1500
Fe (ppm)
3000
Tan/red Tan/gray
100
150
Dark red
150
50
0
50
100 P (ppm)
150
1500
3000
50
100
Core 3 200
4500
150
Core 2 200
0
0
Yellow
Depth (cm)
Orange Red/orange
100
50
4500
Gray
3000
Gray
1500
Gray
0
0
Gray/yellow
Fe (ppm) (b)
200
Core 5 200
0
50
100
150
200
P (ppm)
0
50
100
150
200
P (ppm)
Figure 8 (a) Scientists extruding a sediment core taken through the subterranean estuary of Waquoit Bay, MA. Note the presence of iron oxides within the sediments at the bottom of the core (orange-stained sediments in foreground). (b) Changes in iron and phosphate concentration with depth in three sediment cores similar to the one shown in (a). The red circles indicate Fe concentration (ppm:mg Fe/g dry sediment) while the blue diamonds represent P (ppm:mg P/g dry sediment). Error bars indicate the standard deviation for triplicate leaches performed on a selected number of samples. The dashed lines represent the concentration of Fe and P in ‘off-site’ quartz sand. Also shown is the approximate color stratigraphy for each core. The R2 value for Fe vs. P in cores 2, 3, and 5 is 0.80, 0.91, and 0.16, respectively.
(c) 2011 Elsevier Inc. All Rights Reserved.
96
GROUNDWATER FLOW TO THE COASTAL OCEAN
aquifer. The biogeochemical reactions in such underground estuaries are presumed to be similar to the surface estuary (river water/seawater) counterpart, though few comprehensive studies of chemical cycling in subterranean estuaries have been undertaken. Drivers of biogeochemical reactions in these environments include oxidation–reduction gradients, desorption–sorption processes, and microbially driven diagenesis. In a study of the Waquoit Bay subterranean estuary, a large accumulation of iron (hydr)oxide-coated sediments within the fresh–saline interface was encountered. These iron-oxide-rich sands could act as a geochemical barrier by retaining and accumulating certain dissolved chemical species carried to the subterranean estuary by groundwater and/or coastal seawater. Significant accumulation of phosphorus in the iron oxide zones of the Waquoit cores exemplifies this process (Figure 8). Phosphorous is not the only nutrient that can be retained/removed via reactions in the subterranean estuary. The microbial reduction of nitrate to inert dinitrogen gas, a process known as denitrification, is known to occur in the redox gradients associated with fresh and saline groundwater mixing. Conversely, ammonium, which is more soluble in saline environments, may be released within the subterranean estuary’s mixing zone. While the overall importance of SGD on the ‘global cycle’ of certain chemical species remains to be seen, there is little doubt that SGD is important at the local scale both within the United States and throughout the world.
Summary Groundwater discharge to the coastal ocean can be an important source of fresh water and dissolved solutes. Although a significant amount of research into the role of SGD in solute budgets has only occurred in the past decade or so, there is increasing evidence that solute loading from groundwater can be significant enough to affect solute budgets and even ecosystem health. Proper estimation of solute loads from groundwater requires confident estimates of both the groundwater discharge rate and average solute concentrations in the discharging fluid, neither of which are easily determined. Estimating groundwater discharge rates is complicated by the spatial and temporal variability of groundwater flow. A multitude of time-varying driving mechanisms complicate analysis, as does geologic heterogeneity. Nonetheless, a suite of tools from hydrogeology, geophysics, and geochemistry have been developed for sampling and measuring SGD.
The nature of geochemical reactions within nearshore sediments is not well understood, yet recent studies have shown that important transformations occur over small spatial scales. This is an exciting and critically important area of research that will reveal important process in the near future.
See also Chemical Processes in Estuarine Sediments. Estuarine Circulation. Gas Exchange in Estuaries. Inverse Modeling of Tracers and Nutrients. LongTerm Tracer Changes. Submarine Groundwater Discharge. Trace Element Nutrients. Tracer Release Experiments. Tracers of Ocean Productivity.
Further Reading Burnett WC, Aggarwal PK, Bokuniewicz H, et al. (2006) Quantifying submarine groundwater discharge in the coastal zone via multiple methods. Science of the Total Environment 367: 498--543. Burnett WC, Bokuniewicz H, Huettel M, Moore WS, and Taniguchi M (2003) Groundwater and pore water inputs to the coastal zone. Biogeochemistry 66: 3--33. Cable JE, Burnett WC, Chanton JP, and Weatherly GL (1996) Estimating groundwater discharge into the northeastern Gulf of Mexico using radon-222. Earth and Planetary Science Letters 144: 591--604. Charette MA and Sholkovitz ER (2002) Oxidative precipitation of groundwater-derived ferrous iron in the subterranean estuary of a coastal bay. Geophysical Research Letters 29: 1444. Charette MA and Sholkovitz ER (2006) Trace element cycling in a subterranean estuary. Part 2: Geochemistry of the pore water. Geochimica et Cosmochimica Acta 70: 811--826. Heath R (1998) Basic ground-water hydrology. US Geological Survey Water Supply Paper 2220. Washington, DC: USGS. Kohout F (1960) Cyclic flow of salt water in the Biscayne aquifer of southeastern Florida. Journal of Geophysical Research 65: 2133--2141. Lee DR (1977) A device for measuring seepage in lakes and estuaries. Limnology and Oceanography 22: 140--147. Li L, Barry DA, Stagnitti F, and Parlange J-Y (1999) Submarine groundwater discharge and associated chemical input to a coastal sea. Water Resources Research 35: 3253--3259. Michael HA, Mulligan AE, and Harvey CF (2005) Seasonal water exchange between aquifers and the coastal ocean. Nature 436: 1145--1148 (doi:10.1038/ nature03935). Miller DC and Ullman WJ (2004) Ecological consequences of ground water discharge to Delaware Bay, United States. Ground Water 42: 959--970.
(c) 2011 Elsevier Inc. All Rights Reserved.
GROUNDWATER FLOW TO THE COASTAL OCEAN
Moore WS (1996) Large groundwater inputs to coastal waters revealed by Ra-226 enrichments. Nature 380: 612--614. Moore WS (1999) The subterranean estuary: A reaction zone of ground water and sea water. Marine Chemistry 65: 111--125. Mulligan AE and Charette MA (2006) Intercomparison of submarine ground water discharge estimates from a sandy unconfined aquifer. Journal of Hydrology 327: 411--425. Paulsen RJ, Smith CF, O’Rourke D, and Wong T (2001) Development and evaluation of an ultrasonic ground water seepage meter. Ground Water 39: 904--911. Portnoy JW, Nowicki BL, Roman CT, and Urish DW (1998) The discharge of nitrate-contaminated groundwater from developed shoreline to marsh-fringed estuary. Water Resources Research 34: 3095--3104.
97
Shinn EA, Reich CD, and Hickey TD (2002) Seepage meters and Bernoulli’s revenge. Estuaries 25: 126--132. Sholkovitz ER, Herbold C, and Charette MA (2003) An automated dye-dilution based seepage meter for the timeseries measurement of submarine groundwater discharge. Limnology and Oceanography: Methods 1: 17--29. Slomp CP and Van Cappellen P (2004) Nutrient inputs to the coastal ocean through submarine groundwater discharge: Controls and potential impact. Journal of Hydrology 295: 64--86. Taniguchi M, Burnett WC, Cable JE, and Turner JV (2002) Investigation of submarine groundwater discharge. Hydrological Processes 16: 2115--2129. Valiela I, Foreman K, LaMontagne M, et al. (1992) Couplings of watersheds and coastal waters: Sources and consequences of nutrient enrichment in Waquoit Bay, Massachusetts. Estuaries 15: 443--457.
(c) 2011 Elsevier Inc. All Rights Reserved.
HABITAT MODIFICATION M. J. Kaiser, Bangor University, Bangor, UK & 2009 Elsevier Ltd. All rights reserved.
Introduction This article deals with the causes and consequences of habitat modification in the marine environment, that occur as a result of human activities. The effects of natural forcing (waves, currents, glacial scour, seismic activity, and natural climate change) and how they interact with habitats are dealt with elsewhere. Nevertheless, in order to place the effects of human activities on habitats in a meaningful context, it is necessary to appreciate the different scales at which natural and human-mediated processes occur. Habitat modifications through natural processes can occur as a result of direct physical or chemical processes (e.g., alteration of temperature, seismic activity, and input of freshwater) or through biologically mediated processes (e.g., changes in species composition or dominance). At the largest scale, marine habitats are influenced by the interaction between atmospheric changes and the water column, which in turn affect the biological processes that occur in the habitat. These changes in the watercolumn environment occur over the largest spatial scales (entire oceans) and have periodicities that vary in timescale from millennia to decades (e.g., the North Atlantic Oscillation (NAO)).
Large-Scale Processes: Natural Forcing The NAO is a good example of how physical processes can affect habitat. Decadal changes in the NAO affect wind forcing and precipitation levels according to whether the NAO index is positive or negative. In years when the index is positive, average wind speeds are higher leading to greater physical mixing of the water column, which will affect the timing and persistence of stratification on the continental shelf Northeast Atlantic. When in a negative phase, lower wind speeds are associated with an increased incidence of anoxic events due to prolonged periods of stratification in coastal waters. The latter is particularly pronounced in enclosed bodies of water such as the Baltic Sea. Wind forcing also generates stress at the seabed in shallow water near the coast. These
areas are among the most biologically productive and support important nursery habitats for commercially important fish, and diving and wading birds. Wave stress is a direct controlling factor on benthic production which increases with depth and distance from the shore to a point where it reaches a maximum before decreasing again. Increasing levels of wind stress will decrease near-shore benthic production and lower coastal-carrying capacity for species dependent upon the benthos (Figure 1). Whenever the NAO is positive, increasing precipitation will elevate the amount of sediment discharged from rivers and is associated with increased frequencies of extreme sediment-discharge events linked with flash floods. Such events in large river systems (e.g., the River Amazon) can cause mass mortalities of benthos 10 km2 offshore. Elevated inputs of freshwater discharge will affect the extent of density-driven coastal currents in regions of freshwater influence (ROFIs) leading to extended front systems that run in parallel with the coastline. In addition to these periodic changes in the oceanic regime, we can expect global climate change to greatly exacerbate the extreme nature of many aspects of physical forcing that influences habitat modification in marine systems.
Smaller-Scale Processes: Natural and Biological Forcing The examples given above operate at large scales and influence both water-column and seabed habitat properties. Other more localized natural forcing events can have profound impacts upon the habitat. Examples would include glacial scour that can remove entire habitats on a scale of a few kilometers, localized seismic activity resulting in gas discharge or geothermal activity, localized coastal erosion and daily tidal scour, and carrion falls to the deep seafloor. Apart from the last example these individual processes operate at much larger scales than the modifying processes undertaken by individual biota. Such processes include grazing, predation, and bioturbation. Despite the relatively small scale of individual habitat-modifying events (e.g., an intertidal fiddler crab excavating a burrow), the sum total of many individuals events can equate to a significant disturbance process. However, natural forcing can lead to changes in biological community structure, and habitat changes can arise if ecosystem engineering organisms are affected. Ecosystem engineers
(c) 2011 Elsevier Inc. All Rights Reserved.
99
Number of sightings
(a)
HABITAT MODIFICATION
500
Waves stress−depth
(b)
400
Sea ducks, North Wales
300
Sea ducks, shell flat
200 100 0 1
3
5
7
9
11 13 15 17 19 21 23 25
Wave stress (N m−2)
100
30 25 20 15 10 5 0
Shell flat North Wales Burbo bank
0
5
Depth (m)
20
25
Biomass−Wave stress
(c) log10 bivalve AFDW (g)
10 15 Depth (m)
2 1 0 −0
0
5
10
15
20
25
30
−2 −3 Wave stress (N m−2)
Figure 1 (a) The distribution of shellfish eating diving sea ducks in Liverpool Bay, UK, with mean water depth, (b) the relationship between wave stress and depth, and (c) the relationship between bivalve biomass with depth and wave stress in the same area. Shell flat is more exposed to prevailing winds than the sheltered North Wales coast. Consequently, a peak in bivalve biomass occurs in shallower water off North Wales but in deeper water (where wave stress is lower) on shell flat. The distribution of the diving sea ducks corresponds reasonably well with the observed peak in biomass when it is considered that diving depth limits prey availability. Adapted from Kaiser MJ, Clarke KR, Hinz H, Austen MCV, Somerfield PJ, and Karakassis I (2006) Global analysis and prediction of the response of benthic biota and habitats to fishing. Marine Ecology Progress Series 311: 1–14.
are those taxa that have a strong influence on some component of a marine habitat. A good example of the latter is the fluctuation between kelp and sea urchin-dominated systems in response to periodic natural outbreaks of urchin disease (natural forcing). When urchin population biomass is high, kelp biomass is grazed down, but the latter are released from grazing when urchin populations are reduced by disease. Such systems are also affected by human activities and we will return to this later.
Habitat Modification as an Ecological Process In summary, natural forcing provides a background of habitat modification at generally large scales against which are caste the smaller-scale, more localized natural forcing events and biologically mediated modifying activities. Habitat modification creates disturbance which is a critical process in the maintenance of alpha and beta diversity in terrestrial and aquatic systems. Consequently, the seafloor is a patchwork of communities in different stages of recovery, succession, or climax. It is against this natural background of disturbance and habitat
modification that the impact and significance of human activities needs to be assessed.
Human Forcing The effects of human activities can be split broadly into: direct physical effects on the habitat, and indirect effects through alteration of habitat quality or through changes in biological composition. Direct effects are usually more immediate in terms of their impact and are more readily quantifiable (e.g., aggregate extraction, fishing disturbance, and land reclamation), whereas indirect effects may remain unrecognized for many years or until they interact with some other stressor to induce a habitat change (e.g., increasing levels of sublethal pollutants). Human activities are likely to induce long-lasting habitat changes, if the scale, frequency, or intensity of the human activity exceeds the background natural disturbance regime.
Fishing One of the most widespread agents of direct habitat modification is the use of fishing gears that are towed
(c) 2011 Elsevier Inc. All Rights Reserved.
HABITAT MODIFICATION
Y = loge (1 + (% change from control)/101)
across the seabed. This source of habitat disturbance has been the subject of over 100 experimental manipulations conducted over the last 25 years. In general, as fishing gear is towed across the seabed, parts of the gear dig into the substratum. This causes resuspension of sediments into the water column and, in deeper mud areas, can lead to significant transport of sediment down the continental shelf slope. The consequences of sediment resuspension in more dynamic habitats are likely to be less significant; however, this remains a much understudied topic. What is unequivocal is the direct mortality caused by the direct physical contact of fishing gear with benthic biota. The use of meta-analytic approaches has enabled the general response of benthic biota to different fishing gears in different habitats. What emerges from this analysis is that the direct mortality caused varies among different fishing gears and different habitats. While this result is not surprising, it has important management implications. Scallop (Pectinid) dredges have the most negative effect of all fishing gears across all habitat types, whereas the response of biota to beam trawls varies considerably with habitat type (Figure 2). Perhaps of greater importance is the time taken for disturbed communities to recover to a condition comparable with similar unimpacted areas. These empirical studies have formed the basis for a body-size-based model of the response of benthic communities to fishing disturbance which was validated extensively
with field data. This model has been used to demonstrate that the response of benthic biomass and production to fishing disturbance follows a negative power function such that relatively low levels of fishing intensity cause the greatest proportional loss of virgin biomass and production. Consequently, benthic communities in areas that are fished 2 or 3 times per annum differ little in absolute terms from similar benthic communities that are fished perhaps 10 or more times per annum. This finding is also applicable to other forms of chronic physical disturbance. The preceding observations and predictions apply to seabed habitats on the continental shelf. As yet no one has undertaken experimental manipulations of fishing disturbance in the deep sea. Nevertheless, the life history of deep-sea biota is such that the consequences of physical impacts by fishing gear on the deep-sea benthos and associated habitats will be measured in terms of years to hundreds of years. Recent discoveries of cold-water coral formations on the continental slope in the North Atlantic have coincided with clear evidence of the direct destruction of some of these structures by towed demersal trawl gear. The sensitivity and vulnerability of such longlived biogenic structures subsequently led to the creation of fishing-exclusion areas designed to protect these habitats. Fishing can also modify habitat indirectly through changes in trophic interactions. This is most likely to
Deposit feeders 1
BT
101
Suspension feeders
OT
SD
BT
OT
SD
0
−1
−2
5
4
14
2
38 77
6
28 13
5
25 34
2
35 38
64 11
2
G
S
M
G
S
G
S
G
S
G
S
G
M
M
M
M
M
S
Figure 2 The mean initial response (up to 7 days after impact) of deposit- and suspension-feeding fauna to beam trawling (BT), otter trawling (OT), and scallop dredging (SD) in gravel (G), sand (S), and muddy sand/mud habitats combined (M). Also shown are 95% confidence intervals (dashed lines indicate only two points available for the mean calculation and hence intervals usually extending outside the plotted range). Values above the x-axis denote the number of data points in each mean calculation. An adequate test for a significant initial impact is to note whether the 95% confidence interval crosses the zero-response line (central dashed line). The key point to note from this figure is how the initial response of the biota to different fishing gears varies to a different extent among habitat types. Thus, the response to beam trawling is very variable, whereas scallop dredging appears to have a consistently negative effect in all habitats. Adapted from Kaiser MJ, Galanidi M, Showler DA, et al. (2006) Distribution and behaviour of common scoter Melanitta nigra relative to prey resources and environmental parameters. Ibis 148: 110–128.
(c) 2011 Elsevier Inc. All Rights Reserved.
102
HABITAT MODIFICATION
occur in systems in which keystone species exist that exert a strong influence on other species at a higher or lower trophic level to themselves. Such systems are prone to ‘trophic cascades’. A good example of the latter is the relationship between fishing, sea lions, killer whales, sea otters, sea urchins, and kelp on the west coast of Canada. Fishing activities have depleted populations of forage fish that supported the sea lion population. As sea lions are an important prey species for killer whales, the latter have begun to predate sea otters which in turn feed on sea urchins. The reduction in sea otter population size has released sea urchins from predation pressure such that they have increased in abundance and have grazed down kelp beds such that only urchin barrens dominated by small fleshy algae remain. Similar cascade effects have also been reported for coral reef systems. Cascade effects are manifested relatively quickly due to the strong linkage between each of the components in the system. In other more complex food webs in which the linkages among different species are more complex, the entire removal of one species may have little or no effect on other components. The reason for this lack of response is that multiple species at each trophic level may utilize the same resources; thus, any resources made available after the demise of one species will be utilized by another. This is a phenomenon termed ‘diffuse predation’. Thus, ecosystem complexity provides a buffer against the effects of harvesting specific species. However, a current pressing challenge is to be able to understand when we risk overstepping the threshold beyond which sufficient components of the ecosystem have been removed to alter its functioning. The latter has been thought to occur off the eastern coast of Canada. The removal of demersal fish biomass coupled with a reduction in benthic biomass has resulted in a breakdown in benthopelagic coupling. Consequently, the system has undergone a phase shift such that most of the energy within this ecosystem is recycled within the pelagic compartment.
Aquaculture Wild-capture fishery landings peaked in the early 1990s, while the global contribution of aquaculture to the harvest of marine species has continued to increase at a rate of 10% per year since 1990. Most mariculture practices occur within 10 nautical miles of the coastline due to logistical and technological challenges. Habitat modification can occur through a change in land use through the deforestation of
mangroves for the construction of shrimp ponds or through changes in hydrography and inputs of organic matter by suspended cultivation systems. Most of the important habitat modification processes related to aquaculture relate to pollution from excessive organic inputs from excessive feed and the production of feces. The resulting elevated levels of organic input lead to classic responses in the benthic fauna (Figure 2). Extreme cases of organic-enrichment deoxygenation of the water column and sediment can result in anoxic sediments in the immediate area of cultivation and a benthic community dominated by mats of bacteria. In addition to organic enrichment, the increasing use of antimicrobial products and chemical agents to control sea-lice can alter the microbial composition of the sediment in the immediate vicinity of suspended cages used for fish cultivation. In general, however, the effects of cage cultivation on sediment habitats are relatively localized and can be alleviated by rotating the use of different areas such that areas of the seabed are allowed to recover for a period of time. In enclosed bodies of water, it is possible to modify the water-column habitat by overstocking the biomass of cultivated species such that the carrying capacity of the system is exceeded. Examples of this have occurred in enclosed bays in France (Baie D’Archacon), where oyster culture was allowed to develop in an uncontrolled manner. Overstocking led to poor growth rates (due to competition for food) and changes in benthic community composition. Some bivalve culture systems use direct laying of the stock onto the seabed. This can induce hydrological changes directly above the bivalve bed (e.g., in the case of cockles and mussels) such that turbulence increases. This has the effect of increasing the advection rate of phytoplankton and other particles down to the seabed (and hence enhances food supply) but can also increase the sedimentation rate. These systems are often typified by raised areas of the seabed that are occasionally leveled by the cultivator. The associated community changes that occur with on-bed cultivation are often dramatic, but are confined to a limited area within and around the cultivated plots (Figure 3). Overdevelopment of shrimp culture ponds along the coastal margin has frequently led to degradation of water quality and has exacerbated the spread of disease such that the cultivation activities have become unviable. These environmental problems are well understood and new developments typically involve the use of cultivation practices that are better integrated into natural systems to alleviate and prevent their potential negative effects (e.g., by planting additional mangroves).
(c) 2011 Elsevier Inc. All Rights Reserved.
HABITAT MODIFICATION
Distant control plots (no mussels) Number of species
25
?
20 15 10
Bare mud within mussel bed
5 Mussel covered areas 0
2
4
6
8
10
ln(n + 1) Mussel area cm−2 Figure 3 The response of species richness to the mussel cover in commercially cultivated plots of mussels. At very low mussel cover there is a suggestion that species richness may increase slightly as indicated by the dashed line; however, increasing the extent of mussel cover is associated with a sharp decline in species richness. Nevertheless, these effects are highly restricted to the area of cultivation as indicated by reference to ‘distant control plots’ with no mussel cover. Adapted from Beadman HA, Kaiser MJ, Galanidi M, Shucksmith R, and Willows RI (2004) Changes in species richness with stocking density of marine bivalves. Journal of Applied Ecology 41: 464–475.
Aggregate Extraction and Mining The ecological effects of mining of minerals and aggregates are similar to fishing to the extent that both processes involve the direct removal of the habitat from the seabed. Compared to the spatial area affected by towed bottom fishing gear, the effects of mining and dredging are confined to smaller areas. Nevertheless, the intensity of the impact is much greater (i.e., the depth to which the habitat is affected). Aggregates are often stable substrata that are not easily reworked by natural physical forcing, while mining can occur on the abyssal plain where physical forcing processes are much reduced. Consequently, recovery rates for both the habitats and their associated fauna are likely to exceed those that occur in response to bottom fishing in the same habitat.
River Discharge: Quality and Quantity Agricultural practices such as deforestation have led to a fourfold increase in nutrient loads in river discharge and are associated with phytoplankton blooms that generate mainly diatomaceous phytodetritus. This organic material fuels rapid rates of microbial community respiration and results in depletion of oxygen at the sediment–water interface. Such events are usually periodic and tend to coincide with periods of calm weather when inshore waters undergo stratification. The precise timing and extent of such events will vary interannually according to
103
the exact climatic conditions that prevail and upon the quality and quantity of river discharge. Enclosed water basins such as the Baltic Sea, the Adriatic, and the Gulf of Mexico are well-documented examples of locations where such events occur on a regular basis. These systems are characterized by restricted tidal inundation, poor water exchange, and are highly susceptible to stratification. While anoxic events would appear to be natural events associated with areas of river discharge, the extent and severity of these events would appear to be related to human activities. Nitrogen loading doubled in the River Mississippi catchment between the 1900s and 1980s resulting in eutrophication that has increased the scale of anoxic events in the Gulf of Mexico. This supposition is supported by studies of stable isotope signatures and organic tracers in sediment cores which indicate that nutrient levels started to increase in the 1950s and finally leveled off in the 1980s, and can be linked to the concomitant rise in the use of artificial fertilizers. In addition to concern regarding the effects of elevated nutrient inputs into coastal waters, there are many other contaminants that persist in the marine environment, particularly in sediments. Many of these contaminants, such as radionuclides and organic substances, are derived from industrial and heat-generation sources. Polychlorinated biphenols (PCBs) are particularly persistent and accumulate through the food chain with negative effects expressed in predators at the top of food chains. While the effects of such contaminants on apex predators have been documented in a number of cases, the effects on organisms at lower trophic levels are less well studied. Although invertebrates such as amphipods are used in lethal toxicity tests, the incidence of contaminant mortality associated with bioengineering organisms that affect habitat structure remains unknown. The effects of contaminants may be subtle in that they do not necessarily cause direct mortality but may have negative effects on recruitment processes and larval viability. Sublethal effects may lower the cellular energy activity of organisms such that they are no longer able to cope with more stressful environments. Thus there is a distinct possibility that a combination of factors (increasing temperature, contaminant load) could have lethal effects. The damming of rivers for the purposes of water abstraction has severely lowered the discharge of some river systems with negative consequences for local ecosystems. For example, in the Adriatic Sea, output of the River Po and adjacent river systems has been lowered by 12% in the recent years. The reduction in nutrient inputs has been associated with a
(c) 2011 Elsevier Inc. All Rights Reserved.
104
HABITAT MODIFICATION
decrease in primary production in the local shelf water mass. The structure of many river delta regions is maintained by the supply of suspended river-born sediment. However, human interference with the supply of suspended sediment can lead to alteration in the structure of deltaic regions. For example, the construction of the barrage on the River Nile in 1868 led to the reversal of the accretion of the Nile delta, a situation later worsened by the construction of the Aswan dam. This led to coastal erosion rates of between 5 and 240 m per year. The associated reduction in nutrients transported out to sea was linked with a decline in the catches of sardines. In later years, fishery production increased as the reductions in nutrients supplied through the Nile discharge were replaced by increasing coastal urbanization.
Petroleum Exploration and Production The impacts of oil production on shelf sea ecosystems have been well reviewed. The ecological effects of the contaminants of drilling muds typically lead to a reduction in species diversity, the effects of which ameliorate with increasing distance from the drilling platform. In the North Sea, the chronic use of drilling muds over a period of 6–9 years was associated with community effects at a distance of up to 6 km from the drilling platform, although in cases when water-based muds were used, the ecological effects were reduced. In other areas, the effects may be much more localized as in the case of gas platforms in the Gulf of Mexico where the ecological effects of the drilling muds were confined to a distance of 100 m from the drilling platform. Thus, as with fishing, the impact of hydrocarbon exploration depends upon the nature of the habitat that is affected by this activity.
Discussion The extent of natural and human agents of habitat modification is highly varied both in terms of the temporal and spatial scale at which they affect habitats. Smaller-scale processes should be viewed as being nested with the larger-scale forcing processes that affect global or basin-scale processes. Direct mitigation of human activities that have negative effects on habitat structure that lead to phase shifts or long-term change (such as fishing) are amenable to management at a national or local level. Understanding the interaction between a particular human
activity and the response of different habitat types is a necessary prerequisite for effective spatial management that can lead to sustainable use and conservation of the marine environment.
See also Breaking Waves and Near-Surface Turbulence. Cold-Water Coral Reefs. Groundwater Flow to the Coastal Ocean. Hypoxia. Mangroves. Mariculture Overview.
Further Reading Beadman HA, Kaiser MJ, Galanidi M, Shucksmith R, and Willows RI (2004) Changes in species richness with stocking density of marine bivalves. Journal of Applied Ecology 41: 464--475. Estes JA and Duggins DO (1995) Sea otters and kelp forests in Alaska: Generality and variation in a community ecology paradigm. Ecological Monographs 65: 75--100. Gray JS, Clarke KR, Warwick RM, and Hobbs G (1990) Detection of initial effects of pollution on marine benthos: An examination from the Ekofisk and Eldfisk oilfields, North Sea. Marine Ecology Progress Series 66: 285--299. Hall SJ (1994) Physical disturbance and marine benthic communities: Life in unconsolidated sediments. Oceanography and Marine Biology Annual Review 32: 179--239. Jennings and Kaiser M (1998) The effects of fishing on marine ecosystems. Advances in Marine Biology 34: 201--352. Kaiser MJ, Clarke KR, Hinz H, Austen MCV, Somerfield PJ , and Karakassis I (2006) Global analysis and prediction of the response of benthic biota and habitats to fishing. Marine Ecology Progress Series 311: 1--14. Kaiser MJ, Galanidi M, Showler DA, et al. (2006) Distribution and behaviour of common scoter Melanitta nigra relative to prey resources and environmental parameters. Ibis 148: 110--128. Olsgard F and Gray JF (1995) A comprehensive analysis of the effects of offshore oil and gas exploration and production on the benthic communities of the Norwegian continental shelf. Marine Ecology Progress Series 122: 277--306. Pearson TH and Rosenberg R (1987) Macrobenthic succession in relation to organic enrichment and pollution of the marine environment. Oceanography and Marine Biology Annual Review 16: 229--311. Planques A, Guilen J, and Puig P (2001) Impact of bottom trawling on water turbidity and muddy sediment of an unfished continental shelf. Limnology and Oceanography 46: 1100--1110.
(c) 2011 Elsevier Inc. All Rights Reserved.
HEAT AND MOMENTUM FLUXES AT THE SEA SURFACE P. K. Taylor, Southampton Oceanography Centre, Southampton, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1187–1195, & 2001, Elsevier Ltd.
Introduction The maintenance of the earth’s climate depends on a balance between the absorption of heat from the sun and the loss of heat through radiative cooling to space. For each 100 W of the sun’s radiative energy entering the atmosphere nearly 40 W is absorbed by the ocean – about twice that adsorbed in the atmosphere and three times that falling on land surfaces. Much of this oceanic heat is transferred back to the atmosphere by the local sea to air heat flux. The geographical variation of this atmospheric heating drives the weather systems and their associated winds. The wind transfers momentum to the sea causing waves and the wind-driven currents. Major ocean currents transport heat polewards and at higher latitudes the sea to air heat flux significantly ameliorates the climate. Thus the heat and momentum fluxes through the ocean surface form a crucial component of the earth’s climate system. The total heat transfer through the ocean surface, the net heat flux, is a combination of several components. The heat from the sun is the short-wave radiative flux (wavelength 0.3–3 mm). Around noon on a sunny day this flux may reach about 1000 W m2 but, when averaged over 24 h, a typical value is 100–300 W m2 varying with latitude and season. Part of this flux is reflected from the sea surface – about 6% depending on the solar elevation and the sea state. Most of the remaining short-wave flux is absorbed in the upper few meters of the ocean. In calm weather, with winds less than about 3 m s1, a shallow layer may be formed during the day in which the sea is warmed by a few degrees Celsius (a ‘diurnal thermocline’). However, under stronger winds or at night the absorbed heat becomes mixed down through several tens of metres. Thus, in contrast to land areas, the typical day to night variation in sea surface sea and air temperatures is small, o11C. Both the sea and the sky emit and absorb long-wave radiative energy (wavelength 3–50 mm).
Because, under most circumstances, the radiative temperature of the sky is colder than that of the sea, the downward long-wave flux is usually smaller than the upward flux. Hence the net long-wave flux acts to cool the surface, typically by 30–80 W m2 depending on cloud cover. The turbulent fluxes of sensible and latent heat also typically transfer heat from sea to air. The sensible heat flux is the transfer of heat caused by difference in temperature between the sea and the air. Over much of the ocean this flux cools the sea by perhaps 10–20 W m2. However, where cold wintertime continental air flows over warm ocean currents, for example the Gulf Stream region off the eastern seaboard of North America, the sensible heat flux may reach 100 W m2. Conversely warm winds blowing over a colder ocean region may result in a small sensible heat flux into the ocean – a frequent occurrence over the summertime North Pacific Ocean. The evaporation of water vapor from the sea surface causes the latent heat flux. This is the latent heat of vaporization which is carried by the water vapor and only released to warm the atmosphere when the vapor condenses to form clouds. Usually this flux is significantly greater than the sensible heat flux, being on average 100 W m2 or more over large areas of the ocean. Over regions such as the Gulf Stream latent heat fluxes of several hundred W m2 are observed. In foggy conditions with the air warmer than the sea, the latent heat flux can transfer heat from air to sea. In summertime over the infamous fog-shrouded Grand Banks off Newfoundland the mean monthly latent heat transfer is directed into the ocean, but this is an exceptional case.
Measuring the Fluxes The standard instruments for determining the radiative fluxes measure the voltage generated by a thermopile which is exposed to the incident radiation. Typically the incoming short-wave radiation is measured by a pyranometer which is mounted in gimbals for use on a ship or buoy (Figure 1). For better accuracy the direct and scattered components should be determined separately but, apart from at the Baseline Surface Radiation Network stations which are predominantly situated on land, at present this is rarely done. The reflected short-wave
(c) 2011 Elsevier Inc. All Rights Reserved.
105
106
HEAT AND MOMENTUM FLUXES AT THE SEA SURFACE
Figure 1 A pyranometer used for measuring short-wave radiation. The thermopile is covered by two transparent domes. (Photograph courtesy of Southampton Oceanography Centre.)
radiation is normally determined from the sun’s elevation and lookup tables based on the results of previous experiments. The pyrgeometer used to determine the long-wave radiation is similar to the pyranometer but uses a coated dome to filter out, as far as possible, the effects of the short-wave heating. Because the air close to the sea surface is normally near to the sea temperature, the use of gimbals is less important. However, a clear sky view is required and a number of correction terms have to be calculated for the temperature of the dome and any short-wave leakage. Again, only the downward component is normally measured; the upwards component is calculated from knowledge of the sea temperature and emissivity of the sea surface. The turbulent fluxes may be measured in the nearsurface atmosphere using the eddy correlation method. If upward moving air in an eddy is on average warmer and moister than the downward moving air, then there is an upwards flux of sensible heat and water vapor and hence also an upward latent heat flux. Similarly the momentum flux, or wind stress, may be determined from the correlation between the horizontal and vertical wind fluctuations. Since a large range of eddy sizes may contribute to the flux, fast response sensors capable of sampling at 10 Hz or more must be exposed for periods of the order of 30 min for each flux determination. Three-component ultrasonic anemometers (Figure 2) are relatively robust and, by also determining the speed of sound, can provide an estimate of the sonic temperature flux, a function of the heat and moisture fluxes. The sensors used for determining the fluctuations in temperature and humidity have previously tended to be fragile and prone to contamination by salt particles which are ever-present in the marine atmosphere. However,
Figure 2 The sensing head of a three-component ultrasonic anemometer. The wind components are determined from the different times taken for sound pulses to travel in either direction between the six ceramic transducers. (Photograph courtesy of Southampton Oceanography Centre.)
improved sonic thermometry, and new techniques for water vapor measurement, such as microwave refractometry or differential infrared absorption instruments, are now becoming available. Despite these improvements in instrumentation, obtaining accurate eddy correlation measurements over the sea remains very difficult. If the instrumentation is mounted on a buoy or ship the six components of the wave-induced motion of the measurement platform must be measured and removed from the signal. The distortion both of the turbulence and the mean wind by ship, buoy or fixed tower must be minimized and, as far as possible, corrected for. Thus eddy correlation measurements are not routinely obtained over the ocean, rather they are used in special air–sea interaction experiments to calibrate other less direct methods of flux estimation. For example, in the inertial dissipation method, fluctuations of the wind, temperature, or humidity at a few Hertz are measured and related (through turbulence theory) to the fluxes. This method is less sensitive to flow distortion or platform motion, but relies on various assumptions about the formation and dissipation of turbulent quantities, which may not be valid under some conditions. It has been implemented
(c) 2011 Elsevier Inc. All Rights Reserved.
HEAT AND MOMENTUM FLUXES AT THE SEA SURFACE
on a semi-routine basis on some research ships to increase the range of available flux data. The most commonly used method of flux estimation is variously referred to as the bulk (aerodynamic) formulae. These formulae relate the difference between the value of temperature, humidity or wind (‘x’ in [1]) at some measurement height, z, and the value assumed to exist at the sea surface – respectively the sea surface temperature, 98% saturation humidity (to allow for salinity effects), and zero wind (or any nonwind-induced water current). Thus the flux Fx of some quantity x is: Fx ¼ rUz Cxz ðxz x0 Þ
½1
where r is the air density, and Uz the wind speed at the measurement height. While appearing intuitively correct (for example, blowing over a hot drink will cool it faster) these formulae can also be derived from turbulence theory. The value for the transfer coefficient, Cxz, characterizes both the surface roughness applicable to x and the relationship between Fx and the vertical profile of x. This varies with the atmospheric stability, which itself depends on the momentum, sensible heat, and water vapor fluxes, as well as the measurement height. Thus, although it may appear simple, Eqn [1] must be solved by iteration, initialized using the equivalent neutral value of Cxz at some standard height (normally 10 m), Cx10n. Typical neutral values (determined using eddy correlation or inertial dissipation data) are shown in Table 1. Many research problems remain. For example: CD10n is expected to depend on the state of development of the wave field, but can this be successfully characterized by the ratio of the predominant wave speed to the wind speed (the wave age), or by the wave height and steepness, or is a spectral representation of the wave field required? What are the effects of waves propagating from other regions (i.e., swell waves)? What is the behavior of CD10n in low wind speed conditions? Furthermore CE10n and CH10n are relatively poorly defined by the available experimental data, and Table 1
107
recent bulk algorithms have used theoretical models of the ocean surface (known as surface renewal theory) to predict these quantities from the momentum roughness length.
Sources of Flux Data Until recent years the only source of data for flux calculation routinely available from widespread regions of the world’s oceans was the weather reports from merchant ships. Organized as part of the World Weather Watch system of the World Meteorological Organisation, these ‘Voluntary Observing Ships (VOS)’ are asked to return coded weather messages at 00 00, 06 00, 12 00, and 18 00 h GMT daily, also recording the observation (with further details) in the ship’s weather logbook. The very basic set of instruments provided will normally include a barometer and a means of measuring air temperature and humidity – typically wet and dry bulb thermometers mounted in a hand swung sling psychrometer or a fixed, louvered ‘Stevenson’ screen. Sea temperature is obtained using a thermometer and an insulated bucket, or by reading the temperature gauge for the engine cooling water intake. Depending on which country recruited the VOS an anemometer and wind vane might be provided, or the ship’s officers might be asked to estimate the wind velocity from observations of the sea state using a tabulated ‘Beaufort scale’. Because of the problems of adequately siting an anemometer and maintaining its calibration, these visual estimates are not necessarily inferior to anemometer-based values. Thus the VOS weather reports include all the variables needed for calculating the turbulent fluxes using the bulk formulae. However, in many cases the accuracy of the data is limited both by the instrumentation and its siting. In particular, a large ship can induce significant changes in the local temperature and wind flow, since the VOS are not equipped with radiometers. The short-wave and long-wave fluxes must be estimated from the observer’s estimate
Typical values (with estimated uncertainties) for the transfer coefficientsa
Flux
Transfer coefficients
Typical values
Momentum
Drag coefficient CD10n( 1000) Stanton no., UH10n Dalton no., UE10n
¼ 0.61 (70.05) þ 0.063 (70.005) U10n (U10n43 m s1) ¼ 0.61 þ 0.57/U10no3 m s1 1.1 (70.2) 103 1.2 (70.1) 103
Sensible heat Latent heat
a Neither the low wind speed formula for CD10n, nor the wind speed below which it should be applied, are well defined by the available, very scattered, experimental data. It should be taken simply as an indication that, at low wind speeds, the surface roughness increases as the wind speed decreases due to the dominance of viscous effects.
(c) 2011 Elsevier Inc. All Rights Reserved.
108
HEAT AND MOMENTUM FLUXES AT THE SEA SURFACE
of the cloud amount plus (as appropriate) the solar elevation, or the sea and air temperature and humidity. The unavoidable observational errors and the crude form of the radiative flux formulae imply that large numbers of reports are needed, and correction schemes must be applied, before satisfactory flux estimates can be obtained. While there are presently nearly 7000 VOS, the ships tend to be concentrated in the main shipping lanes. Thus whilst coverage in most of the North Atlantic and North Pacific is adequate to provide monthly mean flux values, elsewhere data is mainly restricted to relatively narrow, major trade routes. For most of the southern hemisphere the VOS data is only capable of providing useful values if averaged over several years, and reports from the Southern Ocean are very few indeed. These shortcomings of VOS-derived fluxes must be borne in mind when studying the flux distribution maps presented below. Satellite-borne sensors offer the potential to overcome these sampling problems. They are of two types, passive sensors which measure the radiation emitted from the sea surface and the intervening atmosphere at visible, infrared, or microwave frequencies, and active sensors which transmit microwave radiation and measure the returned signal. Unfortunately these remotely sensed data do not allow all of the flux components to be adequately estimated. Sea surface temperature has been routinely determined using visible and infrared radiometers since about 1980. Potential errors due, for example, to changes in atmospheric aerosols following volcanic eruptions, mean that these data must be continually checked against ship and buoy data. Algorithms have been developed to estimate the net surface short-wave radiation from top of the atmosphere values; those for estimating the net surface long wave are less successful. The surface wind velocity can be determined to good accuracy by active scatterometer sensors by measuring the microwave radiation backscattered from the sea surface. Unfortunately scatterometers are relatively costly to operate, since they demand significant power from the spacecraft and, to date, few have been flown. The determination of near-surface air temperature and humidity from satellite is hindered by the relatively coarse vertical resolution of the retrieved data. A problem is that the radiation emitted by the near-surface air is dominated by that originating from the sea surface. Statistically based algorithms for determining the near-surface humidity have been successfully demonstrated. More recently neural network techniques have been applied to retrieving both air temperature and humidity; however, at present there is no routinely available product. Thus the satellite flux products for which useful accuracy has
been demonstrated are presently limited to momentum, short-wave radiation, and latent heat flux. Numerical weather prediction (NWP) models (as used in weather forecasting centers) estimate values of the air–sea fluxes as a necessary part of their calculations. Since these models assimilate most of the available data from the World Weather Watch system, including satellite data, radiosonde profiles, and surface observations, it might be expected that NWP models represent the best source of flux data. However, there are other problems. The vertical resolution of these models is relatively poor and many of the near-surface processes which affect the fluxes have to be represented in terms of larger-scale parameters. Improvements to these models are normally judged on the resulting quality of the weather forecasts, not on the accuracy of the surface fluxes; sometimes these may become worse. Indeed, the continual introduction of model changes results in time discontinuities in the output variables. This makes the determination of interannual variations difficult. Because of this, NWP centres such as the European Centre for Medium Range Weather Forecasting (ECMWF) and the US National Centers for Environmental Prediction (NCEP) have reanalyzed the past weather and have gone back several decades. The surface fluxes from these reanalyses are receiving much study. Those presently available appear less accurate than fluxes derived from VOS data in regions where there are many VOS reports; in sparsely sampled regions the model fluxes may be more accurate. There are particular weaknesses in the shortwave radiation and latent heat fluxes. New reanalyses are planned and efforts are being made to improve the flux estimates; eventually these reanalyses will provide the best source of flux data for many purposes.
Regional and Seasonal Variation of the Momentum Flux The main features of the wind regimes over the global oceans have long been recognized and descriptions are available in many books on marine meteorology (see Further Reading). The major features of the wind stress variability derived from ship observations from the period 1980–93 will be summarized here, using plots for January and July to illustrate the seasonal variation. The distribution of the heat fluxes will be discussed in the next section. In northern hemisphere winter (Figure 3A) large wind stresses due to the strong midlatitude westerly winds are obvious in the North Atlantic and the North Pacific west of Japan. To the south of these
(c) 2011 Elsevier Inc. All Rights Reserved.
HEAT AND MOMENTUM FLUXES AT THE SEA SURFACE
109
Figure 3 Monthly vector mean wind stress (N m2) for (A) January and (B) July calculated from Voluntary Observing Ship weather reports for the period 1980–93. (Adapted with permission from Josey SA, Kent EC and Taylor PK (1998) The Southampton Oceanography Centre (SOC) Ocean–Atmosphere Heat, Momentum and Freshwater Flux Atlas. SOC Report no. 6.)
regions the extratropical high pressure zones result in low wind stress values, south of these is the northeast trade wind belt. The Inter-Tropical Convergence Zone (ITCZ) with very light winds is close to the equator in both oceans. In the summertime southern hemisphere the south-east trade wind belt is less well marked. The extratropical high pressure regions are extensive but, despite it being summer, high winds and significant wind stress exist in the midlatitude
southern ocean. The north-east monsoon dominates the wind patterns in the Indian Ocean and the South China Sea (where it is particularly strong). The ITCZ is a diffuse region south of the equator with relatively strong south-east trade winds in the eastern Indian Ocean. In northern hemisphere summer (Figure 3B) the wind stresses in the midlatitude westerlies are very much decreased. Both the north-east and the
(c) 2011 Elsevier Inc. All Rights Reserved.
110
HEAT AND MOMENTUM FLUXES AT THE SEA SURFACE
south-east trade wind zones are evident respectively to the north and south of the ITCZ. This is predominantly north of the equator. The south-east trades are particularly strong in the Indian Ocean and feed into a very strong south-westerly monsoon flow in the Arabian Sea. The ship data indicate very strong winds in the Southern Ocean south west of Australia. These are also evident in satellite scatterometer data, which suggest that the winds in the Pacific sector of the Southern Ocean, while still strong, are somewhat less than those in the Indian Ocean sector. In contrast the ship data appear to show very light winds. The reason is that in wintertime there are practically no VOS observations in the far south Pacific. The analysis technique used to fill in the data gaps has, for want of other information, spread the light winds of the extratropical high pressure region farther south than is realistic; a good example of the care needed in interpreting the flux maps available in many atlases.
Regional and Seasonal Variation of the Heat Fluxes The global distribution of the mean annual net heat flux is shown in Figure 4A. The accuracy and method of determination of such flux distributions will be discussed further below, here they will be used to give a qualitative description. Averaged over the year the ocean is heated in equatorial regions and loses heat in higher latitudes, particularly in the North Atlantic. However, this mean distribution is somewhat misleading, as the plots for January (Figure 4B) and July (Figure 4C) illustrate. The ocean loses heat over most of the extratropical winter hemisphere and gains heat in the extratropical summer hemisphere and in the tropics throughout the year. The relative magnitude of the individual flux components is illustrated in Figure 5 for three representative sites in the North Atlantic Ocean. At the Gulf Stream site (Figure 5A) the large cooling in winter dominates the incoming solar radiation in the annual mean. However, even at this site the mean monthly short-wave flux in summer is greater than the cooling. Indeed the effect of the longer daylight periods increases the mean short-wave radiation to values similar to or larger than those observed in equatorial regions (Figure 5C). The midlatitude site (Figure 5B) is typical of large areas of the ocean. The ocean cools in winter and warms in summer, in each case by around 100 W m2. The annual mean flux is small – around 10 W m2 – but cannot be neglected because of the very large ocean areas involved. At this site, and generally over the ocean, this annual balance is between the sum of the latent heat flux and net
long-wave flux which cool the ocean, and the net short-wave heating. Only in very cold air flows, as over the Gulf Stream in winter, is the sensible heat flux significant. As regards the interannual variation of the surface fluxes, the major large-scale feature over the global ocean is the El Nin˜o-Southern Oscillation system in the equatorial Pacific Ocean. The changes in the net heat flux under El Nin˜o conditions are around 40 W m2 in the eastern equatorial Pacific. For extratropical and midlatitude regions the interannual variability of the summertime net heat flux is typically about 20–30 W m2, being dominated by the variations in latent heat flux. In winter the typical variability increases to about 30–40 W m2, although in particular areas (such as over the Gulf Stream) variations of up to 100 W m2 can occur. The major spatial pattern of interannual variability in the North Atlantic is known as the North Atlantic Oscillation (NAO). This represents a measure of the degree to which mobile depressions, or alternatively near stationary high pressure systems, occur in the midlatitude westerly zone.
Accuracy of Flux Estimates It has been shown that, although the individual flux components are of the order of hundreds of W m2, the net heat flux and its interannual variability over much of the world ocean is around tens of W m2. Furthermore it can be shown that a flux of 10 W m2 over 1 year would, if stored in the top 500 m of the ocean, heat that entire layer by about 0.151C. Temperature changes on a decadal time scale are at most a few tenths of a degree, so the global mean budget must balance to better than a few W m2. For these various reasons there is a need to measure the flux components, which vary on many time and space scales, to an accuracy of a few W m2. Given the available data sources and methods of determining the fluxes described in the previous sections, it is not surprising that this level of accuracy cannot be achieved at present. To take an example, in calculating the flux maps shown in Figure 4 from VOS data many corrections were applied to the VOS observations to attempt to remove biases caused by the methods of observation. For example, air temperature measurements were corrected for the heat island caused by the ship heating up in sunny, low wind conditions. The wind speeds were adjusted depending on the anemometer heights on different ships. Corrections were applied to sea temperatures calculated from engine room intake data. Despite these and other corrections, the
(c) 2011 Elsevier Inc. All Rights Reserved.
HEAT AND MOMENTUM FLUXES AT THE SEA SURFACE
111
Figure 4 Variation of the net heat flux over the ocean, positive values indicate heat entering the ocean: (A) annual mean, (B) January monthly mean, (C) July monthly mean. (Adapted with permission from Josey SA, Kent EC and Taylor PK (1998) The Southampton Oceanography Centre (SOC) Ocean–Atmosphere Heat, Momentum and Freshwater Flux Atlas. SOC Report no. 6.)
global annual mean flux showed about 30 W m2 excess heating of the ocean. Previous climatologies calculated from ship data had shown similar biases and the fluxes had been adjusted to remove the bias,
or to make the fluxes compatible with estimates of the meridional heat transport in the ocean. However, comparison of the unadjusted flux data with accurate data from air–sea interaction buoys showed good
(c) 2011 Elsevier Inc. All Rights Reserved.
112
HEAT AND MOMENTUM FLUXES AT THE SEA SURFACE
600
600
500
500
400
400
300
300
200
200
100
100
0 Annual (A)
January 40°N, 60°W
July
0 Annual (B)
600
January
July
40°N, 20°W SW radn. LW radn.
500
Latent Heat Sensible Heat
400 300 200 100 0 Annual (C)
January 0°N, 20°W
July
Figure 5 Mean heat fluxes at three typical sites in the North Atlantic for the annual mean, and the January and July monthly means. In each case the left-hand column shows the fluxes which act to cool the ocean while the right-hand column shows the solar heating. (A) Gulf Stream site (401N, 601W), (B) midlatitude site (401N, 201W), (C) equatorial site (01N, 201W).
agreement between the two. This suggests that adjusting the fluxes globally is not correct and that regional flux adjustments are required; however, the exact form of these corrections is presently not shown. In the future, computer models are expected to provide a major advance in flux estimation. Recently coupled numerical models of the ocean and of the atmosphere have been run for many simulated years during which the modeled climate has not drifted. This suggests that the air–sea fluxes calculated by the models are in balance with the simulated oceanic and atmospheric heat transports. However, it does not imply that the presently estimated flux values are realistic. Errors in the short-wave and latent heat fluxes may compensate one another; indeed in a typical simulation the sea surface temperature stabilized to a value which was, over large regions of the ocean, a few degrees different from that which is observed. Nevertheless the estimation of flux values using climate or NWP models is a rapidly developing field and improvements will doubtless occur in the next few years. There will be a continued need for in situ and satellite data for assimilation into the models and for model development and verification. However, it seems very likely that in future the most accurate routine source of the air–sea flux data will
be from numerical models of the coupled ocean–atmosphere system.
See also El Nin˜o Southern Oscillation (ENSO). El Nin˜o Southern Oscillation (ENSO) Models. Evaporation and Humidity. Heat Transport and Climate. IR Radiometers. North Atlantic Oscillation (NAO). Satellite Passive-Microwave Measurements of Sea Ice. Satellite Remote Sensing of Sea Surface Temperatures. Sensors for Mean Meteorology. Turbulence Sensors. Upper Ocean Heat and Freshwater Budgets. Wave Energy. Wave Generation by Wind. Wind- and Buoyancy-Forced Upper Ocean.Wind Driven Circulation
Further Reading Browning KA and Gurney RJ (eds.) (1999) Global Energy and Water Cycles. Cambridge: Cambridge University Press. Dobson F, Hasse L, and Davis R (eds.) (1980) Air–Sea Interaction, Instruments and Methods. New York: Plenum Press.
(c) 2011 Elsevier Inc. All Rights Reserved.
HEAT AND MOMENTUM FLUXES AT THE SEA SURFACE
Kraus EB and Businger JA (1994) Atmosphere–Ocean Interaction, 2nd edn. New York: Oxford University Press. Meteorological Office (1978) Meteorology for Mariners, 3rd edn. London: HMSO.
113
Stull RB (1988) An Introduction to Boundary Layer Meteorology. Dordrecht: Kluwer Academic. Wells N (1997) The Atmosphere and Ocean: A Physical Introduction, 2nd edn. London: Taylor and Francis.
(c) 2011 Elsevier Inc. All Rights Reserved.
HEAT TRANSPORT AND CLIMATE H. L. Bryden, University of Southampton, Southampton, UK & 2009 Elsevier Ltd. All rights reserved.
Introduction: The Global Heat Budget The Earth receives energy from the sun (Figure 1) principally in the form of short-wave energy (sunlight). The amount of solar radiation is quantified by the ‘solar constant’ which satellite radiometers have measured since about 1985 to have a mean value of 1366 W m 2, an 11-year sunspot cycle of amplitude 1.5 W m 2, and a maximum at maximum sunspot activity. A fraction of the sunlight is reflected directly back into space and this fraction is termed the albedo. Brighter areas like snow in polar regions have high albedo (0.8) reflecting most of the short-wave radiation back to space, while darker areas like the ocean have small albedo (0.05) and small reflection. Averaged over the Earth’s surface, the albedo is about 0.3. Overall, the net incoming short-wave radiation (incoming minus reflected) peaks in equatorial regions and decreases to small values in polar latitudes (Figure 2). The Earth radiates energy back to space in the form of long-wave, black-body radiation proportional to the fourth power of the absolute temperature at the top of the atmosphere. Because the temperature at the top of the atmosphere is relatively uniform with latitude varying only from 200 to 230 K, there is only a small latitudinal variation in outgoing radiation (Figure 1). Over a year, the net
incoming radiation equals the net outgoing radiation within our ability to measure the radiation, thus maintaining the overall heat balance of the Earth. For the radiation budget as a function of latitude, however, there is more incoming short-wave radiation at equatorial and tropical latitudes and more outgoing long-wave radiation at polar latitudes. To maintain this heating–cooling distribution, the atmosphere and ocean must transport energy poleward from the Tropics toward the Pole and the maximum poleward energy transport in each hemisphere occurs at a latitude of c. 351, where there is a change from net incoming to net outgoing radiation, and the maximum has a magnitude of about 5.8 PW (1 petawatt (PW) ¼ 1015 W). As recently as the mid-1990s, it was controversial whether the ocean or the atmosphere was responsible for the majority of the energy transport. Oceanographers found a maximum ocean heat transport of about 2 PW at 25–301 N, while meteorologists reported a maximum atmospheric transport of about 2.5 PW from the analysis of global radiosonde network observations. Thus, there was a missing petawatt of energy transport in the Northern Hemisphere for the combined ocean–atmosphere system. Recent analyses combining observations and models suggest that the atmosphere does carry the additional petawatt that observational analyses alone could not find Radiation balance Incoming radiation (1– )Q 300
300 E
E
Radiation (W m−2)
Outgoing radiation
Q
200
200
100
100
Net radiation R
0
0 90° N 60°
Q
30°
0°
30°
60° 90° S
−100
−100 R = (1−)Q −E Figure 1 Schematic of the Earth’s radiation budget. Q is incoming, short-wave solar radiation; aQ is reflected solar radiation where a is albedo, and E is outgoing, long-wave black-body radiation.
114
Figure 2 Latitudinal profiles of net incoming short-wave radiation, outgoing long-wave radiation, and the net radiative heating of the Earth. Note the latitudinal scale is stretched so that it is proportional to the surface area of the Earth.
(c) 2011 Elsevier Inc. All Rights Reserved.
HEAT TRANSPORT AND CLIMATE
due to a sparsity of radiosonde observations over the ocean. Now it is generally accepted that the atmosphere carries the majority of energy transport at 351 N, though some oceanographers point out that half of the maximum atmospheric transport is due to latent heat (water vapor) transport that can be considered to be a joint ocean–atmosphere process (Figure 3). Here we will concentrate on ocean heat transport, especially on the mechanisms of ocean heat transport in the Northern Hemisphere. Conservation of energy in the ocean is effectively expressed as conservation of heat, where heat is defined to be rCpY, where r is seawater density, Cp is specific heat of seawater at constant pressure, and Y is potential temperature, the temperature of a water parcel brought adiabatically (without heat exchange) to the sea surface from depth. Because rCp is nearly constant at about 4.08 106 J m 3 1C 1, heat conservation is essentially expressed as conservation of potential temperature. There are many subtleties to the definitions that are described in entries for density, potential temperature, and heat, but here we use traditional definitions of potential temperature, density, and specific heat based on the internationally recognized equation of state for seawater. Ocean heat transport is then the flow of heat through the ocean, rCpYv, where v is the water
115
velocity. Such definition depends on the temperature scale and has little meaning until it is considered for a given volume of the ocean. Because of mass conservation, there is no net mass transport into or out of a fixed ocean volume over long timescales (neglecting the relatively miniscule contributions from evaporation minus precipitation), so it is the ocean heat transport convergence that is meaningful, that is the amount of heat transport into the volume minus the heat transport out of the volume. Since mass is conserved, the heat transport convergence is proportional to the mass transport times the difference in temperature between the inflow and the outflow across the boundaries of the volume. For a complete latitude band like 251 N, where the Atlantic Ocean and Pacific Ocean volume north of 251 N can be considered to be an enclosed ocean, the heat transport convergence is commonly referred to as the ocean heat transport at 251 N. Individually the Atlantic and Pacific Oceans are nearly enclosed with only a small throughflow connecting them in Bering Straits, so the Atlantic heat transport at 251 N is commonly referred to, even though it is formally the heat transport convergence between 251 N and Bering Straits and similarly for Pacific heat transport at 251 N. Such definitions of heat transport are generally used throughout the Atlantic north of 301 S and the Pacific north of about 101 N, where each ocean
Components of global energy balance 6
Northward energy transport (PW)
4
Dry static atmosphere Latent Ocean Total
2
0
−2
−4
−6 90° S 60°
30° S
15° 0° 15° Latitude scaled for spherical area
30° N
60° 90° N
Figure 3 Components of atmosphere and ocean energy transports required to balance the net radiational heating/cooling of the Earth following Figure 2. The standard atmospheric energy transport is here divided into the dry static atmospheric energy transport and the latent heat transport. Latent heat transport is fundamentally a joint atmosphere–ocean process since the atmospheric water vapor transport is balanced by an opposing oceanic freshwater transport. The ocean heat transport is determined by integrating over the oceans the spatial distribution of atmosphere–surface heat exchange calculated by subtracting the atmospheric energy transport divergence from the radiative heating at the top of the atmosphere.
(c) 2011 Elsevier Inc. All Rights Reserved.
116
HEAT TRANSPORT AND CLIMATE
basin is closed except for the small Bering Straits transport.
Air–Sea Heat Exchange Conservation of heat means that any convergence of ocean heat transport is balanced by heat loss to the atmosphere (the amount of heat gain or loss through the ocean bottom is small in comparison to exchanges with the atmosphere). Thus charts of air–sea energy exchange are primary sources for our understanding of ocean heat transport, where it occurs and how big it is. Estimates of air–sea energy exchange have long been made based on measurements of cloud cover, surface air and water temperatures, wind speed, humidity, and bulk formula exchange coefficients to calculate the size of the radiational heating and latent heat cooling of the ocean surface and the sensible heat exchange between the ocean and atmosphere. Combining such shipboard observations on a global scale produces air–sea flux climatologies giving air–sea exchange by month and region. From such climatologies, the global distribution of annual averaged air–sea energy exchange (Figure 4) shows that the ocean gains heat over much of the equatorial and tropical regions and gives up large amounts of heat over the warm poleward flowing western boundary currents like the Gulf Stream, Kuroshio or Agulhas Current, and over open-water subpolar and polar regions. Ocean heat
transport convergence for any arbitrary ocean volume can technically be estimated by summing up the air–sea energy exchange over the surface area of the ocean volume. There is a problem, however, in that the air–sea energy exchange estimates have an uncertainty of about 30 W m 2. One way to determine this uncertainty is to sum the air–sea exchanges globally and to find that there is on average a heat gain by the ocean of 30 W m 2, in each of the two state-of-the-art air–sea exchange climatologies. It is of course possible to remove this bias, either uniformly, by region or by component (radiative, latent, or sensible heat exchange); despite careful comparison with buoy measurements with bulk formula estimates at several oceanic locations, there is no consensus on how to remove the 30 W m 2 uncertainty in air–sea energy exchange.
Distribution of Ocean Heat Transport Estimating ocean heat transport convergence from in situ oceanographic measurements is a second method to quantify the role of the ocean in the global heat balance. The advantage of this direct approach is that the mechanisms of ocean heat transport are examined rather than just their overall effect in terms of the air–sea exchange. It was first reliably applied at 251 N in the Atlantic, a latitude where the warm northward Gulf Stream flow of about 30 Sv (1 Sv ¼ 1 106 m3 s 1) through Florida Straits is regularly
Net heat flux (W m−2), annual 90
Latitude
75 60
150.00
45
125.00
30
100.00 75.00
15
50.00
0
25.00
−15
0.00 −25.00
−30
−50.00
−45
−75.00
−60
−100.00
−75
−125.00
−90 30° E
90° E
150° E
150° W Longitude
90° W
30° W
30° E
Figure 4 Global distribution of the annual mean net heat gain by the ocean as determined from bulk formula calculations. Positive values indicate a gain of heat by the oceans. SOC climatology – Josey SA, Kent EC, and Taylor PK (1999) New insights into the ocean heat budget closure problem from analysis of the SOC air–sea flux climatology. Journal of Climate 12: 2856–2880.
(c) 2011 Elsevier Inc. All Rights Reserved.
HEAT TRANSPORT AND CLIMATE
monitored by submarine telephone cable voltage (the varying flow of conducting seawater through a magnetic field produces varying voltage in the cable which is continuously measured). This is the latitude of the easterly trade winds whose westward wind stress generates a northward surface water transport of about 4 Sv in the Ekman layer. Because the Atlantic is closed to the north except for a small (1 Sv) flow from the Pacific through Bering Straits, a nearly equal amount of southward flow of 35 Sv must cross the mid-ocean section between the Bahamas and Africa. The vertical distribution of this southward return flow (and its temperature) has been estimated from transatlantic hydrographic sections where temperature and salinity profiles are used to derive geostrophic velocity profiles and the reference-level velocity for the geostrophic velocity profiles is set to make the southward geostrophic transport equal to the northward Gulf Stream plus Ekman transport. Analysis of three hydrographic sections along 251 N made in 1957, 1981, and 1992 suggests that the ocean heat transport is 1.3 PW. Estimates of the error in heat transport using this direct method are about 0.3 PW, implying that direct ocean heat transport estimates are more accurate than air–sea flux estimates for areas larger than 301 latitude 301 longitude (approximately, 30 W m 2 3100 km 3100 km ¼ 0.3 PW). In terms of mechanisms, the Gulf Stream carries warm water (transport-weighted temperature of 19 1C) northward and the northward surface wind-driven Ekman transport is also warm (25 1C). The compensating southward mid-ocean flow is about equally divided between a recirculating thermocline flow above 800 m depth with an average temperature close to the 19 1C Gulf Stream flow and a cold deep-water flow with an average temperature less than 3 1C. Overall then, the northward heat transport of 1.3 PW across 251 N is due to a net northward transport of about 18 Sv of warm upper layer waters balanced by a net southward transport of cold deep waters. Deep-water formation in the Nordic and Labrador Seas of the northern Atlantic connects the northward flowing warm waters and southward flowing cold deep waters. This overall vertical circulation is commonly called the Atlantic meridional overturning circulation. During the World Ocean Circulation Experiment 1990–99, transoceanic hydrographic sections were made and ocean heat transport estimated in each ocean basin. The Atlantic heat transport is northward from 401 S to 551 N; it is of the order 0.5 PW at the southern boundary of the Atlantic, increases in the tropical regions as heat is gained at the sea surface, reaches a maximum near 251 N, and then decreases northward as the ocean gives up heat to the
117
atmosphere (Figure 5). At all latitudes, the northward heat transport is due to the meridional overturning circulation where the cold deep-water transport of c. 15–20 Sv persists through the Atlantic and the compensating northward upper water flow changes temperature, warming in tropical regions and then cooling in northern subtropical and subpolar latitudes. The North Pacific, north of 101 N where the basin is essentially closed, exhibits a similar pattern of northward heat transport, but the heat transport achieves a maximum of only 0.8 PW at 251 N, 50% less than the North Atlantic despite the Pacific being more than twice as wide as the Atlantic. The Pacific is different from the Atlantic in that there is no substantial deep meridional overturning circulation, for there is no deep-water formation in the North Pacific. The Pacific heat transport is due to a horizontal circulation where warm waters flow northward in the Kuroshio western boundary current and then recirculate southward over the vast zonal extent of the Pacific, all at depths shallower than 1000 m. The Kuroshio has about the same size and temperature structure as the Gulf Stream, but the zonal temperature distribution in the Pacific exhibits much colder upper water temperatures in mid-ocean and particularly along the eastern boundary of western North America than does the Atlantic along Europe and Africa. Thus, the northward heat transport in the North Pacific is due to the horizontal upper water circulation where warm water flows northward in the Kuroshio, loses heat to the atmosphere at latitudes south of 501 N, and then recirculates southward at colder temperatures in the mid- and eastern Pacific. South of 101 N in the Pacific and throughout the Indian Ocean, ocean heat transports are somewhat ambiguous because of the throughflow from the Pacific to the Indian Ocean through the Indonesian archipelago. The throughflow transport is not yet well defined but there is a substantial observational effort now underway to quantify its mass transport and associated temperature structure. Because mass is not conserved for individual Pacific or Indian zonal sections south of 101 N, heat transports across such zonal sections in the literature generally depend on the temperature scale used as well as on the size of the throughflow assumed. These fluxes should properly be called temperature transports and to be interpreted properly should include both a net mass transport across each section and a transport-weighted temperature. Reported northward temperature transports in the South Pacific are mainly due to a net northward transport (to balance the throughflow) times temperature in degrees Celsius. Similarly, southward temperature transports in the south Indian Ocean
(c) 2011 Elsevier Inc. All Rights Reserved.
118
HEAT TRANSPORT AND CLIMATE
60° N 0.6
30° N
Heat A2
0 P1
1 PW Uncertainty +0.5 PW −0.5 PW
1.3 A5
P3
0.5
A6
0.7
0° 1
A7
0.9
A8
0.8
A10 A11
1.4
1.6
0.2 JADE89
A9
0.3
30° S
I2 I3
P21E
1.8
I4
I10
0.6
1.5
I5
P6
0.7
1.3
60° S
1.7
0.9
1.4
0.7 A21 A12
60° W
0° W
I9S
I6
60° W
P12
120° W
180° W
120° W
Figure 5 Heat transport and temperature transport across hydrographic sections taken during the World Ocean Circulation Experiment. For the Atlantic north of 401 S and the Pacific north of 101 N, heat transport values are presented where there is no net mass transport across the sections. For the remaining sections, temperature transports are presented where there is a net mass transport across the section and the temperature transport includes the net transport multiplied by temperature in degrees Celsius.
tend to be large because they include a substantial net southward mass transport (to balance the throughflow) times an average temperature. Careful ocean heat transport divergence estimates for the Indian Ocean taking into account the throughflow transport and temperature generally suggest that the Indian Ocean gains heat from the atmosphere between 0.4 and 1.2 PW depending on the size and temperature of the throughflow. For the South Pacific, the ocean heat transport convergence or divergence is ambiguous for a normal range of throughflow transport, so it is uncertain whether the South Pacific as a whole gains heat or loses heat to the atmosphere. It is possible to combine the South Pacific and Indian sections to enclose a confined ocean volume north of say 321 S. For such combinations where mass is conserved, the Indo-Pacific Ocean heat transport at 321 S is meaningful and estimates are that it is southward with a magnitude of 0.4–1.2 PW. For the complete latitude band at about 301 S combining Indo-Pacific and Atlantic Ocean heat transports, the ocean heat transport is southward or
poleward but with only a maximum value of about 0.5 PW since the northward Atlantic heat transport partially cancels the southward Indo-Pacific heat transport. Thus, it appears that the ocean contributes much less than the atmosphere to the poleward heat transport in the Southern Hemisphere required to balance the Earth’s radiation budget.
Eddy Heat Transport In addition to heat transport by the steady ocean circulation, temporally and spatially varying currents with scales of 10–100 days and 10–100 km, which are called mesoscale features or eddies, can also transport heat. Correlations between time-varying velocity and temperature fluctuations can transport heat, rCp /Y0 v0 S (where primes denote fluctuations and angular brackets indicate time averages), even when there is no net velocity or mass transport, that is, /v0 S is zero. Such eddy heat transport can be substantial in regions where there are strong lateral temperature gradients like the Gulf Stream or Kuroshio extensions separating
(c) 2011 Elsevier Inc. All Rights Reserved.
HEAT TRANSPORT AND CLIMATE
the subtropical from the subpolar gyre. In the zonally unbounded Southern Ocean where there are several thermal fronts associated with the Antarctic Circumpolar Current, eddy heat transport is the dominant mechanism for transporting heat poleward. Here eddy motions are observed to have colder temperature when they are flowing northward and warmer temperature when following southward though there is no average flow over the eddy scales. The resulting eddy heat transport, typically of the order 4 103 W m 2 for /Y0 v0 S ¼ 0.1 1C cm s 1, is southward, downgradient from high temperature on the northern side toward cold temperature on the southern side of the front, and this downgradient heat flux is a signature of the baroclinic instability process by which eddies form and grow on the potential energy stored in the large-scale lateral temperature distribution. For the 3500 m depth and 20 000 km zonal extent of the Southern Ocean, this poleward eddy heat flux amounts to 0.3 PW across a latitude of 601 S. Individual eddies or rings of isolated water mass properties may also transport heat. For example, Agulhas rings formed with a core of Indian Ocean water properties in the retroflection area south of Africa are observed to transit across the South Atlantic. These eddies have relatively warm water cores and their heat transport is often estimated by multiplying their heat content anomaly by an estimated number of how many such rings are formed each year. Such calculation is somewhat ambiguous because it is not clear how the mass is returned and what its temperature is. Similar estimates have been made with Gulf Stream and Kuroshio rings, both warm core and cold core, and with meddies formed from the outflow of Mediterranean water. While individual rings are impressive, it is not yet clear whether they carry a significant amount of heat compared with the annual averaged air–sea exchange in any region.
Future Developments There is a third method, the residual method, for estimating ocean heat transport that takes the difference between the energy transport required to maintain the Earth’s radiation budget and the atmospheric energy transport to define the ocean heat transport as a residual. In its original implementation, the residual method could only be applied to estimate zonally averaged ocean heat transport into the polar cap north of any given latitude because atmosphere energy transport could only be determined for a complete zonal integral. In a recent development, the atmospheric energy transport divergence is estimated on a grid point basis from a globally consistent model analysis that assimilates atmospheric observations and
119
radiation variables. From such analysis, the surface energy flux can be estimated at each grid point as the difference between radiation input at the top of the atmosphere and atmospheric transport divergence. Presently, the radiation input can only be accurately estimated for the intensive period of Earth Radiation Budget Experiment from 1985 to 1989. Imposing constraints that the annual averaged surface flux over land should be zero and that the net global air–sea flux should be zero leads to realistic charts of air–sea heat exchange from which ocean heat transport divergence can be estimated. Careful comparison of such ocean heat transport convergence with existing air–sea flux climatologies and with direct estimates of ocean heat transport convergence has not yet been done. There are many outstanding questions on how ocean heat transport will change under changing climate conditions. As atmospheric CO2 has increased, the ocean has warmed up by 14 1022 J over the past 40 years. Such warming represents a heat flux of only 0.1 PW or 0.3 W m 2 averaged over the ocean surface area, so it is unlikely to be detectable in local estimates of air–sea exchange that have uncertainties of 30 W m 2 or in direct estimates of ocean heat transport convergence with uncertainties of 0.3 PW. Instead, local estimates of ocean heat content change over decadal timescales provide a sensitive estimate of how the difference between ocean heat transport convergence and air– sea exchange is changing in a changing climate. There may be changes in ocean circulation that will lead to measurable changes in air–sea heat exchange and ocean heat transport. For example, most coupled climate models predict that the Atlantic meridional overturning circulation will slow down by order of 50% over the next century as atmospheric CO2 increases. Because Atlantic heat transport is presently closely related to the strength of the meridional overturning circulation, Atlantic ocean heat transport could reduce measurably. In addition, the absence of a meridional overturning circulation in coupled climate models leads to much colder (10 1C lower) temperatures in the northern Atlantic that greatly reduce the amount of heat given up by the ocean to the atmosphere in northern latitudes. In fact, there has been a recent suggestion that the Atlantic meridional circulation decreased by 30% since 1992. The heat transport decreased by only about 15% from 1.3 to 1.1 PW, however, as the horizontal gyre circulation increased to transport more heat northward. Thus, the Atlantic heat transport may not reduce proportionately with a decreased meridional overturning circulation, because in the absence of a meridional overturning circulation it is possible that the Atlantic will become more like the Pacific with a horizontal
(c) 2011 Elsevier Inc. All Rights Reserved.
120
HEAT TRANSPORT AND CLIMATE
gyre circulation that still transports a substantial amount of heat northward. If Atlantic ocean heat transport reduces under changing climate, will the atmospheric energy transport act to compensate with larger northward energy transport? Presently, the Atlantic Ocean circulation transports more than 20% of the maximum energy transport required to balance the Earth’s radiation budget. The hypothesized Bjerkenes compensation mechanism suggests that a reduction in ocean heat transport will be compensated by increased atmospheric energy transport. For extratropical latitudes, atmospheric transport is primarily effected by eddies, cyclones, and anticyclones. Will a reduction in ocean heat transport then be accompanied by increased mid-latitude storminess and increased atmospheric energy transport? Or will the overall radiation budget for the Earth system be fundamentally altered? Clearly, it is of interest to monitor the changes in ocean circulation and heat transport, most importantly in the Atlantic where there are concerns that increasing atmospheric CO2 may lead relatively quickly to substantial changes in ocean circulation and heat transport. A program to monitor the Atlantic meridional overturning circulation and associated heat transport started in 2004 and such monitoring may provide the first evidence for changes in Atlantic ocean heat transport.
See also Agulhas Current. Ekman Transport and Pumping. Florida Current, Gulf Stream and Labrador Current. Indonesian Throughflow. Kuroshio and Oyashio Currents. Meddies and Sub-Surface Eddies. Mesoscale Eddies. Neutral Surfaces and the Equation of State.
Further Reading Bryden HL (1993) Ocean heat transport across 241 N latitude. In: McBean GA and Hantel M (eds.)
Geophysical Monograph Series, Vol. 75 Interactions between Global Climate Subsystems: The Legacy of Hann, pp. 65--75. Washington, DC: American Geophysical Union. Bryden HL and Beal LM (2001) Role of the Agulhas Current in Indian Ocean circulation and associated heat and freshwater fluxes. Deep-Sea Research I 48: 1821--1845. Bryden HL and Imawaki S (2001) Ocean heat transport. In: Siedler G, Church J, and Gould J (eds.) Ocean Circulation and Climate, pp. 455--474. New York: Academic Press. Bryden HL, Longworth HR, and Cunningham SA (2005) Slowing of the Atlantic meridional overturning circulation at 251 N. Nature 438: 655--657. Cunningham SA, Kanzow T, Rayner D, et al. (2007) Temporal variability of the Atlantic meridional overturning circulation at 26.51 N. Science 317: 935--938. Ganachaud A and Wunsch C (2000) Improved estimates of global ocean circulation, heat transport and mixing from hydrographic data. Nature 408: 453--457. Josey SA, Kent EC, and Taylor PK (1999) New insights into the ocean heat budget closure problem from analysis of the SOC air–sea flux climatology. Journal of Climate 12: 2856--2880. Lavı´n A, Bryden HL, and Parrilla G (1998) Meridional transport and heat flux variations in the subtropical North Atlantic. Global Atmosphere and Ocean System 6: 269--293. Levitus S, Antonov J, and Boyer T (2005) Warming of the World Ocean, 1955–2003. Geophysical Research Letters 32 (doi:10.1029/2004GL021592). Shaffrey L and Sutton R (2006) Bjerknes compensation and the decadal variability of the energy transports in a coupled climate model. Journal of Climate 19(7): 1167--1181. Trenberth KE and Caron JM (2001) Estimates of meridional atmosphere and ocean heat transports. Journal of Climate 14: 3433--3443. Trenberth KE, Caron JM, and Stepanaik DP (2001) The atmospheric energy budget and implications for surface fluxes and ocean heat transports. Climate Dynamics 17: 259--276. Vellinga M and Wood RA (2002) Global climatic impacts of a collapse of the Atlantic thermohaline circulation. Climatic Change 54: 251--267.
(c) 2011 Elsevier Inc. All Rights Reserved.
HISTORY OF OCEAN SCIENCES H. M. Rozwadowski, Georgia Institute of Technology, Atlanta, Georgia, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1206–1210, & 2001, Elsevier Ltd.
Oceanography is the scientific study of the ocean, its inhabitants, and its physical and chemical conditions. Since its emergence as a recognized scientific discipline, oceanography has been characterized less by the intellectual cohesiveness of a traditional academic discipline than by multidisciplinary, often large-scale, investigation of a complex and forbidding environment. The term ‘oceanography’ was not applied until the 1880s, at which point it still competed with such alternatives as ‘thassalography’ and ‘oceanology’. In many countries, ‘oceanography’ now encompasses both biological and physical traditions, but this meaning is not universal. In Russia, for instance, it does not include biological sciences; the umbrella term remains ‘oceanology’. Before scientists studied the ocean as a geographic place with an integrated ecosystem, they addressed individual biological, physical, and chemical questions related to the sea. European expansion promoted the study of marine phenomena, initiating a lasting link between commercial interests and oceanography. Institutional interest in the sea by the Royal Society in London flourished briefly from 1660 to 1675, when members, including Robert Hooke and Robert Boyle, discussed marine research, developed instruments to make observations at sea, and conducted experiments on sea water to discover its physical and chemical properties. Sailors and travelers continued to make scattered observations at sea, working from the research plan and equipment established by the Society. In addition, tidal studies were prosecuted consistently from the late seventeenth century to well into the nineteenth. Following decades of quiescence, a renewal of interest in marine phenomena occurred in the mideighteenth century. As part of that century’s growth of astronomy, geophysics, chemistry, geology, and meteorology, investigators began making observations at sea as part of their scientific pursuit of other fields. Initially, observers were traveling gentlemen, naturalists, or eclipse-expedition astronomers; later they were scientific explorers. Beginning with the voyages of Captain James Cook in the last third of the century, British exploration became
characterized by attention to scientific observations. The amount of energy devoted to marine science depended on the level of interest of expedition leaders or individual members, but, in the last quarter of the century, the volume of experiments and observations made at sea increased. Virtually all ocean investigation during this time focused on temperature and salinity of water, reflecting the growth of chemical sciences. The concept of oceanic circulation sustained by differences in density, which was widely discussed in the early nineteenth century, emerged at this time. The study of waves became more important while marine natural history remained at a modest level. Early in the nineteenth century the emphasis on physical and chemical studies continued. Curtailed exploration due to the American Revolutionary War and the Napoleonic Wars slowed work in marine science until another period of rapid expansion between 1815 and 1830. Work then focused largely on currents and salinity, reflecting the fact that marine science was conducted on Arctic expeditions which searched for sperm whaling grounds and the Northwest Passage. Navigation and the pursuit of whales prompted Arctic explorers to investigate water temperature and pressure at depths. Enthusiastic individuals, most frequently ships’ captains such as William Scorseby, continued to carry on observation programs at sea, but the 1830s saw a decrease of interest in marine science by physical scientists, who turned to rival fields of discovery including meteorology and terrestrial magnetism. At this time, however, British zoologists directed particular attention to marine fauna, embarking on small boats and yachts to collect with modified oyster dredges. Pursuit of new species as well as living relatives of fossilized ones inspired naturalists to reach deeper and deeper into the sea. Beginning in the 1850s, both British and American hydrographic institutions began deep sea sounding experiments which were guided by the promise of submarine telegraphy, although initiated in support of whaling and navigational concerns. The decades after 1840 saw the gradual awareness by scientists, sailors, entrepreneurs, and governments that the ocean’s depths were commercially and intellectually important places to investigate. An unprecedented increase in popular awareness of the sea accompanied this trend. Prefaced by the rage for seaside vacations which was initiated by railroad access to beaches, cultural interest in the ocean was
(c) 2011 Elsevier Inc. All Rights Reserved.
121
122
HISTORY OF OCEAN SCIENCES
manifested by the popularity of marine natural history collecting and the new maritime novels as well as the vogue for yachting and the personal experiences of growing numbers of ocean travelers and emigrants. Submarine telegraphy and public interest in maritime issues helped scientists argue successfully for government funding of oceanographic voyages. Declining fisheries emphasized the need to know more about the biology of the seas. Until 1840, Britain led marine science, although important work was conducted by investigators in other countries, especially Scandinavian countries. In Norway, for example, G. O. Sars carried out important research for fisheries and also studied unusual creatures dredged from very deep water. American marine science began to threaten British dominance after mid-century. Matthew Fontaine Maury was the first investigator to compile wind, current, and whale charts to improve navigation and commerce. He also dispatched the first trans-Atlantic sounding voyages in the early 1850s. Under Alexander Dallas Bache’s tenure, the US Coast Survey began to include, alongside routine charting work, special studies such as detailed surveys of the Gulf Stream, microscopic examinations of bottom sediments, and dredging cruises with the renowned zoologist Louis Agassiz and his son Alexander. Spencer F. Baird, Secretary of the Smithsonian Institution and United States Fish Commissioner, oversaw American efforts to study marine fauna. In Britain, while physical work was undertaken by the Hydrographic Office and Admiralty exploring expeditions, marine biological science centered around the British Association for the Advancement of Science Dredging Committee from 1839 until the mid-1860s. After that time, a Royal Society–Admiralty partnership took the lead and dispatched a series of summer expeditions. These culminated in the famous fouryear circumnavigation of HMS Challenger (1872– 1876), the first expedition sent out with a mandate to study the world’s oceans. The United States, Britain, and other nations as well, quickly followed the example of the Naples Zoological Station and set up coastal marine biological laboratories in the 1880s and 1890s. In the last quarter of the nineteenth century, many nations sponsored oceanographic voyages modeled after that of Challenger. The United States, Russia, Germany, Norway, France, and Italy contributed to the effort to define the limits and contents of the oceans. Late in the nineteenth century, however, the Scandinavian countries promoted a new style of ocean science to replace the great voyage tradition. Mounting concerns about depleted fisheries inspired national efforts in many countries to study the
biology of fish species as well as their migration. Sweden initiated the formation of what became in 1902 the International Council for the Exploration of the Seas (ICES), which coordinated research undertaken by eight northern European member nations. Although not a member, the United States also continued active biological research. Victor Henson’s discovery of plankton and efforts to quantify its distribution and the subsequent realization of how to use physical oceanography to investigate the movements of fish populations gave oceanographers confidence with which to study the ocean as an undivided system. World War I disrupted the international community of oceanographers but provided the impetus for developing echo-sounding technology for submarine detection, which had been pioneered for ice detection partly in response to the Titanic disaster. By the late 1920s, echo-sounders revolutionized the study of underwater topography and helped scientists to recognize the rift valleys of midocean ridges, showing them to be active, unstable regions. Fishermen took up this technology for locating schools of pelagic fish; later scientists adapted echo-sounding to create fishery-independent surveying tools. Although government funding of oceanographic research dropped back almost to pre-war levels, the late 1920s saw the appearance of the first oceanographic institutions, sponsored mostly by foundations and private individuals. In the United States, the Scripps Institution changed its mission from biological research to oceanography in 1925 and, five years later, the Woods Hole Oceanographic Institution was established on the Atlantic. The availability of oceangoing vessels in the 1930s spurred development of American oceanography on a new scale. Fisheries research, which blossomed in Europe in the 1930s, included important theoretical advances which provided the foundation for later fish population dynamics modeling as well as open ocean research such as Sir Alister Hardy’s Continuous Plankton Recorder surveys and Johannes Schmidt’s publicly acclaimed search for the mid-Atlantic spawning ground of the European eel. Economic depression affected practical government science, such as fisheries research, as well as private projects, such as several of Schmidt’s voyages, which were sponsored by the Carlsberg Foundation. As a consequence, oceanographic work, which was particularly expensive, slowed down dramatically until preparations began for World War II. As with other sciences, World War II partnerships helped forge a new relationship between governments and oceanography. Oceanographic work during and after the war carried the imprint of wartime
(c) 2011 Elsevier Inc. All Rights Reserved.
HISTORY OF OCEAN SCIENCES
government support and policy in both its problem selection and its scale. Physical studies gained and maintained precedence over biological ones. The areas of inquiry promoted by wartime efforts related to submarine and antisubmarine tactics. New research began on underwater acoustics, and ocean floor sediment charts were compiled from existing data. Wave studies also received precedence for their value in predicting surf conditions for landings. After the war, oceanographers and academic institutions, newly accustomed to generous funding, learned to accept and even encourage government support. The foundation of the National Science Foundation, which became the major federal supporter of marine science in the USA, exemplified the new high levels of support for this technology-intensive science. Oceanography became characterized by large-scale, expensive research projects such as the 1960s deepsea drilling by proponents of the theory of seafloor spreading. Major international projects also became an integral part of ocean sciences, in conjunction with postwar internationalization of science and other sectors. The International Indian Ocean Expedition, for example, targeted the least well-studied of the world’s oceans. By 1950 physical oceanographers were aware of the wanderings of the Gulf Stream and pressed urgently for systematic exploration of ocean variability. Development of deep ocean mooring technology provided the opportunity to study currents and temperatures continuously. This work showed that mesoscale variability was important and was incorporated into general circulation models. Progress in postwar physical oceanography in general proceeded in step with technological improvements, particularly the development of computing power. Another example, that of radioactive tracers and ‘experiments’ of atomic bomb-testing at sea, led to an intensification of research into biogeochemical cycling from the 1950s. Through programs such as the Geochemical Ocean Sections Study (GEOSECS), tracer use permitted estimates of mixing and circulation times by the late 1960s and led to programs such as the World Ocean Circulation Experiment (WOCE). Circulation and diffusion studies took on renewed importance as public concern about pollution rose. The idea that it was necessary to understand how the ocean functioned in order to avoid inadvertently destroying it fanned development of biological and chemical oceanography from the late 1960s onward. Much effort was funneled into baseline surveying and monitoring programs, initially focusing on potential threats to human health through eating poisoned seafood but soon broadening to investigations of biological effects of
123
contaminants. The fast growth of mariculture from this period also encouraged these studies. Although many oceanographers became interested in open ocean research and theoretical modeling in the postwar period, some in Europe resuscitated efforts, begun within ICES in the 1930s, to integrate hydrographic knowledge with fisheries biology. This work encompassed the practical attempt to guide fishermen to catch more fish as well as the scientific project of trying to relate recruitment fluctuations to environmental factors. Parallel efforts in the United States bore fruit in the California Cooperative Oceanic Fisheries Investigations (CalCOFI), a cooperative government and academic program which investigated causes for the decline of the California coastal sardine fishery. CalCOFI, though, was an exceptional project; in general through the 1960s little significant intellectual exchange occurred between biological oceanography and fisheries science. Within biological oceanography, the legacy of E. Steeman-Nielsen’s radioactive carbon tracer method led to active work measuring primary productivity, with global estimates by J. H. Ryther and others in the late 1960s. Discovery of high biological diversity in the deep sea in the late 1960s led to work which was considerably enhanced by deep-diving submersibles, whose use dramatically changed our perception of the ocean’s depths. In the area of marine geology and geophysics, the two greatest achievements since World War II have been the development of the theory of plate tectonics and the deciphering of Earth’s paleoclimate record from deep-sea sediments. The first of these was set in motion by the Vine and Matthews hypothesis of 1963, which rocked the scientific community but was accepted with remarkable speed. Research submersibles played a central role in proving the theory of seafloor spreading. The 1979 discovery of hydrothermal vents, made in the process of these geophysical investigations, led to the surprise discovery of chemosynthetic ecosystems. Reconstruction of Earth’s paleoclimates, which became possible due to the large accumulation of data and the availability of deep sea cores from the ocean drilling projects, was fueled by growing scientific concerns about global climate change. This work began during the 1960s with C. David Keeling’s launch of time-series measurements of carbon dioxide, which provided the data documenting the increase of atmospheric carbon dioxide attributable to human activities and prompted inquiry into the magnitude of exchange of carbon dioxide between the atmosphere and oceans. Studies of the ocean’s role in global warming and weather production have grown in importance in recent years.
(c) 2011 Elsevier Inc. All Rights Reserved.
124
HISTORY OF OCEAN SCIENCES
Accompanying the rise of the environmental movement, oceanography’s multi-faceted attempt to understand the oceans as integrated biological and physical environments made it a compelling discipline to ecologically and environmentally minded scientists from the 1970s onward. Food web investigations became prominent, including those which used mesoscale enclosures such as one designed and run at Canada’s Pacific Biological Station in Nanaimo. Discovery of the microbial character of the pelagic food web added a new dimension to biological oceanography. New understanding of production, including a distinction between recycled nutrients and new production made by Dugdale and Goering in 1967, provided biological oceanography with the mathematical formalism for rigorous, quantitative modeling of ocean productivity and biogeochemical fluxes. Modeling capability encouraged a link to be forged between physical and biological oceanography, because new concepts required that physical processes of mixing and upwelling be integrated into ecosystem models dealing with new production, fish production, or export of organic material from the surface layer. The international Global Ocean Ecosystem Dynamics program and the recent resurgence of fisheries oceanography exemplify relatively successful efforts to bridge trophic levels, from plankton to marine mammals and sea birds, and thereby study the ecosystem as a whole.
Distribution. Primary Production Primary Production Processes.
Methods.
Further Reading Deacon M (1971) Scientists and the Sea, 1650–1900: A Study of Marine Science. London: Academic Press. Herdman W (1923) Founders of Oceanography and their Work: An Introduction to the Science of the Sea. London: Edward Arnold. Idyll CP (1969) Exploring the Ocean World: A History of Oceanography. New York: Thomas Y. Crowell. Lee AJ (1992) The Directorate of Fisheries Research: Its Origins and Development. London: Ministry of Agriculture, Fisheries, and Food. Mills EL (1989) Biological Oceanography: An Early History, 1870–1960. Ithaca: Cornell University Press. Schlee S (1973) The Edge of an Unfamiliar World: A History of Oceanography. New York: E.P. Dutton. Smith TD (1994) Scaling Fisheries: The Science of Measuring the Effects of Fishing, 1855–1955. Cambridge: Cambridge University Press. Went AEJ (1972) Seventy years agrowing: a history of the International Council for the Exploration of the Sea, 1902–1972. Rapport et Proce`s-Verbaux des Re´unions, Conseil Permanent International pour l’Exploration de la Mer 165: 1--252. National Research Council 50 (2000) Years of Ocean Discovery: National Science Foundation, 1950–2000. Washington, DC: National Academy Press, 2000.
See also Continuous Plankton Recorders. Fisheries and Climate. Indonesian Throughflow. International Organizations. Law of the Sea. Leeuwin Current. Maritime Archaeology. Primary Production
(c) 2011 Elsevier Inc. All Rights Reserved.
HOLOCENE CLIMATE VARIABILITY M. Maslin, C. Stickley, and V. Ettwein, University College London, London, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1210–1217, & 2001, Elsevier Ltd.
Introduction Until a few decades ago it was generally thought that significant large-scale global and regional climate changes occurred at a gradual pace within a timescale of many centuries or millennia. Climate change was assumed to be scarcely perceptible during a human lifetime. The tendency for climate to change abruptly has been one of the most surprising outcomes of the study of Earth history. In particular, paleoceanographic records demonstrate that our present interglacial, the Holocene (the last B10 000 years), has not been as climatically stable as first thought. It has been suggested that Holocene climate is dominated by millennial-scale variability, with some authorities suggesting that this is a 1500 year cyclicity. These pronounced Holocene climate changes can occur extremely rapidly, within a few centuries or even within a few decades, and involve regional-scale changes in mean annual temperature of several degrees Celsius. In addition, many of these Holocene climate changes are stepwise in nature and may be due to thresholds in the climate system. Holocene decadal-scale transitions would presumably have been quite noticeable to ancient civilizations. For instance, the emergence of crop agriculture in the Middle East corresponds very closely with a sudden warming event marking the beginning of the Holocene, and the widespread collapse of the first urban civilizations, such as the Old Kingdom in Egypt and the Akkadian Empire, coincided with a cooling event at around 4300 BP. In addition, paleo-records from the late Holocene demonstrate the possible influence of climate change on the collapse of the Mayan civilization (Classic Period), while Andean ice core records suggests that alternating wet and dry periods influenced the rise and fall of coastal and highland cultures of Ecuador and Peru. It would be foolhardy not to bear in mind such sudden stepwise climate transitions when considering the effects that humans might have upon the present climate system, via the rapid generation of greenhouse gases for instance. Judging by what we
have already learnt from Holocene records, it is not improbable that the system may gradually build up over hundreds of years to a ‘breaking point’ or threshold, after which some dramatic change in the system occurs over just a decade or two. At the threshold point, the climate system is in a delicate and somewhat critical state. It may take only a relatively minor ‘adjustment’ to trigger the transition and tip the system into abrupt change. This article summarizes the current paleoceanographic records of Holocene climate variability and the current theories for their causes. Concentrating on records that cover a significant portion of the Holocene. The discussion is limited to centennial– millennial-scale variations. Figure 1 illustrates the Holocene and its climate variability in context of the major global climatic changes that have occurred during the last 2.5 million years. Short-term variations such as the North Atlantic Oscillation and the El Nin˜o-Southern Oscillation will not be discussed.
The Importance of the Oceans and Holocene Paleoceanography Climate is created from the effects of differential latitudinal solar heating. Energy is constantly transferred from the equator (relatively hot) toward the poles (relatively cold). There are two transporters of such energy – the atmosphere and the oceans. The atmosphere responds to an internal or external change in a matter of days, months, or may be a few years. The oceans, however, have a longer response time. The surface ocean can change over months to a few years, but the deep ocean takes decades to centuries. From a physical point of view, in terms of volume, heat capacity and inertia, the deep ocean is the only viable contender for driving and sustaining long-term climate change on centennial to millennial timescales. Since the process of oceanic heat transfer largely regulates climate change on longer time-scales and historic records are too short to provide any record of the ocean system prior to human intervention, we turn to marine sediment archives to provide information about ocean-driven climate change. Such archives can often provide a continuous record on a variety of timescales. They are the primary means for the study and reconstruction of the stability and natural variability of the ocean system prior to anthropogenic influences.
(c) 2011 Elsevier Inc. All Rights Reserved.
125
126
HOLOCENE CLIMATE VARIABILITY
Future 0
+100 a
Quaternary Period Holocene Pleistocene
Recent 100 a
1 ka
6 NOW Little Ice Age
El Niño and the North Atlantic Oscillation
Dansgaard _ Oeschger cycles
10 Ma
2 3
Deglaciation and the Younger Dryas
Global warming
1 Ma
4
5
7 8
100 ka
10 ka
Plio-Miocene
Heinrich events and Dansgaard _ Oeschger cycles
Development of NADW Mid-Pleistocene revolution
Interglacials: 410 ka and 125 ka
1
Onset of Northern Hemisphere glaciation
Figure 1 Log time scale cartoon, illustrating the most important climate events in the Quaternary Period. (a, ka, Ma, refer to years ago.) (1) Onset of Northern Hemisphere Glaciation (3.2–2.5 Ma), ushering in the strong glacial–interglacial cycles which are characteristic of the Quaternary Period. (2) Mid-Pleistocene Revolution when the dominant periodicity of glacial–interglacial cycles switched from 41 000 y, to every 100 000 y. The external forcing of the climate did not change; thus, the internal climate feedback’s must have altered. (3) The two closest analogues to the present climate are the interglacial periods at 420 000 to 390 000 years ago (oxygen isotope stage 11) and 130 000 to 115 000 years ago (oxygen isotope stage 5e, also known as the Eemian). (4) Heinrich events and Dansgaard–Oeschger cycles (see text). (5) Deglaciation and the Younger Dryas events. (6) Holocene Dansgaard– Oeschger cycles (see text). (7) Little Ice Age (AD 1700), the most recent climate event that seems to have occurred throughout the Northern Hemisphere. (8) Anthropogenic global warming.
One advantage of marine sediments is that they can provide long, continuous records of Holocene climate at annual (sometimes intraannual) to centennial time-resolutions. However, there is commonly a trade-off between temporal and spatial resolution. Deep-ocean sediments usually represent a large spatial area, but sedimentation rates in the deep-ocean are on average between 0.002 and 0.005 cm y 1, with very productive areas producing a maximum of 0.02 cm y 1. This limits the temporal resolution to a maximum of 200 years per cm (50 y cm 1 for productive areas). Mixing by the process of bioturbation will reduce the resolution further. On continental shelves and in bays and other specialized sediment traps such as anoxic basins and fiords, sedimentation rates can exceed 1 cm y 1 providing temporal resolution of over 1 y cm 1. More local conditions are recorded in laminated marine sediments formed in anoxic environments, where biological activity can not disturb the sediments. For example, Pike and Kemp (1997) analysed annual and intraannual variability within the Gulf of California from laminated sediments containing a record of diatom-mat accumulation. Time series analysis highlighted a decadal-scale variability in mat-deposition associated with Pacific-wide changes in surface water circulation, suggested to be influenced by solar-cycles. In addition, anoxic sediments from the Mediterranean Ridge (ODP Site 971) reveal
seasonal-scale variability during the late Quaternary from a laminated diatom-ooze sapropel. Pearce et al., (1998) inferred changes in the monsoon-related nutrient input to the Mediterranean Basin via the Nile River as the main cause of the variations in the laminated sediments, which suggests a wide influence of changes in seasonality. Other potentially extremely high-resolution studies will come from Saanich Inlet, a Canadian fiord and Prydz Bay in Antarctica, sites recently drilled by the Ocean Drilling Program. However, the main drawback to such high-resolution locations is that they contain highly localized environmental and climate information. An additional problem associated mainly with continental margins is reworking, erosion, and redistribution of the sediment by mass density flows such as turbidities and slumps. Hence we concentrate on wider-scale records of Holocene climate change.
Holocene Climatic Variability Initial studies of the Greenland ice core records concluded the absence of major climate variation within the Holocene. This view is being progressively eroded, particular in the light of new information being obtained from marine sediments (Figure 2). Long-term trends indicate an early to mid-Holocene climatic optimum with a cooling trend in the late
(c) 2011 Elsevier Inc. All Rights Reserved.
127
HOLOCENE CLIMATE VARIABILITY
Northwest African Climate
North Atlantic Climate
ODP 658C
V29-191
Summer radiation _ Terrigenous (%) SST anomaly (˚C) (W m 2 ) 450
470
490
40
50
60
_8_
6 _4 _2 0
2
_1
(lithics g ) 1000 100 10
1
NEAP 15K Sortable silt mean size (10_63 µm) 11 12 13 14 15
0
Calendar age (ka BP)
_ 34
1
0
2
2
2
3
3
3 4
4
4 5
5
5
6
6
10000
7 8
7
5000
Calendar age (BP)
African Humid Period
10000
_ 35
Holocene
5000
δ18O (ppt) _ 36
LIA 1
LIA 1
GISP2
FASTER
8
COLD
WARM YD
(A)
(B)
(C)
SSTwarm SSTcold
(D)
(E)
(F)
Figure 2 Comparison of summer insolation for 651N with north-west African climate (deMenocal et al., 2000) and North Atlantic climate (V29-191, Bond et al., 1997; NEAP 15 K, Bianchi and McCave, 1999; GISP2, O’Brien et al., 1996). Note the similarity of events labeled 1 to 8 and the Little Ice Age (LIA).
Holocene. Superimposed on this trend are several distinct oscillations or climatic cooling steps that appear to be of widespread significance (see Figure 2), the most dramatic of which occurred 8200, 5500, and 4400 years ago and between AD 1200 and AD 1650. The event 8200 years ago is the most striking and abrupt, leading to widespread cool and dry conditions lasting perhaps 200 years, before a rapid return to climates warmer and generally moister than at present. This event is noticeably present in the GISP2 Greenland ice cores, from which it appears to have been about half as severe as the Younger Dryas to Holocene transition. Marine records of North African to Southern Asian climate suggest more arid conditions involving a failure of the summer monsoon rains. Cold and/or arid conditions also seem to have occurred in northernmost South America, eastern North America and parts of north-west Europe. In the middle Holocene approximately 5500–5300 years ago there was a sudden and widespread shift in precipitation, causing many regions to become either noticeably drier or moister. The dust and sea surface temperature records off north-west Africa show that the African Humid Period, when much of subtropical West Africa was vegetated, lasted from 14 800 to 5500 years ago and was followed by a 300year transition to much drier conditions (de Menocal et al., 2000). This shift also corresponds to the decline of the elm (Ulmus) in Europe about 5700, and of hemlock (Tsuga) in North America about 5300
years ago. Both vegetation changes were initially attributed to specific pathogen attacks, but it is now thought they may have been related to climate deterioration. The step to colder and drier conditions in the middle of an interglacial period is analogous to a similar change that is observed in records of the last interglacial period referred to as Marine Oxygen Isotope Stage 5e (Eemian). There is also evidence for a strong cold and arid event occurring about 4400 years ago across the North Atlantic, northern Africa, and southern Asia. This cold, and arid event coincides with the collapse of a large number of major urban civilizations, including the Old Kingdom in Egypt, the Akkadian Empire in Mesopotamia, the Early Bronze Age societies of Anatolia, Greece, Israel, the Indus Valley civilization in India, the Hilmand civilization in Afganistan, and the Hongshan culture of China.
Little Ice Age (LIA) The most recent Holocene cold event is the Little Ice Age (see Figures 2 and 3). This event really consists of two cold periods, the first of which followed the Medieval Warm Period (MWP) that ended B1000 years ago. This first cold period is often referred to as the Medieval Cold Period (MCP) or LIAb. The MCP played a role in extinguishing Norse colonies on Greenland and caused famine and mass migration in Europe. It started gradually before AD 1200 and ended at about AD 1650. This second cold period,
(c) 2011 Elsevier Inc. All Rights Reserved.
128
HOLOCENE CLIMATE VARIABILITY
Greenland temperature
Bermuda Rise SST
(˚C)
(˚C)
_ 32
_ 31
21
22
23
24
West African SST 25
_6
(˚C, anomaly) _4 _2 0
2
0 a 500
LIA
Age (BP)
b 1000 MWP 1500 Mid-Neoglacial
2000 2500
Figure 3 Comparison of Greenland temperatures, the Bermuda Rise sea surface temperatures (SST) (Keigwin, 1996), and west African and a sea surface temperature (deMenocal et al., 2000) for the last 2500 years. LIALittle Ice Age; MWPMedieval Warm Period. Solid triangles indicate radiocarbon dates.
may have been the most rapid and the largest change in the North Atlantic during the Holocene, as suggested from ice-core and deep-sea sediment records. The Little Ice Age events are characterized by a drop in temperature of 0.5–11C in Greenland and a sea surface temperature falls of 41C off the coast of west Africa and 21C off the Bermuda Rise (see Figure 3).
Holocene Dansgaard–Oeschger Cycles The above events are now regarded as part of the millennial-scale quasiperiodic climate changes characteristic of the Holocene (see Figure 2) and are thought to be similar to glacial Dansgaard–Oeschger (D/O) cycles. The periodicity of these Holocene D/O cycles is a subject of much debate. Initial analysis of the GISP2 Greenland ice core and North Atlantic sediment records revealed cycles at approximately the same 1500 (7500)-year rhythm as that found within the last glacial period. Subsequent analyses have also found a strong 1000-year cycle and a 550year cycle. These shorter cycles have also been recorded in the residual d14 C data derived from dendrochronologically calibrated bidecadal tree-ring measurements spanning the last 11 500 years. In general, during the coldest point of each of the millennial-scale cycles shown in Figure 2, surface water temperatures of the North Atlantic were about 2– 41C cooler than during the warmest part. One cautionary note is that Wunsch has suggested a more radical explanation for the pervasive 1500year cycle seen in both deep-sea and ice core, glacial and interglacial records. Wunsch suggests that the
extremely narrow spectral lines (less than two bandwidths) that have been found at about 1500 years in many paleo-records may be due to aliasing. The 1500-year peak appears precisely at the period predicted for a simple alias of the seasonal cycle sampled inadequately (under the Nyquist criterion) at integer multiples of the common year. When Wunsch removes this peak from the Greenland ice core data and deep-sea spectral records, the climate variability appears as expected to be a continuum process in the millennial band. This work suggests that finding a cyclicity of 1500 years in a dataset may not represent the true periodicity of the millennialscale events. The Holocene Dansgaard–Oeschger events are quasi periodic, with different and possibly stochastic influences.
Causes of Millennial Climate Fluctuation during the Holocene As we have already suggested, deep water circulation plays a key role in the regulation of global climate. In the North Atlantic, the north-east-trending Gulf Stream carries warm and relatively salty surface water from the Gulf of Mexico up to the Nordic seas. Upon reaching this region, the surface water has cooled sufficiently that it becomes dense enough to sink, forming the North Atlantic Deep Water (NADW). The ‘pull’ exerted by this dense sinking maintains the strength of the warm Gulf Stream, ensuring a current of warm tropical water into the North Atlantic that sends mild air masses across to the European continent. Formation of the NADW can be weakened by two processes. (1) The presence
(c) 2011 Elsevier Inc. All Rights Reserved.
HOLOCENE CLIMATE VARIABILITY
of huge ice sheets over North America and Europe changes the position of the atmospheric polar front, preventing the Gulf Stream from traveling so far north. This reduces the amount of cooling and the capacity of the surface water to sink. Such a reduction of formation occurred during the last glacial period. (2) The input of fresh water forms a lens of less-dense water, preventing sinking. If NADW formation is reduced, the weakening of the warm Gulf Stream causes colder conditions within the entire North Atlantic region and has a major impact on global climate. Bianchi and McCave, using deep-sea sediments from the North Atlantic, have shown that during the Holocene there have been regular reductions in the intensity of NADW (Figure 2E), which they link to the 1500-year D/O cycles identified by O’Brien and by Bond (1997). There are two possible causes for the millennial-scale changes observed in the intensity of the NADW: (1) instability in the North Atlantic region caused by varying
129
freshwater input into the surface waters; and (2) the ‘bipolar seesaw’. There are a number of possible reasons for the instability in the North Atlantic region caused by varying fresh water input into the surface waters:
• • • •
Internal instability of the Greenland ice sheet, causing increased meltwater in the Nordic Seas that reduces deep water formation. Cyclic changes in sea ice formation forced by solar variations. Increased precipitation in the Nordic Seas due to more northerly penetration of North Atlantic storm tracks. Changes in surface currents, allowing a larger import of fresher water from the Pacific, possibly due to reduction in sea ice in the Arctic Ocean.
The other possible cause for the millennial-scale changes is an extension of the suggested glacial intrinsic millennial-scale ‘bipolar seesaw’ to the
0.8
Northward heat flux (1015 W)
0.4 N. Atlantic heat piracy
0.0 (1) Modern
S. Atlantic counter heat piracy
(2) CLIMAP LGM (3) Cold tropics LGM (4) Heinrich event (HL)
_ 0.4
(5) Heinrich event (HB) (6) D/O event
_ 0.8 _ 40 N ˚
0˚N
40˚S Latitude
80˚S
120˚S
Figure 4 Atlantic Ocean poleward heat transport (positive indicates a northward movement) as given by the ocean circulation model (Seidov and Maslin, 1999) for the following scenarios: (1) present-day (warm interglacial) climate; (2) last glacial maximum (LGM) with generic CLIMAP data; (3) ‘Cold tropics’ LGM scenario; (4) a Heinrich-type event driven by the meltwater delivered by icebergs from decaying Laurentide ice sheet; (5) a Heinrich-type event driven by meltwater delivered by icebergs from decaying Barents Shelf ice sheet or Scandinavian ice sheet; (6) a general Holocene or glacial Dansgaard–Oeschger (D/O) meltwater confined to the Nordic Seas. Note that the total meridional heat transport can only be correctly mathematically computed in the cases of cyclic boundary conditions (as in Drake Passage for the global ocean) or between meridional boundaries, as in the Atlantic Ocean to the north of the tip of Africa. Therefore the northward heat transport in the Atlantic ocean is shown to the north of 301S only.
(c) 2011 Elsevier Inc. All Rights Reserved.
130
HOLOCENE CLIMATE VARIABILITY
Holocene. One of the most important finds in the study of glacial millennial-scale events is the apparent out-of-phase climate response of the two hemispheres seen in the ice core climate records from Greenland and Antarctica. It has been suggested that this bipolar seesaw can be explained by variations in the relative amount of deep water formation in the two hemispheres and heat piracy (Figure 4). This mechanism of altering dominance of the NADW and the Antarctic Bottom Water (AABW) can also be applied to the Holocene. The important difference with this theory is that the trigger for a sudden ‘switching off’ or a strong decrease in rate of deep water formation could occur either in the North Atlantic or in the Southern Ocean. AABW forms in a different way than NADW, in two general areas around the Antarctic continent: (1) near-shore at the shelf–ice, sea-ice interface and (2) in open ocean areas. In near-shore areas, coastal polynya are formed where katabatic winds push sea ice away from the shelf edge, creating further opportunity for sea ice formation. As ice forms, the surface water becomes saltier (owing to salt rejection by the ice) and colder (owing to loss of heat via latent heat of freezing). This density instability causes sinking of surface waters to form AABW, the coldest and
saltiest water in the world. AABW can also form in open-ocean Antarctic waters; particularly in the Weddell and Ross seas; AABW flows around Antarctica and penetrates the North Atlantic, flowing under the less dense NADW. It also flows into the Indian and Pacific Oceans, but the most significant gateway to deep ocean flow is in the south-west Pacific, where 40% of the world’s deep water enters the Pacific. Interestingly, Seidov and colleagues have shown that the Southern Ocean is twice as sensitive to meltwater input as is the North Atlantic, and that the Southern Ocean can no longer be seen as a passive player in global climate change. The bipolar seesaw model may also be self sustaining, with meltwater events in either hemisphere, triggering a train of climate changes that causes a meltwater event in the opposite hemisphere, thus switching the direction of heat piracy (Figure 5).
Conclusion The Holocene, or the last 10 000 years, was once thought to be climatically stable. Recent evidence, including that from marine sediments, have altered this view, showing that there are millennial-scale
Internal ice sheet oscillation ?
North Atlantic
Heinrich Events
Reduced NADW
Antarctic Warm Events Antartica melting or surging?
Southern Ocean warms up
S. Hemisphere heat piracy
1500 year D/O cycle
Meltwater
Warmer southern ocean
S. Hemisphere heat piracy
Reduced NADW
Southern Ocean
Laurentide ice sheet collapse
Sea level
Meltwater
Deep water oscillating system
Greenland / Iceland melting or surging?
Warmer N. Atlantic
N. Hemisphere heat piracy
Increased NADW
Reduced AABW
? Internal ice sheet oscillation
Figure 5 Possible deep water oscillatory system explaining the glacial and interglacial Dansgaard–Oeschger cycles. Additional loop demonstrates the possible link between interglacial Dansgaard–Oeschger cycles and Heinrich events.
(c) 2011 Elsevier Inc. All Rights Reserved.
HOLOCENE CLIMATE VARIABILITY
Solar cycles North Atlantic Oscillation
100
D/O cycles
Heinrich events
0.1
Obliquity
1
Precession
10 Eccentricity
Relative variance
Annual cycle
Millennial-scale quasiperiodicities?
0.01
El Nino_Southern Oscillation
Astronomical forcing
1000
131
? T h e _6
10
_5
10
_
‘ g a p ’ _
_
_1
10 4 10 3 10 2 10 _ Frequency (cycles y 1)
100
Figure 6 Spectrum of climate variance showing the climatic cycles for which we have good understanding and the ‘gap’ between hundreds and thousands of years for which we still do not have adequate understanding of the causes.
climate cycles throughout the Holocene. In fact we are still in a period of recovery from the last of these cycles, the Little Ice Age. It is still widely debated whether these cycles are quasiperiodic or have a regular cyclicity of 1500 years. It is also still widely debated whether these Holocene Dansgaard–Oeschger cycles are similar in time and characteristic to those observed during the last glacial period. A number of different theories have been put forward for the causes of these Holocene climate cycles, most suggesting variations in the deep water circulation system. One suggestion is that these cycles are caused by the oscillating relative dominance of North Atlantic Deep Water and Antarctic Bottom Water. Holocene climate variability still has no adequate explanation and falls in the ‘gap’ of our knowledge between Milankovitch forcing of ice ages and rapid variations such as El Nin˜o and the North Atlantic Oscillation (Figure 6). Future research is essential for understanding these climate cycles so that we can better predict the climate response to anthropogenic ‘global warming.’
See also Antarctic Circumpolar Current. Calcium Carbonates. Cenozoic Oceans – Carbon Cycle Models. Current Systems in the Atlantic Ocean. Deep Convection. Deep-Sea Drilling Methodology. Deep-Sea Drilling Results. MillennialScale Climate Variability. Ocean Circulation. Ocean Subduction. Ocean Margin Sediments. Paleoceanography. Sediment Chronologies. Water Types and Water Masses.
Further Reading Adams J, Maslin MA, and Thomas E (1999) Sudden climate transitions during the Quaternary. Progress in Physical Geography 23(1): 1--36. Alley RB and Clark PU (1999) The deglaciation of the Northern hemisphere: A global perspective. Annual Review of Earth and Planetary Science 27: 149--182. Bianchi GG and McCave IN (1999) Holocene periodicity in North Atlantic climate and deep-ocean flow south of Iceland. Nature 397: 515--523. Broecker W (1998) Paleocean circulation during the last deglaciation: a bipolar seesaw? Paleoceanography 13: 119--121. Bond G, Showers W, Cheseby M, et al. (1997) A pervasive millenial-scale cycle in North Atlantic Holocene and glacial climates. Science 278: 1257--1265. Chapman MR and Shackleton NJ (2000) Evidence of 550 year and 1000 years cyclicities in North Atlantic pattern during the Holocene. Holocene 10: 287--291. Cullen HM, et al. (2000) Climate change and the collapse of the Akkadian Empire: evidence from the deep sea. Geology 28: 379--382. Dansgaard W, Johnson SJ, and Clausen HB (1993) Evidence for general instability of past climate from a 250-kyr ice-core record. Nature 364: 218--220. deMenocal P, Ortiz J, Guilderson T, and Sarnthein M (2000) Coherent high- and low-latitude climate variability during the Holocene warm period. Science 288: 2198--2202. Keigwin LD (1996) The Little Ice Age and Medieval warm period in the Sargasso sea. Science 274: 1504--1507. Maslin MA, Seidov D, and Lowe J (2001) Synthesis of the nature and causes of sudden climate transitions during the Quaternary. AGU Monograph: Oceans and Rapid Past and Future Climate Changes: North–South Connections. Washington DC: American Geophysical. Union No. 119.
(c) 2011 Elsevier Inc. All Rights Reserved.
132
HOLOCENE CLIMATE VARIABILITY
O’Brien SR, Mayewski A, and Meeker LD (1996) Complexity of Holocene climate as reconstructed from a Greenland ice core. Science 270: 1962--1964. Pearce RB, Kemp AES, Koizumi I, Pike J, Cramp A, and Rowland SJ (1998) A lamina-scale, SEM-based study of a late quaternary diatom-ooze sapropel from the Mediterranean ridge, Site 971. In: Robertson AHF, Emeis K-C, Richter C, and Camerlenghi A (eds.) Proceedings of the Ocean Drilling Program, Scientific Results 160, 349--363. Peiser BJ (1998) Comparative analysis of late Holocene environmental and social upheaval: evidence for a disaster around 4000 BP. In: Peiser BJ, Palmer T, and Bailey M (eds.) Natural Catastrophes during Bronze Age Civilisations, 117--139.
Pike J and Kemp AES (1997) Early Holocene decadal-scale ocean variability recorded in Gulf of California laminated sediments. Paleoceanography 12: 227--238. Seidov D and Maslin M (1999) North Atlantic Deep Water circulation collapse during the Heinrich events. Geology 27: 23--26. Seidov D, Barron E, Haupt BJ, and Maslin MA (2001) Meltwater and the ocean conveyor: past, present and future of the ocean bi-polar seesaw. AGU Monograph: Oceans and Rapid Past and Future Climate Changes: North–South Connections. Washington DC: American Geophysical. Union No. 119.. Wunsch C (2000) On sharp spectral lines in the climate record and millennial peak. Paleoceanography 15: 417--424.
(c) 2011 Elsevier Inc. All Rights Reserved.
HYDROTHERMAL VENT BIOTA R. A. Lutz, Rutgers University, New Brunswick, NJ, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1217–1227, & 2001, Elsevier Ltd.
On 17 February 1977, the deep-submergence vehicle Alvin descended 2500 m to the crest of the Galapagos Rift spreading center to first visit an ecosystem that would forever change our view of life in the deep sea. Cracks and crevices in the ocean floor were emanating fluids with temperatures up to 171C. None of the bizarre organisms clustering around these ‘hydrothermal vents’ had ever been encountered; they comprised new species, genera, families, superfamilies, and bizarre ‘tubeworms,’ up to 2 m long, which were subsequently placed in a new phylum (Vestimentifera) (Figure 1). Since the Galapagos Rift discovery, numerous hydrothermal vent sites have been found throughout the world’s oceans and over 500 new species have been described from these regions. Figure 2 depicts many of the major hydrothermal systems from which organisms have been collected to date. Fluids with temperatures as high as 4031C exit from polymetallic sulfide chimneys in many of these regions (Figure 3). Most ecosystems on earth ultimately rely on photosynthesis, with the energy source being solar. In marked contrast, deep-sea hydrothermal ecosystems
are based predominantly on chemosynthesis, with the energy source being geothermal. Many of the chemosynthetic microbes are fueled by hydrogen sulfide, which is present at low-temperature vents in concentrations up to several hundred micromoles per liter and at high-temperature vents in concentrations up to 100 milimoles per liter. These microbial organisms can be either ‘free-living’ (in the water or on the surface of various substrates) or symbiotic in association with certain vent organisms. The vestimentiferan tubeworms Riftia pachyptila and Tevnia jerichonana (Figures 1, 4, 5, and 6), for example, each have a specialized ‘tissue,’ known as the trophosome, which is comprised entirely of chemosynthetic bacteria. The tubeworms have no mouth, no digestive system, and no anus; in short, no opening to the external environment. Hydrogen sulfide diffuses across cell membranes and is transported via the hemoglobin-containing circulatory system to the trophosome, where it is utilized by the associated symbionts. Mussels (Bathymodiolus thermophilus) (Figures 7 and 8) and vesicomyid clams (Calyptogena magnifica) (Figure 9), common along both the Galapagos Rift and East Pacific Rise (EPR), represent two of the other dominant members of the vent megafauna that house chemosynthetic symbionts. In the case of each of these bivalves, the symbionts are associated with the gills and both species have modified feeding apparatuses relative to those of shallow-water related species (likely a result of their
Figure 1 A cluster of vestimentiferan tubeworms (Riftia pachyptila and Tevnia jerichonana), together with a zoarcid fish (Thermarces andersoni) and brachyuran crabs (Bythograea thermydron) inhabiting low-temperature hydrothermal vents at 91500 N along the East Pacific Rise (depth 2500 m).
(c) 2011 Elsevier Inc. All Rights Reserved.
133
134
HYDROTHERMAL VENT BIOTA
Lucky Strike
Southern Explorer Ridge (SEXP) Juan de Fuca Ridge (JDFR) Gorda Ridge (GORD) Sunrise Okinawa Trough
Mariana Trough Manus Basin
Rainbow Broken Spur
Guaymas Basin
TAG
Galapagos′ Rift (GAR)
Carlsberg Ridge Central Indian Ridge
Ri
se
e
dg
A R)
Ch
ile
Red Sea
Snake Pit Logatchev
ic Ridge (M
21° 13° Northern East 11° Pacific Rise (NEPR) 9° 7° Lau Basin 11° 14° 16° 17° 18° 20° 21° Southern East 22° Pacific Rise (SEPR) 31° 32°
Mid Atlant
North Fiji Basin
Menez Gwen
n dia
Ri
t In
as
e uth
Southeast Indian Ridge So
ge
arctic Rid
nt Pacific A
Figure 2 Deep-sea hydrothermal vent sites along mid-oceanic and back-arc ridge systems from which vent organisms have been collected to date. Numbers indicate approximate latitude of site along the East Pacific Rise.
predominant reliance on the associated symbionts for nutrition). Closely related mussels and clams within the families Mytilidae and Vesicomyidae are common constituents of the fauna associated with vents along mid-oceanic ridge and back-arc spreading centers (as well as at many cold-water hydrocarbon seeps) throughout the world’s oceans. All of
Figure 3 A ‘black smoker’ polymetallic sulfide chimney at 201500 N along the East Pacific Rise (depth 2615 m). Temperatures as high as 4031C have been recorded at the orifice of such edifices from which mineral-rich fluids emanate violently in many deep-sea hydrothermal systems.
these bivalve mollusks retrieved to date appear to contain thiotrophic (‘sulfur-feeding’) or methanotrophic (‘methane-feeding’) chemosynthetic symbionts. It should be mentioned that, while chemosynthetic bacteria play a critical role in food chains associated with many vent systems, there are numerous other microbes at the vents belonging to two other recognized kingdoms. Both eukaryotes and Archaea occupy specialized niches within various vent ecosystems and certain Archaea, which have been isolated from environments associated with hightemperature sulfide chimneys, have been reported to occupy the most primitive ‘node’ on the phylogenetic tree of life. Such reports have led to considerable speculation as to whether or not life itself may have originated in hydrothermal vent environments. At various high-temperature vent sites, numerous organisms colonize the sides of active sulfide edifices. Figure 4 depicts a sulfide edifice (named ‘Tubeworm Pillar’) that is 11 m high, the sides of which are covered with tubeworms (both R. pachyptila and T. jerichonana), crabs, zoarcid fish, and a spectrum of smaller, associated vent fauna. At other sites, the sides of ‘black smoker’ chimneys are frequently covered with the tubes of polychaetes, such as Alvinella pompejana, a ‘Pompeii’ worm named after the
(c) 2011 Elsevier Inc. All Rights Reserved.
HYDROTHERMAL VENT BIOTA
135
Figure 4 A portion of an 11 m high polymetallic sulfide edifice known as Tubeworm Pillar at 91500 N along the East Pacific Rise (depth 2500 m). The top two-thirds of the edifice is covered with vestimentiferan tubeworms, both Riftia pachyptila (larger organisms) and Tevnia jerichonana (smaller organisms), as well as numerous brachyuran crabs (Bythograea thermydron) and zoarcid fish (Thermarces andersoni).
submersible Alvin (Figure 10). This organism has been reported to routinely withstand long-term exposure to temperatures ranging from 21C to 351C and short-term exposure to temperatures in excess of 1001C, rendering this annelid perhaps the most eurythermal organism on the planet. In addition to the numerous species of bivalves, such as mussels and clams mentioned earlier, there are myriad other common mollusks at vents, including numerous gastropods, particularly archeogastropod limpets. Thirteen different gastropod species have been reported from vents along the narrow range of the East Pacific Rise from 91170 N to
91540 N. These organisms can achieve high population densities and are found on a wide variety of substrates including the tubes of Riftia pachyptila and the shells of Bathymodiolus thermophilus (Figures 8 and 11). Limpets graze on free-living microbes that coat the majority of surfaces associated with both low-temperature and high-temperature (e.g., sulfide chimney) vents. Over 100 species of mollusks have been collected to date from the mid-oceanic and back-arc spreading centers visited to date. Virtually all of the invertebrates inhabiting deepsea hydrothermal vents have planktonic larval stages. These free-swimming stages serve as the primary
Figure 5 Higher magnification of the side of Tubeworm Pillar (depicted in Figure 4), showing tubeworms (Riftia pachyptila and Tevnia jerichonana) and scavenging brachyuran crabs (Bythograea thermydron).
(c) 2011 Elsevier Inc. All Rights Reserved.
136
HYDROTHERMAL VENT BIOTA
Figure 6 Close-up image of a cluster of Tevnia jerichonana, together with a brachyuran crab (Bythograea thermydron) and a zoarcid fish (Thermarces andersoni). Note the ‘accordion-like’ morphology of the tube of this species of vestimentiferan tubeworm.
Figure 7 A dense population of mussels (Bathymodiolus thermophilus) inhabiting a low-temperature hydrothermal vent field along the East Pacific Rise. Associated fauna in the field of view include tubeworms (Riftia pachyptila), brachyuran crabs (Bythograea thermydron), zoarcid fish (Thermarces andersoni), and a galatheid crab (Munidopsis subsquamosa) (lower left).
means of dispersal between isolated vent systems, which can be separated by hundreds of kilometers. The larval stages can be either planktotrophic (feeding within the water column) or lecithotrophic (utilizing yolk reserves for nutrition). In either case, it appears that the early life history stages of the vast majority of vent organisms are capable of staying in the water column for considerable lengths of time (likely months in the case of many vent species). Such extended planktonic durations facilitate passive transport over vast distances via ocean currents (velocities of ocean currents between 15 and 30 cm s1 are commonly encountered along the crest of ridge systems). Numerous crustaceans inhabit the vent environment and represent perhaps the dominant scavengers
Figure 8 Close-up of mussels (Bathymodiolus thermophilus) attached to the tubes of the tubeworm Riftia pachyptila. Limpets (Lepetodrilus elevatus) are seen attached to the external surfaces of both the mussel shells and tubeworm tubes.
(c) 2011 Elsevier Inc. All Rights Reserved.
HYDROTHERMAL VENT BIOTA
137
Figure 9 Vesicomyid clams (Calyptogena magnifica) line cracks and crevices from which low-temperature fluids are venting in an area known as Clam Acres at 201500 N along the East Pacific Rise (depth 2615 m). This species is common at many vent sites that are believed to be in relatively late stages of succession along the Galapagos Rift and the entire stretch of the East Pacific Rise from 211N to 181S.
of the ecosystem. Brachyuran crabs (e.g., Bythograea thermydron on the Galapagos Rift and the East Pacific Rise) (Figures 1, 5, 6, and 7) appear to be one of the earliest colonizers of hydrothermal vent environments. When new vents were formed during the April, 1991 volcanic eruption along the crest of the East Pacific Rise at 91500 N, several regions were referred to as ‘crab nurseries’ owing to the relatively large abundance of crab larvae (megalopae) in areas
of low-temperature discharge. Eleven months after the eruptive event, the region was populated by tremendous numbers of large crabs. These were frequently observed to be holding with their claws various substrates, such as empty tubeworm tubes or pieces of basalt covered with microbial mats, and voraciously scraping the surfaces with their mouth parts, suggesting that microbes and/or their products may represent an important source of nutrition for
Figure 10 The polychaete Alvinella pompejana emerging from its tube on the side of a black smoker chimney at 91500 N along the East Pacific Rise. It has been suggested that the organism may be the most eurythermal invertebrate on the planet, capable of withstanding short-term exposures to temperatures ranging from less than 21C to in excess of 1001C.
(c) 2011 Elsevier Inc. All Rights Reserved.
138
HYDROTHERMAL VENT BIOTA
Figure 11 Limpets (Lepetodrilus elevatus) coating the surfaces of the tubes of Riftia pachyptila and the shell of Bathymodiolus thermophilus (left). A zoarcid fish (Thermarces andersoni) is seen emerging from the tubeworm tubes.
these organisms. They are also frequently seen extending their claws into the tubes of Riftia pachyptila and will readily devour the tissues of any tubeworm or mussel damaged during routine submersible activities. They are easily captured in ‘crab traps’ baited with a wide variety of dead fish. Intraspecific attacks appear to be relatively common occurrences and crabs with only one claw or a missing leg are frequently seen on the bottom, presumably reflecting an encounter with another crab. Galatheid crabs (e.g., Munidopsis subsquamosa) (Figure 12) inhabit peripheral areas of the vents during earlier stages of succession and, as hydrogen sulfide levels gradually decrease over time, they are
commonly encountered in central areas as well, often among tubeworms and mussels. While these ‘squat lobsters’ may also be scavengers, they are seldom, if ever, caught in traps. One chance encounter with a relatively large dead octopod in the peripheral area of a vent revealed a large quantity of galatheid crabs on the carcass and a noticeable absence of brachyuran crabs or other vent organisms (Figure 13). Close-up imagery revealed that the crabs were actively feeding on the dead tissues of the carcass. Galatheids that have been collected to date have generally been ‘transported’ back to the surface in random regions of the submersible itself. The carapaces of this species are frequently covered with fine, filamentous bacteria (Figure 12).
Figure 12 A galatheid crab (Munidopsis subsquamosa) with filamentous bacteria on its carapace. This species is a common inhabitant of peripheral vent environments and is also frequently observed in ambient deep-sea environments.
(c) 2011 Elsevier Inc. All Rights Reserved.
HYDROTHERMAL VENT BIOTA
Figure 13 Numerous galatheid crabs (Munidopsis subsquamosa) on and in the immediate vicinity of the carcass of a cirrate octopod approximately 50 m from an active hydrothermal vent at 91500 N along the East Pacific Rise. Closeup imagery revealed that the crabs were actively feeding on the dead tissues of the carcass.
Numerous species of shrimp (Figure 14) have been encountered at vent sites visited to date along ridge axes in all ocean basins. At several sites (e.g., the TAG hydrothermal vent field) along the Mid-Atlantic Ridge, thousands of shrimp are frequently seen ‘swarming’ on top of one another, completely carpeting the sides of large sulfide chimneys. It has been suggested that a large ‘eye spot’ on the back of the carapace of the shrimp is capable of sensing longwavelength light emitted by high-temperature smokers. Shrimp are also relatively common inhabitants
139
of vents visited to date in the western Pacific (e.g., at the Alice Springs vent field along the back-arc spreading center in the Mariana Trough). Other smaller, common vent-endemic crustaceans include numerous species of amphipods, copepods, and leptostracans. Halice hesmonectes, an amphipod common at various vent sites along the East Pacific Rise is frequently seen ‘swarming’ above mussel and tubeworm colonies in regions of active low-temperature venting. It has been reported that these dense swarms represent the highest concentration of planktonic invertebrates in the ocean. Serpulid polychaetes (Figure 15) are common inhabitants of the peripheral area of many vent fields. Commonly referred to as ‘feather dusters,’ they were seen extending over large expanses of lava during the early expeditions to the Galapagos Rift in 1977 and 1979 and were subsequently reported at numerous vents along the East Pacific Rise. Their small tubes, which generally reach lengths of about 5 cm, consist of calcite. When the tentacular plumes are withdrawn into the tube, a small ‘plug’ seals the tube from the external environment. Numerous species of apparently vent-endemic fish have been reported from hydrothermal systems along ridge systems throughout the world’s oceans. Bythites hollisi is a bythitid that was encountered on the first dive to the Galapagos Rift vent field in 1977. Bythitids have been observed in large numbers at several other vent sites, such as the hydrothermal field at 91500 N along the East Pacific Rise (Figure 16). One large ‘pit,’ with a diameter of several meters, from which cloudy, shimmering water was emanating, had a concentration of over 20 bythitids,
Figure 14 The shrimp Alvinocaris lusca perched atop a tube of Riftia pachyptila at a low temperature vent along the East Pacific Rise.
(c) 2011 Elsevier Inc. All Rights Reserved.
140
HYDROTHERMAL VENT BIOTA
Figure 15 Serpulid polychaetes in the peripheral area of a vent field at 91500 N along the East Pacific Rise.
which were frequently observed with their heads projecting downward into the cloudiest portions of the water at the base of the pit. Zoarcids (eel pouts) are common members of the vent fauna at various sites along the East Pacific Rise and the Mid-Atlantic Ridge, with several different species having been encountered at the various vent fields visited to date. The common species that inhabits many East Pacific Rise vents is Thermarces andersoni (Figures 1, 4, 6, and 11). From analyses of extensive video footage and stomach contents of retrieved specimens of this species, it appears to commonly feed on a wide range of organisms, including bacteria, amphipods, leptostracans, and shrimp. It has also been observed scavenging on the plumes of specimens of vestimentiferan tubeworms that have
been damaged in the process of sampling or maneuvering with the submersible. Enteroptneusts, commonly referred to as ‘spaghetti worms’ were first observed in 1977 draped over pillow lava in peripheral regions of the Galapagos Rift vent field. Few individuals were observed at any of the other many vent fields visited over the next 20 years throughout the world. In 1997, six years after the volcanic eruption at 91500 N along the East Pacific Rise, numerous enteroptneusts were observed at distances ranging from a few meters to a few hundred meters from active vent sites within the region. No individual organisms were observed living in direct association with venting fluids or at distances in excess of a kilometer from active hydrothermal systems. This unusual, soft-bodied
Figure 16 The vent fish Bythites hollisi emerging from a high-temperature vent region known as Hole-to-Hell at 91500 N along the East Pacific Rise. The species is common at various vent fields along the Galapagos Rift and East Pacific Rise.
(c) 2011 Elsevier Inc. All Rights Reserved.
HYDROTHERMAL VENT BIOTA
Figure 17 The enteroptneust Saxipendium coronatum (commonly known as a ‘spaghetti worm’) on the surface of basalt about 30 m from an active low-temperature vent field at 91500 N along the East Pacific Rise. The image was taken with a prototype high-resolution video camera system equipped with a macro-lens.
invertebrate may represent an organism that is uniquely adapted to an ecotone between active vents and the ambient deep sea. Analyses of video images of numerous individuals taken with a high-resolution video camera system equipped with a macro-lens (Figure 17) have revealed behavioral patterns suggesting that the organism may be ‘grazing’ directly on basaltic surfaces, potentially consuming microbes or organic substances ultimately originating in the vent environment. One of the most unusual and spectacular organisms inhabiting vent ecosystems is the vestimentiferan tubeworm Riftia pachyptila which thrives at numerous vent fields along the Galapagos Rift and East Pacific Rise (Figures 1, 4, and 8). As mentioned earlier, it lives in a symbiotic relationship with chemosynthetic bacteria concentrated within its body. Such a relationship provides an internal, hydrogen sulfide-nourished ‘garden’ that, in turn, nourishes the tubeworm. Although the mechanism by which the host obtains energy from the bacteria is unclear, the energy transfer appears remarkably efficient. An unique opportunity to determine the growth rate potential of R. pachyptila arose as a result of the April, 1991 volcanic eruptive event at 91500 N along the East Pacific Rise mentioned above. In March,
Figure 18 (Right)Temporal sequence of vent community development at a low-temperature vent in a region known as Hole-to-Hell at 91500 N along the East Pacific Rise. This was the site of a volcanic eruption in April, 1991 that entirely decimated previously existing communities within the region. The field of view of each image and the heading of the camera system utilized are approximately the same for each image taken at the following times: (A) April, 1991; (B) March, 1992; (C) December, 1993; (D) October, 1994; and (E) November, 1995. (Shank et al., 1998.)
(c) 2011 Elsevier Inc. All Rights Reserved.
141
142
HYDROTHERMAL VENT BIOTA
1992, no R. pachyptila were present within the region, which had been devastated by the eruption. By December, 1993, less than 2 years later, huge colonies of this tubeworm had colonized the active lowtemperature vents; the tube lengths of many individuals were in excess of 1.5 m (Figure 18 and 19). Such growth rates of more than 85 cm y1 increase in tube length represent the fastest rates of growth documented for any marine invertebrate. It is interesting to note that one of the other fastest-growing marine invertebrates is the giant clam Tridacna squamosa, a bivalve that has a symbiotic association with photosynthetic algae (zooxanthellae). The efficiency with which energy is transferred from the symbiont to the host in certain invertebrates may well be a contributing factor to the remarkable growth rates of these organisms. The rapid succession of a tubeworm-dominated vent community over the 5-year period following the April, 1991 eruption is dramatically illustrated in Figures 18 and 19. Inhabitants of deep-sea hydrothermal vents are among the most spectacular and unusual organisms on the planet. Given the relatively small number of vent ecosystems found to date, many questions have been raised concerning whether conservation measures need to be taken to protect vent communities from anthropogenic disturbances. Any assessment of the potential consequences of anthropogenic impacts needs to consider that the organisms inhabiting deepsea hydrothermal vents thrive in an environment that is constantly being altered radically by geological process. Periodic devastation of entire biological communities is a relatively common occurrence as volcanic and tectonic processes proceed along active ridge axes. Over the past two decades we have learned that many vent communities are remarkably resilient. Populations of essentially all vent organisms indigenous to regions decimated by massive volcanic eruptions have, like a Phoenix rising from the ashes, reestablished themselves in less than a decade. This remarkable resilience in the face of huge natural disasters has profound implications as one considers the potential impacts of exploitation of
Figure 19 (Left)Temporal sequence of vent community development at a low-temperature vent located approximately 500 m from the community depicted in Figure 18. This was also a region that was buried with fresh lava by the April, 1991 volcanic eruptive event, decimating previously existing communities within the area. The field of view of each image and the heading of the camera system utilized are approximately the same for each image taken at the following times: (A) April, 1991; (B) March, 1992; (C) December, 1993; (D) October, 1994; and (E) November, 1995. (Shank et al., 1998.)
(c) 2011 Elsevier Inc. All Rights Reserved.
HYDROTHERMAL VENT BIOTA
precious mineral and biological resources associated with active hydrothermal systems throughout the world’s oceans.
Acknowledgments I thank the pilots and crew of the DSV Alvin and the R/V Atlantis for their expertise, assistance, and patience over the years; W. Lange and the Woods Hole Oceanographic Institution for technical expertise and the provision of the camera and recording systems critical to the generation of the majority of images presented in this article; Emory Kristof, Stephen Low, and Michael V. DeGruy for inspiration and assistance with the procurement of video images using a variety of camera systems; and Matt Tieger for assistance in the generation of video prints. Supported by National Science Foundation Grants OCE-95-29819 and OCE-96-33131.
See also Deep-Sea Ridges, Microbiology. Hydrothermal Vent Ecology Hydrothermal Vent Fauna, Physiology of
143
Further Reading Childress JJ and Fisher CR (1992) The biology of hydrothermal vent animals: physiology, biochemistry, and autotrophic symbioses. Oceanography and Marine Biology Annual Review 30: 337--441. Gage JD and Tyler PA (1991) Deep-Sea Biology: A Natural History of Organisms at the Deep-Sea Floor. Cambridge: Cambridge University Press. Lutz RA (2000) Deep sea vents. National Geographic 198(4): 116--127. Shank TM, Fornari DJ, and Von Damm KL (1998) Temporal and spatial patterns of biological community development at nascent deep-sea hydrothermal vents (91N, East Pacific Rise). Deep-Sea Research. 45: 465– 515. Jones ML (ed.) (1985) The Hydrothermal Vents of the Eastern Pacific: An Overview. Bulletin of the Biological Society of Washington, vol. 6. Washington, DC: Biological Society of Washington. Rona PA, Bostrom K, Laubier L, and Smith KL Jr (eds.) (1983) Hydrothermal Processes at Seafloor Spreading Centers. New York: Plenum Press. Tunnicliffe V (1991) The biology of hydrothermal vents: ecology and evolution. Oceanography and Marine Biology Annual Review 29: 319--417. Van Dover CL (2000) The Ecology of Deep-Sea Hydrothermal Vents. Princeton, NJ: Princeton University Press, Princeton.
(c) 2011 Elsevier Inc. All Rights Reserved.
HYDROTHERMAL VENT DEPOSITS R. M. Haymon, University of California, CA, USA Copyright & 2001 Elsevier Ltd. All rights reserved. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1228 –1234, & 2001, Elsevier Ltd.
Where Deposits Form: Geologic Controls
Introduction In April 1979, submersible divers exploring the midocean ridge crest at latitude 211N on the East Pacific Rise discovered superheated (3807301C) fluids, blackened by tiny metal-sulfide mineral crystals, spewing from the seafloor through tall mineral conduits (see Hydrothermal Vent Biota, Hydrothermal Vent Fluids, Chemistry of). The crystalline conduits at these ‘black smoker’ hydrothermal vents were made of minerals rich in copper, iron, zinc, and other metals. Since 1979, hundreds of similar hydrothermal deposits have been located along the midocean ridge. It is now clear that deposition of hydrothermal mineral deposits is a common process, and is integrally linked to cracking, magmatism, and cooling of new seafloor as it accretes and spreads away from the ridge (see Propagating Rifts and Microplates, Mid-Ocean Ridge Geochemistry and Petrology, Seamounts and Off-Ridge Volcanism, Mid-Ocean Ridge Seismic Structure). For thousands of years before mid-ocean ridge hot springs were discovered in the oceans, people mined copper from mineral deposits that were originally formed on oceanic spreading ridges. These fossil deposits are embedded in old fragments of seafloor called ‘ophiolites’ that have been uplifted and emplaced onto land by fault movements. The copperrich mineral deposits in the Troodos ophiolite of Cyprus are well-known examples of fossil oceanridge deposits that have been mined for at least 2500 years; in fact, the word ‘copper’ is derived from the Latin word ‘cyprium’ which means ‘from Cyprus.’ The mineral deposits accumulating today at hot springs along the mid-ocean ridge are habitats for a variety of remarkable organisms ranging in size from tiny microbes to large worms (see Hydrothermal Vent Biota, Deep-Sea Ridges, Microbiology, Hydrothermal Vent Fauna, Physiology of). The properties of the mineral deposits are inextricably linked to the organisms that inhabit them. The mineral deposits contain important clues about the physical–chemical environments in which some of these organisms live, and also preserve fossils of some organisms, creating a geologic record of their existence.
144
Hydrothermal vent deposits are thus a renewable source of metals and a record of the physical, chemical, biological, and geological processes at modern and ancient submarine vents.
Less than 2% of the total area of the mid-ocean ridge crest has been studied at a resolution sufficient to reveal the spatial distribution of hydrothermal vents, mineral deposits, and other significant small-scale geologic features. Nevertheless, because study areas have been carefully selected and strategically surveyed, much has been learned about where vents and deposits form, and about the geologic controls on their distribution. The basic requirements for hydrothermal systems include heat to drive fluid circulation, and high-permeability pathways to facilitate fluid flow through crustal rocks. On midocean ridges, vents and deposits are forming at sites where ascending magma intrusions introduce heat into the permeable shallow crust, and at sites where deep cracks provide permeability and fluid access to heat sources at depth.
Fast-spreading Ridges Near- and on-bottom studies along the fast-spreading East Pacific Rise suggest that most hydrothermal mineral deposits form along the summit of the ridge crest within a narrow ‘axial zone’ less than 500 m wide. Only a few active sites of mineral deposition have been located outside this zone; however, more exploration of the vast area outside the axial zone is needed to establish unequivocally whether or not mineral deposition is uncommon in this region. The overall spatial distribution of hydrothermal vents and mineral deposits along fast-spreading ridges traces the segmented configuration of cracks and magma sources along the ridge crest (see Propagating Rifts and Microplates, Mid-Ocean Ridge Geochemistry and Petrology, Seamounts and Off-Ridge Volcanism, Mid-Ocean Ridge Seismic Structure). Within the axial zone, mineral deposition is concentrated along the floors and walls of axial troughs created by volcanic collapse and/or faulting along the summit of the ridge crest. The majority of the deposits are located along fissures that have opened above magmatic dike intrusions, and along collapsed
(c) 2011 Elsevier Inc. All Rights Reserved.
HYDROTHERMAL VENT DEPOSITS
lava ponds formed above these fissures by pooling and drainage of erupted lava. Where fault-bounded troughs have formed along the summit of the ridge crest, mineral deposition is focused along the bounding faults and also along fissures and collapsed lava ponds in the trough floor. Hydrothermal vents appear to be most abundant along magmatically inflated segments of fast-spreading ridges; however, the mineral deposits precipitated on the seafloor on magmatically active segments are often buried beneath frequent eruptions of new lava flows. The greatest number of deposits, therefore, are observed on inflated ridge segments that are surfaced by somewhat older flows, i.e., along segments where: (1) much heat is available to power hydrothermal vents; and (2) mineral deposits have had time to develop but have not yet been buried by renewed eruptions.
Intermediate- and Slow-spreading Ridges Most hydrothermal deposits that have been found on intermediate- and slow-spreading ridge crests are focused along faults, fissures, and volcanic structures within large rift valleys that are several kilometers wide. The fault scarps along the margins of rift valleys are common sites for hydrothermal venting and mineral deposition. Fault intersections are thought to be particularly favorable sites for hydrothermal mineral deposition because they are zones of high permeability that can focus fluid flow. Mineral deposition on rift valley floors is observed along fissures above dike intrusions, along eruptive fissures and
Dead edifice
Flanges
volcanic collapse troughs, and on top of volcanic mounds, cones and other constructions. In general at slower-spreading ridges, faults appear to play a greater role in controlling the distribution of hydrothermal vents and mineral deposits than they do at fast-spreading ridges, where magmatic fissures are clearly a dominant geologic control on where vents and deposits are forming.
Structures, Morphologies, and Sizes of Deposits A typical hydrothermal mineral deposit on an unsedimented mid-ocean ridge accumulates directly on top of the volcanic flows covering the ridge crest. On sedimented ridges, minerals are deposited within and on top of the sediments. Beneath seafloor mineral deposits are networks of feeder cracks through which fluids travel to the seafloor. Precipitation of hydrothermal minerals in these cracks and in the surrounding rocks or sediments creates a subseafloor zone of mineralization called a ‘stockwork’. In hydrothermal systems where fluid flow is weak, unfocused, or where the fluids mix extensively with sea water beneath the seafloor, most of the minerals will precipitate in the stockwork rather than on the seafloor. Hydrothermal deposits on mid-ocean ridges are composed of: (1) vertical structures, including individual conduits known as ‘chimneys’ (Figure 1) and larger structures of coalesced conduits that are often called ‘edifices’; (2) horizontal ‘flange’ structures that extend outward from chimneys and edifices
Black smoke plume
Fossil worm tubes
145
Black smoker chimneys
T decreasing, mixing increasing
White smoke plume White smoker edifices Live alvinelline worm Basal mound
Basalt Figure 1 Composite sketch of the mineral structures and zones in hydrothermal mineral deposits on unsedimented ridge crests (modified after Haymon, 1989). Although mound interiors are seldom observed on the seafloor, the simplified sketch of mineral zoning within the mound is predicted by analogy with chimneys and massive sulfide deposits exposed in ophiolites. An outer peripheral zone (unshaded) of anhydrite þ amorphous silica þ Zn-rich sulfide, dominantly ZnS þ FeS2, is replaced in the interior by an inner zone (hatched) of Cu-rich sulfide (CuFeS2 þ FeS2) þ minor anhydrite and amorphous silica. The inner zone may be replaced by a basal zone (cross-pattern) of Cu-rich sulfide (CuFeS2 þ FeS2) þ quartz. Zones migrate as thermochemical conditions within the mound evolve. Although not shown here, it is expected that zoning around individual fractures cutting through the mound will be superimposed on the simplified zone structure in this sketch.
(c) 2011 Elsevier Inc. All Rights Reserved.
146
HYDROTHERMAL VENT DEPOSITS
at a particular site, which depends on the nature of the heat source and plumbing system, and the rate of seafloor spreading; frequency with which deposits are buried beneath lava flows; and the compositions of the vent fluids and minerals. The large deposits found on slower-spreading ridge crests are located on faults that have moved slowly away from the ridge axis and have experienced repeated episodes of venting and accumulated mineral deposition over thousands of years, without being buried by lava flows. The tall Endeavor Segment edifices are formed because ammonia-enriched fluid compositions favor precipitation of silica in the edifice walls. The silica is strong enough to stabilize these structures so that they do not collapse as they grow taller.
How Do Chimneys Grow? A relatively simple two-stage inorganic growth model has been advanced to explain the basic characteristics of black smoker chimneys (Figure 2). In this model, a chimney wall composed largely of anhydrite (calcium sulfate) precipitates initially from sea water that is heated around discharging jets of hydrothermal fluid. The anhydrite-rich chimney wall precipitated during stage I contains only a small Cross-section chimney wall +
+
+
+
+
+
+ +
Hydrothermal fluid
+ +
+ +
+ +
+ +
+ +
+
Pillow basalt (A)
Arrows indicate directions of growth
+ + +
Anhydrite ± (caminite) + FeS2 and Zn(Fe)S Fine-grained pyrrhotite + pyrite + sphalerite intergrowths in an anhydrite matrix +
+
+
+
+
+
+
+
Hydrothermal fluid
+
+
+ +
+ +
(B)
+
+
Pillow basalt + + +
(Figure 1); (3) mounds of accumulated mineral precipitates (Figure 1); and (4) horizontal layers of hydrothermal sediments, debris, and encrustations. Chimneys are initially built directly on top of the seabed around focused jets of high-temperature effluents. Chimneys and edifices are physically unstable and often break or collapse into pieces that accumulate into piles of debris. The debris piles are cemented into consolidated mounds by precipitation of minerals from solutions percolating through the piles. New chimneys are constructed on top of the mounds as the mounds grow in size. Hydrothermal plume particles and particulate debris from chimneys settle around the periphery of the mounds to form layers of hydrothermal sediment. Diffuse seepage of fluids also precipitates mineral encrustations on mound surfaces, on volcanic flows and sediments, and on biological substrates, such as microbial mats or the shells and tubes of sessile macrofauna. The morphologies of chimneys are highly variable and evolve as the chimneys grow, becoming more complex with time. Black smoker chimneys are often colonized by organisms and evolve into ‘white smokers’ that emit diffuse, diluted vent fluids through a porous carapace of worm tubes (Figure 1). Fluid compositions and temperatures, flow dynamics, and biota are all factors that influence the development of chimney morphology. The complexity of the interactions between these factors, and the high degree of spatial–temporal heterogeneity in the physical, chemical, biological and geological conditions influencing chimney growth, account for the diverse morphologies exhibited by chimneys, and present a challenge to researchers attempting to unravel the processes producing these morphologies. The sizes of hydrothermal mineral deposits on ridges also vary widely. It has been suggested that the largest deposits are accumulating on sedimented ridges, where almost all of the metals in the fluids are deposited within the sediments rather than being dispersed into the oceans by hydrothermal plumes. On unsedimented ridges, the structures deposited on the seabed at fast spreading rates are usually relatively small in dimension (mounds are typically less than a few meters in thickness and less than tens of meters in length, and vertical structures are o15 m high). On intermediate- and slow-spreading ridges, mounds are sometimes much larger (up to tens of meters in thickness, and up to 300 m in length). On the Endeavour Segment of the Juan de Fuca Ridge, vertical structures reach heights of 45 m. The size of a deposit depends on many factors, including: magnitude of the heat source, which influences the duration of venting and mineral deposition; tendency of venting and mineral deposition to recur episodically
Chalcopyrite, or cubanite + (pyrrhotite) Cu_Fe sulfides in an anhydrite matrix As in stage I:higher sulfide:sulfate ratio
Figure 2 Two-stage model of black smoker chimney growth. (A) Stage I, sulfate-dominated stage; (B) stage II, sulfide replacement stage. During stage II, several different sulfide mineral zonation sequences develop, depending on permeability and thickness of chimney walls, hydrodynamic variables, and hydrothermal fluid composition.
(c) 2011 Elsevier Inc. All Rights Reserved.
HYDROTHERMAL VENT DEPOSITS
component of metal sulfide mineral particles that crystallize because of rapid chilling of the hydrothermal fluids. In stage II, the anhydrite-rich wall continues to grow upward and to thicken radially,
Table 1 Minerals occurring in ocean ridge hydrothermal mineral deposits
147
protecting the fluid flowing through the chimney from very rapid chilling and dilution by sea water. This allows metal sulfide minerals to precipitate into the central conduit of the chimney from the hydrothermal fluid. The hydrothermal fluid percolates outward through the chimney wall, gradually replacing anhydrite and filling voids with metal sulfide minerals. During stage II, the chimney increases in height, girth and wall thickness, and both the calcium sulfate/metal sulfide ratio and permeability of the walls decrease. Equilibration of minerals with pore fluid in the walls occurs continuously along steep, time–variant temperature and chemical gradients between fluids in the central conduit and sea water surrounding the chimney. This equilibration produces sequences of concentric mineral zones across chimney walls that evolve with changes in thermal and chemical gradients and wall permeability. The model of chimney growth described above is accurate but incomplete, as it does not include the effects on chimney development of fluid phase separation, biological activity, or variations in fluid composition. Augmented models that address these
Table 2 Ranges of elemental compositions in bulk midocean ridge hydrothermal mineral deposits
(c) 2011 Elsevier Inc. All Rights Reserved.
148
HYDROTHERMAL VENT DEPOSITS
Morphological and Mineralogical Evolution of Chimneys on the East Pacific Rise at 9˚-10˚N Stage 2
1991 "Proto-chimney" Anhydrite-dominated T = 389˚-403˚C
Stage 2
1992-1995 CuFe-sulfide-dominated T = 340˚-392˚C
1992-1995 Zn-sulfide-dominated T = 264˚-340˚C
KEY an _ anhydrite cp _ chalcopyrite po _ pyrrhotite py _ pyrite
(c) 2011 Elsevier Inc. All Rights Reserved.
HYDROTHERMAL VENT DEPOSITS
complexities are needed to fully characterize the processes governing chimney growth.
Elemental and Mineral Compositions of Deposits Ridge crest hydrothermal deposits are composed predominantly of iron-, copper- and zinc sulfide minerals, calcium- and barium-sulfate minerals, iron oxide and iron oxyhydroxide minerals, and silicate minerals (Table 1). These minerals precipitate from diverse processes, including: heating of sea water; cooling of hydrothermal fluid; mixing between sea water and hydrothermal fluid; reaction of hydrothermal minerals with fluid, sea water, or fluid–sea water mixtures; reaction between hydrothermal fluid and seafloor rocks and sediments; and reactions that are mediated or catalyzed biologically. This diversity in the processes and environments of mineral precipitation results in the deposition of many different minerals and elements (Tables 1 and 2). High concentrations of strategic and precious metals are found in some deposits (Table 2). The deposits are potentially valuable, if economic and environmentally safe methods of mining them can be developed. Chimneys can be classified broadly by composition into four groups: sulfate-rich, copper-rich, zinc-rich and silica-rich structures. Copper-rich chimney compositions are indicative of formation at temperatures above 3001C. Sulfate-rich compositions are characteristic of active and immature chimneys. Many chimneys are mineralogically zoned, with hot interior regions enriched in copper, and cooler exterior zones enriched in iron, zinc, and sulfate (Figures 1 and 2). Mounds exhibit a similar gross mineral zoning, and those which are exposed by erosion in ophiolites often have silicified (quartzrich) interiors (Figure 2). Seafloor weathering of deposits after active venting ceases results in dissolution of anhydrite, and oxidation and dissolution of metal-sulfide minerals. Small deposits that are not sealed by silicification or buried by lava flows will
149
not be well preserved in the geologic record (Figure 3).
Chimneys as Habitats Chimney and mound surfaces are substrates populated by microbial colonies and sessile organisms such as vestimentiferan and polychaete worms, limpets, mussels, and clams. It is likely that pore spaces in exterior regions of chimney walls are also inhabited by microbes. All of these organisms that are dependent on chemosynthesis benefit from the seepage of hydrothermal fluid through active mineral structures, and from the thermal and chemical gradients across mineral structures. The structures provide an interface between sea water and hydrothermal fluid that maintains tolerable temperatures for biota, and allows organisms simultaneous access to the chemical constituents in both sea water and hydrothermal fluid. However, organisms attached to active mineral structures must cope with changes in fluid flow across chimney walls (which sometimes occur rapidly), and with ongoing engulfment by mineral precipitation. Some organisms actively participate in the precipitation of minerals; for example, sulfide-oxidizing microbes mediate the crystallization of native sulfur crystals, and microbes are also thought to participate in the precipitation of marcasite and iron oxide minerals. Additionally, the surfaces of organisms provide favorable sites for nucleation and growth of amorphous silica, metal sulfide and metal oxide crystals, and this facilitates mineral precipitation and fossilization of vent fauna (Figure 3).
Fossil Record of Hydrothermal Vent Organisms Fossil molds and casts of worm tubes, mollusc shells, and microbial filaments have been identified in both modern ridge hydrothermal deposits and in Cretaceous, Jurassic, Devonian, and Silurian deposits. This fossil record establishes the antiquity of vent
Figure 3 (Left)On left: a time series of seafloor photographs showing the morphological development of a chimney that grew on top of lava flows erupted in 1991 on the crest of the East Pacific Rise near 9150.30 N (Haymon et al., 1993). Within a few days-to-weeks after the eruption, anhydrite-rich ‘Stage 1 Protochimneys’ a few cm high had formed where hot fluids emerged from volcanic outcrops covered with white microbial mats (top left). Eleven months later, the chimney consisted of cylindrical ‘Stage 2’ anhydrite-sulfide mineral spires approximately one meter in height, and as-yet unpopulated by macrofauna (middle left). Three and a half years after the eruption, the cylindrical conduits had coalesced into a 7 m-high chimneys structure that was covered with inhabited Alvinelline worms tubes (bottom left). On right: photomicrographs of chimney samples from the eruption area that show how the chimneys evolved from Stage 1 (anhydrite-dominated; top right) to Stage 2 (metal-sulfide dominated) mineral compositions (see text). As the fluids passing through the chimneys cooled below B3301C during Stage 2, the CuFe-sulfide minerals in the chimney walls (middle right) were replaced by Zn- and Fe-sulfide minerals (bottom right).
(c) 2011 Elsevier Inc. All Rights Reserved.
150
HYDROTHERMAL VENT DEPOSITS
communities and the long evolutionary history of specific faunal groups. The singular Jurassic fossil assemblage preserved in a small ophiolite-hosted deposit in central California is particularly interesting because it contains fossils of vestimentiferan worms, gastropods and brachiopods, but no clam or mussel fossils. In contrast, modern and Paleozoic faunal assemblages described thus far include clams, mussels and gastropods, but no brachiopods. Does this mean that brachiopods have competed with molluscs for ecological niches at vents, and have moved in and out of the hydrothermal vent environment over time? Fossilization of organisms is a selective process that does not preserve all the fauna that are present at vents. Identification of fossils at the species level is often difficult, especially where microbes are concerned. Notwithstanding, it is important to search for more examples of ancient fossil assemblages and to trace the fossil record of life at hydrothermal vents back as far as possible to shed light on how vent communities have evolved, and whether life on earth might have originated at submarine hydrothermal vents.
Summary Formation of hydrothermal deposits is an integral aspect of seafloor accretion at mid-ocean ridges. These deposits are valuable for their metals, for the role that they play in fostering hydrothermal vent ecosystems, for the clues that they hold to understanding spatial–temporal variability in hydrothermal vent systems, and as geologic records of how life at hydrothermal vents has evolved. From these deposits we may gain insights about biogeochemical processes at high temperatures and pressures that can be applied to understanding life in inaccessible realms within the earth’s crust or on other planetary bodies. We are only beginning to unravel the complexities of ridge hydrothermal vent deposits. Much exploration and interdisciplinary study remains to be done to obtain the valuable information that they contain.
See also Deep-Sea Ridges, Microbiology. Hydrothermal Vent Fluids, Chemistry of. Hydrothermal Vent Biota. Hydrothermal Vent Fauna, Physiology of. Mid-Ocean Ridge Geochemistry and Petrology. Mid-Ocean Ridge Seismic Structure. Propagating Rifts and Microplates. Seamounts and Off-Ridge Volcanism.
Further Reading Dilek Y, Moores E, Elthon D, and Nicolas A (eds.) (2000) Ophiolites and Oceanic Crust: New Insights from Field Studies and the Ocean Drilling Program. Geological Society of America Memoir. Boulder: Geological Society of America. Haymon RM (1989) Hydrothermal processes and products on the Galapagos Rift and East Pacific Rise, 1989. In: Winterer EL, Hussong DM, and Decker RW (eds.) The Geology of North America: The Eastern Pacific Ocean and Hawaii, vol. N, pp. 125--144. Boulder: Geological Society of America. Haymon RM (1996) The response of ridge crest hydrothermal systems to segmented, episodic magma supply. In: MacLeod CJ, Tyler P, and Walker CL (eds.) Tectonic, Magmatic, Hydrothermal, and Biological Segmentation of Mid-Ocean Ridges, vol. Special Publication 118, pp. 157--168. London: Geological Society. Humphris SE, Zierenberg RA, Mullineaux LS, and Thomson RE (eds.) (1995) Seafloor Hydrothermal Systems: Physical, Chemical, Biological, and Geological Interactions, Geophysical Monograph, vol. 91. Washington, DC: American Geophysical Union. Little CTS, Herrington RJ, Haymon RM, and Danelian T (1999) Early Jurassic hydrothermal vent community from the Franciscan Complex, San Rafael Mountains, California. Geology 27: 167--170. Tivey MK, Stakes DS, Cook TL, Hannington MD, and Petersen S (1999) A model for growth of steep-sided vent structures on the Endeavour Segment of the Juan de Fuca Ridge: results of a petrological and geochemical study. Journal of Geophysical Research 104: 22859--22883.
(c) 2011 Elsevier Inc. All Rights Reserved.
HYDROTHERMAL VENT ECOLOGY C. L. Van Dover, The College of William and Mary, Williamsburg, VA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1234–1241, & 2001, Elsevier Ltd.
Introduction Most of the ocean floor is covered with a thick layer of sediment and is populated by sparse and minute, mud-dwelling and mud-consuming invertebrates. In striking contrast, the volcanic basalt pavement of mid-ocean ridges hosts hydrothermal vents and their attendant lush communities of large invertebrates that ultimately rely on inorganic chemicals for their nutrition. Vents themselves are sustained by tectonic forces that fracture the basalt and allow sea water to penetrate deep within the ocean crust, and by volcanism, which generates the hot rock at depth that strips sea water of oxygen and magnesium. The hot rock gives up to the nascent vent fluid a variety of metals, especially copper, iron, and zinc, as well as reduced compounds such as hydrogen sulfide and methane. The vent fluid, thermally buoyant, rises to exit as hot springs on the seafloor. Discovered first by geologists in 1977 along a stretch of mountain range known as the Galapagos Rift, near the equator in the eastern Pacific Ocean, hydrothermal vents are now known to occur along every major ridge system on the planet. Several of these ridge systems – in the Arctic and Antarctic, in the Indian Ocean, in the southern Atlantic – are only just beginning to be explored. Regional species composition of the vent fauna differs between ocean basins and, sometimes, even within a basin. Most of the species that occur at vents have never been found in the adjacent, non-vent deep sea and are considered to be endemic, adapted to the chemical milieu of the vent environment. Reduced compounds carried in hydrothermal fluids, together with oxygen from the surrounding sea water, fuel the microbial fixation of inorganic carbon into organic carbon that forms the chemosynthetic base of the vent food web. Hydrothermal vents on midocean ridges are thus globally distributed, insular ecosystems that support endemic faunas through chemosynthetic processes rather than through photosynthesis. They are effectively decoupled by depth (typically 41000 m) from climatic variations and anthropogenic activities, but are tightly coupled to geophysical processes
of tectonism and volcanism. Vents thus offer unique opportunities for biologists to study adaptations that allow life to persist in this extreme environment and to explore planetary controls on biodiversity and biogeography along submarine, hydrothermal ‘archipelagoes’, where propagules are water-borne and subject to dispersal in an open system. Further, because vents are thought to have been a primordial component of the oceans and because there is increasing speculation that early life on this planet may have thrived in hot environments and on chemicals rather than on an organic soup for nourishment, extant hydrothermal systems are thought by many to be analogues for sites where early life may have evolved on this and other planets or planetary bodies in our solar system.
Microorganisms and the Chemosynthetic Basis for Life at Vents The terms ‘chemosynthesis’ and ‘photosynthesis’ are imprecise. While a voluminous nomenclature is available to differentiate among variations in these processes, for simplicity, chemosynthesis and photosynthesis are used here. In photosynthesis, sunlight captured by proteins provides energy for the conversion of inorganic carbon (carbon dioxide, CO2) and water (H2O) into organic carbon (carbohydrates, [CH2O] and oxygen (O2) (eqn [1]). light
CO2 þ H2 O - ½CH2 O þ O2
½1
Photosynthesis by plants is the basis for consumer and degradative food webs both on land and, as a rain of organic detritus derived from surface phytoplankton productivity, on the seabed. In the deep sea, detrital inputs of organic carbon are exceedingly small, accounting for the paucity of consumer biomass in abyssal muds. At hydrothermal vents, the supply of surface-derived organic material is overwhelmed by the supply of new organic carbon generated through chemical oxidation of hydrogen sulfide (H2S) (eqn [2]). CO2 þ H2 O þ H2 S þ O2 -½CH2 O þ H2 SO4
½2
Metabolic fixation pathways for carbon can be identical in photosynthetic plants and chemosynthetic
(c) 2011 Elsevier Inc. All Rights Reserved.
151
152
HYDROTHERMAL VENT ECOLOGY
microorganisms, namely the Calvin–Benson cycle, but the energy-yielding processes that fuel the Calvin– Benson cycle (photon capture versus chemical oxidation) are distinctive. High biomass at hydrothermal vents is in part a consequence of the aerobic nature of the process described in eqn [2]. Oxygen is used to oxidize the hydrogen sulfide, generating a large energy yield that in turn can fuel the production of large
amounts of organic carbon (Figure 1). Nonaerobic chemical reactions, such as oxidation of vent-supplied hydrogen (H2) by carbon dioxide (CO2), can also support chemosynthesis at vents, but energy yields under such anaerobic conditions are much lower than from aerobic oxidation. Microorganisms using these anaerobic reactions cannot by themselves support complex food webs and large invertebrates.
Symbiosis and the Host–Symbiont Relationship One of the hallmarks of many hydrothermal vent communities is the dominance of the biomass by invertebrate species that host chemosynthetic microorganisms within their tissues. Giant, redplumed, vestimentiferan tubeworms (Riftia pachyptila; Figure 2) so far provide the ultimate in host accommodation of endosymbiotic bacteria. These worms live in white, chitinous tubes, with their plumes extended into the zone of turbulent mixing of warm (B201C), sulfide-rich, hydrothermal fluid and cold (21C), oxygenated sea water. When discovered
CO2+H2O solar energy [CH2O] + O2 Photosynthesis
Chemical energy CO2 + H2O + H2S + O2 [CH2O] + H2SO4
Chemosynthesis
Figure 1 Photosynthetic and chemosynthetic processes in the ocean. Sunlight fuels the generation of organic material (CH2O) from inorganic carbon dioxide (CO2) and water (H2O) by phytoplankton in surface, illuminated waters. At depths where hydrothermal vents exist (typically>2000 m), no sunlight penetrates. In place of sunlight, the chemical oxidation of sulfide (H2S) by oxygen (O2) fuels the conversion of carbon dioxide to organic carbon by chemosynthetic bacteria.
Figure 2 The giant tubeworm Riftia pachyptila. The red plume of the tubeworm acts as a gill for uptake of dissolved gases. The trunk of the worm is found inside the white, chitinous tube. (Photograph by Dudley Foster, Woods Hole Oceanographic Institution.)
(c) 2011 Elsevier Inc. All Rights Reserved.
HYDROTHERMAL VENT ECOLOGY
in 1977, vestimentiferan tubeworms were remarkable for their size (up to several meters in length) and the complete absence of a digestive system in adults. In fact, the digestive system has been replaced by the trophosome, which is a specialized, paired organ derived from the larval gut. The trophosome is richly infiltrated with blood capillaries and each of its lobes lies within a blood-filled body cavity. The blood itself is rich in hemoglobin. Host bacteriocyte cells in the trophosome house chemosynthetic bacteria that use hydrogen sulfide and oxygen to fuel the production of organic carbon, as described above. The metabolic requirements of the tubeworm endosymbiotic bacteria place some remarkable burdens on the host. First, there is a novel requirement for delivery of sulfide to the bacteria, which reside at a location remote from the site of gas exchange (the plume). Sulfide is normally a potent toxin to animals, poisoning the cellular enzyme system that generates ATP (adenosine triphosphate), the currency of metabolism. Sulfide also competes with oxygen for binding sites on hemoglobin. Tubeworm hemoglobin has separate binding sites for oxygen and sulfide, so that both can be transported throughout the worm in the circulatory system without competition. When bound to hemoglobin, the sulfide is not reactive and so enzyme systems remain unchallenged. Once delivered to the bacteria in the trophosome, the sulfide is quickly oxidized and loses its toxic potential. Novel requirements for carbon dioxide are also found in tubeworms. The usual flow of CO2 is out of an animal, as the end-product of metabolism, but the resident bacteria of the trophosome require a net uptake of CO2. Maintenance of high concentrations of inorganic carbon in the blood of the tubeworm is facilitated by the high partial pressure of CO2 in the water surrounding the site of uptake (the plume) and by the alkaline internal pH of the blood (7.3–7.4), which favors the bicarbonate form (HCO 3 ) of carbon dioxide and thus maintains a steep concentration gradient for diffusion of CO2 from the environment into the blood. As described above, the anatomy of the tubeworm is well adapted for life in sulfide- and CO2-rich vent fluids and for supporting its endosymbiotic, chemosynthetic bacteria. The bacteria provide nearly all of the nutrition for the host, with the exception, perhaps, of small amounts of dissolved organic materials taken up across the tissues of the plume. In turn, the bacteria are provided with a chemically rich and stable environment for growth. Other large invertebrates at vents also derive much of their nutrition from endosymbiotic, chemosynthetic bacteria, including 20 to 30-cm long to vesicomyid clams and bathymodiolid mussels. While
153
Figure 3 Swarming shrimp (Rimicaris exoculata) at a hydrothermal vent on the Mid-Atlantic Ridge. (Photograph by C.L. Van Dover.)
vent mussels have a fairly normal digestive system and are capable of filter-feeding just as shallow-water mussels do, vent clams lack a functional digestive system. Both types of bivalves have enlarged gills and it is within these gills that the endosymbiotic bacteria are found. Clams are thought to take up hydrogen sulfide via their highly vascularized foot, with which they probe cracks in the basalt where vent fluids emanate. Not all chemosynthetic bacteria that nourish vent invertebrates are endosymbiotic. Shrimp that dominate vents in the Atlantic (Figure 3) host chemosynthetic bacteria on their carapace (i.e., the bacteria are episymbiotic) and seem to depend on these bacteria for a significant portion of their diet. Other chemosynthetic bacteria are free-living, suspended in the water column, providing nourishment to suspension-feeding invertebrates such as barnacles, or grow as mats or films on surfaces, where grazers such as limpets and polychaetes forage. Heterotrophic bacteria (using organic rather than inorganic compounds) may also be important for consumers within the vent invertebrate food web, but this has yet to be examined carefully.
Thermal Adaptations While hydrothermal vent communities live at temperatures slightly elevated above ambient sea water temperature, the existence of highly productive communities at vents is a consequence of fluid chemistry rather than thermal input. Nevertheless, there are some invertebrates, notably the large, thumb-sized polychaetes in the family Alvinellidae (Figure 4), that are especially tolerant of high temperatures and that compete with desert ants for the
(c) 2011 Elsevier Inc. All Rights Reserved.
154
HYDROTHERMAL VENT ECOLOGY
Guinness Book of World Records in the category of multicellular animal living at a thermal extreme. Alvinellids live on the sides of black smoker sulfide chimneys, and are reported to survive brief exposures to temperatures as high as 1051C. They routinely experience a thermal gradient of 50–601C over the length of their bodies. Temperature tolerance in these worms is not completely understood, but thermal stability of enzymes has been shown to increase in alvinellid species that occupy increasingly warmer habitats. Membranes may also be adapted for thermal stability through an increase in the degree of double bonding in fatty acids in species living in warmer portions of the vent habitat. Shrimp (Rimicaris exoculata) that swarm on black smoker chimneys in close proximity to extremely high temperature (3501C) fluids are not bathed in excessively hot water, but they do have paired photoreceptive organs that may function to detect the glow emitted by the hot water. The organs are derived from ordinary shrimp eyes, but the photoreceptive surfaces are hypertrophied and rich in visual pigment (rhodopsin). The eyestalks are lost and the derived eyes extend back along the dorsal surface of the shrimp, beneath the transparent carapace. These ‘eyes’ have also lost their optic (lens) systems, so they cannot form an image, but they are
Figure 4 The alvinellid polychaete Alvinella caudata, beside its fragile tube. (Photograph by J. Porteous, Woods Hole Oceanographic Institution.)
optimized to detect gradients of dim light. While light from black smokers has now been well documented, the behavioral response of the shrimp to this light has not been studied.
Community Dynamics There could hardly be a greater dynamic contrast than that found between the cold, food-limited, relentlessly stable and vast deep sea environment, and the thermally complex, trophically rich, ephemeral, and insular deep-sea hydrothermal vent fields. Since the discovery of vents, ecologists have attempted to predict the cycle of community development over the life span of a vent field by interpolation and extrapolation from snapshot observations. The 1991 seafloor volcanic eruption at the Venture Hydrothermal Field on the East Pacific Rise was witnessed within days to weeks of the event, providing biologists with the first submarine equivalent of Krakatau. At the time of the Venture eruption, existing vent communities were obliterated by fresh lava flows, pervasive warm-water venting was observed along B1.5 km of ridge axis, and a dense ‘bloom’ of flocculent material poured from cracks between lobes of lava, obscuring visual navigation (Figure 5). The flocs were determined to be the mineral by-product of microbial production. Within one year, venting became focused at numerous sites along the ridge axis and the first colonists had arrived, including a small tubeworm species (Tevnia jerichonana) and dense aggregations of several species of limpets. Mobile vent organisms, including vent crabs and squat lobsters, zoarcid fish, and swarming amphipods were also well represented. After 2.5 years, some vents had shut off but, where venting persisted, mature colonies of the giant tubeworm (Riftia
Figure 5 Flocculent material suspended in the water column 1 m above new ocean crust. The floc is a mineral by-product of microbial production and emanated from cracks in the seafloor. (Photograph by Alvin, Woods Hole Oceanographic Institution.)
(c) 2011 Elsevier Inc. All Rights Reserved.
HYDROTHERMAL VENT ECOLOGY
pachyptila) were established, along with a variety of smaller invertebrates that live among the tubeworm tubes, including shrimp, limpets, amphipods, and polychaetes. Growth rates of R. pachyptila were measured to be among the most rapid of any aquatic invertebrate. Within 5 years, 75% of the regional species pool could be found at the new vents and mussel beds were well-established and beginning to overwhelm the R. pachyptila thickets. In this example, vesicoymid clams were the last of the big megafaunal species to arrive at the site, despite the presence of adult populations within several kilometers of the area overrun by fresh lava. Mussels appear to have a competitive edge among the larger taxa at vents, in part because they are mobile and can relocate as necessary to cope with changing flow patterns of vent fluids, while tubeworms are stationary and have few options for tracking vent flow. In addition, because mussels can filter-feed as well as derive nutrition from their endosymbiotic bacteria, they are among the last species to disappear as a vent shuts down. Ultimately, it is the mobile scavengers – the crabs and fish and octopus – that witness the final demise of a vent field. The eruption at the Venture Hydrothermal Field was not entirely unexpected – geologists were studying this region of the ridge axis because it was so inflated and appeared to be ripe for an eruptive event. Localization of seafloor eruptions on the Juan de Fuca Ridge in the northeast Pacific (off the coast of Vancouver Island) now takes place in real-time, facilitated by a legacy of the cold war era, namely through acoustic signals received by underwater sound-surveillance systems originally designed to track enemy submarines. As eruptions take place, T-waves (low-frequency, tertiary waves characteristic of lava in motion) are transmitted into the water column. Navy hydrophone networks allow the T-waves to be placed in a geographical context and traced from start to finish. One migration of lava in June 1993 was traced for about 40 km below the ocean crust over a two-day period before it finally erupted onto the seafloor. As at the Venture Field, fresh lava on the seafloor was observed, along with venting of flocculent material derived from microbial production. Sites of persistent venting were colonized by populations of vent invertebrates within one year, some of which were reproductively mature. The two examples of community dynamics cited here suggest that the species composition of all vent communities is constantly changing. But some vent sites are long-lived, and repeat visits to these longlived sites over a 15-year period document essentially no change in the nature of the fauna. The best
155
example of this to date is the TAG site on the MidAtlantic Ridge. TAG is a large sulfide mound (100 m diameter) that has occupied the junction of crosscutting faults and fissures for more than 100 000 years. Swarming shrimp (Rimicaris exoculata) and anemones (Maractis rimicarivora) dominated the site when it was first discovered in 1985 and continued to dominate through the most recent set of observations in 1998, despite massive disruptions of hydrothermal flow caused by drilling in 1994. There have been shifts in the precise location of the primary masses of shrimp as they track local natural and anthropogenic changes in vent flow on the mound, and we assume that there has been replacement of generations by recruitment. But there has been no succession observed, no invasion by other taxa. The absence of change does little to attract ecologists, who thrive on dynamic systems, but the stability of the species composition at TAG and other sites has profound implications regarding the selective pressures encountered by the species that inhabit these sites compared to sites that are constantly threatened by lava overruns or tectonic shifts in plumbing.
Origins of Vent Faunas, Biogeography, and Biodiversity Evolutionary paths that brought invertebrate taxa to vents are varied. Several major taxa, including the galatheid squat lobsters, pycnogonid sea spiders, and echinoderms are likely to be immigrants from the surrounding deep sea. Some species are closely related to (and presumably derived from) shallowwater genera. Many of the most familiar vent taxa – the vesicomyid clams, bathymodiolid mussels, vestimentiferan tubeworms, alvinocarid shrimp – are allied to genera and families found in a variety of deepsea chemosynthetic ecosystems (i.e., seeps and whale-falls as well as vents). The direction of invasion (seep-to-vent or vice versa) can be inferred using molecular techniques. For example, molecular phylogenetics suggests that the bathymodiolid mussel group invaded vents from seeps, with several seep species resulting from reinvasion of the seep habitat by vent ancestors. There are also specialized taxa so far known only from hydrothermal vents. The most conspicuous of these is the alvinellid polychaete family, whose members often occupy the warmest habitable waters of a vent site. Still other taxa appear to be relicts of ancient lineages that have found refuge in the vent environment. These relict taxa include the stalked
(c) 2011 Elsevier Inc. All Rights Reserved.
156
HYDROTHERMAL VENT ECOLOGY
1 cm
Figure 6 The ancient barnacle Neolepas zevinae. (Photograph by C.L. Van Dover.)
Present
barnacle, Neolepas zevinae (Figure 6) and the archaeogastropod limpet, Neomphalus fretterae. Fossil vent communities found on land provide a glimpse of Silurian vent assemblages. Vestimentiferans are reported from some of the oldest vent deposits known, accompanied by brachiopods and monoplacophorans. Brachiopods and monoplacophorans are so far poorly represented in modern vent communities, if at all. When vents were first discovered, some biologists hypothesized that the vent fauna would be globally cosmopolitan. Subsequent explorations have shown this hypothesis to be false, and mechanisms that allow isolation and differentiation of faunas have been postulated. One of the best examples comes from a comparison of faunas from the Juan de Fuca Ridge and East Pacific Rise. At one time (56 Ma), these ridges were one continuous ridge system, but around 37 Ma the North American Plate began to
56 Ma
Kula North America
Pacific Pacific
65 Ma
Farallon
110 Ma
Kula Farallon
Izanagi Farallon Pacific
Pacific
Figure 7 Bisection of a mid-ocean ridge by the North American Plate. At one time (56 Ma), there was a continuous mid-ocean ridge in the eastern Pacific basin but, as the North American Plate overrode the ridge system, it was bisected to form the northeast Pacific ridge system (Juan de Fuca, Explorer, Endeavour, and Gorda Ridges) and the East Pacific Rise (emerging in the Gulf of California and running south toward Antarctica). (Reproduced from Tunnicliffe et al. 1996.)
(c) 2011 Elsevier Inc. All Rights Reserved.
HYDROTHERMAL VENT ECOLOGY
override this ridge (Figure 7), splitting the parent faunal assemblage into two daughter assemblages that are distinct yet closely allied at the generic level. Similarities among faunas at the species level may be reduced to nearly zero at vent sites that are in separate ocean basins. This observation suggests that major additions to the global vent faunal inventory await us in the unexplored ocean basins; the potential for discovery of novel adaptations to the vent ecosystem is extremely high. Vent biologists are just beginning to examine global patterns in biodiversity. Preliminary measures show a strong correlation between spreading rate and species richness, raising the hypothesis that patterns of volcanism may be an ultimate control on species diversity in hydrothermal vent ecosystems. Where ridges are fast-spreading, the magma budget is high and the temporal and spatial frequency of vents along the ridge axis is high. In contrast, on slow-spreading ridges, the magma budget is low and vents are far apart. Because distances between vents are short on fast-spreading ridges, species are less likely to go extinct. At slow-spreading ridges, allopatric speciation by isolation may be favored, but species are more likely to go extinct because of the distance between vents. Many of the species that dominate vents on slow-spreading systems have mobile rather than sessile adults, suggesting that distance may act as a selective filter for dispersal capability in these systems.
Vent Systems and the Origin of Life Without doubt, the most provocative consequence of the discovery of seafloor hydrothermal vents is the suggestion that vents may have been the site where life originated on Earth. In contrast to the heterotrophic hypothesis of the origin of life, with its nourishing ‘organic soup’, the vent theory suggests that the earliest life was chemosynthetic, taking biochemical advantage of the large degree of chemical disequilibrium associated with mixing zones of low- and high-temperature portions of hydrothermal systems. A case has been made that the thermophilic nature of the most ancient known lineages of life is indicative of an origin of life at hot springs, but this argument finds limited support, since these most ancient cell types are still very complex and far removed from the progenitors of life. One model for the origin of life at vents suggests the following acellular precursor: a monomolecular, negatively-charged organic layer bonded to positively charged mineral surfaces at the interface of hot water. In this model, pyrite, which forms exergonically from
157
iron monosulfide and hydrogen sulfide (both components of vent fluids), serves as the mineral surface. The pyrite-forming reaction yields two free electrons that can be used for building biochemical constituents; simple organic molecules interacting with the pyrite could be reduced to more complex organic molecules. The simple organics also derive from the vent fluids, through abiogenic synthesis. In a secondary stage of development, the precursor evolves to being semicellular, still supported by minerals, but with a lipid membrane and internal broth, with increasing metabolic capabilities. In the final stage of origin, the pyrite support is abandoned and true cellular organisms arise. The elegance of this model is that it uses an energetically realistic inorganic chemical reaction to create a cationic substrate that can bind with organic compounds, all within a single setting. Discovery of vent ecosystems and the appreciation of their chemosynthetic basis has influenced the search for life elsewhere in the solar system. When the Viking Mission to Mars took place, the emphasis was on a search for photoautotrophic processes, but now the search for evidence of past or extant life on other planets highlights environments where chemosynthetic processes may take place, including hydrothermal areas.
Closing Remarks Hydrothermal vent ecology remains a field ripe for discovery of novel faunas and adaptations, as vents in new ocean basins are explored. Because access to deep-sea ecosystems is continually improving, biologists can now undertake quantitative sampling and time-series investigations that are certain to reshape our understanding of the physiological ecology, population biology, community dynamics, and biogeography of vent faunas in the near future.
See also Hydrothermal Vent Biota. Hydrothermal Vent Fluids, Chemistry of. Hydrothermal Vent Deposits. Hydrothermal Vent Fauna, Physiology of. MidOcean Ridge Tectonics, Volcanism, and Geomorphology.
Further Reading Bock GR and Goode JA (eds.) (1996) Evolution of Hydrothermal Ecosystems on Earth (and Mars?). New York: Ciba Foundation.
(c) 2011 Elsevier Inc. All Rights Reserved.
158
HYDROTHERMAL VENT ECOLOGY
Childress JJ and Fisher CR (1992) The biology of hydrothermal vent animals: physiology, biochemistry, and autotrophic symbioses. Oceanography and Marine Biology Annual Review 30: 337--441. Desbruye`res D and Segonzac M (eds.) (1997) Handbook of Deep-Sea Hydrothermal Vent Fauna. Brest: IFREMER. Fisher CR (1990) Chemoautotrophic and methanotrophic symbioses in marine invertebrates. Critical Reviews in Aquatic Science 2: 399--436. Humphris SE, Zierenberg RA, Mullineaux LS, and Thomson RE (eds.) (1995) Seafloor Hydrothermal Systems: Physical, Chemical, Biological and Geological Interactions. Washington, DC: American Geophysical Union. Karl DM (ed.) (1995) The Microbiology of Deep-Sea Hydrothermal Vents. New York: CRC Press. Shank TM, Fornari DJ, Von Damm KL et al. (1998) Temporal and spatial patterns of biological community
development at nascent deep-sea hydrothermal vents (91N, East Pacific Rise). Deep-Sea Research 45: 465--516. Tunnicliffe V (1991) The biology of hydrothermal vents: ecology and evolution. Ocenography and Marine Biology Annual Review 29: 319--407. Tunnicliffe V, Fowler CMR, and McArthur AG (1996) Plate tectonic history and hot vent biogeography. In: MacLeod CJ, Tyler PA, and Walker CL (eds.) Tectonic, Magmatic, Hydrothermal and Biological Segmentation of Mid-Ocean Ridges, Geological Society Special Publication 118, pp. 225–238. Tyler PA and Young CM (1999) Reproduction and dispersal at vents and cold seeps: a review. Journal of the Marine Biology Association of the UK 79: 193--208. Van Dover CL (2000) The Ecology of Deep-Sea Hydrothermal Vents. Princeton: Princeton University Press.
(c) 2011 Elsevier Inc. All Rights Reserved.
HYDROTHERMAL VENT FAUNA, PHYSIOLOGY OF A. J. Arp, Romberg Tiburon Center for Environment Studies, Tiburon, CA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1242–1246, & 2001, Elsevier Ltd.
Hydrothermal Vent Environments are Dynamic, Hot, and Toxic The hydrothermal vent environments, lying at the bottom of the ocean at depths of 2.5 km or more, were discovered in 1977 by a group of geologists exploring spreading centers at midocean ridges on the sea floor. As fissures open up in the earth’s surface, lava is extruded onto the ocean floor and sea water is pulled towards the center of the earth so deeply that it comes into contact with hot, molten magma. The sea water is superheated and then discharged back into the environment through fissures in the ocean floor. As sea water moves from the center of the earth and into the vent habitat on the seafloor, it becomes laden with inorganic chemicals. In particular, hydrogen sulfide, an essential chemical in this unique environment, leaches into the water at depth, and water discharging into the environment can be highly enriched in this toxic but energy-rich chemical. There are two types of hydrothermal vents. In those characterized by diffuse venting, sea water percolates out at a moderate rate and is approximately 10–201C in temperature. As the bottom water temperature of the majority of the earth’s deep ocean is about 21C, these hydrothermal fluids are elevated in temperature, but rapidly mix with the surrounding sea water. One of the first hydrothermal vent environments discovered, Rose Garden at the hydrothermal vent environment near the Galapagos Islands, is an example of a diffuse vent habitat. In the black smoker environment of the hydrothermal vents, things are a lot hotter, such as at those on the Juan de Fuca Ridge off the coast from the state of Washington. This is a very dynamic, high-temperature environment where water issuing forth from large chimneylike structures can be as hot as 4001C in temperature. In spite of the hydrogen sulfide-enriched environment, elevated temperatures, and the dynamic volcanic activity, numerous and varied animals cluster around these sites, taking advantage of the hard substrate provided by the extruded pillow lava. A
typical vent environment begins with a stretch of pillow lava in the foreground, with clams wedged into the numerous fissures. At the heart of the vent environment are diffuse venting hydrothermal fluids or actively spewing white smoker chimneys and black smoker chimneys, teeming with life. Typical inhabitants include dense clusters of tubeworms and many free-ranging animals roaming in and out of the vent environment such as brachyuran crabs, galatheid crabs, numerous amphipods, a few species of fish, and a host of other smaller animals.
Chemosynthesis — The Basis of All Life in the Vent Environment Possibly the most revolutionary outcome of the discovery of hydrothermal vents is the story of how life exists in this challenging habitat and the unique nature of the food chain and the source of basic energy in this remote location. Prior to the discovery of the hydrothermal vents, most biologists believed that all life depended upon the energy of sunlight and that the basis of all food chains was photosynthesis. When the hydrothermal vents were discovered, it was immediately clear that they represent a very enriched biological environment very remote from the surface sunlight. It was difficult to imagine that organic material could drift down in large enough quantities to provide the energy to fuel this environment. Other interesting data materializing rapidly after the discovery of the vents indicated that most of the animals, especially the large invertebrates, have no digestive systems. For example, the large tubeworm, Riftia pachyptila, has no mouth and no intestine; however, it is a large animal approximately 1.5 m in length and up to 2 cm in diameter, and large colonies of these animals flourish in the remote vent habitat. The question of how animal life is supported in the deep sea communities of the hydrothermal vents became, and remains to this day, a focus of intense research. Several lines of evidence lead to the realization that some of the major invertebrates endemic to the vent environment harbor bacteria within their body cavities. Free-living bacteria in this environment and in our own backyard have been known for years to be able to use chemical energy as a basis of their metabolism. In the case of the free-living bacteria, there are many hydrogen sulfide-oxidizing bacteria that can use this chemical as the basis of
(c) 2011 Elsevier Inc. All Rights Reserved.
159
160
HYDROTHERMAL VENT FAUNA, PHYSIOLOGY OF
their metabolic pathways and produce organic compounds that form the basis of their nutrition. Hydrothermal vent animals harboring these chemical-utilizing bacteria within their body tissues as symbionts include the large tubeworm Riftia pachyptila, which has dense aggregations of bacteria in a residual gutlike organ, and the clam Calyptogena magnifica, which harbors bacteria in the gills. These symbiotic bacteria are able to utilize the inorganic chemical hydrogen sulfide, so plentiful in this environment, in a manner analogous to what plants do with energy from the sun. This process is therefore termed chemosynthesis rather than photosynthesis. The symbiotic bacteria living within the bodies of the larger invertebrate animals have been demonstrated to oxidize hydrogen sulfide, and the energy released from this biochemical process is used to power the fixation of carbon dioxide into small organic compounds – just as in free-living bacterial sulfide oxidation. The Calvin–Benson cycle employed in both cases is the same metabolic pathway that is utilized by plants in photosynthesis to transform inorganic carbon dioxide into organic compounds that are then utilized as food higher up in the food chain. The critical difference with chemosynthetic metabolism is that, rather than using sunlight, these animals and bacteria utilize chemical energy to power that reaction and are completely independent of sunlight (Figure 1). The net result is that freeliving bacteria in the environment and symbiotic bacteria living within animal tissues are able to live independently of sunlight by utilizing chemicals from the core of the earth, thus forming a very different basis for the food chain in the hydrothermal vent environment. The discovery of the hydrothermal vent environment was a fundamental discovery of a well-defined ecosystem that is completely independent of sunlight at any level of the food chain.
Figure 1 Chemosynthetic pathways in Riftia pachyptila.
The Ecophysiology of the Giant Tubeworm Riftia pachyptila One of the most dramatic and best-known of the animals endemic to the hydrothermal vent environment is the giant tubeworm Riftia pachyptila. Colonies of these worms are clumped together around effluent points in the hydrothermal vent habitat, growing toward and into the water that is percolating out from the seafloor. An individual animal lives inside a single, unbranched chitinous tube and the red structure protruding out of the end of the tube is the respiratory plume. The animal can retract the plume back into the tube if disturbed by a roaming predator. There is a collarlike vestimentum organ that positions the animal within the tube, and a large trunk region of the animal is filled with an organ termed the trophosome. This organ is believed to be the vestigial gut of the worm and is composed, literally, of masses of bacteria. A pool of coelomic fluid bathes the trophosome and contains a largemolecular-weight extracellular respiratory hemoglobin (Figure 2). Riftia pachyptila also has a separate pool of blood that contains high concentrations of an extracellular hemoglobin that circulates in an elaborate closed circulatory system powered by a heartlike structure in the vestimentum region. The blood is pumped in a complete circuit
Figure 2 The external anatomy of Riftia pachyptila.
(c) 2011 Elsevier Inc. All Rights Reserved.
HYDROTHERMAL VENT FAUNA, PHYSIOLOGY OF
161
from the respiratory plume to body tissues, and on to elaborate capillary beds in the region of the trophosome and bacteria (Figure 3). The red color of the blood is due to the high concentration of hemoglobin and gives the characteristic red color to the plumes. The respiratory plume and the circulating hemoglobin are essential for the transport of the key metabolites oxygen, hydrogen sulfide, and carbon dioxide, which are the principal components of the metabolism of the symbiotic bacteria. The respiratory hemoglobins present in the plume and the coelomic fluid of the animal bind oxygen with a very high affinity. The binding is reversible and cooperative, such that oxygen uptake is enhanced at the respiratory plume, and oxygen delivery is augmented at the tissues and trophosome organ. Hydrogen sulfide is a highly toxic molecule that typically acts in a similar manner to cyanide by binding at the iron center of cytochrome molecules and hemoglobin molecules, thus arresting aerobic metabolism. Although Riftia pachyptila and other hydrothermal vent animals utilize hydrogen sulfide for their metabolism, they also have tissues that are highly sensitive to sulfide poisoning. Detoxification of hydrogen sulfide is essential for aerobic life in this
dynamic, chemically enriched environment. The key to the simultaneous needs for transportation and detoxification is the respiratory hemoglobin present in the plume and coelomic fluid. Riftia pachyptila hemoglobin binds hydrogen sulfide with a very high affinity. The toxic hydrogen sulfide is transported to the trophosome region in the center of the worm’s body as a tightly bound molecule that cannot chemically interact with sulfide-sensitive tissues. Oxygen and sulfide are simultaneously bound to the hemoglobin at separate binding sites and are transported to the trophosome, where they are believed to be delivered to the symbiotic bacteria for metabolism. In this way hydrogen sulfide is taken up from the surrounding sea water and transported to the site of bacterial metabolism while interaction is prevented with other tissues, such as the body wall, that are highly aerobic and sensitive to the toxic effects of hydrogen sulfide. These unusual adaptations function for respiratory gas transport and metabolism as well as for detoxification and tolerance of toxic chemicals in what would be a very inhospitable environment for most animals. Dense colonies of Riftia pachyptila flourish in a specialized microhabitat within the vent environment. The worms anchor themselves on the rocks where the hydrothermal vent fluid is issuing out into the seafloor. The base of the tube is bathed in hydrothermal fluid enriched in hydrogen sulfide and carbon dioxide, but de
