Long-term Environmental Change in Arctic and Antarctic Lakes, Vol. 8 (Developments in Paleoenvironmental Research)

Long-term Environmental Change in Arctic and Antarctic Lakes

Developments in Paleoenvironmental Research VOLUME 8

Long-term Environmental Change in Arctic and Antarctic Lakes Edited by

Reinhard Pienitz Université Laval, Québec, Canada and

Marianne S.V. Douglas University of Toronto, Toronto, Canada and

John P. Smol Queen's University, Kingston, Canada

A C.I.P. Catalogue record for this book is available from the Library of Congress.

ISBN-10 1-4020-2125-9 (HB) Springer Dordrecht, Berlin, Heidelberg, New York ISBN-10 1-4020-2126-7 (e-book) Springer Dordrecht, Berlin, Heidelberg, New York ISBN-13 978-1-4020-2125-1 (HB) Springer Dordrecht, Berlin, Heidelberg, New York ISBN-13 978-1-4020-2126-8 (e-book) Springer Dordrecht, Berlin, Heidelberg, New York

Published by Springer, P.O. Box 17, 3300 AA Dordrecht, The Netherlands. Cover Photo: Hanging Lake, Yukon Territory, Canada (photo taken by Julien Racca). Inset: Subarctic Lake near Umiujaq, Northern Québec, Canada (photo taken by Reinhard Pienitz).

Printed on acid-free paper

All Rights Reserved © 2004 Springer No part of this work may be reproduced, stored in a retrieval system, or transmitted in any form or by any means, electronic, mechanical, photocopying, microfilming, recording or otherwise, without written permission from the Publisher, with the exception of any material supplied specifically for the purpose of being entered and executed on a computer system, for exclusive use by the purchaser of the work. Printed in the Netherlands.

DEDICATION This book is dedicated to our friend and colleague Dr. John C. Kingston (1949-2004), whose passion for diatoms and life affected us all.

CONTENTS Acknowledgements ................................................................................................. xiii The Editors .............................................................................................................. xv List of Contributors .............................................................................................. xvii Editors and Board of Advisors of Developments in Paleoenvironmental Research Book Series...........................................................................................xxvii

1. Paleolimnological research in polar regions: An introduction. Reinhard Pienitz, Marianne S.V. Douglas and John P. Smol..................................... 1 Introduction Observational/instrumental evidence for rapid climate change in the circumpolar regions Why do we need paleolimnological data from arctic and antarctic regions? Geographic scope of the book The focus and the structure of this volume Acknowledgements References Part I: Major Indicators and Approaches 2. Geochronology of high latitude lake sediments. Alexander P. Wolfe, Gifford H. Miller, Carrie A. Olsen, Steven L. Forman, Peter T. Doran and Sofia U. Holmgren.................................................................... 19 Introduction Dating recent high latitude lake sediments using 210Pb and 137Cs Radiocarbon dating Optically stimulated luminescence Future directions in high latitude lake sediment geochronology Summary Acknowledgements References 3. Physical and chemical properties and proxies of high latitude lake sediments. Scott F. Lamoureux and Robert Gilbert ................................................................... 53 Introduction Polar environmental systems High latitude lake systems Physical and chemical proxy records in high latitude lakes vii

viii Framework for interpreting the environmental significance of physical and biogeochemical sedimentary records in high latitude lakes Summary Acknowledgements References 4. Palynology of North American arctic lakes. Konrad Gajewski and Glen M. MacDonald ............................................................. 89 Introduction Synthesis of methodological aspects Select North American studies Conclusion – Outlook Summary Acknowledgements References 5. Algal indicators of environmental change in arctic and antarctic lakes and ponds. Marianne S.V. Douglas, Paul B. Hamilton, Reinhard Pienitz and John P. Smol... 117 Introduction Historical overview of algal research in the Arctic and Antarctic Algal indicators Ecological classifications of algae Paleolimnological reconstructions Other applications Summary Acknowledgements References 6. Aquatic invertebrates and high latitude paleolimnology. Ole Bennike, Klaus P. Brodersen, Erik Jeppesen and Ian R. Walker..................... 159 Introduction Notes on different zoological indicators present in lake sediments Discussion Summary Acknowledgements References

ix 7. Use of water isotope tracers in high latitude hydrology and paleohydrology. Thomas W.D. Edwards, Brent B. Wolfe, John J. Gibson and Dan Hammarlund... 187 Introduction Isotopic labelling in the hydrological cycle Isotope hydrology at high latitudes Water isotope tracers in paleolimnology Summary and future perspectives Acknowledgements References 8. Lake sediments as records of arctic and antarctic pollution. Derek C.G. Muir and Neil L. Rose ......................................................................... 209 Introduction Challenges in the study of high latitude lake sediment cores Spatial and temporal trends of metals, persistent organic pollutants and anthropogenic particles Summary Acknowledgements References Part II: Regional Syntheses 9. Paleolimnology of the middle and high Canadian Arctic. Alexander P. Wolfe and I. Rod Smith ..................................................................... 241 Introduction Environmental background Wisconsinan glacial history The limnological legacy Historical development Pre-Holocene lake sediment records Holocene climatic evolution and paleolimnology The latest Holocene: a time of unprecedented change Problems, recommendations and conclusions Summary Acknowledgements References

x 10. Paleolimnology of the North American Subarctic. Bruce P. Finney, Kathleen Rühland, John P. Smol and Marie-Andrée Fallu ........ 269 Introduction Description of study region Paleoindicators of the North American Subarctic Regional syntheses Challenges and future directions Summary Acknowledgements References 11. Holocene paleolimnology of Greenland and the North Atlantic islands (north of 60°N). N. John Anderson, David B. Ryves, Marianne Grauert and Suzanne McGowan... 319 Introduction Paleolimnological themes across the northern North Atlantic Synthesis and areas for further research Summary Acknowledgements References 12. Paleolimnological research from northern Russian Eurasia. Glen M. MacDonald, Thomas W.D. Edwards, Bruce Gervais, Tamsin E. Laing, Michael F.J. Pisaric, David F. Porinchu, Jeffrey A. Snyder, Nadia Solovieva, Pavel Tarasov and Brent B. Wolfe ......................................................................... 349 Introduction Recent analyses using biological evidence Stable isotope studies of lake sediments from across northern Russian Eurasia Regional lake status data bases and lake-level records from northern Russian Eurasia Summary Acknowledgements References 13. Paleolimnological studies in arctic Fennoscandia and the Kola Peninsula (Russia). Atte Korhola and Jan Weckström........................................................................... 381 Introduction Origin and sedimentological characteristics of lakes Holocene changes in physical attributes Holocene trends in chemical attributes

xi Holocene trends in aquatic communities Recent limnological changes Summary Acknowledgements References 14. Paleolimnological studies from the Antarctic and subantarctic islands. Dominic A. Hodgson, Peter T. Doran, Donna Roberts and Andrew McMinn ....... 419 Introduction The role of the Antarctic in Earth system science Antarctic limnology Antarctic paleolimnology Antarctic paleolimnology and Earth system science – case studies Discussion Outlook Conclusions Summary Acknowledgements References 15. Paleolimnology of extreme cold terrestrial and extraterrestrial environments. Peter T. Doran, John C. Priscu, W. Berry Lyons, Ross D. Powell, Dale T. Andersen and Robert J. Poreda................................................................. 475 Introduction Perennially ice-covered lakes Perennially ice-sealed lakes Subglacial lakes Exopaleolimnology Summary Acknowledgements References 16. Epilogue: Paleolimnological research from arctic and antarctic regions. Reinhard Pienitz, Marianne S.V. Douglas and John P. Smol................................. 509 Glossary, Acronyms and Abbreviations ................................................................. 513 Index....................................................................................................................... 541

ACKNOWLEDGEMENTS We are very grateful to our many colleagues who have helped in the preparation and completion of this volume. In addition to the dedicated work provided by the chapter authors, we would also like to acknowledge the important contributions made by our chapter reviewers, whose constructive and detailed comments greatly improved the quality of many contributions. We would like to thank Claudia Zimmermann for her help with the final preparation and revision of the chapters, as well as Fawn Ginn and Carol Nöel for technical assistance. Thanks are also due to Ms. Judith Terpos and Dr. Anna Besse-Lototskaya at Kluwer Academic Publishers, who helped bring this project to its conclusion. Finally, the book series co-editor, Prof. William M. Last, is thanked for his continued encouragement, enthusiasm, and editorial advice.

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THE EDITORS Reinhard Pienitz is a professor in the Department of Geography at Université Laval (Québec City, Québec, Canada). He directs the Paleolimnology-Paleoecology Laboratory at the Centre d’Études Nordiques (CEN). His research focuses on the use of modern and fossil algae and insects as indicators of environmental change in lakes and rivers of arctic and temperate regions, as well as marine coastal environments. Marianne S.V. Douglas is an associate professor in the Department of Geology at the University of Toronto (Toronto, Ontario, Canada). Her laboratory, the Paleoenvironmental Assessment Laboratory (PAL), conducts work primarily in the Arctic and Antarctic. John P. Smol is a professor in the Biology Department at Queen’s University (Kingston, Ontario, Canada), with a cross-appointment at the School of Environmental Studies. He co-directs the Paleoecological Environmental Assessment and Research Lab (PEARL). Professor Smol is co-editor of the Journal of Paleolimnology and holds the Canada Research Chair in Environmental Change.

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LIST OF CONTRIBUTORS

DALE T. ANDERSEN SETI Institute Center for the Study of Life in the Universe 2035 Landings Drive Mountain View, California 94043, USA ([email protected]) N. JOHN ANDERSON Department of Geography Loughborough University Loughborough, Leicestershire LE11 3TU United Kingdom ([email protected]) OLE BENNIKE Geological Survey of Denmark and Greenland Øster Voldgade 10 1350 Copenhagen K Denmark ([email protected]) KLAUS P. BRODERSEN Freshwater Biological Laboratory University of Copenhagen 51 Helsingørsgade 3400 Hillerød Denmark ([email protected]) PETER T. DORAN Department of Earth and Environmental Sciences University of Illinois at Chicago 845 W. Taylor St. Chicago, Illinois 60607-7059, USA ([email protected])

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xviii MARIANNE S.V. DOUGLAS Paleoecological Assessment Laboratory Department of Geology University of Toronto Toronto, Ontario M5S 3B1, Canada ([email protected]) THOMAS W.D. EDWARDS Department of Earth Sciences University of Waterloo Waterloo, Ontario N2L 3G1, Canada ([email protected]) MARIE-ANDRÉE FALLU Département de Chimie-Biologie Université du Québec à Trois-Rivières C.P. 500 Trois-Rivières, Québec G9A 5H7, Canada ([email protected]) BRUCE P. FINNEY Institute of Marine Science University of Alaska Fairbanks Fairbanks, Alaska 99775, USA ([email protected]) STEVEN L. FORMAN Department of Earth and Environmental Sciences University of Illinois at Chicago 845 W. Taylor St. Chicago, Illinois 60607-7059, USA ([email protected]) KONRAD GAJEWSKI Laboratory of Paleoclimatology and Climatology Department of Geography University of Ottawa Ottawa, Ontario K1N 6N5, Canada ([email protected])

xix BRUCE GERVAIS Department of Geography California State University Sacramento, California 95819-6003, USA ([email protected]) JOHN J. GIBSON Water and Climate Impacts Research Centre Department of Geography University of Victoria Victoria, British Columbia V8W 3P5, Canada ([email protected]) ROBERT GILBERT Department of Geography Queen’s University Kingston, Ontario K7L 3N6, Canada ([email protected]) MARIANNE GRAUERT Department of Geography University of Copenhagen Øster Voldgade 10 DK-1350 Copenhagen K Denmark ([email protected]) PAUL B. HAMILTON Canadian Museum of Nature Life Sciences Section, Research Division P.O. Box 3443, Station D Ottawa, Ontario K1P 6P4, Canada ([email protected]) DAN HAMMARLUND GeoBiosphere Science Centre Quaternary Sciences Lund University Sölvegatan 12 SE-223 62 Lund Sweden ([email protected])

xx DOMINIC A. HODGSON British Antarctic Survey Natural Environment Research Council High Cross, Madingley Road Cambridge, CB3 0ET United Kingdom ([email protected]) SOFIA U. HOLMGREN Department of Earth Sciences Geovetarcentrum Göteborgs Universitet SE-405 30 Göteborg Sweden ([email protected]) ERIK JEPPESEN Department of Freshwater Ecology National Environmental Research Institute Vejlsøvej 25 8600 Silkeborg Denmark and Department of Plant Biology University of Aarhus Nordlandsvej 68 8240 Risskov Denmark ([email protected]) ATTE KORHOLA Environmental Change Research Unit (ECRU) Department of Biological and Environmental Sciences University of Helsinki P.O. Box 65 (Viikinkaari 1) FIN-00014 Helsinki Finland ([email protected]) TAMSIN E. LAING Environmental Sciences Group Royal Military College of Canada PO Box 17000 Stn. Forces Kingston, Ontario K7K 7B4, Canada ([email protected])

xxi SCOTT F. LAMOUREUX Department of Geography Queen’s University Kingston, Ontario K7L 3N6, Canada ([email protected]) W. BERRY LYONS Byrd Polar Research Center Ohio State University Columbus, Ohio 43210, USA ([email protected]) GLEN M. MACDONALD Departments of Geography and Organismic Biology, Ecology and Evolution University of California Los Angeles, California 90095-1524, USA ([email protected]) SUZANNE MCGOWAN School of Geography The University of Nottingham University Park Nottingham, NG7 2RD United Kingdom ([email protected]) ANDREW MCMINN Institute of Antarctic and Southern Ocean Studies University of Tasmania Private Bag 77 Hobart, Tasmania 7001 Australia ([email protected]) GIFFORD H. MILLER Institute of Arctic and Alpine Research and Department of Geological Sciences University of Colorado Boulder, Colorado 80309-0450, USA ([email protected])

xxii DEREK C.G. MUIR National Water Research Institute Environment Canada Burlington, Ontario L7R 4A6, Canada ([email protected]) CARRIE A. OLSEN Department of Earth and Environmental Sciences University of Illinois at Chicago 845 W. Taylor St. Chicago, Illinois 60607-7059, USA ([email protected]) REINHARD PIENITZ Paleolimnology-Paleoecology Laboratory Centre d’études nordiques Département de Géographie Université Laval Québec, Québec G1K 7P4, Canada ([email protected]) MICHAEL F.J. PISARIC Department of Geography and Environmental Studies Carleton University Ottawa, Ontario K1S 5B6, Canada ([email protected]) ROBERT J. POREDA Earth and Environmental Sciences University of Rochester Rochester, New York 14627, USA ([email protected]) DAVID F. PORINCHU Department of Geography Ohio State University Columbus, Ohio 43210-1002, USA ([email protected])

xxiii ROSS D. POWELL Department of Geology and Environmental Geosciences Northern Illinois University DeKalb, Illinois 60115, USA ([email protected]) JOHN C. PRISCU Land Resources and Environmental Sciences Montana State University Bozeman, Montana 59717, USA ([email protected]) DONNA ROBERTS Institute of Antarctic and Southern Ocean Studies University of Tasmania Private Bag 77 Hobart, Tasmania 7001 Australia ([email protected]) NEIL L. ROSE Environmental Change Research Centre University College London 26 Bedford Way London, WC1H 0AP United Kingdom ([email protected]) KATHLEEN RÜHLAND Paleoecological Environmental Assessment and Research Laboratory Department of Biology Queen’s University Kingston, Ontario K7L 3N6, Canada ([email protected]) DAVID B. RYVES Department of Geography Loughborough University Loughborough, Leicestershire LE11 3TU United Kingdom ([email protected])

xxiv I. ROD SMITH Geological Survey of Canada Terrain Sciences Division 3303 - 33rd St. NW Calgary, Alberta T2L 2A7, Canada ([email protected]) JOHN P. SMOL Paleoecological Environmental Assessment and Research Laboratory Department of Biology Queen’s University Kingston, Ontario K7L 3N6, Canada ([email protected]) JEFFREY A. SNYDER Department of Geology Bowling Green State University Bowling Green, Ohio 43403, USA ([email protected]) NADIA SOLOVIEVA Environmental Change Research Centre (ECRC) University College London London, WC1H 0AP United Kingdom ([email protected]) PAVEL TARASOV Department of Geography Moscow State University Moscow, 119899 Russia ([email protected]) IAN R. WALKER Departments of Biology and Earth and Environmental Sciences Okanagan University College 3333 University Way Kelowna, British Columbia V1V 1V7, Canada ([email protected])

xxv JAN WECKSTRÖM Environmental Change Research Unit (ECRU) Department of Biological and Environmental Sciences University of Helsinki P.O. Box 65 (Viikinkaari 1) FIN-00014 Helsinki Finland ([email protected]) ALEXANDER P. WOLFE Department of Earth and Atmospheric Sciences University of Alberta Edmonton, Alberta T6G 2E3, Canada ([email protected]) BRENT B. WOLFE Department of Geography and Environmental Studies Wilfrid Laurier University Waterloo, Ontario N2L 3C5, Canada ([email protected])

EDITORS AND BOARD OF ADVISORS OF DEVELOPMENTS IN PALEOENVIRONMENTAL RESEARCH BOOK SERIES Series Editors: John P. Smol Paleoecological Environmental Assessment and Research Lab (PEARL) Department of Biology Queen’s University Kingston, Ontario, K7L 3N6, Canada e-mail: [email protected] William M. Last Department of Geological Sciences University of Manitoba Winnipeg, Manitoba R3T 2N2, Canada e-mail: [email protected]

Advisory Board : Professor Raymond S. Bradley Department of Geosciences University of Massachusetts Amherst, MA 01003-5820 USA e-mail: [email protected] Professor H. John B. Birks Botanical Institute University of Bergen Allégaten 41 N-5007 Bergen Norway e-mail: [email protected] Dr. Keith Alverson Director, GOOS Project Office Intergovernmental Oceanographic Commission (IOC) UNESCO 1, rue Miollis 75732 Paris Cedex 15 France Tel: +33 (0)1-45-68-40-42 Fax: +33 (0)1-45-68-58-13 (or 12) e-mail: [email protected]

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DEVELOPMENTS IN PALEOENVIRONMENTAL RESEARCH BOOK SERIES http://www.springeronline.com/sgw/cda/frontpage/0,,4-40109-69-33113470-0,00.html http://home.cc.umanitoba.ca/~mlast/paleolim/dper.html

Series Editors: John P. Smol, Department of Biology, Queen's University Kingston, Ontario, Canada William M. Last, Department of Geological Sciences, University of Manitoba, Winnipeg,

Manitoba, Canada Volume 9:

Earth Paleoenvironments: Records Preserved in Mid-and Low-Latitude Glaciers Edited by L. D. Cecil, J. R. Green and L. G. Thompson Hardbound, ISBN 1-4020-2145-3, July 2004

Volume 8:

Long-term Environmental Change in Arctic and Antarctic Lakes Edited by R. Pienitz, M. S. V. Douglas and J. P. Smol Hardbound, ISBN 1-4020-2125-9, 2004

Volume 7:

Image Analysis, Sediments and Paleoenvironments Edited by P. Francus Hardbound, ISBN 1-4020-2061-9, forthcoming

Volume 6:

Past Climate Variability through Europe and Africa Edited by R. W. Battarbee, F. Gasse and C. E. Stickley Hardbound, ISBN 1-4020-2120-8, 2004

Volume 5:

Tracking Environmental Change Using Lake Sediments. Volume 5: Data Handling and Numerical Techniques Edited by H. J. B. Birks et al. Hardbound, forthcoming

Volume 4:

Tracking Environmental Change Using Lake Sediments. Volume 4: Zoological Indicators Edited by J. P. Smol, H. J. B. Birks and W. M. Last Hardbound, ISBN 1-4020-0658-6, June 2001

Volume 3:

Tracking Environmental Change Using Lake Sediments. Volume 3: Terrestrial, Algal, and Siliceous Indicators Edited by J. P. Smol, H. J. B. Birks and W. M. Last Hardbound, ISBN 1-4020-0681-0, June 2001

Volume 2:

Tracking Environmental Change Using Lake Sediments. Volume 2: Physical and Geochemical Methods Edited by W. M. Last and J. P. Smol Hardbound, ISBN 1-4020-0628-4, June 2001

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xxx Volume 1:

Tracking Environmental Change Using Lake Sediments. Volume 1: Basin Analysis, Coring, and Chronological Techniques Edited by W. M. Last and J. P. Smol Hardbound, ISBN 0-7923-6482-1, June 2001

A Continuation Order Plan is available for this series. A continuation order will bring delivery of each new volume immediately upon publication. Volumes are billed only upon actual shipment. For further information please contact the publisher.

1. PALEOLIMNOLOGICAL RESEARCH IN POLAR REGIONS: AN INTRODUCTION

REINHARD PIENITZ ([email protected]) Paleolimnology-Paleoecology Laboratory Centre d’études nordiques Département de Géographie Université Laval Québec, Québec G1K 7P4, Canada MARIANNE S.V. DOUGLAS ([email protected]) Paleoecological Assessment Laboratory Department of Geology University of Toronto Toronto, Ontario M5S 3B1, Canada and JOHN P. SMOL ([email protected]) Paleoecological Environmental Assessment and Research Laboratory Department of Biology Queen’s University Kingston, Ontario K7L 3N6, Canada

The world can tell us everything we want to know. The only problem for the w orld is that it doesn’t have a v oice. But the world’s indicators are there. They are always talking to us. Quitsak Tarkiasuk, Ivujivik (1997 Canadian Arctic Resources Committee)

Key words: Arctic, Antarctic, Polar, Lakes, Paleolimnology, Climate change, Environmental change

1 R. Pienitz, M.S.V. Douglas and J.P. Smol (eds), 2004. Long-term Environmental Change in Arctic and Antarctic Lakes. Springer. Printed in the Netherlands.

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Introduction There is growing consensus that we have entered the "Anthropocene" (Crutzen 2002), a time period during which the major changes in the global biosphere are primarily the result of human actions and their impacts on the environment. Anthropogenic activities have a history that, in some areas, stretches back over thousands of years. Yet it is the exceptional warming of the last decade - the warmest decade since instrumental records began - that has provoked widespread concern over human influences on climate. Both indigenous observations of environmental change (e.g., Canadian Arctic Resources Committee (CARC) 1997; Krupnik and Jolly 2002; Fox 2003) and instrumental data (e.g., Magnuson et al. 2000; Houghton et al. 2001; Moritz et al. 2002) provide evidence that the Earth’s circumpolar regions are especially sensitive to rapid biophysical and social changes. However, the nature and characteristics of 20th-century climate variability are difficult to define, as the period for which we have instrumental records coincides with the time during which the atmosphere has been increasingly enriched by greenhouse gases, and so we now live in a “no-analogue” biosphere. To put present-day conditions into context and to more clearly understand future global changes, we require an appreciation of long-term climatic and environmental changes. General circulation model (GCM) simulations of future climatic changes, which may be expected with higher levels of carbon dioxide (CO2) and other greenhouse gases, show a remarkable consensus that the effects of anthropogenic greenhouse warming will be particularly intense in polar regions (e.g., Nicholls et al. 1996; Kattenberg et al. 1996; ACIA 2001; Houghton et al. 2001). For example, while the most recent report by the Intergovernmental Panel on Climate Change (IPCC; Houghton et al. 2001) predicts an average global temperature rise of between 1.5 and 6°C by the end of the 21st century, increases are expected to be much more significant in the Arctic and Antarctic (Kattenberg et al. 1996; Houghton et al. 2001). GCMs also predict a complete loss of sea ice across the Arctic Ocean basin by the end of this century (Vincent et al. 2001 and references therein). This warming will especially influence the late-autumn and winter temperatures and precipitation, mainly due to the increased sea ice – albedo feedbacks during the winter period, with the thermal inertia of the mixed layer of the open ocean preventing substantial warming during the short summer period (Nicholls et al. 1996). Collectively, the interactions or feedback mechanisms between high latitude oceans (sea-surface temperature), the cryosphere (continental ice masses, sea ice, and snow cover), and the atmosphere amplify the effects of global change (Kattenberg et al. 1996; Nicholls et al. 1996) and are dominant mechanisms controlling circumpolar climate on decadal to centennial timescales (Dickinson et al. 1996). Not surprisingly, the circumpolar regions have been the foci of scientific debates regarding global environmental and climatic change, as changes will likely first be discernible in these regions (Anderson and Willebrand 1996). The amplification of the global warming signal and increased precipitation in high latitude regions are characteristic features of model predictions, and they are believed to be largely due to feedback mechanisms in which key roles are played by variations in snow and sea ice extent (albedo), the stability of the lower troposphere, and (in the Arctic) thawing of permafrost. However, discrepancies in GCM scenarios arise concerning the magnitude and spatial structure of high latitude warming (reviewed in Serreze et al. 2000). The projections of future changes are complicated, for example, by

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uncertainties about the effects of changes in cloud cover, albedo, vegetation cover, topography, as well as possible interactions involving stratospheric temperature, stratospheric ozone, and changes in other parts of the system. For these reasons, current estimates of future changes vary significantly. Regardless of the mechanisms, northern populations are already experiencing the consequences of rapid climatic change in the Arctic, such as increased variability and unpredictability of the weather, fewer extended periods of extreme cold, more frequent extreme weather events, such as high wind (summer storm) events and lightning, changes in the seasonal extent and distribution of sea ice cover, different lake ice conditions, the melting of permafrost leading to unstable soils, as well as modified conditions for hunting of wildlife and waterfowl (CARC 1997; Krupnik and Jolly 2002; Fox 2003). Observational/instrumental evidence for rapid climate change in the circumpolar regions Although there are pronounced differences between the Arctic and Antarctica (reviewed in Grémillet and Le Maho 2003), both high latitude regions have many similarities arising from a common response to Earth’s orbital parameters. For example, the cryosphere is the dominant feature in both regions, and they share common temperature and radiation regimes, ozone characteristics, sea ice ecosystems and polynyas, as well as pronounced aridity. Moreover, atmospheric and oceanographic circulation processes link the polar regions, and also significantly influence world climates. For these reasons, this book describes long-term paleolimnological research from both polar regions, as they are intrinsically connected to the global environment. Monitoring data from both hemispheres have recorded the warming of surface air temperatures in recent decades, with the largest temperature increases occurring over northern hemisphere land areas from about 40 to 70°N (Serreze et al. 2000). Temperature data from arctic stations over the period 1966-1995 indicate a general warming trend, with the greatest effects in the western Arctic (up to 0.7°C per decade; Weller 1998). In the Antarctic, warming is largely focused in the Peninsula region. The 3.4 ± 1.6°C per century (weighted by length of record) warming of the Antarctic Peninsula over the past 40 to 50 years is amongst the most rapid air temperature increases recorded on Earth (Quayle et al. 2002; Vaughan et al. 2003). Increased open sea conditions will be accompanied by increased evaporation into the overlying air masses that, at warmer temperatures, can hold more moisture for subsequent cloud formation and precipitation. There is wide regional variability in precipitation trends, from significant increases at certain locations (e.g., from the 1960s to 1990s at Spitsbergen (Svalbard); Hanssenbauer and Forland 1998) to significant decreases at other sites (e.g., Alaska; Curtis et al. 1998). There is ample evidence that polar regions are already experiencing major ecosystem changes related to recent climate change. Amongst the most striking evidence has been from records of the declining extent and thickness of multi-year sea ice across the Arctic Ocean. Over the period 1978 to 1998, sea ice cover diminished in area by 14% in winter (Johannessen et al. 1999), and by 44% in average thickness during the past three decades (Rothrock et al. 1999). The largest arctic ice shelf, the Ward Hunt Ice Shelf (83°N, 74°W; > 10 m-thick sea ice) along the northern coast of Ellesmere Island,

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retreated by 79% over the course of the 20th century, with evidence for substantial thinning of the remaining landfast ice in the 1980s and 1990s (Vincent et al. 2001) and break-up and associated loss of an ice-dammed lake in 2002 (Mueller et al. 2003). Likewise, the collapse of extensive ice shelves on both sides of the Antarctic Peninsula during the last two decades has been attributed to rapid regional warming in that region (Vaughan and Doake 1996; Vaughan et al. 2003). In contrast, the ecosystem of the Dry Valleys in eastern Antarctica responded to a recent 14-year cooling with a doubling in lake ice thickness and resulting in 6 to 9% per year decrease in primary productivity in the lakes (Doran et al. 2002). Arctic and antarctic regions are also experiencing changes in incident ultraviolet (UV) radiation flux as a result of declining concentrations of stratospheric ozone in spring which have been linked to the production of CFCs. Alarms concerning this relatively recent environmental problem were first raised in the Antarctic (Farman et al. 1985), and some of the biological effects were soon realized (e.g., Vincent and Roy 1993). An integrated overview of UV radiation and its effects on terrestrial, freshwater and marine arctic biota is presented in Hessen (2002) and on polar lakes in Helbling and Zagarese (2003). During the last decade, marked stratospheric ozone declines have been observed in the Arctic during winter (Staehelin et al. 2001). During spring, the resulting increases in surface erythemal UV radiation are estimated to be about 22% relative to the values in the 1970s (Madronich et al. 1998). This increase is likely to worsen in duration and severity in the future, in part associated with greenhouse gas effects on stratospheric cooling (Shindell et al. 1998). In the Antarctic, ozone depletion is most significant during spring when total atmospheric concentrations can decline up to 40%. Fluxes of UV-B (280-315 nm) have increased 6 to 14% since 1980 and are expected to persist until at least 2050 (WMO 2002). UV radiation has a broad range of photobiological and photochemical effects on aquatic ecosystems (Moran and Zepp 1997; Miller 1998; Vincent and Neale 2000), emphasizing the need to better understand the implications of changes in snow and ice cover for UV exposure in polar lakes. Why do we need paleolimnological data from arctic and antarctic regions? Despite clear signs of marked recent environmental changes in the circumpolar regions, we as yet have only a limited perspective on how arctic and antarctic climates and environmental conditions have varied in the past. To better understand and anticipate the magnitude, nature, and direction of future changes, it is essential to compare currentday conditions to records of past environmental changes. As long-term monitoring programmes in high latitude regions have only been established for the past several decades, indirect proxy methods must be used to infer these past conditions. Several international and national research initiatives have recognized the need for obtaining long-term, high-resolution paleoenvironmental records from which any recent changes can be compared. These include, for example, projects under the auspices of the IGBPPAGES (International Geosphere-Biosphere Programme - Past Global Changes), CAPE (Circum-Arctic PaleoEnvironments), NSF-PALE (National Science Foundation Paleoclimates of Arctic Lakes and Estuaries), NSF-PARCS (Paleoenvironmental Arctic Sciences), and SCAR (Scientific Committee on Antarctic Research). The two fundamental goals common to many of these programmes are: (1) to gain a more

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complete understanding of the role and impact of the polar regions in global change issues; and (2) to provide a longer temporal baseline against which to measure ongoing changes in polar environments. As summarized by Bradley (1999), a large number of paleoenvironmental approaches are available to reconstruct climatic and other environmental changes from natural archives (e.g., ice cores, dendroecology, corals, pack rat middens). Whilst some of these records, such as ice cores, have provided important paleoenvironmental information from high latitude regions, many of these natural archives are more applicable in temperate and tropical regions. Fortunately, a characteristic feature of most arctic and antarctic landscapes is the large number of lakes and ponds. For example, although 750,000 km2 of Canada’s land mass is covered by freshwater, 18% of these surface waters are north of latitude 60°N (Prowse and Ommanney 1990). Arctic lakes and ponds, and the biota they support, are sentinels of ecosystem change (Douglas and Smol 1994, 1999; Vincent and Pienitz 1996; Rouse et al. 1997). The information contained in the sediments accumulating in each of these lakes and ponds potentially offers a sensitive record of past environmental changes (Smol and Cumming 2000). Similar lakes are present on many antarctic ice-free coastlines. Paleolimnology, which is the study of the physical, chemical and biological information stored in lake and river sediments (Smol 2002; Cohen 2003), offers considerable potential for reconstructing the long-term trends in environmental and climatic conditions. Paleolimnological records can be determined for periods well before any marked human influences on the global environment, and can thereby provide a unique guide to natural baseline conditions, as well as to the rate of environmental changes up to the modern day. The abundance of lakes and ponds throughout the circumpolar regions makes paleolimnological approaches especially powerful tools to assist in the reconstruction and interpretation of long-term environmental changes. The overall goal of this book is to provide the reader with an appreciation of the broad spectrum of techniques available for generating historical records, their respective potentials and limits, as well as to provide an overview of the geographic extent of paleolimnological work completed thus far in circumpolar regions. As noted earlier, the Arctic and Antarctica are likely to experience profound changes in the coming decades. However, due to the lack of direct monitoring data some of the most fundamental environmental questions cannot be addressed. For example: (1) Have these ecosystems changed over time, and if so, when and by how much? (2) What is the range of natural variability in these regions? (3) If systems have changed, were these shifts the result of natural phenomena or related to human activities? Although direct monitoring data are not available, paleoenvironmental data can be used in lieu of these missing data sets. As evidenced by the contents of this volume, the Arctic and Antarctica display a great diversity of paleoenvironmental records. Polar biotic and abiotic processes are strongly governed by climate and the large seasonal differences between the relatively productive summers and non-productive winters. Extreme seasonality is therefore a dominant factor for the geographical distribution and adaptation of polar life (e.g., Billings 1987). This enhances the role of climate forcing over non-climatic processes in these depositional environments and makes the identification of climate forcing in sedimentary records more likely (Smol 1988). There is already widespread and accumulating paleolimnological evidence for dramatic changes in terrestrial and aquatic

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communities across northern circumpolar regions within the last two centuries (e.g., Douglas et al. 1994; Overpeck et al. 1997; Rouse et al. 1997; Joynt and Wolfe 2001; Wolfe and Perren 2001; Korhola et al. 2002; Sorvari et al. 2002; Perren et al. 2003, Rühland et al. 2003), which has been attributed at least partly to anthropogenicallyinduced climatic warming. However, as expected, some studies have also shown that the so-called “global climate change” may have regionally varying impacts (e.g., Laing et al. 2002; Ponader et al. 2002; Paterson et al. 2003; Wrona et al. 2004), as has also been shown for Antarctica (Thompson and Solomon 2002; Gillett and Thompson 2003). Terrestrial records with decadal to centennial resolution exist in the form of lake sediments and peat deposits, whereas annually resolved records can be found in varved lake records (Lamoureux and Gilbert, this volume) and ice cores (e.g., Reeh 1989), which have been a central focus of high latitude paleoenvironmental research over the last few decades (e.g., PARCS 1999; CAPE 2001). Other terrestrial records, such as aeolian deposits (mainly sand sheets and thin loess deposits), which are widespread along proglacial outwash, may provide important complementary information on similar or longer timescales (Eisner et al. 1995; Lamoureux and Gilbert, this volume). Many of the lakes that are located close to present-day ice sheets and glaciers receive considerable amounts of aeolian sediments, and their records may provide a continuous archive of both paleolimnological changes and aeolian activity. Paleolimnological studies may also help elucidate the origins and pathways of atmospheric pollutants in the polar regions (Muir et al. 2002; Muir and Rose, this volume). Many scientists now characterize the Arctic and Antarctica as “pollution sinks”, the final resting places for many contaminants used in industry and agriculture thousands of kilometres away. These contaminants, in particular organochlorines, are persistent, entering the food chain and bioaccumulating at each trophic level. Inuit and other aboriginal peoples may be exposed to health risks as they ingest such contaminants when eating traditional food, including freshwater organisms (e.g., arctic char). Lake sediments form ideal archives for studying the patterns of paleoenvironmental change and pollution in circumpolar regions (PARCS 1999) for many reasons, including: (1) Sources of high-resolution paleoenvironmental records, such as ice cores and tree rings, often provide important environmental information at high temporal resolution, though their availability may often be limited spatially (e.g., trees are absent in more extreme regions, ice caps are limited in geographic coverage). Lakes and ponds, on the other hand, have excellent spatial coverage in arctic and coastal antarctic regions. (2) Many lake basins contain reliable sedimentary records which, given sufficient dating control, allow for continuous reconstructions of environmental change extending back thousands of years. (3) Paleoecological reconstructions of climate change are likely to be more reliable if derived from areas with little impact from local human activities such as forestry, agriculture and local pollution. Such impacts can impair our ability to estimate the relationship between individual proxies and climate, as the community composition also responds to changes unrelated to climatic conditions (e.g., nutrients, pH and erosion). The circumpolar regions are the least densely

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populated areas of the planet and therefore contain some of the most pristine environments, or at least almost always unaffected by local disturbances. (4) The biological, chemical and physical dynamics of high latitude lakes and ponds are closely linked to climate and meteorological variability at a variety of timescales (e.g., Welch et al. 1987; Hostetler 1995; Doran et al. 1996; Hardy et al. 1996; Schindler et al. 1996). (5) Polar lake ecosystems have simplified food webs relative to lower latitudes, and so even minor climatic shifts are expected to generate relatively large changes in biota and depositional processes. Geographic scope of the book There is considerable debate as to how the polar and subpolar regions should be circumscribed and defined. The terms Arctic, Antarctic, Subarctic and Subantarctic are commonly used, yet ecologists, climatologists, administrators and politicians often use them in different ways (discussed in Osherenko and Young 1989; AMAP 1998; Hansom and Gordon 1998; Hamelin 2000; Nuttall and Callaghan 2000; Przybylak 2003; Sater 2003). Boundaries exist on maps primarily for political convenience, but in reality their positions reflect a continuum or gradients of environmental variables. In addition, these boundaries may also be time-transgressive due to changes, in space or in time, of environmental patterns, such as temperature or sea ice extent. Despite the lack of conformity concerning boundaries and definitions, it seems helpful to sketch the broad limits of the regions covered in this book by referring to the most commonly used delineations and definitions (Figures 1 and 2). In this volume we refer mainly to ecoclimatic regions and boundaries, which integrate ecological and climatological aspects, such as the position of treeline and the 10°C summer isotherm (corresponding respectively to the July and February isotherms in the northern and southern hemispheres). Some ecologists prefer the summer isotherms to other polar boundaries as they are equally valid over land and ocean, and on land correspond fairly closely with the position of northern or southern treeline (Larsen 1989). Further, the same criteria can be used for both hemispheres and can provide a more solid basis for comparison, although this comprises a much wider area in the south than in the north. In fact, winters are colder over much of the Subarctic in continental northern Canada and Siberia than in the equivalent subantarctic zone, as the latter covers a wide expanse of ocean and scattered small islands where maritime influences stabilize temperatures. In the Arctic, the position of treeline (Figure 1) is often a more appealing boundary to ecologists, as treeless tundra is an easily recognized biome, characterized by particular kinds of soil, vegetation and fauna. Also, the presence or absence of trees can be plotted accurately over large areas, whereas isotherms are drawn from records of stations that are generally few and far between. Although the position of treeline often matches closely the 10°C July isotherm for much of its length (Bryson 1966), reflecting the strong influence of summer temperatures on tree growth in both hemispheres, departures arise because treeline location is also determined by many other physical, geological, biological, and social factors, including climate (winds), topography, soil characteristics (permafrost presence or absence), biological interactions such as grazing,

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and human pressure such as farming or logging (e.g., in Norway the treeline-forming species mountain birch shifted northward and upward because of both climate change and grazing by reindeer). Thus, for the terrestrial Arctic, the position of treeline is a useful practical boundary, whilst the 10°C July isotherm is a useful conceptual one. There is no clearly defined southern (biological) boundary for the Subarctic, whereas in the southern hemisphere, treeline is influenced by oceanographic (water mass) boundaries on the southern tip of South America and the subantarctic islands.

Figure 1 . The Arctic as defined by different boundaries discussed in the text (after Stonehouse 1989).

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Figure 2. The Antarctic as defined by different boundaries discussed in the text (after Stonehouse 1989).

Antarctica is defined as the Antarctic Continent, whereas the Antarctic has a variety of definitions but is loosely defined as the area south of 60°S. An alternative definition is that the Antarctic lies south of the natural boundary formed by the Polar Frontal Zone or Antarctic Convergence (Figure 2), the region where the cold and dense Antarctic waters meet and sink beneath the warmer and less dense waters of the Pacific, Atlantic and Indian oceans. This oceanographic boundary, which is defined by the convergence of oceanic waters, encircles the continent between latitudes 50°S and 60°S (on average at about 58°S), and roughly corresponds to the 10°C February isotherm (Hansom and Gordon 1998). The subantarctic islands, including Bouvet Island, Îles Crozet, Heard

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Island, Îles Kerguelen, Macdonald Islands, Macquarie Island, Marion Island, Prince Edward Island, South Georgia, and South Sandwich Islands, lie close to, but mostly north of, the Polar Frontal Zone. There is also a legal political definition, embodied in the Antarctic Treaty, which defines Antarctica as south of 60°S. Though the terms “Antarctica” and “Antarctic” are often used interchangeably, the former generally refers to the continent itself, whereas the latter denotes the region that includes both ocean and the continent. Due to the different ways of delineating the boundaries of the Earth’s polar regions, we refrained from imposing rigid definitions for the book chapters. Instead, authors defined geographic boundaries in their respective chapters, which in fact were very consistent between contributions. The focus and the structure of this volume Aims and scope As is evident from our title “Long-Term Environmental Change in Arctic and Antarctic Lakes”, the over-arching theme of this volume is the paleolimnology of high latitude regions, and not long-term environmental change in general. The time frame covered by the chapters emphasizes late Pleistocene and Holocene events, although other time frames are discussed when appropriate. Each chapter represents a peer-reviewed, up-todate overview of key aspects of high latitude paleolimnological research by specialists in their fields. Our overall goal is to present an accessible introduction to a nonspecialist readership, but also provide a relevant source for paleolimnologists working in high latitude regions. Despite their importance, the human and social aspects of rapidly changing environments in the circumpolar regions are beyond the scope of this book. Those who are interested in these aspects may refer to other textbooks (e.g., Osherenko and Young 1989; Nuttall and Callaghan 2000). Also, a comprehensive presentation (and overview) of modern arctic and antarctic environments, climate, biota, permafrost and vegetation zonation is beyond the scope of this volume and is not included here. Reviews of these and other diverse themes are provided in Sugden (1982), Stonehouse (1989), Chaturvedi (1996), King and Turner (1997), Hansom and Gordon (1998), Nuttall and Callaghan (2000), Huiskes et al. (2003), Przybylak (2003) and Wonders (2003), to mention just a few sources. Structure The general structure of this book is divided into two parts: More specific methodological aspects related to different indicators or approaches are presented in “theme” chapters in Part One, with a focus on their strengths, weaknesses, and challenges. As a series of chapters summarizing paleolimnological techniques has recently been published (Last and Smol 2001a,b; Smol et al. 2001a,b), the focus of the introductory chapters in this volume is on the specialized literature dealing with high latitude regions.

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A.P. Wolfe, G.H. Miller, C.A. Olsen, S.L. Forman, P.T. Doran and S.U. Holmgren begin with an introduction to the approaches and problems of dating lake sediment cores in high latitude regions, such as low sedimentation rates, low 210Pb concentrations and the paucity of terrestrial macrofossils (Chapter 2). In their chapter on the physical and chemical properties and proxies of high latitude lake sediments (Chapter 3), S.F. Lamoureux and R. Gilbert provide an overview of present-day processes that govern sediment formation in glacial and periglacial lake systems, and how these processes might be reflected in sedimentary records. Their chapter also introduces methodologies to investigate sediments from these remote sites, including dating, mineralogy, sediment texture, organic matter and trace element analyses. K. Gajewski and G.M. MacDonald discuss pollen and charcoal studies of ice cores and lacustrine records in arctic regions, including a synthesis of some of the palynological methods used in polar regions (Chapter 4). Algal indicators and their potential as proxies of long-term environmental change in high latitude lakes are summarized in Chapter 5 by M.S.V. Douglas, P. Hamilton, R. Pienitz and J.P. Smol. A variety of morphological and biogeochemical indicators and their applications in studies of high latitude environmental and climatic changes (e.g., tracking past changes in the underwater penetration of ultraviolet radiation; the effects of local human and animal populations on water bodies) are discussed. Chapter 6 by O. Bennike, K.P. Brodersen, E. Jeppesen and I.R. Walker reviews the use of aquatic invertebrates as proxy indicators in high latitude paleolimnological studies. The chapter focuses primarily on the usefulness of the remains of testate amoebae, chironomids and cladocerans as paleoclimate proxies recovered from Holocene lake sediments, with reference mainly to studies from Greenland and northern Europe. The potential of stable isotopes as tracers of environmental change in polar lakes is explored by T.W.D. Edwards, B.B. Wolfe, J.J. Gibson and D. Hammarlund (Chapter 7). Their chapter reviews major processes controlling isotopic fractionation, as well as other aspects concerning the use of stable water isotope tracers in lake sediment records. This overview is followed by selected case studies of modern isotopic hydrology at regional to watershed scales and paleohydrological applications at sites in northern Canada and Sweden. In the final chapter of Part One, the pollution of arctic and antarctic lakes, as revealed by their sedimentary archives, is summarized by D.C.G. Muir and N.L. Rose (Chapter 8). The authors critically examine studies on heavy metals, persistent organic pollutants (POPs), and anthropogenic particles in sediment cores, while demonstrating their value and potential for the assessment of large-scale spatial and temporal trends of atmospheric contaminants and the pathways of exposure to wildlife and humans in polar regions. Part Two of this volume consists of a series of regional syntheses that provide overviews of paleolimnological work completed in high latitude regions from both hemispheres. In their review of paleolimnological studies from the Canadian Mid- and High Arctic (Chapter 9), A.P. Wolfe and I.R. Smith summarize investigations into late Quaternary and Holocene climatic and environmental changes in regions north of mainland Canada

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(including the Arctic Archipelago and Boothia and Melville peninsulas), with special emphasis on case studies from Ellesmere and Baffin islands. B. Finney, K. Rühland, J.P. Smol and M.-A. Fallu provide an overview of paleolimnological work completed in the North American Subarctic (Chapter 10), including studies from Alaska, Yukon, Northwest Territories, southern Nunavut, northern Québec, and Labrador. The importance of the treeline ecotone and permafrost is also emphasized in this chapter. The Holocene paleolimnology of Greenland and the North Atlantic Islands, including Iceland, Svalbard, Jan Mayen, Bjørnøya, and the Faeroe Islands is surveyed in Chapter 11, authored by N.J. Anderson, D.B. Ryves, M. Grauert and S. McGowan. They elaborate on the potential role of paleolimnological research from the North Atlantic region in studies addressing questions related to climate variability, including the North Atlantic Oscillation (NAO), as well as post-glacial refugia and pathways of species migration. In their chapter on paleolimnological research from northern Russian Eurasia, G.M. MacDonald, T.W.D. Edwards, B. Gervais, T.E. Laing, M.F.J. Pisaric, D.F. Porinchu, J.A. Snyder, N. Solovieva, P. Tarasov, and B.B. Wolfe present examples of recent biological and stable isotope records from lakes in the far north of European Russia (Kola Peninsula) and Siberia (Chapter 12). They also provide a synthesis of Holocene paleohydrological changes based on Russian lake-level data sets. A. Korhola and J. Weckström complete the geographical coverage of work done in the northern hemisphere with their review of paleolimnological studies completed in the northwestern European Arctic, including arctic Fennoscandia and the Kola Peninsula (Chapter 13). Aspects of long-term environmental change in the Earth’s southern hemisphere are dealt with by D.A. Hodgson, P.T. Doran, D. Roberts, and A. McMinn (Chapter 14). Their comprehensive review discusses many fundamental and practical aspects of paleolimnological studies in Antarctica and the Subantarctic Islands, links to other paleo-disciplines (e.g., ice core studies), and the relevance of this work to Earth system science and global change research. In the final chapter of Part Two, P.T. Doran, J.C. Priscu, W.B. Lyons, R.D. Powell, D.T. Andersen, and R.J. Poreda provide insights into the historical records preserved in extreme cold lake habitats, including Lake Vostok and the Dry Valley lakes in Antarctica, and their relevance for a better understanding of the origin and evolution of life on Earth (Chapter 15). They also discuss the connection between these extreme polar environments and water bodies that may exist elsewhere in our solar system, notably on Mars and Jupiter’s moon Europa. Finally, a short epilogue presents a brief synthesis on the strengths and weaknesses of different paleolimnological methods and approaches used in arctic and antarctic regions. It briefly defines some of the most important scientific problems and provides an outlook on future paleolimnological challenges. The volume concludes with a Glossary and Index. Paleolimnological research in polar regions has seen tremendous advances and progress over the last two decades. We hope this volume will provide a useful synthesis of research completed to date, and serve as a point of departure for the many exciting new projects to be completed in arctic and antarctic regions.

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Acknowledgements Polar research in our laboratories is primarily funded by the Natural Sciences and Engineering Research Council of Canada and the Polar Continental Shelf Project. We thank P.T. Doran, D.A. Hodgson, K. Rühland, and W.F. Vincent for helpful comments on this chapter. References ACIA 2001. Report of the Arctic Climate Impact Assessment Scoping Workshop, February 28March 1, 2000, Washington, DC, USA, 54 pp. AMAP 1998. AMAP Assessment Report: Arctic Pollution Issues. Arctic Monitoring and Assessment Programme (AMAP), Oslo, Norway, 859 pp. Anderson D.L.T. and Willebrand J. (eds) 1996. Decadal climate variability, dynamics and predictability. NATO Advanced Science Institute Series I44, Springer-Verlag, Berlin. Billings W.D. 1987. Constraints to plant growth, reproduction and establishment in arctic environments. Arct. Alp. Res. 19: 357-365. Bradley R.S. 1999. Quaternary Paleoclimatology. Academic Press, San Diego, 610 pp. Bryson R.A. 1966. Air masses, streamlines and the boreal forest. Geographical Bull. 8: 228-269. Cohen A.S. 2003. Paleolimnology: The History and Evolution of Lake Systems. Oxford University Press, New York, 500 pp. CAPE (Circumpolar Arctic PaleoEnvironments) Project Members 2001. Holocene paleoclimate data from the Arctic: testing models of global climate change. Quat. Sci. Rev. 20: 1275-1287. CARC (Canadian Arctic Resources Committee) 1997. Voices from the bay: traditional ecological knowledge of Inuit and Cree in the Hudson Bay bioregion. CARC Ottawa, Canada, 98 pp. Chaturvedi S. 1996. The Polar Regions. A Political Geography. John Wiley, Chichester, 806 pp. Crutzen P.J. 2002. Geology of mankind. Nature 415: 23. Curtis J., Wendler G., Stone R. and Dutton E. 1998. Precipitation decrease in the western Arctic, with special emphasis on Barrow and Barter Island, Alaska. Int. J. Climatol. 18: 1687-1707. Dickinson R.E., Meleshko V., Randall D., Sarachik E., Silva-Dias P. and Slingo A. 1996. Climate Processes. In: Houghton J.T., Jenkins G.J. and Ephraums J.J. (eds), Climate Change, the IPCC Scientific Assessment. Cambridge, Cambridge University Press, pp. 193-227. Doran P.T., McKay C.P., Adams W.P., English M.C., Wharton R.A. and Meyer M.A. 1996. Climate forcing and thermal feedback of residual lake-ice cover in the high Arctic. Limnol. Oceanogr. 41: 839-848. Doran P.T., Priscu J., Lyons W., Walsh J., Fountain A., McKnight D., Moorhead D., Virginia R., Wall D., Clow G., Fritsen C., McKay C. and Parsons A. 2002. Antarctic climate cooling and terrestrial ecosystem response. Nature 415: 517-520. Douglas M.S.V. and Smol J.P. 1994. Limnology of high arctic ponds (Cape Herschel, Ellesmere Island, N.W.T.). Arch. Hydrobiologie 131: 401-434. Douglas M.S.V. and Smol J.P. 1999. Freshwater diatoms as indicators of environmental change in the High Arctic. In: Stoermer E.F. and Smol J.P. (eds), The Diatoms: Applications for the Environmental and Earth Sciences. Cambridge University Press, Cambridge, pp. 227-244. Eisner W.E., TĘrnqvist T.E., Koster E.A., Bennike O. and Van Leeuwen J.F.N. 1995. Paleoecological studies of a Holocene lacustrine record from the Kangerlussuaq (Søndre Strømfjord) region of West Greenland. Quat. Res. 43: 55-66. Farman J.C., Gardiner B.G. and Shanklin J.D. 1985. Large losses of total ozone in Antarctica reveal seasonal ClOx/NOx interaction. Nature 315: 207-210. Fox S. 2003. When the Weather is Uggianaqtuq: Inuit Observations of Environmental Change. Boulder, Colorado: National Snow and Ice Data Center. Digital media.

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Ponader K., Pienitz R., Vincent W.F. and Gajewski K. 2002. Limnological conditions in a subarctic lake (northern Québec, Canada) during the late Holocene: Analyses based on fossil diatoms. J. Paleolim. 27: 353-366. Prowse T.D. and Ommanney C.S.L. (eds) 1990. Northern Hydrology: Canadian Perspectives. NHRI Science Report No. 1. Environment Canada, Ottawa, 308 pp. Przybylak R. 2003. The Climate of the Arctic. Atmospheric and Oceanographic Sciences Library vol. 26, Kluwer Academic Publishers, Dordrecht, The Netherlands, 270 pp. Quayle W.C., Peck L.S., Peat H., Ellis-Evans J.C. and Harrigan P.R. 2002. Extreme responses to climate change in Antarctic lakes. Science 295: 645. Reeh N. 1989. Dynamic and climatic history of the Greenland Ice Sheet. In: Fulton R.J. (ed.), Quaternary Geology of Canada and Greenland. Geological Survey of Canada, Geology of Canada 1: 793-822. Rothrock D.A., Yu Y. and Maykut G.A. 1999. Thinning of Arctic sea ice cover. Geophys. Res. Lett. 26: 3469-3472. Rouse W., Douglas M., Hecky R., Kling G., Lesack L., Marsh P., McDonald M., Nicholson B., Roulet N. and Smol J. 1997. Effects of climate change on fresh waters of Region 2: Arctic and Sub-Arctic North America. Hydrologic Proc. 11: 873-902. Rühland K., Priesnitz A. and Smol J.P. 2003. Evidence for recent environmental changes in 50 lakes the across Canadian arctic treeline. Arct. Ant. Alp. Res. 35: 110-123. Sater J.E. 2003. The Arctic Basin and the Arctic: Some definitions. In: Wonders W.C. (ed.), Canada's Changing North. McGill-Queen's University Press, Montreal & Kingston, pp. 3-7. SCAR 1993. The role of Antarctica in Global Change. An International Plan for a regional Research Programme. Scientific Committee on Antarctic Research, Cambridge. Schindler D.W., Bayley S.E., Parker B.R., Beaty K.G. and Cruikshank D.R. 1996. The effects of climate warming on the properties of boreal lakes and streams at the Experimental Lakes Area, northwestern Ontario. Limnol. Oceanogr. 41: 1004-1017. Serreze M.C., Walsh J.E., Chapin F.S. III, Osterkamp T., Dyurgerov M., Romanovsky V., Oechel W.C., Morison J., Zhang T. and Barry R.G. 2000. Observational evidence of recent change in the northern high-latitude environment. Climatic Change 46: 159-207. Shindell D., Rind D. and Lonergan P. 1998. Increased stratospheric ozone losses and delayed eventual recovery owing to increasing greenhouse-gas concentrations. Nature 392: 589-592. Smol J.P. 1988. Paleoclimatic proxy data from freshwater Arctic diatoms. Verh. Int. Ver. Limnol. 23: 37-44. Smol J.P. 2002. Pollution of Lakes and Rivers: A Paleoenvironmental Perspective. Arnold Publishers, London, 280 pp. Smol J.P. and Cumming B.F. 2000. Tracking long-term changes in climate using algal indicators in lake sediments. J. Phycology 36: 986-1011. Smol J.P., Birks H.J.B. and Last W.M. (eds) 2001a. Tracking Environmental Change Using Lake Sediments. Volume 3: Terrestrial, Algal, and Siliceous Indicators. Kluwer Academic Publishers, Dordrecht, The Netherlands, 371 pp. Smol J.P., Birks H.J.B. and Last W.M. (eds) 2001b. Tracking Environmental Change Using Lake Sediments. Volume 4: Zoological Indicators. Kluwer Academic Publishers, Dordrecht, The Netherlands, 217 pp. Sorvari S., Korhola A. and Thompson R. 2002. Lake diatom response to recent Arctic warming in Finnish Lapland. Global Change Biol. 8: 171-181. Staehelin J., Harris N.R.P., Appenzeller C. and Eberhard J. 2001. Ozone trends: a review. Rev. Geophysics 39: 231-290. Stonehouse B. 1989. Polar Ecology. Blackie and Son Ltd., Glasgow/London, 222 pp. Sugden D.E. 1982. Arctic and Antarctic: A Modern Geographical Synthesis. Blackwell, Oxford, 472 pp. Thompson D.W.J. and Solomon S. 2002. Interpretation of recent Southern Hemisphere climate change. Science 296: 895-899.

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Vaughan D.G. and Doake C.S.M. 1996. Recent atmospheric warming and retreat of ice shelves on the Antarctic Peninsula. Nature 379: 328-331. Vaughan D.G., Marshall G., Connolley W.M., Parkinson C., Mulvaney R., Hodgson D.A., King J.C., Pudsey C.J., Turner J. and Wolff E. 2003. Recent rapid regional climate warming on the Antarctic Peninsula. Climatic Change 60: 243-274. Vincent W.F. and Roy S. 1993. Solar ultraviolet-B radiation and aquatic primary production: damage, protection, and recovery. Environ. Rev. 1: 1-12. Vincent W.F. and Pienitz R. 1996. Sensitivity of high latitude freshwater ecosystems to global change: temperature and solar ultraviolet radiation. Geoscience Canada 23: 231-236. Vincent W.F. and Neale P.J. 2000. Mechanisms of UV damage to aquatic organisms. In: de Mora S.J., Demers S. and Vernet M. (eds), The Effects of UV Radiation in the Marine Environment. Cambridge University Press, United Kingdom, pp. 149-176. Vincent W.F., Gibson J.A.E. and Jeffries M.O. 2001. Ice shelf collapse, climate change and habitat loss in the Canadian High Arctic. Polar Rec. 37: 133-142. Welch H.E., Legault J.A. and Bergmann M.A. 1987. Effects of snow and ice on the annual cycles of heat and light in Saqvaqjuac lakes. Can. J. Fish. Aquat. Sci. 44: 1451-1461. Weller G. 1998. Regional impacts of climate change in the Arctic and Antarctic. Ann. Glaciol. 27: 543-552. WMO 2002. Scientific Assessment of Ozone Depletion: 2002, Global Ozone Research and Monitoring Project – Report No. 47. World Meteorological Organization, Geneva, Switzerland. Wolfe A.P. and Perren B.B. 2001. Chrysophyte microfossils record marked responses to recent environmental changes in high- and mid-arctic lakes. Can. J. Bot. 79: 747-752. Wonders W.C. (ed.) 2003. Canada's Changing North. Revised edition, McGill-Queen's University Press, Montreal & Kingston, 449 pp. Wrona F. et al. 2004. Arctic Climate Change Impact Assessment (ACIA), Freshwater Ecosystems Chapter (in press).

2. GEOCHRONOLOGY OF HIGH LATITUDE LAKE SEDIMENTS

ALEXANDER P. WOLFE ([email protected]) Department of Earth and Atmospheric Sciences University of Alberta Edmonton, Alberta T6G 2E3, Canada GIFFORD H. MILLER ([email protected]) Institute of Arctic and Alpine Research and Department of Geological Sciences University of Colorado Boulder, Colorado 80309-0450, USA CARRIE A. OLSEN ([email protected]) STEVEN L. FORMAN ([email protected]) PETER T. DORAN ([email protected]) Department of Earth and Environmental Sciences University of Illinois at Chicago 845 W. Taylor St. Chicago, Illinois 60607-7059, USA and SOFIA U. HOLMGREN ([email protected]) Department of Earth Sciences Geovetarcentrum Göteborgs Universitet SE-405 30 Göteborg Sweden

Key words: Arctic, Antarctica, Lake sediment chronology, Baffin Island, Optically stimulated luminescence

210

Pb,

137

Cs, 14C, Accelerator mass spectrometry,

19 R. Pienitz, M.S.V. Douglas and J.P. Smol (eds), 2004. Long-term Environmental Change in Arctic and Antarctic Lakes. Springer. Printed in the Netherlands.

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Introduction This chapter reviews the most relevant techniques for dating lake sediments from the Arctic and Antarctica, with specific reference to the challenges encountered in environments characterized by prolonged lake ice cover, low terrestrial and aquatic biological production, slow rates of organic matter decomposition, and periglaciation. We focus on applications of three general techniques: short-lived radioisotopes (210Pb and 137Cs), radiocarbon (14C), and optically stimulated luminescence (OSL), which collectively offer the potential for dating sediments ranging from decades to 105 years in age. Our focus is placed on the strengths and weaknesses of these techniques in establishing a chronostratigraphic framework for reconstructing environmental change in high latitude environments, with only brief theoretical and methodological considerations. The underlying principles of each of these methods are covered extensively by complementary reviews in the DPER book series (Appleby 2001; Björck and Wohlfarth 2001; Lian and Huntley 2001), as are additional dating techniques including varved sediments (Lamoureux 2001) and tephrochronology (Turney and Lowe 2001). In general, high latitude lakes have significantly lower sediment accumulation rates relative to temperate counterparts. Cold climates (mean annual temperatures < 0°C) result in the prolongation of seasonal lake ice cover, coupled to a shortening of the hydrological season. Lake ice strongly limits biological production and hence autochthonous sedimentation. Lake hydrology is typically dominated by summer snowmelt and remains largely inactive during winter, thus narrowing the window available for clastic (allochthonous) sedimentation. Although these unique environmental attributes present several challenges for the successful development of arctic and antarctic lake sediment chronologies, recent progress suggests that most high latitude lake sediments can be successfully dated over a broad range of late Quaternary timescales. Dating recent high latitude lake sediments using 210Pb and 137Cs Sediment accumulations of fallout 210Pb and 137Cs have been used extensively to date recent events in lacustrine environments. 210Pb (half-life, t1/2 = 22.3 yr) is a natural radioactive daughter in the decay series of 238U and is produced both in the atmosphere and the subsurface. The decay of 226Ra (t1/2 = 1602 yr) in soil and rock produces 222Rn (t1/2 = 3.82 days), a portion of which diffuses through particle interstices and escapes to the atmosphere. Fallout 210Pb is produced by the decay of emanated 222Rn and resides in the atmosphere in the order of 10 days. This is termed unsupported or excess 210Pb, which provides the basis for 210Pb geochronology. Terrestrial, or supported 210Pb, is produced by the in situ decay of particulate 226Ra and its daughter 222Rn, both associated with the presence of U-bearing mineral phases. In an ideal sediment profile, total (supported + unsupported) 210Pb activity decreases exponentially with depth to a point where it becomes constant, marking the depth at which only supported 210Pb exists. Given its relatively short half-life, 210Pb dating is only applicable to sediments younger than ca. 150 yr. Several assumptions must be made when applying the 210Pb methodology as a recent geochronometer: (1) fallout of 210Pb has been constant over the

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interval studied; (2) deposition of 210Pb in sediment occurs within a closed system; (3) 210 Pb is immobile within the sediment column; and (4) the sediment does not undergo extensive mixing once 210Pb is deposited. Although the first two assumptions generally appear valid, the last two are more frequently violated and may potentially introduce error to age determinations. Confidence in the 210Pb method is enhanced when coupled measurements of 137Cs activity are available, or when cross-checked with annually-laminated sediments. 137Cs is an anthropogenic radionuclide first introduced into the atmosphere in measurable quantities with the onset of nuclear weapons testing ca. 1954. The maximum atmospheric fallout of 137Cs occurred in 1963, immediately prior to the ban of nuclear testing. In some northern regions, the 1986 Chernobyl reactor meltdown has created a detectable secondary peak (Johnson-Pyrtle et al. 2000). Thus, 137Cs activity profiles within sediment columns may ideally encompass all three of these known marker horizons. 137Cs has a short half-life (t1/2 = 30.1 yr) and displays similar chemical and physical behaviour to 210Pb. Since the pioneering applications of 210Pb and 137Cs geochronology to lake sediments (Krishnaswamy et al. 1971), both methods have been widely used in temperate and, more recently, high latitude lakes. Fundamental differences between these environments have important implications for their successful application. Physical and chemical processes associated with 210Pb and 137Cs Successful dating by 210Pb requires a progressive decrease in activity with increasing depth in the sediment, permitting the modeling of sedimentation rates using various approaches, most commonly the CRS (constant rate of supply) and CIC (constant initial concentration) models (Appleby and Oldfield 1992; Appleby 2001). However, lake sediments may experience various degrees of mixing by wave action, turbidity currents, turnover, bioturbation, and mass movement, which can homogenize or dilute the stratigraphic distribution of short-lived radionuclides. Transport of fine-grained littoral sediments into deeper regions of a lake can result in significant inter-basin variability of 210 Pb and 137Cs inventories in temperate lakes (Crusius and Anderson 1995a,b). In general, non-glacial high latitude lakes are less affected by these processes for several reasons. First, prolonged ice cover (8-12 months per year) greatly reduces, or even eliminates, sediment redistribution by wave and current energy. Second, burrowing communities attain lower densities in the benthos of most high latitude lakes, minimizing the potential of bioturbation. On the other hand, sparse vegetation, often accompanied by significant local relief, increases the potential for mass movements which are typically expressed in adjacent deep-water sediments as turbidites. Periglacial processes associated with seasonal freeze-thaw cycles also have the potential to remobilize hillslope sediments surrounding lake basins (e.g., French and Guglielmin 2000). It is therefore imperative that the disruptive effects of low-frequency geomorphic events on radionuclide inventories be carefully assessed through critical visual or microscopic (thin-section) examinations of sediment lithology. Where these deposits are identified, the possibility must also be considered that they have eroded underlying sediments, thereby removing (and redepositing elsewhere) a portion of the nuclide inventory (Francus et al. 1998).

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In addition to physical processes, post-depositional resolubilization of 210Pb and 137Cs may influence their distribution in lake sediments. This is especially the case for 137Cs, which is more soluble and less reactive than 210Pb. Four pathways exist for the scavenging of excess 210Pb to sediments: (1) cation exchange and surface complexation with clay minerals; (2) binding with negatively-charged functional groups in organic matter; (3) adsorption to Fe and Mn oxides; and (4) formation of PbS. In the case of 137 Cs, only (1) and (2) apply. Thus, there are several conditions under which particulate 210 Pb and 137Cs can be remobilized. The presence of anions such as Cl-, HCO3- and SO4-2 can induce desorption of Pb from clays, whereas elevated NH4+ concentrations have been shown to displace 137Cs from clay mineral surfaces (Evans et al. 1983). Although 210 Pb bound to particulate organic matter generally remains immobile once deposited, complexation of 210Pb and 137Cs with dissolved organic matter has been documented in Alaskan lakes and ponds (Cooper et al. 1995). 210Pb bound to Fe and Mn oxides may also be released under changing sediment redox conditions (Benoit and Hemond 1990; von Guten and Moser 1993; Balistrieri et al. 1995; Canfield et al. 1995). Furthermore, the oxidation of 210PbS can revert Pb to more soluble minerals such as anglesite (PbSO4) and cerussite (PbCO3) (Vile et al. 1999). Clearly, any significant remobilization of sedimentary 210Pb and 137Cs through combinations of physical or chemical processes will hamper or preclude accurate age determinations. Unfortunately, at present there is insufficient quantification of the relative importance of these processes in high latitude lakes.

Atmospheric 210Pb distribution 90°N Arctic Circle

60°N

Latitude

30°N Equator 30°S Antarctic Circle 0

1

2

5 10 20 -3 (•10 dpm•m -3)

50

60°S 90°S

Figure 1. Latitudinal distribution and range (shaded area) of 210Pb activity in surface air (redrawn after Robbins 1978).

GEOCHRONOLOGY OF HIGH LATITUDE LAKE SEDIMENTS 210

23

Pb and 137Cs behaviour in high latitude environments

The largest single problem associated with 210Pb dating in polar regions is that of low (Douglas et al. 1994; Appleby et al. 1995) to undetectable (Doran et al. 1999; Hodgson et al. 2001) excess 210Pb activities in lake sediments. Two factors appear responsible for this. First, the release of 222Rn from soil decreases with latitude (Table 1), due to the influence of permafrost on retarding its diffusion to the atmosphere (Wilkening et al. 1975; Baskaran and Naidu 1995). This results in low ambient 210Pb fluxes at high increasing latitudes, especially in the southern hemisphere where the ice-free landmass is already limited (Figure 1). Secondly, the persistence of lake ice for all but a few weeks of the year limits the efficiency with which atmospheric 210Pb is transferred to lake sediments. For this reason, the successful development of 210Pb and 137Cs chronologies is compromised in perennially ice-covered lakes of the High Arctic and Antarctic, with the exceptions of sites that effectively integrate melt from largely glacierized catchments, or for which ice-marginal moats are sustained through the summer season (see Doran et al. 1999, 2000). However, such ecosystems occur in only the most extreme polar climates, implying that, while the caveats above still apply, the majority of high latitude lakes have sufficient exchange with the atmosphere for adequate 210Pb and 137Cs incorporation into sediments. While prolonged ice cover may indeed influence the transfer of radionuclides to lake sediments, it also reduces or even eliminates sediment redistribution by physical processes, while restricting benthic communities and hence bioturbation. Furthermore, most high latitude lakes are far removed from direct anthropogenic influences on sedimentation. In these ways, many Table 1. Estimates of 222Rn emanation and 210Pb production rates, with standard deviations shown in parentheses, where available. Location

222

Rn flux (Bq m-2 d-1)

Temperate land surfacesa Western Europe 1505 (429) Central Europe 1100 (427) Former U.S.S.R. (European) 587 (219) Former U.S.S.R. (Central Asia) 1029 (414) U.S.A. (central) 2200 (1158) U.S.A. (east) 2054 (1450) Hawaii 962 (1243) Arctic land surfaces 600 (112) Alaskaa Alaskab 599 Russiac 327 Canadad 671 a adapted from Appleby and Oldfield (1992) b adapted from Kirichenko (1970) c adapted from Anderson and Larson (1974) d adapted from Wilkening et al. (1975)

210

Pb production rate (Bq m-2 d-1) 259 (74) 189 (73) 101 (38) 177 (71) 378 (199) 353 (249) 165 (214) 103 (19) n/a n/a n/a

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WOLFE, MILLER, OLSEN, FORMAN, DORAN AND HOLMGREN

high latitude lacustrine environments remain highly promising for 210Pb and 137Cs geochronology, provided that sufficiently high-resolution sampling equipment (e.g., Glew 1988, 1989) is available to resolve radionuclide inventories in lakes typically characterized by extremely low sediment accumulation rates (i.e., < 10 to ca. 50 cm ka-1). For example, in non-glacial arctic lakes the entire unsupported 210Pb inventory is frequently contained in the uppermost 5 cm or less of the core (Douglas et al. 1994; Bindler et al. 2001), mandating great care in sample collection and core sectioning. Case studies Varves, or annually-deposited couplets of sediment laminae, provide the strongest independent chronometer against which 210Pb- and 137Cs-based ages for recent sediments can be rigorously assessed. Only recently have varved sediments been identified from lakes in the Canadian Arctic, and for the most part the confirmation of their annually-laminated character has relied on short-lived radioisotopes (Lamoureux 2001). In cores from non-glacial lakes on Ellesmere (82°50’N, 78°00’W) and Devon (75°34’N, 89°19’W) islands, respectively, Zolitschka (1996) and Gajewski et al. (1997) were the first to produce sufficiently close correspondence between varve and 210Pb chronologies to ascertain that couplets are in fact annual. Hughen et al. (2000) have documented even tighter correlations between 210Pb ages and varve counts in sediments from Upper Soper Lake on southern Baffin Island (62°55’N, 69°53’W), which have further been supported by correspondence between peak 239+240Pu activity and the 1962-1964 varve years in a second core (Figures 2A and B). 239+240Pu is another anthropogenic by-product of nuclear testing, with a stratigraphic behaviour that mirrors that of 137Cs (Ketterer et al. 2002). But more complicated results have also emerged. Varve, 210Pb, and 137Cs chronologies from Nicolay Lake on Cornwall Island (Lamoureux 1999a,b) exhibit three distinct stages of concordance: (1) excellent accord between varves and isotopes from the surface to ca. 10 cm, which includes the 1963 137Cs peak; (2) progressive divergence of varve and 210Pb chronologies to a depth of ca. 17 cm; and (3) a consistent 35-year offset below 15 cm (Figure 2C). Given that the 1963 and 1986 (Chernobyl) 137Cs peaks are separated by 23 couplets, and that the slopes of 210Pb and varve-based depth-age curves are identical below 17 cm, a compelling case can nonetheless be made that the sediments are indeed varves. The interval of pronounced divergence, in which 210Pb dates appear younger than the varves, is believed to reflect reduced clastic dilution of sediment 210Pb and hence higher 210Pb activity relative to adjacent levels, given that this interval is characterized by a reduction in couplet thickness (Lamoureux 1999b). In addition to strictly geochronological applications, 210Pb and 137Cs distributions may provide broader environmental inferences concerning sedimentary processes in high latitude lakes. For example, Hermanson (1990) measured 210Pb and 137Cs in multiple sediment cores from Imitavik and Annak lakes on Flaherty Island, the largest of the Belcher Islands in southeastern Hudson Bay (56°53’N, 79°25’W). Although each of seven cores analysed from Imitavik Lake showed good correspondence between 137Cs and unsupported 210Pb, the three shallowest of these have significantly greater inventories than the four cores from deeper water, contrary to the predicted radionuclide

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distribution if sediment focussing was preferentially moving flocculent material to the deepest portions of the lake. Instead, the anomalously large radionuclide inventories originate from regions of the lake that undergo periodic bottom-freezing. Hermanson (1990) suggests that significant transport of sediments frozen to the base of floating ice pans may effectively redistribute 210Pb, leading to preferential accumulations in shallow zones of the lake. Although this type of defocusing may be common in some shallow arctic lakes and ponds, it need not disrupt the overall stratigraphic integrity of 210Pb profiles. For example, shallow ponds on Ellesmere Island that freeze solid each winter nonetheless preserve intact 210Pb inventories, despite the influence of bottom-freezing (Douglas et al. 1994). Annak Lake, a small pond used to treat sewage from the nearby Inuit hamlet of Sanikiluaq, was also cored by Hermanson (1990) to investigate potential anthropogenic influences on sediment 210Pb distribution. Although the 137Cs profile from this core verifies its overall stratigraphic integrity, 210Pb activities decline sharply above 5 cm (Figure 3A). This closely matches the establishment of Sanikiluaq in 1968. Although the flux of 210Pb to the lake has remained constant since the community’s inception, as indicated by the results from adjacent Imitavik Lake, 210Pb has progressively become diluted by organic matter associated with sewage treatment. Annak Lake provides one of the first records of point-source pollution in an arctic lake, as well as a cautionary notice regarding anthropogenic effects on sediment 210Pb profiles. But dilution of sediment 210Pb activity is not solely associated with anthropogenic activity. For example, at Skardtjørna, a small coastal lake on western Svalbard (Spitsbergen; 78°00’N, 13°40’E), episodic reductions of sediment 210Pb activity preclude straight-forward modeling of isotopic decay (Figure 3B; Holmgren et al. unpublished data). However, the profile of 137Cs remains robust and interpretable, implying that the core retains stratigraphic integrity despite these 210Pb reversals. This implies that the 137Cs signal is less susceptible than 210Pb to short-term variations of sediment input at this site, likely because of efficient 137Cs sorption to mineral phases in these largely inorganic (< 3% total C) sediments. An abrupt decline in diatom concentrations closely matches the 210Pb anomaly between 9 and 12 cm, confirming the dilution of autochthonous sediment constituents by inputs of inorganic material that incompletely scavenges 210Pb. This is attributed to accelerated rates of allochthonous sediment transfer during these episodes. Indeed, sediment accumulation rates since ca. 1960 at Skardtjørna are several times higher than the Holocene mean, even when corrected for the higher water content of surficial sediments. Possibly, increased erosion and mass movement may be associated with recent pan-Arctic warming (Overpeck et al. 1997). The results from Skardtjørna, as well as additional sites on Svalbard (Appleby 2004) illustrate the opposite phenomenon to that depicted in Figure 2C, where a reduction in mass sedimentation has led to relative 210Pb enrichment and, consequently, an underestimation of true (varve) age. The recommendation is that, in all investigations, the sensitivity of 210Pb activity and derived chronologies to rapid changes in sedimenation must be verified. Although measurements of 210Pb and 137Cs are most commonly performed on single cores that are assumed to represent average depositional conditions in a given lake basin, the few studies where multiple cores have been analysed generally reveal that sedimentation is far more complex than what may be inferred from a single core. For example, the distribution of 137Cs in multiple cores from Toolik Lake, Alaska (68°38’N,

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Figure 2 . Comparisons of short-lived radionuclide activities and their derived chronologies to varve counts. (A) 210Pb activity versus depth for Upper Soper Lake core 93-7 showing the excellent agreement between model 210Pb age and varve age-depth relationships. (B) 239+240Pu activities versus depth for adjacent core 93-6, showing the degree of concordance between peak 139+140 Pu activity and the 1963 varve year. Both cores indicate average sediment accumlation rates of ca. 0.17 cm yr-1. (A) and (B) are adapted and redrawn after Hughen et al. (2000). (C) 137Cs activity (left panel) and comparison of model 210Pb and varve ages (right panel) for Nicolay Lake core NL-2. Peak 137Cs activity provides a closer match to the varve chronology than does the CRS 210 Pb model, suggesting sensitivity of the latter to sedimentation rates, while confirming the annual cyclicity of couplets. Further divergence occurs below 8 cm, where varve thicknesses decrease, resulting in a fairly consistent ca. 35-year offset between chronometers below 15 cm. Adapted and redrawn from Lamoureux (1999a,b).

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Figure 3 . (A) Loss-on-ignition (LOI, an estimate of organic matter content), 137Cs activity, and excess 210Pb from Annak Lake core A-2 in the Belcher Islands, Canada (redrawn after Hermanson 1990). The arrow indicates the year 1968 according to CRS 210Pb age modeling. This is also the year the hamlet of Sanikiluaq was established, resulting in effluent that has diluted recent sediment 210Pb activities above 5 cm in the core. The dashed line approximates decay of the unsupported 210Pb inventory and highlights the marked deviation from natural behaviour in nearsurface sediments. (B) Down-core 137Cs and 210Pb activities from Skardtjørna on Nordenskiöldkysten (Spitsbergen, Svalbard), shown alongside diatom valve concentrations. The shaded areas indicate dilution of the 210Pb signal due to increased siliciclastic sedimentation, as indicated most notably by the striking decline of diatom numbers between 9 and 12 cm. However, this dilution does not mask the prominence of the 137Cs peak. Dilution of 210Pb is also evident in the uppermost 4 cm of the core, but appears less extreme than the lower event. Note how removal of the shaded areas results in a more typical profile of down-core 210Pb activity.

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149°39’W), indicate preferential radionuclide accumulation near the lake’s inflow (Cooper et al. 1995). This pattern is associated with enhanced fluxes of dissolved organic matter near the inflow, and its efficient scavenging of 137Cs from adjacent tundra. However, considerable spatial variability in 137Cs activities coupled to relatively large inventories imply that this is not the sole process involved, and that some degree of post-depositional Cs redistribution occurs in Toolik Lake. One possibility is that 137 Cs distributions are modulated by high soil ammonia concentrations under reducing conditions, resulting in competition between ammonium and Cs ion binding sites (e.g., Evans et al. 1983). More detailed studies of the nature of Cs binding to organic matter in arctic ecosystems are clearly needed to resolve questions concerning post-depositional mobility. Although few studies have addressed sources and fluxes of minerogenic sediments entering high-energy proglacial lakes using 210Pb and 137Cs, the methodology appears to have considerable promise. By comparing sediment 137Cs inventories between an icecontact lake and another one 5 km downstream of the Mittivakkat Glacier, southeast Greenland (65°40’N, 38°11’W), Hasholt et al. (2000) have quantified the direct influences of ice-proximal processes such as calving and ice-rafting on recent sedimentation. Although sedimentation rates are one to two orders of magnitude greater in the ice-contact lake, well-defined subsurface 137Cs peaks are nonetheless identifiable in six of eight cores. Predictably, 137Cs inventories nearest to the calving margin are greatly reduced by comparison to the main basin, demonstrating the effective dilution of 137 Cs by subglacially-derived sediments characterized by limited exchange with recent atmospheric fallout. This study illustrates that applications of 210Pb and 137Cs geochronology need not be restricted to quiescent, low-energy lacustrine environments. Additional considerations High latitude lakes present several challenges to successful applications of 210Pb and Cs dating. Permafrost reduces 222Rn exhalation from arctic and antarctic land surfaces so that most atmospheric 210Pb originates from lower latitudes, resulting in reduced 210 Pb fallout over polar regions. Furthermore, stratospheric air masses may provide additional transport mechanisms for 210Pb to both the Arctic (Dibb et al. 1992) and the Antarctic (Pourchet et al. 1997), implying that a substantial (but as yet not quantified) fraction of the 210Pb flux to lake sediments depends on stratosphere-troposphere exchange processes. The additive effects of low atmospheric fluxes and severe lake ice regimes may, at their most extreme, preclude 210Pb and 137Cs from reaching lake sediments. However, in most cases where measurable 210Pb and 137Cs activities occur, their profiles may allow precise quantifications of sedimentation rates, and a variety of inferences concerning sedimentary processes. Ice cover may substantially reduce wave and current activity, resulting in more accurate 210Pb and 137Cs ages and accumulation histories. These conditions should prove conducive to future investigations of the chemical remobilization of these radionuclides. A final effect of ice cover is its role in the creation of complex and spatially heterogeneous sedimentary environments that integrate aeolian, colluvial, and/or ice-rafted materials, together complicating the evaluation of basin-wide sedimentation. Unlike sediment focussing, which is fairly well constrained (Rowan et al. 1995), deposition of sediments entrained on or within lake ice 137

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29

remains largely unknown. By collecting and analysing multiple cores from both shallow and deep locations within the lake, 210Pb and 137Cs methods may provide greater insight into the sedimentation processes and into the large scale dynamics of high latitude lacustrine environments. Radiocarbon dating Radiocarbon (14C, t1/2 = 5730 yr) is the primary geochronometer for lake sediments beyond the range of 210Pb and younger than ca. 40,000 yr. Radiocarbon is produced by cosmic ray bombardment of 14N in the upper atmosphere, where it is rapidly oxidized to 14 CO2. A poleward reduction of geomagnetic shielding, coupled to higher cosmic ray neutron densities near the geomagnetic poles, results in some degree of latitudinal variability in 14C production rates. However, 14CO2 is sufficiently well mixed and has a long enough atmospheric residence time to minimize this effect, particularly in the lower atmosphere where it enters the planetary carbon cycle. The fundamental assumption in radiocarbon dating is that metabolic uptake of 14C, primarily by photosynthesis, occurs in equilibrium with atmospheric 14C while the organism is alive. Once metabolism ceases upon death, decay of 14C in organic tissues is initiated. The residual 14C content in organic matter is the basis for its radiocarbon age, assuming that C isotopic composition has not been altered subsequent to death. The history and development of 14C dating can be read from many sources (e.g., Polach 1988; Taylor et al. 1992), and succinct reviews of the technique exist (e.g., Taylor 1992). Conventions for the expression of sample 14C activities as radiocarbon ages are provided by Stuiver and Polach (1977). The advent and refinement of direct 14C measurement by accelerator mass spectrometry (AMS) has revolutionized radiocarbon dating over the last two decades (Linick et al. 1989; Stafford et al. 1991). In paleolimnology, AMS has essentially supplanted conventional dates obtained by E-decay counting. This is because AMS enables much smaller samples (by as much as 6 orders of magnitude) to be analysed with precision. This is especially relevant for high latitude lakes characterized by low sediment accumulation rates, because the advantage of dating small samples is directly conveyed to the possibility of attaining high resolution chronostratigraphy. Furthermore, AMS has created new horizons with the ability to target discrete fractions of a given sediment, such as chemically extracted and purified organic macromolecules (Eglinton et al. 1997; Hwang and Druffel 2003) and particulate nanno-fractions including pollen (Mensing and Southon 1999) and insect chitin (Jones et al. 1993; Hodgins et al. 2001; Fallu et al. 2004). AMS has also refined studies of modern carbon cycling at the ecosystem scale, largely because anthropogenic 14C from thermonuclear testing in the 1960s provides a useful signature for tracking the movement of newly accrued C in soil and biomass (Richter et al. 1999; Neff et al. 2002). Unfortunately, detailed studies of this type are still lacking for most high latitude terrestrial and aquatic ecosystems, the exceptions being Abbott and Stafford’s (1996) research on Baffin Island, and Doran et al.’s (1999) work in the McMurdo Dry Valleys of Antarctica. As a result of the expanded analytical capability afforded by AMS, it is an increasingly common strategy to date several individual fractions from the same depth in a sediment core (e.g., MacDonald et al. 1987; Lowe et al. 1988; Björck et al. 1998;

30

WOLFE, MILLER, OLSEN, FORMAN, DORAN AND HOLMGREN

Turney et al. 2000; Nilsson et al. 2001; Fallu et al. 2004). This approach is of great assistance in the critical selection of fractions most likely to produce accurate ages. In addition to being conducive to more robust sediment chronologies, in many instances the dating of multiple fractions also provides useful information on changes in lake basin carbon dynamics over time. For high latitude lake sediments, the following general guidelines have emerged. In regions underlain by crystalline bedrock lithologies, where hardwater effects are negligible, autochthonous aquatic macrofossils are an appropriate target for AMS 14C dating, because lake waters are generally equilibrated with the atmosphere with respect to their 14C activities. For hardwater lakes, terrestrial plant macrofossils or pollen are the most desirable targets for AMS 14C dating, although the possibility of macrofossil redeposition from older sediments certainly exists (Lozhkin et al. 2001). Of all datable organic fractions, terrestrial plant material is clearly the most directly linked to atmospheric CO2 in the metabolic sense. However, slow rates of organic matter decomposition result in a reservoir of plant macrofossils that may yield apparent 14C ages hundreds to thousands of years too old (Oswald et al. 2004). In almost all cases, the least reliable fraction for 14C dating is undifferentiated organic matter from bulk sediment. This is because most bulk sediments comprise highly heterogeneous admixtures, in age and composition, of organic carbon. Clear warnings have been issued concerning problems associated with dating bulk sediments (Nelson et al. 1988; Colman et al. 1996), including specific examples from oligotrophic arctic lakes (Snyder et al. 1994; Child and Werner 1999; Gajewski et al. 2000). The practice of dating bulk sediments, whether by AMS or conventional counting, cannot be recommended under most circumstances. The targeting of optimal carbon-bearing targets for AMS 14C of arctic lake sediment has transformed the temporal control from an inherently uncertain effort with typical uncertainties in core chronologies of more than 1000 years, to a more reliable and more accurate science. Arctic and antarctic lake sediments have provided several challenging complications for dating. First, terrestrial plant production is low, so that terrestrial macrofossils are usually scarce. Moreover, the presence of permafrost and low temperatures, even in the active layer, retards the decomposition of what little terrestrial organic carbon is present, so that fluxes of old organic carbon are present in most catchments. Because low primary production is also the rule for high latitude lakes (e.g., Flanagan et al. 2003), the autochthonous carbon flux to sediments is also low, which increases the potential magnitude of contamination effects by ancient carbon derived from a variety of allochthonous sources. These sources include (1) dissolved organic carbon (DOC) sequestered from lake catchments; (2) dissolved inorganic carbon (DIC) from the weathering of carbonate-rich bedrock and surficial deposits; (3) particulate organic carbon (POC) including undecomposed plant tissues, as well as older materials such as pre-Quaternary palynomorphs and coal. The latter may be important in regions where poorly consolidated outcrops of Mesozoic-Cenozoic strata occur. In some cases, the input of DIC that is influenced by ancient CO2 in glacial ice must also be considered. Most of these problems are discussed in greater detail elsewhere (Björck et al. 1991; Snyder et al. 1994; Gajewski et al. 1995; Hall and Henderson 2001). We emphasize that, despite abundant sources of carbon depleted in 14C, radiocarbon dating of high latitude lake sediments can be optimized through careful selection of a subset of the total carbon pool preserved in the sedimentary record. Only polar lakes that are completely isolated from the atmosphere by perennial ice cover and thus do not

GEOCHRONOLOGY OF HIGH LATITUDE LAKE SEDIMENTS

31

exchange 14CO2 (Blake 1989), and extreme hardwater lakes in carbonate terrain lacking terrestrial macrofossils, produce essentially intractable settings for radiocarbon dating. Rather than providing an exhaustive review of the evolving body of literature concerning applications of 14C dating to high latitude lakes, much of which has been treated in detail by Björck and Wohlfarth (2001), we provide instead a compilation of published and unpublished results from our ongoing research that serves to illustrate several facets of the application of AMS 14C to dating arctic and antarctic lake sediments. Unless indicated, all previously unpublished radiocarbon measurements are reported as uncalibrated 14C years using the Libby half-life of 5568 years, and corrected for fractionation effects. Radiocarbon geochemistry of arctic and antarctic lakes The geochemistry of radiocarbon in modern arctic and antarctic lake systems is illustrated schematically in Figure 4. Whereas lakes on Baffin Island are seasonally open to the atmosphere, surrounded by herbaceous heath tundra, and underlain by crystalline bedrock, the McMurdo Dry Valley lakes have permanent ice cover, are glacier-fed or glacier dammed, and are surrounded by polar desert. The comparative radiocarbon geochemistry of these modern systems pertains directly to the dating of their sediment records. On Baffin Island, soils have near surface pools of 14C-depleted organic material at depths well above the permafrost table (i.e., 4 cm; Figure 4A). Individual soil organic fractions increase in age with depth without exception, indicating that permafrost restricts their vertical migration in the soil profile, in sharp contrast to the mobility of organic C observed in temperate peatlands (Charman et al. 1999). Because the decomposition of terrestrial organic matter is comparatively much slower on Baffin Island, there is evidence of 14C depletion in both POC and DOC pools in modern streams and lakes by as much as 20% relative to the atmosphere. Despite this, photosynthetic organisms in Baffin Island lakes (moss and algae) remain equilibrated with the atmosphere (Table 2). Only one of eight modern plant samples from six different lakes (all on crystalline bedrock) resulted in a 14C activity significantly different from the atmosphere at the time they were collected. These data suggest that, despite the short ice-free season, most arctic lakes exchange 14C with the atmosphere efficiently. The single depleted sample of modern photosynthate suggests that algae may occasionally incorporate some old C. This is most likely the result of DOC metabolism, resulting in a depletion of up to 10% relative to the atmosphere. However, relative to living plants and algae, 14C depletion is consistently more pervasive in detrital and dissolved organic C pools. This necessarily reflects photosynthetic assimilation of DIC, which is modulated by wind mixing of surface waters. Metabolism of DOC must be either secondary or insignificant as a carbon source to the modern plant samples in our data set. Summarily, despite the presence of old carbon in soils and its potential influence on several organic C pools in streams and lakes, there is acceptable overall equilibration with respect to atmospheric 14C, as the grand mean 14C activity of modern stream, lake and snowpack POC and DOC is Fm (fraction modern) = 1.06 ± 0.078 (n = 11). When algae and moss are added to this total, the mean increases to Fm = 1.08 ± 0.084 (n = 18).

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Figure 4 . Synthesis of radiocarbon geochemistry for lakes, lake inputs, and catchment soils on southern and east-central Baffin Island (A) and the McMurdo Dry Valley lakes of Antarctica (B). Most results are reported as Fm (fraction modern) with radiocarbon ages in parentheses when different from modern. Soil data in (A) are also in radiocarbon years before present. For pools with several individual measurements, the range of results is shown, which is more informative than the mean in terms of assessing degrees of potential contamination. Data for Baffin Island are from Abbott and Stafford (1996), Miller et al. (1999), and previously unpublished results, integrating measurements from seven lake basins on Meta Incognita, Hall, and Cumberland peninsulas (Avataq, Brevoort Water, Dyer Lower Water, Meech, Robinson, Totnes, and Tuktu lakes). Because all of these lakes are located on Precambrian crystalline bedrock, only DIC values from soft-water lakes (Robinson and Dyer Lower Water) are used (Table 3). Data from Antarctica are mostly from Doran et al. (1999, 2003), and include results from lakes Bonney, Fryxell, Hoare, and Vida. DOC values are from Aiken et al. (1996) and are limited to Lake Fryxell. Bottom water DIC 14C activity from this lake is not shown due to apparent accidental contamination (Doran et al. 1999).

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33

Table 2 . Radiocarbon activities of modern plants and algae from several Baffin Island lakes (Abbott and Stafford 1996; Miller et al. 1999). All samples have Fm > 1, and therefore contain bomb radiocarbon. The most common aquatic moss species is Warnstorfia exannulata. Lake

Year collected

Material

14

C activity (Fm)

Lab. accession no.

Totnes

1993

moss

1.1514 ± 0.0112

CAMS-7780

Robinson

1993

moss

1.1788 ± 0.0077

CAMS-11334

Brevoort Water

1990

moss

1.1867 ± 0.0094

AA-6532

Avataq

1989

moss

1.2157 ± 0.0074

AA-5476

Tuktu

1989

moss

1.1651 ± 0.0070

AA-5252

Tuktu

1989

rooted aquatic vascular plant 1.1841 ± 0.0078

AA-5253

Tuktu

1989

algal filaments

1.1536 ± 0.0096

AA-7006

Meech

1989

algal filaments

1.0363 ± 0.0095

AA-7003

Mean ± 1ı

1.1590 ± 0.0537

The 14C activities of lake water DIC and mosses living in this water are available for three lakes from Arctic Canada, two in crystalline terrain, and one in carbonate bedrock (Table 3). In both lakes underlain by crystalline lithologies, lake water DIC and living moss 14C activities have reasonable concordance, which suggests the mosses represent appropriate targets for 14C dating in these systems. In our sole carbonate system (Hall Beach #1), which was sampled in early June near the end of the ice cover season, DIC is clearly influenced by contributions from 14C-depleted carbonate bedrock, resulting in an apparent age for late-winter water of almost 9000 years. However, moss collected at the same time from the same lake has no apparent hardwater effect and is nearly equilibrated with the atmosphere. The only tenable explanation for this result is that DIC from carbonate weathering accumulates beneath the ice during winter, when lake overflow ceases, but is rapidly flushed from the lake during snowmelt, when the lake’s water is recirculated and replaced. By the time seasonal moss growth begins, the DIC pool that was exploited metabolically had become largely re-equilibrated with the atmosphere. Of course, such findings must not be considered universal. For example, in hardwater seepage lakes, aquatic mosses are likely to remain susceptible to hardwater effects (MacDonald et al. 1991). Thus, our results suggest that the presence of carbonate bedrock does not necessarily preclude aquatic plants as potential 14C targets. On the contrary, our results from Baffin Island suggest that aquatic mosses are optimal dating targets irrespective of catchment lithology, provided that the lakes are well-mixed during the growing season. This conclusion can be generalized, given that aquatic plants have been shown to be suitable for 14C dating in lake sediments from western (Willemse 2002) and southern (Björck et al. 2002) Greenland, as well as soft-water lakes on the Antarctic Peninsula (Björck et al. 1991). In the McMurdo Dry Valleys, the situation is considerably more complicated (Figure 4B). First, there are two additional sources of old DIC: glacial meltwater and pedogenic carbonate. Both of these pools are several millennia in 14C age. Although stream DIC is

34

WOLFE, MILLER, OLSEN, FORMAN, DORAN AND HOLMGREN

clearly influenced by these sources, especially since all lakes are glacially-fed, there is strong evidence for the progressive re-equilibration of stream DIC with distance from the glacier (Doran et al. 1999 and unpublished data). As a consequence, and perhaps surprisingly, several pools have produced modern 14C activities: stream, lake margin, and lake ice microbial mats, as well as DIC and DOC samples collected immediately beneath lake ice. Given the sheer abundance of 14C-depleted materials in these catchments, these results are quite remarkable, as they suggest that: (1) soil carbonate dissolution in stream beds does not significantly affect DIC 14C activity; and (2) steep localized gradients in aquatic 14C activity are present. Table 3 . Radiocarbon activities of modern moss and late-winter water column DIC (collected under ice in May-June, 1993) from three Baffin Island lakes. Only the Hall Beach site is underlain by carbonate bedrock (Paleozoic limestone). Lake

Material

14

C activity (Fm)

C age

Lab. accession no.

14

Robinson

moss

1.1788 ± 0.0077

modern

CAMS-11334

Robinson

DIC

1.1125 ± 0.0074

modern

CAMS-12258

Dyer Lower Water

moss

1.0607 ± 0.0072

modern

CAMS-12293

Dyer Lower Water

DIC

1.1079 ± 0.0073

modern

CAMS-12259

Hall Beach #1

moss

1.1298 ± 0.0076

modern

CAMS-12295

Hall Beach #1

DIC

0.3278 ± 0.0115

8960 ± 290

CAMS-12260

The second major complication is that these lakes have residence times in the order of millennia. Thus, DIC, DOC and POC pools are all severely depleted in 14C, especially in water depths > 5 m. However, a microbial mat at 12.5 m in the ice cover over Lake Vida, Victoria Valley (77°23’S, 161°45’E) has a 14C age of 2770 yr BP while preserving cyanobacterial cells that are viable upon thawing (Doran et al. 2003). In Lake Fryxell (77°37’S, 163°09’E), the upwards diffusion of relict DOC from ancient bottom waters (ca. 3000 yr) has been invoked to reconcile the gradient in water column 14C activity (Aiken et al. 1996). However, there is evidence for periodic incursions of younger waters into the hypolimnion, perhaps associated with brine exsolution events when littoral moats refreeze at the end of summer. Despite these factors, the simplest explanation for hypolimnetic DIC and DOC dates is that they approximate the actual age of the water. Thus, water ages can be used to infer minimum ages for the last major disruption of water column stability associated with lake-level lowering. For example, in lakes Hoare (77°38’S, 162°55’E) and East and West Bonney (77°43’S, 162°20’E), the last major drawdowns appear no younger than 1100, 8000, and 23,000 yr BP, respectively (Doran et al. 1999 and unpublished data). Can cores from the Dry Valley lakes of Antarctica be dated with radiocarbon? Eight dates have been obtained on discrete microbial mat horizons in a core from Lake Hoare (Figure 5). Because the core was retrieved by SCUBA with an intact sediment-water interface, the surface mat is presumed modern, despite producing a 14C age of 2600 yr BP. The remaining dates progressively increase downcore to 4650 yr BP, with a single

GEOCHRONOLOGY OF HIGH LATITUDE LAKE SEDIMENTS

35

reversal near the base (Doran et al. 1999). Collectively, these dates result in a linear sediment accumulation rate of 15 cm ka-1 (r2 = 0.97; n = 8), providing an estimated basal age for the core of ca. 2400 yr BP. This accumulation rate is comparable to earlier estimates based on fewer dates (Squyres et al. 1991). Clearly, the radiocarbon-inferred sedimentation history of Lake Hoare must be viewed as a first approximation. Indeed, it is somewhat disquieting that the 14C-based sediment accumulation rate produces a basal age that is somewhat younger (ca. 2400 yr) than the 14C age of the sediment-water interface (2600 yr BP). On the other hand, the possibility that the lake’s sediments have maintained a more or less constant old carbon effect of ca. 2600 yr during the late Holocene cannot be dismissed, given that hypolimnetic DIC has an apparent age of 2670 yr BP (Doran et al. 1999).

Figure 5. Radiocarbon age determinations from microbial mats in a SCUBA-collected core from Lake Hoare, McMurdo Dry Valley, Antarctica (Doran et al. 1999). Although the dates can be fitted satisfactorily by linear interpolation, this assumes that the large 14C age offset of modern microbial mat (ca. 2600 yr) has remained constant over the late Holocene.

Humic acid 14C dating: panacea or further complication? As outlined above, macrofossils are rare in many high latitude lake sediments, yet the dating of bulk sediment is clearly undesirable due to the compounded effects of long residence times for terrestrial carbon and low accumulation rates for autochthonous organic matter. Based on a series of paired dates (Table 4), Abbott and Stafford (1996) have shown that lake sediment humic acid extracts (HA: the base-soluble, acidprecipitated fraction of organic matter) produce AMS 14C dates that are consistently

36

WOLFE, MILLER, OLSEN, FORMAN, DORAN AND HOLMGREN

younger than other fractions, including bulk sediment, fulvic acids (the unprecipitated base-soluble fraction), and humins (acid-soluble). They consider that, next to terrestrial macrofossils, HA provide the most reliable dating target for the following reason. Although a proportion of HA is undeniably terrestrial in origin, there are strong geochemical indications that much of the HA in lake sediments have aquatic precursors (Nissenbaum and Kaplin 1972; Ishiwatari 1985), especially at high latitudes where terrestrial plant biomass is minimal (McKnight et al. 2001). One consequence of this ancestry has direct implications for radiocarbon dating: the aquatic HA pool can be rapidly delivered to sediments given a simplified depositional trajectory that involves degradation of organic matter from and in the water column, followed by the in situ polymerization of HA. Humic acids are also sufficiently abundant for AMS 14C dating in almost all lake sediments, which may also reflect their predominantly aquatic origin. Table 4 . Radiocarbon ages of discrete sediment fractions in a core from Baby Tuktu Lake, southern Baffin Island (Abbott and Stafford 1996). Depth (cm) 1.0-2.0

Humic acid 14 C age

Lab. accession no.

Bulk sediment (< 63µm) 14 C age

Lab. accession no.

Humin C age

14

Lab. accession no.

1045 ± 50 AA-6026 1065 ± 50 AA-6045 1085 ± 60 AA-6038

Fulvic acid 14 C age

Lab. accession no.

n.a.

n.a.

40.0-41.0

3015 ± 50 AA-6027 3420 ± 50 AA-6046 3230 ± 70 AA-6039 3385 ± 50 AA-6031

71.0-72.0

5675 ± 90 AA-6028 6040 ± 70 AA-6047 6260 ± 80 AA-6040

n.a.

n.a.

76.0-77.0

6160 ± 90 AA-6029 6445 ± 90 AA-6048 6950 ± 70 AA-6041

n.a.

n.a.

We have evaluated the precision of HA dates with a series of dates on stratigraphically adjacent HA extracts from Robinson Lake, Baffin Island (63°24’N, 64°16’W). In all four pairs, the lower of the two is always older (Figure 6A). Indeed, in Baffin Island lake sediments we have observed fewer reversals in HA 14C dates than in macrofossil dates, irrespective of sediment accumulation rate. We have also noted that in older sediments approaching the limit of 14C dating, both HA and macrofossils become nonfinite (i.e., 14C depleted beyond instrument detection) at the same stratigraphic levels, lending further support to the contention that HA is immobile within the sediment column (Table 5). The higher molecular weight of HA relative to other humic fractions is likely to both promote its deposition and restrict its postdepositional mobility. In resolving the question of accuracy, however, HA chronologies must be anchored to dates based on materials that are more directly associated with atmospheric 14CO2, namely plant macrofossils. We have accumulated a set of paired dates from HA and macrofossils (Table 5, Figure 6B). Macrofossils are on average younger than HA, which reflects the lag time associated with the sequestration of the terrestrial fraction of HA to lake sediments. The magnitude of this lag depends on several factors, including the status of catchment vegetation, soils, and hydrology, as well as climate and basin topography. Furthermore, as the ontogeny of a given lake continues, so too will the relationship between autochthonous and allochthonous organic matter contributions to

GEOCHRONOLOGY OF HIGH LATITUDE LAKE SEDIMENTS

37

lake sediments. For example, it is only in the early postglacial (> 9000 14C yr BP) that we have encountered sediments from which macrofossils produce older 14C dates than HA (Figure 6B), likely resulting from unresolved combinations of: (1) plants metabolizing DIC that is influenced by glacial melt; (2) redepositon of macrofossils from older deposits; and (3) the absence of an accumulated soil carbon pool during initial stages of ecosystem development. There is strong independent support for the importance of the latter phenomenon: at Kråkenes in Norway (62°02’N, 5°00’E); for example, there is no detectable offset between plant macrofossil and HA 14C dates during the Younger Dryas chron (Gulliksen et al. 1998). In contrast, the age offset

Figure 6. Radiocarbon age determinations on humic acid extracts from adjacent 1 and 2 cm slices in Robinson Lake core 91-RL4 (A). In every case the age of HA increases with depth, indicating little or no migration within the sediment column and testifying to the technique’s high level of precision. Original data are from Miller et al. (1999). In (B), the temporal evolution in the offset between paired humic acid and macrofossil 14C dates is represented, using data from Table 5. The correction factor of 300 ± 300 years is shaded.

38

WOLFE, MILLER, OLSEN, FORMAN, DORAN AND HOLMGREN

Table 5 . Paired radiocarbon ages of macrofossils and humic acid extracts from five lakes on Baffin Island and one (Qipisarqo) in southwestern Greenland, arranged by increasing macrofossil age. The italicized mean and SD are for the subset that excludes pairs where humic acids are younger than macrofossils. Lake

Depth in core (cm)

Macrofossil radiocarbon age

Lab. accession no.

Humic acid radiocarbon age

Lab. accession no.

Humic acid macrofossil offset

Robinson

0.0-2.0

modern1

CAMS-11334

505 ± 70

CAMS-11127

505

Birch

8.0-9.0

1000 ± 50

OS-25908

1700 ± 45

OS-25910

700

Birch

28.0-29.0

2270 ± 60

OS-25911

2840 ± 55

OS-25912

570

Qipisarqo

130.0-132.0

5610 ± 60

OS-25736

5840 ± 60

OS-25735

230

Fog

45.5-46.5

6150 ± 70

CAMS-30535

6320 ± 50

CAMS-30534

170

Robinson

64.0-66.0

7670 ± 85

AA-7589

8245 ± 85

AA-9002

575

Birch

79.5-80.5

7670 ± 60

OS-17680

7330 ± 95

OS-20406

340

Amakuttalik

208.0-209.0

9000 ± 60

CAMS-27574

9620 ± 60

CAMS-27264

620

Robinson

94.0-96.0

9335 ± 80

AA-9012

9145 ± 65

AA-9005

-190

Robinson

96.0-98.0

9575 ± 95

AA-9011

9235 ± 80

AA-9006

-340

Robinson

110.0-112.0

10,160 ± 145

AA-9013

9955 ± 130

AA-9007

-205

Donard

352.0-353.0

12,600 ± 60

CAMS-23555

11,970 ± 60

CAMS-23554

-630

Fog

110.0-111.0

> 52,200

CAMS-28652

> 44,400

CAMS-31808

n/a

1

mean:

195 (464)

SD:

438 (193)

Fm = 1.1788

between HA and macrofossil dates increases with age in late-Wisconsinan sediments from Arolik Lake in Alaska (59°28’N, 161°07’W), implying the presence of a significant pool of ancient soil HA (Kaufman et al. 2003). As a first-order solution, we correct for the average influence of aged terrestrial HA in Baffin Island lakes by subtracting 300 years from HA 14C dates (Miller et al. 1999; Kaplan et al. 2002; Wolfe 2002, 2003). Two examples of this practice are illustrated on Figure 7, demonstrating its apparent suitability in this region. For example, the parabolic curves fitted to the corrected and calibrated dates on Figure 7 have not been forced through the origin, but intersect it naturally well within error. This is reassuring, since both cores were retrieved with undisturbed mud-water interfaces (Glew 1989). Despite the apparent successes in dating HA in Holocene sediments (Figure 7), some HA dates are highly misleading. In Fog Lake on Cumberland Peninsula, Baffin Island (67°11’N, 63°15’W), which escaped glacial erosion during the Late-Wisconsinan and consequently preserved organic-rich sediments of last interglacial age, we have obtained a series of HA ages that suggest, at face value, organic production in lakes during the 15-33 ka BP interval. Although the possibility has been advanced that these dates indicate the existence of genuine full-glacial refugia (Steig et al. 1998; Wolfe et al. 2000), it is now clear that these HA results integrate both contemporaneous and ancient carbon. In one example of this (Figure 8), dates on an individual moss fragment and on chironomid head capsules produce ages of 8530 ± 70 and 7590 ± 240 yr BP, respectively, in sediments for which HA are in the 14-16 ka BP range. Simple mixing

GEOCHRONOLOGY OF HIGH LATITUDE LAKE SEDIMENTS

39

Figure 7 . Age models for the upper 35 cm of gravity cores from (A) Kekerturnak (67°53’N, 64°50’W) and (B) Fog (67°11’N, 63°15’W) lakes on Baffin Island separated by ca. 100 km. Humic acid extracts have been dated by AMS, a 300 year correction has been subtracted from these results, and the corrected dates then calibrated (Stuiver et al. 1998) to provide a timescale in calendar years (from Wolfe 2002).

model calculations imply that, at these times, 60% of HA in Fog Lake sediments is reworked from interglacial carbon assumed to be exhausted in 14C. But this contamination is evidently short-lived: in sediments less than 5 cm above the highest erroneous date, the behaviour of HA appears to return to one dominated by contemporaneous sources, resulting in a suite of early Holocene ages that do not contradict the moss and chironomid ages (Figure 8). Our revised paleoenvironmental reconstruction is that the Fog Lake basin was geomorphically inactive and biologically sterile for tens of millennia prior to the early Holocene, under cold full-glacial climates (Dahl-Jensen et al. 1998). The return to conditions of summer thaw in the early Holocene remobilized organic matter of last interglacial age, leading to anomalously old HA ages. This example illustrates how, even in the absence of carbonates, ancient carbon may still be present. In resolving the problem of old apparent HA ages in early Holocene sediments, the ability to date chitin from chironomid head capsules has been instrumental. Despite representing extremely small samples (down to ca. 100 µg C), this material appears suitable for AMS 14C dating of lake sediments from the Arctic, and further tests of the isotopic integrity of lacustrine chitin should be undertaken. However, certain chironomids are detritivores, so that the incorporation of old carbon in chitin remains possible. For example, Snyder et al. (1994) demonstrated that chironimids are 1200 years older than terrestrial plant macrofossils in hardwater Linnévatnet, Svalbard (78°03’N, 13°50’E). Our initial data from Baffin Island, as well as results from northern

40

WOLFE, MILLER, OLSEN, FORMAN, DORAN AND HOLMGREN

Québec (Fallu et al. 2004) suggest a somewhat smaller effect, implying that chironomids in these instances consume well-equilibrated materials such as large benthic algae, as a higher-quality dietary source relative to detritus (e.g., Edlund and Francis 1999). Although the HA results presented on Figures 6B and 8 illustrate some of the complexities inherent to the development of lake sediment 14C chronologies in the Arctic, they also demonstrate how the dating of distinct organic fractions may contribute to the interpretation of ecological, biogeochemical, and sedimentological changes. A holistic approach to radiocarbon geochemistry in high latitude lakes is clearly desirable, as exemplified by recent successes in mid-latitude terrestrial (Neff et al. 2002) and riverine (Raymond and Bauer 2001) ecosystems.

Figure 8 . Compilation of (uncalibrated) radiocarbon and luminescence ages from Fog Lake percussion cores 96-05 and 04. The site presents significant dating challenges because reworked carbon of interglacial age has contaminated the HA pool in sediments of early Holocene age, resulting in apparent ages of 14-16 ka BP (box: 65-75 cm). Key dates from moss and insect chitin extracted from the same core depths (bold numbers) provide more accurate ages for these sediments. Dates in bolded italics are interpreted as minimum ages only, because luminescence dating of the equivalent lithostratigraphic unit in adjacent core 96-4 (correlated by magnetic susceptibility) are ca. 90 ka BP. Luminescence results are shown on the far right, derived from either infrared (IRSL: bold italics) or thermal (TL: plain type) stimulation. Modified from Wolfe et al. (2000).

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Pushing the limits Because 14C-free alanine blanks produce finite ages in the 48-56 ka range (Wolfe et al. 2001), this must be considered beyond the range of interpretable results from most AMS laboratories. In reality, however, the upper limit may be somewhat younger. Humic acid dates from sediments with luminescence ages > 85 ka have produced finite ages in the 36-39 ka range, even though these sediments should have no measurable 14C (Figure 8). Similar finite ages from lakes in other regions of the Canadian Arctic (e.g., Wolfe and Härtling 1996; Wolfe and King 1999) should therefore be reconsidered as minimum ages. It has been demonstrated that the contamination of macrofossil tissues by de novo microbial synthesis can occur during prolonged core storage, leading to anomalously young 14C dates (Wohlfarth et al. 1998). This is perhaps the case for some, but not all, macrofossil samples from Robinson Lake, where non-finite moss remains occur adjacent to specimens producing finite ages in the 29-38 ka range, all of which occur in a lithostratigraphic unit independently dated by luminescence to the last interglacial (Miller et al. 1999). We recommend that all pre-Holocene lake sediment 14C dates should be confirmed using a second method, of which optically stimulated luminescence appears the most promising. The same recommendation has been reached for Holocene sediments in antarctic Dry Valley lakes (Doran et al. 1999, 2000). Optically stimulated luminescence Luminescence geochronology is the most promising technique for dating aquatic sediments beyond the limit of 14C and potentially spanning the past ca. 200 ka. Luminescence dating is focus of many reviews (e.g., Aitken 1985; Berger 1988; Wintle 1990; Stokes 1999; Lian and Huntley 2001). Optically stimulated luminescence (OSL) appears more sensitive than thermoluminescence (TL) for dating waterlain sediments, in part because effective zeroing of the OSL signal prior to deposition can be accomplished rapidly during transport over relatively short distances (Godfrey-Smith et al. 1988). A detailed examination of the theory and applications of OSL is provided by Aitken (1998), whereas applications specific to subaqueous sediments include Forman (1999) and Wallinga et al. (2001). Here, we emphasize aspects of OSL dating most relevant to high latitude lake sediments, based on a limited array of presently available results (Krause et al. 1997; Doran et al. 1999; Miller et al. 1999; Wolfe et al. 2000). The full potential of OSL dating is realized only when knowledge of sedimentary processes, efficiency of solar resetting, and environmental radioactivity are combined. Luminescence and geochronology Silicate minerals (e.g., quartz and feldspar) contain lattice charge defects formed during crystallization or from exposure to nuclear radiation. These defects are potential sites of electron capture over geological time and therefore a source of time-dependent luminescence. Luminescence reflects exposure to radiation from the radioactive decay of 238U, 235U, 232Th (and daughter isotopes), and 40K. The energy associated with the release of D, E and J particles from isotopic decay displaces lattice-bound electrons,

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which are then captured in charge defects called electron traps. A population of these electrons is stored in deep traps that are theoretically stable for > 106 years (Aitken 1985). Ideally, luminescence emitted by silicate minerals increases with time, reflecting progressively longer exposures to environmental radiation. Heating or light exposure of these minerals causes eviction of stored electrons, which are redistributed to luminescence centers where they recombine and emit light. This luminescence emission (at a particular wavelength) is proportional to the total environmental dose and is a function of time. The time-dependent luminescence of silicate minerals can be liberated either by heating (thermoluminescence, TL) or exposure to wavelengths from 400 to 900 nm (optically stimulated luminescence, OSL). In dating sediments, a particular wavelength (e.g., 880 ± 80 nm) is used to excite mineral grains and the resultant emission is measured for a particular band (e.g., 400-500 nm). Luminescence laboratories often select blue emissions for measurement, reflecting the dominant response of quartz and feldspar (Krebetschek et al. 1997), and avoid ultraviolet emissions, which may yield age underestimates (Balescu and Lamothe 1992). Thus, silicate minerals are long-term radiation dosimeters for which naturally acquired luminescence is a measure of radiation exposure during the period of burial. The radiation measured in the laboratory approximates natural luminescence and is termed equivalent dose (DE, measured in grays (Gy), where 1 Gy = 1 J kg-1). DE is the numerator in the luminescence age equation: Luminescence age =

DE wD a  wD b  wD g  wD c

(1)

where Da = D dose including compensation for reduced efficiency of D radiation (Gy ka-1), Db = E dose (Gy ka-1), Dg = J dose (Gy ka-1), Dc = cosmic dose (Gy ka-1), and w = moisture content correction factor. The denominator in equation (1) is termed dose rate, and approximates the sample’s environmental radioactivity. Dose rate is typically assessed by measuring U, Th, and K concentrations and calculating a small contribution from cosmic sources. Water content is an important parameter because water strongly attenuates environmental radiation, reducing dose rate as a near-linear function (for fine grains). For many lacustrine sediments, moisture content is measured, reflecting variability in sediment particle size and compaction. There are two principal methodologies for translating stored luminescence to equivalent dose. The multiple aliquot additive dose method (MAAD) was initially developed for TL and is used primarily on fine-grained polymineral or quartz fractions (Singhvi et al. 1982). This method applies additional doses (E or J) to the natural luminescence of separate sample aliquots in order to build a dose-response curve, from which equivalent dose can be extrapolated. When applied to loess, MAAD yields ages concordant with other geochronometers for at least 50 ka (e.g., Berger et al. 1992; Forman et al. 1995; Forman and Pierson 2002). The second method is single aliquot regeneration (SAR), which determines an age for each aliquot by matching the regenerated signal to original natural luminescence (Murray and Wintle 2000). This technique can be applied to both fine and coarse fractions of either mono- or polyminerals. The SAR method is particularly useful for sediments that are < 15 ka old, as

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there is evidence that it may yield underestimates for sediments > 50 ka old (e.g., Mangerud et al. 2001; Wallinga et al. 2001). Although solar resetting of luminescence is in principle rapid in aquatic environments, zeroing may be incomplete due to spectral attenuation associated with suspended sediment and water depth (Berger 1990). Turbid glacial meltwaters may also result in ineffective zeroing, especially if minerals are only transported short distances prior to deposition (Doran et al. 1999). Fine-grained (4-11 µm) lake sediments in non-glacial catchments are more likely to be fully reset, due to prolonged sunlight exposure during particle transport and deposition. In addition, most sediment transport occurs during summer when sunlight is abundant, and deposition occurs by hypopycnal flows (Retelle and Child 1996) that are also exposed to sunlight. On the other hand, fine-grained sediments may occasionally be deposited rapidly by turbidity currents or hyperpycnal flows with suspended sediment loads > 20 mg l-1, thus restricting sunlight exposure, elevating inherited luminescence, and yielding age overestimates. Variability in natural luminescence underscores the necessity of understanding lake sediment transport and deposition. On the other hand, dating coarse-grained fractions (ca. 100-200 µm) permits greater mineral specificity. With the advent of the SAR technique, analysis can extended to smaller samples with improved precision and speed (Murray and Wintle 2000). Several studies have defined the OSL response of coarse-grained quartz and feldspar fractions from waterlain sediments (e.g., Duller et al. 1995; Wallinga et al. 2001). The ability of OSL (514 nm) to date the coarse quartz fraction in very young sediments is well demonstrated in aeolian and fluvial environments (Stokes 1999). These sediments yield DE values of 0.05 to 0.3 Gy, equivalent to < 200 years. However, coarse fractions from high energy aquatic environments may also produce apparent OSL ages that are too old (Duller et al. 1995). This is because saltation of sand in turbid flow severely limits light exposure and hence the effective solar resetting (Ditlefsen 1992). Single-grain OSL provides a means of identifying incompletely reset grains, and restricting age determinations to the fully reset population (Roberts et al. 2000). In organic-rich silty sediments from Robinson and Fog lakes on Baffin Island, the concordance between TL, OSL, and 14C dates for Holocene sediments implies that effective zeroing of the luminescence signal occurs prior to mineral deposition in these lakes (Miller et al. 1999; Wolfe et al. 2000). This increases our confidence in luminescence dates obtained on lithologically similar sediments of pre-Holocene age, which yield, in both lakes, dates in the 85-95 ka range. Examples of these results are provided in Figure 8, and are discussed elsewhere (Wolfe and Smith; this volume). To date there are no applications of OSL to dating the coarse mineral fraction of arctic lake sediments. Time range and limitations The temporal limitation of luminescence dating is affected by several factors that influence the acquisition, retention, and measurement of luminescence. Dating precision is somewhat dependent on analytical approach, with MAAD yielding typical errors in the 5-15% range, and SAR potentially reducing error to as little as 3%. Furthermore, the development of quartz SAR has greatly improved the resolution of OSL dating over the

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last 1000 years (e.g., Bailey et al. 2001). This advance implies that OSL may now provide dating control for the interval when radiocarbon dating becomes problematic, and 210Pb activities approach purely supported levels (i.e., 150-300 yr BP). When luminescence is acquired at a low rate, as in quartz in aeolian dune sands (environmental dose rate of ca. 0.5 Gy ka-1 and DE of ca. 350 Gy; Huntley and Prescott 2001), the temporal range for dating beyond 400 ka and potentially to as much as ca. 700 ka. In contrast, Baffin Island lake sediments have dose rates over four times higher (ca. 2-3 Gy ka-1), projecting an upper age limit of ca. 200 ka (Miller et al. 1999, Wolfe et al. 2000). The temporal limit of OSL is thus largely related to environmental dose rate, in conjunction with the density and stability of the charge stored in lattice defects. Sites with dose rates < 2.0 Gy ka-1 may extend the datable range beyond 200 ka, but will exhibit limited sensitivity over recent millennia. Conversely, sites with higher dose rates (e.g., > 4 Gy ka-1) approach saturation ca. 200 ka, but remain sensitive over the past ca. 100 ka. This assessment underscores several fundamental aspects of luminescence geochronology: the ability of silicate minerals to store and retain charge over geological time, the spatial and temporal variability of environmental dose rates, and the range of analytical approaches for measuring the stored charge and translating luminescence into an equivalent dose. OSL dating achieves the greatest utility for lacustrine systems when facies analysis is used to identify sediments having maximum potential for full solar resetting, and multiple luminescence sensors are used to extract credible ages. Future directions in high latitude lake sediment geochronology The need for accurate post-1950 sediment chronologies has increased sharply with pressing scientific agendas concerning high latitude climate change and air-borne pollution. Furthermore, as existing anthropogenic radioisotope marker horizons such as 137 Cs progressively decay, there are strong incentives to develop and apply additional isotopic suites that may be applicable to the development of chronologies for recent lake sediments (e.g., Am, Np, Pb, Po, Pu and U). Advances in rapid and precise isotope quantification by inductively coupled plasma mass spectrometry (ICP-MS) is an integral component of this research (Ketterer et al. 2002). Atmospheric radionuclide fluxes are insufficiently constrained at high latitudes. Thus, precipitation sampling and measurements, coupled to more detailed coupled soil and sediment inventories (e.g., Appleby et al. 2003), are still needed to adequately quantify nuclide mass balances at the basin scale. With respect to radiocarbon dating of humic acids and other fractions isolated from sediment, there remains a need for more detailed characterizations of organic matter in high latitude lakes, work that has only recently been initiated (Sauer et al. 2001). For example, fractions such as HA are operationally and not functionally defined, which complicates the task of reliably assigning provenance. The powerful allegiance of organic geochemistry and AMS has had a significant impact in oceanography (Ohkouchi et al. 2002; Hwang and Druffel 2003; Smittenberg et al. 2004). Applications of techniques developed in the marine sciences can only benefit our understanding of high latitude lacustrine environments.

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Although OSL has tremendous promise for dating lake sediments both within and beyond the useful range of 14C, additional calibration using independent chronometers is still needed, particularly in the upper age range (i.e., > 50 ka). Possible candidates for the testing of OSL dates in ancient aquatic sediments include U-Th where carbonates are present (Hall and Henderson 2001), and amino acid racemization of the protein fraction within diatom silica (Harada et al. 2002). Finally, considerable potential exists for the application of micro-tephra techniques to faciliate correlations between lake sediment records from high latitude regions (e.g., Turney et al. 1997). These techniques exploit the widespread, and potentially inter-hemispheric (Palais et al. 1992), distribution of volcanic ash particles that occur distally to their source in such low concentrations that they are not visually apparent in sediments. We are confident that high latitude lake sediment geochronology will directly profit from rigorous applications of multiple dating techniques, each of which are constantly being refined. Summary Providing accurate chronologies for high latitude lake sediments requires a degree of understanding concerning the environmental and limnological uniqueness of these systems. Biological production is typically low both on land and in water, resulting in generally low sediment accumulation rates. Slow organic matter decomposition, severe lake ice regimes, and permafrost strongly influence the transfer of radioisotopes to lake sediments used for dating purposes. On the other hand, cold temperatures and prolonged lake ice cover serve to buffer lake sediments from a variety of physical disturbances, so that most high latitude lakes have sedimentary sequences that preserve stratigraphic integrity, and can be dated with accuracy and precision. The various applications of short-lived radioisotopes (137Cs and 210Pb), radiocarbon (14C), and optically stimulated luminescence (OSL) presented in this chapter verify this to be the case. In most instances, the development of sediment chronologies provides complementary information concerning lacustrine sedimentary processes, which may assist in refining paleoenvironmental reconstruction. Examples of this, among others, include the dilution of excess 210Pb inventories in recent lake sediments, and the influence of organic matter provenance on 14C dating results. Acknowledgements Our research on dating high latitude lake sediments has been supported by various grants from the U.S. National Science Foundation and the Natural Sciences and Engineering Research Council of Canada. We thank Art Dyke and the volume editors for their constructive comments.

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Hodgson D.A., Noon P.E., Vyverman W., Bryant C.L., Gore D.B., Appleby P., Gilmour M., Verleyen E., Sabbe K., Jones V.J., Ellis-Evans J.C. and Wood P.B. 2001. Were parts of Larsemann Hills, Antarctica, ice free through the Last Glacial Maximum? Antarctic Science 13: 440-454. Hughen K.A., Overpeck J.T. and Anderson R.F. 2000. Recent warming in a 500-year palaeotemperature record from varved sediments, Upper Soper Lake, Baffin Island, Canada. The Holocene 10: 9-19. Huntley D.J. and Prescott J.R. 2001. Improved methodology and new thermoluminescence ages for the dune sequence in southeast Australia. Quat. Sci. Rev. 20: 687-699. Hwang J. and Druffel E.R.M. 2003. Lipid-like material as the source of uncharacterized organic carbon in the ocean? Science 299: 881-884. Ishiwatari R. 1985. Geochemistry of humic substances in lake sediments. In: Aiken G.A., McKnight D.M. and Wershaw R.L. (eds), Humic Substance in Soil, Sediment, and Water. Wiley, New York, pp. 147-180. Johnson-Pyrtle A., Scott M., Laing T. and Smol J.P. 2000. 137Cs distribution and geochemistry of Lena River (Siberia) drainage basin sediments. Sci. Tot. Env. 255: 145-159. Jones V.J., Battarbee R.W. and Hedges R.E.M. 1993. The use of chironomid remains for AMS 14 C dating of lake sediments. The Holocene 3: 161-163. Kaplan M.R., Wolfe A.P. and Miller G.H. 2002. Holocene environmental variability in southern Greenland inferred from lake sediments. Quat. Res. 58: 149-159. Kaufman D.S., Hu F.S., Briner J.P., Werner A., Finney B.P. and Gregory-Eaves I. 2003. A ~33,000 year record of environmental change from Arolik Lake, Ahklun Mountains, Alaska, USA. J. Paleolim. 30: 343-362. Ketterer M.E., Watson B.R., Matisoff G. and Wilson C.G. 2002. Rapid dating of recent aquatic sediments using Pu activities and 239Pu/240Pu as determined by quadrupole inductively coupled plasma mass spectrometry. Environ. Sci. Technol. 36: 1307-1311. Kirichenko L.V. 1970. Radon exhalation from vast areas according to vertical distribution of its short-lived decay products. J. Geophys. Res. 75D: 3539-3549. Krause W.E., Krbetschek M.R. and Stolz W. 1997. Dating of Quaternary lake sediments from the Schirmacher Oasis (east Antarctica) by infra-red stimulated luminescence (IRSL) detected at the waavelegth of 560 nm. Quat. Sci. Rev. 16: 387-392. Krbetschek M.R., Gotze J., Dietrich A. and Trautmann T. 1997. Spectral information from minerals relevant for luminescence dating. Radiation Measurements 27: 695-748. Krishnaswamy S., Lal D., Martin J. and Meybeck M. 1971. Geochronology of lake sediments. Earth Planet. Sci. Let. 11: 407-414. Lamoureux S.F. 1999a. Catchment and lake controls over the formation of varves in monomictic Nicolay Lake, Cornwall Island, Nunavut. Can. J. Earth Sci. 36: 1533-1546. Lamoureux S.F. 1999b. Spatial and interannual variations in sedimentation patterns recorded in nonglacial varved sediments from the Canadian High Arctic. J. Paleolim. 21: 73-84. Lamoureux S.F. 2001. Varve chronological techniques. In: Last W.M. and Smol J.P. (eds), Tracking Environmental Change Using Lake Sediments. Volume 1: Basin Analysis, Coring, and Chronological Techniques. Kluwer, Dordrecht, The Netherlands, pp. 247-260. Lian O.B. and Huntley D.J. 2001. Luminescence dating. In: Last, W.M. and Smol J.P. (eds), Tracking Environmental Change Using Lake Sediments. Volume 1: Basin Analysis, Coring, and Chronological Techniques. Kluwer, Dordrecht, The Netherlands, pp. 261-282. Linick T.W., Damon P.E., Donahue D.J. and Jull A.J.T. 1989. Accelerator mass spectrometry: the new revolution in radiocarbon dating. Quat. Internat. 1: 1-6. Lowe J.J, Lowe S., Fowler A.J., Hedges R.E.M. and Austin T.J.F. 1988. Comparison of accelerator and radiometric radiocarbon measurements obtained from Late Devensian Lateglacial lake sediments from Llyn Gwernan, North Wales, UK. Boreas 17: 355-369.

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3. PHYSICAL AND CHEMICAL PROPERTIES AND PROXIES OF HIGH LATITUDE LAKE SEDIMENTS

SCOTT F. LAMOUREUX ([email protected]) ROBERT GILBERT ([email protected]) Department of Geography Queen’s University Kingston, Ontario K7L 3N6, Canada

Key words : Arctic, Antarctic, Clastic, Geochemical, Mineral, Catchment, Lacustrine processes, Lake ice, Sedimentology

Introduction With increasing interest in the sensitivity and potential changes to high latitude environments, considerable attention has been focused on reconstructing past environmental conditions using lacustrine sediments from these regions. Although largely driven by concerns about amplified, human-induced high latitude climate change (Intergovernmental Panel on Climate Change 2001), these records also provide important information to answer other environmental issues (e.g., ecosystem function, geomorphology, hydrology). Collectively, this research has provided the first indications of how the polar terrestrial environmental systems respond to climatic and other sources of variability. Much of this paleoenvironmental information has been summarized in several regional and circumpolar composite records (e.g., Bradley 1990; Overpeck et al. 1997) and demonstrates the valuable prospective that lacustrine sedimentary records can provide. The diversity of possible proxy records contained in lacustrine sediments has led to considerable progress in several different research areas. Subfossil and other biogenic indicators are widely utilized to document ecosystem changes (see other chapters in this volume). However, the relatively low productivity found in many high latitude environments frequently results in lacustrine deposits that are dominated by inorganic material, primarily derived from around the lake or throughout the watershed. The organic component of sedimentary material is further limited by poor soil development, limited vegetation cover in the catchment and typically low autochthonous production in the lake. Indeed, the widespread occurrence of glacial, raised marine and other sediments commonly creates situations where allochthonous inorganic fluxes dominate the sedimentary record. Thus, lacustrine sedimentary records that use physical and 53 R. Pienitz, M.S.V. Douglas and J.P. Smol (eds), 2004. Long-term Environmental Change in Arctic and Antarctic Lakes. Springer. Printed in the Netherlands.

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chemical proxy information are important research tools; they hold the potential to reconstruct environmental conditions throughout the Holocene, and in select areas (e.g., Wolfe and Härtling 1996), through the last glaciation and beyond. This chapter reviews the various physical and chemical archives that are commonly found in high latitude lacustrine sediments. A review of the characteristics of the polar environments is followed by a discussion of the sedimentary characteristics associated with a range of lake-watershed environments. Our examples are drawn largely from the Canadian Arctic, although references to pertinent work in other polar environments are included. Detailed discussion of the paleoenvironmental records from different polar regions can be found in regional summaries that follow in this volume. Further, although some discussion is devoted to the unique chronological issues associated with these sedimentary records, the reader is directed to a companion chapter for more detailed information (Wolfe et al., this volume). Similarly, considerable benefits can arise from multi-proxy studies from polar lakes (e.g., Pienitz et al. 2000). The reader is directed to several companion chapters that describe the characteristics of high latitude biogenic records (Bennike et al.; Douglas et al.; Gajewski and MacDonald; all this volume). For additional detailed discussion of proxy records and paleoenvironmental methodologies, recent compilations in this series are recommended (Last and Smol 2001a). Polar environmental systems The physical sedimentary environment of a lake is a response to the composite of the processes in the lake and catchment. These include the geology which determines source materials, and (through relief) the potential and kinetic energy of the basin, the geomorphic processes which shape landscape, the hydrologic and fluvial conditions that move water and sediment to the lake, the limnic processes especially of circulation in the lake, and the overarching climate with its regional to global controls which influences most of the others. For arctic and antarctic lakes, the interaction between these components is unique and particularly important when sediments are considered as paleoenvironmental proxy because: (1) low temperatures and especially the presence of ice dominates the lake and basin; (2) seasonal variations in processes are more extreme than anywhere on Earth, with the possible exception of the monsoon environment; (3) biological and, to some extent, chemical processes are reduced and simplified, commonly rendering physical processes more evident in the sedimentary record; (4) direct human impact is limited in most cases; and (5) both natural and human-induced climate change are estimated to be greater in high latitude regions than elsewhere (Intergovernmental Panel on Climate Change 2001). Of these factors, the occurrence of ice is perhaps the most important. The cold polar climate dictates the dominance of ice as glaciers, in the fluvial system, in the ground within both permafrost and the seasonally frozen active layer on top, as snow on the land surface, and on the lake itself. Glaciers presently occupy 1.73 x 106 km2 on Greenland, 0.147 x 106 km2 in arctic North America, 0.0038 x 106 km2 in arctic Europe, and 12.5 x 106 km2 in Antarctica (Andrews 1975). Glacial erosion has and is continuing to create many lakes as scouring by glaciers overdeepens lake basins. The hydrologic

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regime of glaciers (Church 1974; Woo 1993) determines that flow: (1) is delayed and sustained throughout the summer due to ice melt; (2) continues during some or all of the winter period from water stored deep within the glacier (Benn and Evans 1998); and (3) is commonly catastrophically augmented by floods (jökulhlaups) of water dammed by glacial ice (Röthlisberger and Lang 1987; Braun et al. 2000). Sediment output from glacial erosion is normally significantly greater than in non-glacial settings (Fenn 1987) and dominates lakes where glaciers occur in the drainage basin (Lewis et al. 2002). As well, glacial retreat exposes soft sediments, including tills and glacio-fluvial deposits, which are easily eroded and which continue to produce sediment for extended periods after deglaciation in what has been termed a “paraglacial regime” (Church and Ryder 1972; Ballantyne 2001). Finally, glaciers exert a significant local effect on climate, including temperature, wind and precipitation, all of which influence the processes that deliver sediment to the lake, and the limnic processes that distribute it in the lake. In a periglacial environment, especially where permafrost prevails, ice occurs abundantly in the ground due to freezing of groundwater, migration of water to the freezing plane to form segregated ice, burial of glacial ice, and as icings in rivers, and from seepage to the land surface (French 1993). In each case, this ice may exert a significant influence on sediment production and delivery to the lake. Colluvial processes are accelerated by freeze-burst-thaw on exposed steep bedrock surfaces (Church et al. 1979) and as a result of gelifluction associated with heaving and creep in the active layer (Benedict 1976). The melting of ice-rich permafrost produces abundant water and sediment, and the subsequent mass failures, even on low slopes (Harris and Lewkowicz 1993, 2000), deliver this sediment to the rivers and lakes in the drainage basin. This thermokarsting is also important in the formation of lake basins in ice-rich environments (Carsen and Heusey 1962; French 1996). Although the presence of snow and associated nival melt in spring is not restricted to high latitude settings, the impact in the polar desert is commonly more significant where most of the annual discharge may occur during a brief period in spring (Chinn 1993; Woo 1993). As well, the bulk of the melt and runoff occurs before the active layer has formed. Thus, there is limited opportunity for recharge of soil and groundwater, increasing the magnitude of these short-lived floods. The extreme seasonality of runoff means that, in all but the larger drainage basins, especially those with large lakes and glaciers, water and sediment input to the lake during winter is close to zero. Although above freezing periods with associated runoff may be common in winter in some regions (Gilbert and McKenna Neuman 1988), they tend to be brief, and the impact on the lacustrine system is normally minimal. Ice on the surface of arctic lakes commonly lasts for 8 to 12 months of each year, although, unlike the polar lakes of Antarctica (Nedell et al. 1987), normally at least a moat of open water develops every year in the nearshore area of the most polar lakes (Figure 1) that do not become ice-free (Doran et al. 1996). The effects of ice on the sedimentary environment are several: (1) Thermal expansion during early winter and wind-driven push at break-up modify shorelines (Ashton 1986). (2) Sediments may be frozen into the ice during formation or delivered to the ice surface by fluvial, colluvial or aeolian processes, and subsequently rafted away from the sources (Luckman 1975). Ice-rafted gravel is particularly noticeable in lacustrine sedimentary deposits where water-borne sediment is normally limited to fine sand, silt and clay size (e.g., Squyres et al. 1991). (3) Ice reduces the erosional effect of waves in nearshore regions, although it

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has been observed to accelerate shoreline currents and the erosion and elongation of the oriented lakes of Alaska (Carson and Hussey 1962). (4) Ice provides a barrier to mass and energy transport across the air-water interface. The importance of ice is that its presence, and growth or decay seasonally and over longer time periods, strongly affects the sedimentary environment of the drainage basin and, therefore, the lake, and in many cases responds significantly and quickly to changes in climate. The effect is enhanced by positive feedback between the presence of snow and ice cover and the short-wave radiation balance responsible for much of the annual heating of the arctic landscape: a snow or ice cover raises the albedo, thus reducing the amount of incident radiation absorbed and the associated heating of the surface, and so preserving the ice cover, reducing the runoff of water and sediment, and the subsequent depositional record in the lake. Aeolian processes are important in the Arctic where an incomplete cover of tundra vegetation is less effective in preventing erosion, and blowing snow is an agent in initiating transport. Significant erosion occurs in winter associated with sublimation of binding ice from sedimentary deposits such as sandurs (McKenna Neuman 1993). Lake ice provides a platform, allowing the traction and saltating load of sand to be distributed widely over the lake surface, especially to areas of thicker snow on the ice where the sediment is trapped. Accelerated melting associated with the darkened surface causes earlier than normal break-up at these sites (Simmons et al. 1986; Anderson et al. 1993; Lewis et al. 2002). Thus, even where ice rafting does not occur during break-up, distribution of aeolian sediment may be widespread through lacustrine sedimentary deposits (Figure 1). High latitude lake systems Detailed description of the sedimentary structures represents an important proxy record in high latitude lacustrine sediments. Interpretation of lacustrine sedimentology provides information regarding the changes in the depositional environment through time and the relative proportion of allochthonous and autochthonous materials delivered to the lake bottom (Figure 1). In high latitude environments, lacustrine biogenic productivity is typically low by temperate measures, and therefore the sedimentology of lacustrine deposits is dominated by allochthonous material. Sedimentological descriptions are available from a limited number of high latitude sites and interpretations draw from the extensive sedimentology literature. Several reviews provide in-depth discussion of lacustrine sedimentary processes and deposits (e.g., Jopling and McDonald 1975; Davidson-Arnott et al. 1982; Håkanson and Jansson 1983; Ashley et al. 1985), although they do not provide information specific to high latitude lakes. Due to the wide range of high latitude lacustrine environments and the unique climate, hydrology and land surface characteristics of these regions (see previous section), sedimentary units (facies) vary substantially from most mid- and low latitude lakes. Indeed, the seasonal shift in the dominance of environmental processes within high latitude regions generates considerable local variation in sedimentary facies

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Figure 1. Sedimentray processes and patterns in proglacial Bear Lake, Devon Island (75°N, 85°W). In addition to sediment distribution via underflows and homopycnal flow, the presence of a persistent ice cover acts as a platform for aeolian processes. Note the relative importance of infrequent, high magnitude snowmelt and rainfall events in the distal portion of the lake, compared to the proximal zone that is dominated by glacial and aeolian influxes. Drawing by J. Glew (reprinted with permission from Lamoureux et al. 2002).

between lakes. Therefore, the facies contained in a lake are sensitive to environmental processes in the catchment and are an important environmental proxy. Notwithstanding these local differences, careful investigation of sedimentary facies provides crucial information regarding the environmental processes that generate the sedimentary record. Additionally, the facies provide evidence for the characteristic timescales for specific deposits and, hence, whether the sedimentary proxy record is episodic or essentially continuous. For example, deposits produced by floods and other extreme processes in the lake or catchment can produce thick lacustrine deposits over the course of days, hours or less (Doran 1993; Lamoureux 2000). Yet, in the context of Holocene paleoenvironmental analysis, these depositional episodes may constitute a disproportionately large component in the sedimentary sequence. Recognizing these temporal disparities is crucial for generating appropriate paleoenvironmental

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interpretations and has important ramifications for subsampling of different proxy materials, particularly subfossils such as pollen and diatoms. Therefore, as a starting point for any sedimentological investigation, it is appropriate to consider the typical catchment and lacustrine environments in high latitude regions. Even though we argue that arctic and antarctic catchments and lakes are distinct from more temperate settings where most of the sedimentological literature has focused, high latitude environments retain many common sedimentological facies associations (e.g., Wolfe and Härtling 1996; Lamoureux et al. 2002). The major lake types and their characteristic sedimentary properties are discussed below, including examples drawn from different high latitude lakes. As well, these high latitude sediments can be used to interpret the immediate postglacial sedimentary record of many temperate lakes, when high latitude conditions may have prevailed. Glacial lakes Studies of contemporary ice-proximal glacio-lacustrine deposits are comparatively rare, although they are identified in exposures during geological investigations. Proximal glacio-lacustrine settings are high-energy environments typified by dynamic ice margins and the interface between supraglacial, englacial, and subglacial meltwater channels (Lawson 1993; Benn and Evans 1998). The location of the ice margin is controlled by water depth, ice thickness, and ice flow and calving rates (Holdsworth 1973). Over annual to decadal scales, the ice margin may appear to be relatively stationary, although it may be subject to rapid adjustments on the sub-seasonal timescale. At multi-decadal and longer timescales, glacier mass balance may vary substantially, resulting in significant changes to the location of the ice margin, and imparting an ephemeral nature to many glacio-lacustrine deposits (Lawson 1993). Over Holocene and Pleistocene timescales, many proximal glacio-lacustrine deposits appear to have been relatively short-lived, often resulting from temporary ice damming (Figure 2) or other unstable geomorphic conditions (Benn and Evans 1998). The sediment carried by glaciers is controlled by a variety of bedrock and glaciological conditions. Typically, abrasion of bedrock at the ice bed produces abundant fine-grained sediment, frequently referred to as rock flour (Benn and Evans 1998). Glacier ice can transport sediment in a wide range of sizes, up to and including large boulders. Sources for these materials range from the glacier bed to sediments delivered by mass wasting from adjacent slopes. Finer sediments are frequently carried via meltwater pathways to the ice margin, compared with the slow movement of large clasts carried by the relatively slow moving ice (Lawson 1993). Ice calving releases sediments directly into the lake with minimal redistribution, frequently resulting in poorly sorted deposits at the calving front. Due to the wide range of clast sizes commonly contained in ice and the lack of sorting mechanisms, deposits at the ice front tend to resemble glacial diamicton composed of coarse, angular clasts in a matrix of finer sediment (Gravenor et al. 1984). Deposition in deep lake water may remove appreciable amounts of the clay and silt that remain in suspension and are transported elsewhere in the lake. Sediment is commonly concentrated in crevasses and, thus, deposition as rainout at the calving ice front is highly localized on the lake bed (Figure 2).

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By contrast, inflow and distribution of sediment-laden meltwater beyond the ice margin is controlled by the density of the meltwater compared to the lake water. Water density is controlled primarily by suspended sediment concentration, temperature, and salinity. Freshwater reaches maximum density at ca. 4ºC, but density may be substantially increased by suspended or dissolved constituents. Given that lake water temperatures in contact with glacier ice are already close to the maximum density, the inflow density is often controlled by the suspended sediment load (Figure 3). Due to the abundance of fine rock flour found in many glacial environments, suspended loads in rivers are frequently high, at least several g·L-1 but uncommonly ranging up to 80 g·L-1 (Gilbert 2000). Therefore, turbid glacial meltwater is frequently denser than the lake water and results in underflows in the lake. During conditions of lower suspended sediment discharge, inflow may not be sufficiently dense to sink to the lake bottom and distributes as overflows or interflows. In cases where the lake water is not stratified, the meltwater is frequently distributed throughout the water column as homopycnal flow (Smith and Ashley 1985).

Figure 2 . View in September 1983 of the retreating glacier separating glacier-dammed North (left) and South (right) Stewart lakes, eastern Baffin Island (70°40’N, 71°30’W), Nunavut. The Little Ice Age terminal moraine indicates the source of sediments formerly deposited proximally at the ice front and by rafting distally in the Stewart lakes. Ice-proximal deposition in the lake within the moraine has been occurring since it began to form in the mid-1950s (Gilbert et al. 1985).

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Figure 3. Nomograph showing how the variables of temperature and concentration of suspended sediment (SSC) with density ȡs = 2.7 g·mL-1 determine the density of freshwater at atmospheric pressure, where ȡw refers to the density of pure water. For example, if water at 10°C is flowing into a lake at 3°C having SSC of 10 mg·L-1, the inflow must have SSC of 500 mg·L-1 to be equally dense as the lake water. A somewhat greater density would be required to generate a turbidity current. This ignores the minor effect of cabbling (Hamblin and Carmack 1978). However, because most high latitude rivers remain close to 4°C through the summer, relatively low SSC may be sufficient to induce turbidity current flow in the lakes into which they flow.

Unlike the sediments that rain out at the ice front, turbid meltwater plumes travel considerable distances in the lake, gradually losing energy to turbulence. With decreasing competence, coarser sediments begin to settle from the meltwater. Settling rate, as defined by Stoke’s Law, is largely a function of particle diameter (Dietrich 1982). Hence, sand settles in most lakes rapidly (minutes to hours), while silt and clay may remain in the water column for weeks to years. This range of settling times leads to substantial sorting of suspended sediments in the water column which is reflected in the stratigraphy of the bottom sediments. Sorting is especially pronounced from overflows, interflows and homopycnal flow, where the settling distances keep the finest sediments suspended during the entire melt season and well into the winter (Smith and Ashley 1985). Deposition from underflows is more rapid due to the proximity of the lake bottom. Moreover, many underflows carry coarse silt and sand which settle out rapidly, leading to relatively uniform (massive) or graded coarse deposits interspersed with fines.

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Estimating settling times on the basis of Stoke’s Law is an unrealistic assumption in many cases. Not only are most sedimentary particles substantially non-spherical, but chemical and biological processes in lakes alter the effective particle size and settling times, particularly for clay (Dietrich 1982). Electrical charges associated with clay minerals cause attraction of individual particles and the formation of flocs that have shorter settling times (Droppo et al. 1998). Additionally, fine particles digested by zooplankton are incorporated into fecal pellets which have rapid settling times (Smith and Syvitski 1982). Therefore, even though arguments have been made that settling times for fine clay may take years (Lemmen et al. 1988), the water column appears to clear over the winter in some relatively deep arctic lakes (e.g., Lamoureux 1999). Sediment accumulation rates generally decrease distally, leading to thick deposits in the proximal inflow regions of the lake (Smith 1978). The exception to this is the case of turbidity currents, where deposition can be higher distally due to proximal bypassing (Gilbert et al. 1997). The spatial pattern of deposition also typically varies with depth, particularly where deposition is associated with underflows and either interflows or homopycnal flow. Ashley (1975) identified three categories of glacio-lacustrine sedimentation, depending on the proportion of coarse to fine sediment. In near-proximal locations, the coarse (silty) layers were thicker than the clay layers, indicating increased dominance by underflows and more energetic sediment plumes. In contrast, clay layers were thicker than the silts in distal locations due to the reduced importance of underflows compared to lower energy interflows and suspension settling (Ashley 1975). Underflows concentrate in the deepest part of the lake compared to the lake-wide distribution of the shallower flows (Smith and Ashley 1985). As a result, underflow deposits are localized in the deepest portions of the lake. In cases where the lake bottom is comparatively flat (often resulting from frequent high bottom sedimentation by underflows) underflows spread laterally and down-lake (Gilbert 1975; Smith et al. 1982; Smith and Ashley 1985; Chikita et al. 1999; Lamoureux et al. 2002). Minor topographic obstacles on the lake floor may divert underflows (Chikita et al. 1999; Lewis et al. 2002), although high sills can halt the forward movement of the plume, potentially causing the flow to reverse (Gilbert et al. 1997). While the processes and characteristics of glacio-lacustrine sediments associated with warm-based glaciers are comparatively well understood, much of the ice cover in the High Arctic and Antarctica is cold-based and frozen to the bed (Benn and Evans 1998). Sediment delivery and deposition associated with cold-based ice are poorly documented, as the geomorphic processes operating at the base do not appear to be conducive to rock abrasion and sediment transport (Lawson 1993; Benn and Evans 1998). Moreover, the absence of basal sliding and subglacial channels further reduces potential rock abrasion and erosion, lowering available sediment for transport by meltwater. In most situations, calving at the margin and drainage of supraglacial channels are the primary ablation mechanisms. Therefore, although some studies have interpreted coarse, poorly sorted lacustrine sediments to be generated in ice-proximal conditions, potentially with cold-based ice (Smith 2000), this combination remains poorly understood. Ice-distal lacustrine environments are more common, although, like more temperate examples, they are limited primarily to alpine regions. Distal glacio-lacustrine environments have been the subject of considerable research concerned with Holocene and recent glacier activity (Østrem and Olson 1987; Lemmen et al. 1988; Mangerud and

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Svendsen 1990; Svendsen and Mangerud 1997; Hasholt et al. 2000; Snyder et al. 2000; Wagner et al. 2000; Noon et al. 2001; Lamoureux et al. 2002; Verkulich et al. 2002). These deposits are also frequently found in lake sediments of deglacial age (late Pleistocene and early Holocene). Therefore, even though Holocene records are commonly composed of organic-rich gyttja, many high latitude lake records have some postglacial glacio-lacustrine sediments. The key difference between proximal and distal glacio-lacustrine deposits is the absence of the poorly sorted, coarse sediment found in the ice-proximal deposits. Intermediate fluvial processes in the proglacial environment introduce the potential for sediment storage, release and sorting. Sediment storage through channel aggradation may operate over a wide range of timescales, acting as a buffer for glacio-fluvial sediment transfer. Conversely, degradation of postglacial channel deposits and erosion of glacially-placed sediment on hillslopes generates additional complexity to long-term sediment yield and accumulation in downstream lakes. This paraglacial sedimentation (Church and Ryder 1972; Ballantyne 2001) appears to operate over timescales of thousands of years and can result in an indirect glacial influence on lacustrine sedimentation long after ice has retreated or disappeared from the catchment (Church and Slaymaker 1989; Lamoureux et al. 2002). Facies found in lakes dominated by glacial meltwater are controlled by the same inflow density and lake stratification considerations as detailed in the previous section. The distance from the ice margin increases water temperature in both the meltwater and especially in the lake. However, warming of the lake may not result in a greater density contrast with the cold inflowing water because many high latitude lakes do not warm significantly above 4-6ºC due to the persistence of lake ice (Lewis et al. 2002). Regardless, turbid inflows can lead to underflows and associated sediment facies (Hasholt et al. 2000). When sediment loads decrease due to exhaustion or reduced meltwater production, interflows tend to dominate only if the lake is stratified (Smith and Ashley 1985). The lack of stratification is an important difference between high latitude and temperate proglacial lakes. For example, in monomictic Bear Lake in the Canadian High Arctic, Lewis et al. (2002) observed underflows on warm days that were generated by high melt and runoff from the adjacent Devon Island Ice Cap. Sediment accumulation rates determined from cores and acoustic data suggested that, over the Holocene, underflows have been an important sediment delivery mechanism (Lamoureux et al. 2002). On days with lower melt intensity, underflows were not observed, suggesting that the sediment load in the meltwater was not sufficient to generate a density contrast with the cold water column (Lewis et al. 2002). Thus, homopycnal flow occurred through much of the melt season, leading to suspension settling across much of the lake basin (Lamoureux et al. 2002). Due to the high sediment loads and strong diurnal and seasonal variability in sediment delivery, glacio-lacustrine sediments are commonly composed of rhythmic laminations representing phases of relatively rapid deposition of coarse sand and silt during high inflows alternating with fine silt and clay that settle out during quiescent periods (Hasholt et al. 2000; Lamoureux et al. 2002) (Figure 4). Depending on the location in the lake and the hydrological regime, the laminae may represent annual deposits (varves) that are especially useful in paleoenvironmental reconstructions. However, varves are not produced in every case, even though the sediments may mimic the classic

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glacial varve structure of a coarse and fine couplet (de Geer 1912). While many proglacial lake sediments in high latitude regions do appear to be varved (Smith 1997; Hasholt et al. 2000; Lamoureux et al. 2002), some occur as rhythmites without a recognizable annual signal (Lemmen et al. 1988).

Figure 4. Thin section illustrating laminated sediments in the proximal basin of Bear Lake, Nunavut (75ºN, 85ºW). Deposition in this proglacial lake is dominated by irregular underflows and aeolian deposition of isolated coarse grains (A). Individual laminae represent sediment input events and are likely deposited from turbidity currents. Groups of laminae capped by darker, thin clay layers (marked) are interpreted as varves. Black marks on edge of photograph indicate the upper boundary of each varve.

Smith (1978) noted that the sedimentary structures varied systematically along a proximal to distal transect within Bow Lake, Alberta (51ºN, 116ºW). In the proximal zone, frequent inflows of sediment-laden meltwater produced frequent laminae, primarily from underflows. In progressively more distal sites, the dominance of underflows decreased and the frequency and magnitude of sediment inputs were attenuated, resulting in fewer laminae in the sedimentary record. He suggested that the reduced number of laminae represented progressively longer term sediment delivery variations, ranging from diurnal in the proximal zone to synoptic (weekly) in the distal zone. In the most distal zone, the sedimentary structures appeared as a simple varve couplet. Smith’s conclusions have been demonstrated in other environments and similar proximal-distal progressions can be observed in high latitude sites (e.g., Lamoureux

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1999; Hasholt et al. 2000). Therefore, it is important to consider the spatial context for the sedimentary facies along with the limnological processes. It is interesting to consider why varved lakes are not more common in high latitude regions. Based on the increased seasonality of these environments, one would expect that annual signal would be dominant in sedimentary sequences. However, recent work has shown that short-lived environmental events (e.g., storms) in high latitude environments can dominate the sedimentary record (Lamoureux 2000; Lewis et al. 2002). Hence, sedimentary rhythms are frequently driven by irregular sediment delivery mechanisms. Where lake conditions are suitable (deep, high sediment supply), varves have been documented (Hughen et al. 2000), but other proglacial lakes do not contain varves (Lemmen et al. 1988). Similarly, in a study including several hundred northern lakes, Larsen et al. (1998) noted that laminated sediments were infrequent as most lakes were too shallow to consistently preserve sedimentary structures. Therefore, even while laminae may be present in a wide range of polar lacustrine settings, care must be taken to avoid incorrectly assigning a varve interpretation. However, the sensitivity of the high latitude environment to short-lived storm events provides the opportunity to identify the changing frequency of these important events in response to other environmental changes (e.g., Lamoureux 2000). Periglacial lakes The vast majority of the catchments in high latitudes do not contain glaciers (Figure 5) and several processes typical only of the periglacial environment play an important role in generating the lacustrine record. Sediments delivered by fluvial processes may be sufficient to be important for generating lake sedimentary deposits (Figure 6). More commonly, high latitude lake sediments are dominated by modest or low fluvial sediment delivery due to the relative absence of autochthonous biological productivity. Other processes in the periglacial environment that may influence the sedimentary record include aeolian and ice rafting deposition (e.g., Retelle 1986; Squyres et al. 1991). Taken collectively, these processes often lead to lacustrine sequences that are substantially different from low latitude sites. Fluvial processes and facies Like glacierized catchments, the hydrological regime of most polar catchments is strongly seasonal and dominated by the spring nival freshet (Woo 1993). Exposed surface and soil materials are frequently highly erodible due to limited vegetation cover, leading to pronounced erosion and transport of sediment to lakes during the freshet (e.g., Braun et al. 2000). Initially, the active layer is impermeable due to frost, leading to higher surface runoff. As the summer progresses and the active layer thickens, the potential infiltration and water storage increases. As the spring flood recedes, the amount of sediment transported also tends to decrease substantially. In most cases, sediment fluxes to the lake are minimal for the remaining hydrological season. Summer precipitation can be an important mechanism for sediment transport, although active

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Figure 5. Examples of periglacial lakes in the Canadian High Arctic Archipelago, including: (A) small lakes on northeastern Ellesmere Island (July 20, 1988) (82ºN, 70ºW), and (B) Nicolay Lake, Cornwall Island (July 15, 1997) (78ºN, 95ºW). Note the prominent river delta built into Nicolay Lake resulting from sediment delivery from a large catchment. By comparison, the lakes shown on Ellesmere Island have limited watersheds and sediment influxes.

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Figure 6 . Laminated sediment in thin section from a distal location in Nicolay Lake, Cornwall Island (78ºN, 95ºW). The sediments are made of couplets of silt and clay laminae with frequent aeolian sand and silt grains (A). These sediments are also interpreted as varved, based on the sedimentology, the catchment hydrology, and independent radioisotopic dating. Black marks on edge of photograph indicate the upper boundary of each varve. Unlike the varves from a proglacial environment (Figure 4), these varves do not have evident subannual laminae due to their distal location, although subannual structures are common in proximal locations.

layer storage must be exceeded before a major runoff response will occur. If the active layer is saturated, rainfall-induced runoff is routed without temporary storage in melt ponds. This runoff has access to thawed channel and slope sediments and can be an effective denudation agent and be important for lake sedimentation (Lamoureux 2000). Sedimentary facies associated with fluvial processes vary substantially but are strongly affected by the catchment sediment yield, the ratio of lake to watershed area, and limnological processes, particularly sediment focusing and stratification. The pronounced seasonality of sediment delivery found in most high latitude regions is highly conducive to the formation of laminated sedimentary structures in lakes (Figure 6). As discussed above, the periglacial environment is especially conducive to an amplified geomorphic response to modest hydrometeorological events. Typically, laminae are well-sorted and commonly normally graded. Sedimentary contacts are usually conformable but may be indistinct, depending on the range of particle sizes present. Where yield is comparatively high, individual millimetre-scale lamina may be produced by fluvial processes, generating an overall laminated sequence (e.g., Coakley and Rust 1968). Catchments with lower sediment yields punctuated by occasional

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increased sediment delivery events are more likely to produce an inconsistent laminated sequence, containing sections of massive sediment (e.g., Doran 1993). In cases of consistently high sediment yield found in catchments with readily erodible surface materials, laminae may be produced by major hydrometeorological events (Lamoureux 2000). Lake and catchment areas are also important factors that control the appearance of lacustrine facies. In the case of a small lake to catchment ratio, sediment delivered by streams is deposited in a smaller area, leading to thicker deposits and greater potential for discernable structures. By contrast, relatively large lake to catchment ratios result in less concentrated sediment deposition. Moreover, larger lakes provide greater opportunities for inflow sorting and dispersion processes, frequently leading to decreased particle size variability in central and distal depositional sites. Similarly, lake stratification can play an important role in the distribution, deposition and preservation of the sediment. While stratification of the water column may be important for directing sediment inflows to different depths (e.g., Retelle and Child 1996), suspended sediment concentrations are frequently too low to generate persistent underflows. More commonly, thermal or chemical stratification leads to interflows that disperse sediment throughout the lake basin. Preservation of the original sedimentary structures is enhanced by stratification by minimizing wave-induced resuspension and, in some cases, by the formation of an anoxic hypolimnion (Figure 7). The latter appears to be relatively uncommon, due the low levels of productivity in most polar lakes. The exception is meromictic lakes found in coastal areas that contain trapped sea water or brine expelled during permafrost growth that lead to permanent hypolimnetic anoxia (Gallagher et al. 1989; Bradley et al. 1996; Verkulich et al. 2002). Biogenic processes and facies Low levels of available energy and the short ice-free season limit biogenic production in most polar lakes (Smol 1988). Productivity is further limited by the oligotrophy of many polar lakes (Douglas and Smol 1999). Therefore, biogenic facies are rarely described, and what authigenic deposition occurs is frequently in conjunction with a substantial quantity of clastic sediment. Arctic environments are capable of supporting diverse ecosystems (Douglas and Smol 1999) but most sediment contains relatively little material that can be identified at the core level. Most frequently, detrital organics composed of material eroded from the catchment or littoral zone of the lake constitute identifiable materials in the sediments. In some cases, these materials may form individual layers in the sedimentary sequence. Microflora including siliceous algae (diatoms) and chironomids are ubiquitous in polar water bodies (Bennike et al.; Douglas et al.; all this volume) and represent a rich paleoenvironmental archive. Only rarely, however, do these subfossils generate discernable sedimentary structures. In some cases, particularly meromictic lakes where the preservation potential of initial sedimentary structures is high, biogenic laminae have been documented (Retelle et al. 1996; Hughen et al. 2000). Indeed, the presence of clear biogenic layers is largely due to the absence of substantial co-deposition of clastic material during the productive summer months.

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Figure 7. Microphotograph of thin-sectioned varves found in meromictic Lake C2, northern Ellesmere Island, Nunavut (83ºN, 76ºW). Black marks on edge of photograph indicate the upper boundary of each varve. These sub-millimetre varves are preserved due to persistent hypolimnetic anoxia in the lake, despite a very low sediment yield.

Aeolian processes and facies The limited vegetation cover and strong winds frequently found in high latitude regions are important preconditions for widespread aeolian erosion and transport (McKenna Neumann 1993). Deposition of aeolian sediments has been widely documented in high latitude regions (McKenna Neuman 1990; Squyres et al. 1991; Edlund and Woo 1992; Lewis et al. 2002). In extreme cases, winds are sufficiently strong to transport gravel (Edlund and Woo 1992). Recognition of aeolian deposits in lacustrine deposits has been infrequent, in part due to the lack of detailed sedimentology necessary to observe these features. Where they have been identified, aeolian deposits occur as isolated sand and silt grains supported by a finer matrix of sediment (Retelle 1986; Lamoureux 1999; Lamoureux et al. 2002). These observations were not apparent from core lithostratigraphy and required identification using embedded sediments (e.g., Pike and Kemp 1996; Lamoureux 2001) under low magnification (Figure 6). Although the descriptions focus on sand grains ranging to 500 µm in diameter, clearly finer aeolian sediments may be present but are indistinguishable from the matrix sediment. In most cases, the sand grains are isolated and appear in low concentrations. However, layers of discrete aeolian sediment have been documented in lakes that are prone to frequent aeolian deposition on the lake ice (Lamoureux et al. 2002). The high density of aeolian sediment required to produce contiguous layers in the lake bottom probably results from concentration on the lake ice surface in snow concavities and during spring snow melt. Concentrated sediments lower the surface albedo and locally accelerate the decay of the ice cover (Lewis et al. 2002).

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Once the sediment melts through ice, the grains are subject to radial dispersal due to currents and turbulence in the water column. As well, grain to grain interactions and random lateral movement cause clouds of sediment to spread from a point source to diameters of up to several hundred metres, depending on the volume of sediment, the grain size and the water depth (Krishnappan 1975; Gilbert 1990). Therefore, the production of a dense, contiguous aeolian deposit in a deep lake requires relatively minimal currents and is aided by the high density of the sediment released from the ice that tends to promote sinking rather than lateral dispersal. In relatively shallow water, Squyres et al. (1991) observed localized mounds of sediment released from the permanent lake ice of an antarctic Dry Valley lake. Therefore, the sedimentary evidence to date suggests that aeolian deposits may be important in polar environments, particularly where they are concentrated on the lake ice. Ice rafting processes and facies The presence of thick lake ice during part of the active deposition season can also form unique sedimentary deposits in lakes. Moreover, in the High Arctic and Antarctica, lake ice may persist through most of the summer and become progressively more mobile as the initial shoreline moat widens by melt and edge attrition. Coarse, isolated gravel clasts found in lake sediments are frequently attributed to sediment captured or frozen to the ice edge and transported to other areas of the lake where they are released into the water column (Figure 8) (Gilbert 1990). Similarly, strong winds can force the mobile ice pan onto the shoreline resulting in the entrainment of sediment that is subsequently released elsewhere in the lake. Finally, early season fluvial sediment can be prevented from entering the lake water by ice cover. This sediment, much like aeolian sediment, is rafted on the lake ice until it is released by ice melt or disintegration. Mixed lake environments Late Pleistocene and Holocene deglaciation has led to many substantial landscape changes, most particularly the relatively slow glacio-isostatic adjustment of the land due to removal of Late-Wisconsinan ice sheets. As a result of glacio-isostatic submergence and subsequent emergence (up to several hundred metres), many lakes in high latitude regions were or continue to be influenced by salt water (Figure 9) (McLaren 1967; Retelle 1986; Ouellet et al. 1987; Gallagher et al. 1989; Svendsen et al. 1989; SaulnierTalbot and Pienitz 2001; Verkulich et al. 2002). The impact on sedimentary sequences is two-fold: (1) the early sedimentary sequence may actually be marine sediments; and (2) residual saline water can lead to permanent stratification or meromixis. Prior to emergence, lake basins that were subject to marine processes contain an important paleoenvironmental archive (e.g., Gallagher et al. 1989; Gilbert 2000). Key differences that differentiate most polar lacustrine and marine sedimentary systems include water salinity, proximity to sediment supplies, and bioturbation. The high

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Figure 8. An ice-rafted stone in a section of a late Pleistocene glacial lake at the head of Itirbilung Fiord, eastern Baffin Island (69ºN, 67ºW). Silty sediments in the vicinity of the stone have been penetrated or down-warped by its impact, while subsequent deposits conformably overlie it.

salinity of marine waters is important for both the vertical density structure of the water column and for fine-grained sediment rapidly flocculating into larger aggregates. Depending on the location of the marine basin, sedimentation rates may be much lower than comparable lacustrine rates, due to the proximity of river inflows (Svendsen et al. 1989). Lamoureux (1999) noted that a pre-emergent High Arctic lake received significantly less sediment due to inland marine basins that captured terrestrial sediment for much of the Holocene. As emergence progressed, the inland basins were isolated, locating the zone of active fluvial sedimentation closer to the modern coast. Finally, biological activity, particularly by benthic fauna, has important implications for the preservation of sedimentary structures of marine sediments (Dale et al. 1989). Burrowing organisms are common in marine environments and frequently mix to a depth of several decimetres. If sedimentation rates are low, primary sedimentary structures can be destroyed by the pervasive action of marine organisms. For this reason, pre-emergent marine sediments are commonly massive and contain fossils of molluscs and other burrowing organisms (e.g., Retelle 1986; Lamoureux 1999). In coastal areas subject to glacio-isostatic emergence, saline bottom waters can remain during and following isolation of the lake and result in a meromictic (chemically stratified) water column. In some cases, a lake basin may remain linked to the sea through tidal flushing (e.g., Hughen et al. 2000) that replenishes the saline water and maintains meromixis (Figure 9). Alternatively, salinity may be increased by brine rejection due to permafrost growth in the newly exposed land (Ouellet et al. 1987).

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These types of meromixis persist after isolation (Bradley et al. 1996), leading to permanent hypolimnetic anoxia. In both of these situations, the absence of bioturbation and wave resuspension results in the preservation of sub-millimetre laminae and varved sedimentary structures (Figure 7) (Zolitschka 1996; Hughen et al. 2000). Therefore, coastal lakes represent a special situation with the potential to preserve many sedimentary features that are often obscured by bioturbation of the primary structures.

Figure 9 . An oblique aerial photograph of the north coast of Cornwall Island, Nunavut (78ºN, 95ºW), showing several lakes that have been isolated during Holocene glacio-isostatic emergence. Emergence of the coastline transforms embayments into lakes. In some cases, sea water becomes trapped, thus generating meromixis. In the early stages of isolation, intermittent replenishment of sea water may occur during spring tides and storm surges. In addition, as permafrost grows beneath the newly exposed sea floor, brine rejection at the freezing front forces saline groundwater into the lake, commonly raising its salinity to several times that of sea water (Ouellet et al. 1987). Photo T440R-124, copyright Her Majesty the Queen in Right of Canada.

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Physical and chemical proxy records in high latitude lakes The techniques employed by lacustrine sedimentologists fall into three general categories: (1) measurement of the physical and chemical processes in lakes to understand the factors that control the delivery and deposition of sediment; (2) geophysical characterization of the lacustrine sedimentary environment; and (3) direct analyses of sediments recovered from the lake. Discussion of the first category is beyond the scope of this chapter (see Last and Smol 2001a for detailed coverage); however, it is important to note that comparatively few process studies have been carried out in high latitude lake environments (Lemmen et al. 1988; Retelle and Child 1996; Lewis et al. 2002). In this section we illustrate some of the applications of these techniques to high latitude lacustrine records and provide examples of how they instruct environmental and paleoenvironmental assessment from sedimentary records. Geophysical characterization of lacustrine sediments (acoustics) Following the long tradition of marine sedimentologists, those studying lakes are increasingly using geophysical techniques to understand aspects ranging from the morphology of the lake to the spatial and temporal pattern of sedimentation in the lake (e.g., Gilbert et al. 1997; Eyles et al. 2000). The technique is analogous to remote sensing as applied to the terrestrial environment with sensors in the electromagnetic spectrum (including visible light). However, because much electromagnetic energy propagates poorly in water, acoustic energy is normally substituted (cf. Moorman and Michel 1997), and the techniques involve active emission of sound. Some of the range of equipment and analyses are described by Gilbert (1999); they include mapping of bathymetry and classification of the type of bottom by high-frequency echo sounding, assessment of the structure and content of the water column using high-frequency sound sources, high-resolution mapping of thin, soft sediments using sources of intermediate energy and frequency, and high energy, low-frequency seismic techniques to penetrate entire sedimentary records. Developments with digital equipment have greatly expanded the versatility of these techniques including digital chirp systems which allow resolution of internal reflectors on a sub-decimetre scale (Fang 1996) and swath mapping which yields high-resolution bathymetry, approaching the definition provided by air photographs on land (Matula 1992). Because most acoustic techniques are not applicable to ice-covered water and require a boat or float-equipped aircraft to operate, they have not been widely applied in studies of high latitude lakes, despite extensive use in similar polar marine environments (Hay 1984; Hill et al. 1999; Gilbert 2000). Svendsen et al. (1989) distinguished lacustrine and marine sediment in a lake on Svalbard using a 3.5 kHz profiler, and assessed the rates of accumulation and thus terrestrial denudation from the records. Figure 10 illustrates a sub-bottom acoustic survey of Bear Lake, Devon Island using chirp technology (Lamoureux et al. 2002). From these results we may assess: (1) bathymetry and morphometry of the lake floor; (2) spatial variability in the thickness of sediment across the lake floor; (3) rates of accumulation based on thickness; (4) texture and water content of the sediment from the degree of penetration and reflectivity of the materials (for example, multiple acoustic returns indicate a high degree of reflectivity

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from consolidated sediments or rock); (5) morphometry of the glacially eroded bedrock beneath the lake and thus the origin of the lake; (6) stratigraphy of the sedimentary record based on the acoustic layering (reflectors are generated by changes in acoustic impedance between water and sediment or within sediments, which are, in turn, defined by the physical properties of the sediment); and (7) sedimentary processes (for example, conformable sediments are deposited from suspension in the water column, while sediments filling depressions with non-conformable, flat-lying reflectors are deposited by gravity flows, especially turbidity currents, and sediment focusing by creep and slope failure is clearly evident). This information is valuable in planning a coring program (to avoid anomalous records, for example) and in interpreting the sedimentary evidence from cores. Features such as dewatering structures are generally not evident in cores. They are caused by the vertical migration of water through the sediment in response to consolidation, and by upwelling of groundwater from beneath the lake. Beneath arctic lakes a talik (unfrozen zone) is maintained by the perpetually above-freezing water (Burn 2002); these are important conduits of water between surface and depth in a landscape where the substrate is dominated by ice-bonded permafrost which prevents the transmission of groundwater. The dewatering structures indicated in Figure 10 may relate to a fault in the bedrock at the lowest point which serves as a conduit for groundwater and which represents a zone of weakness more deeply eroded by Pleistocene glaciation. Sediment cores Core studies are the most common technique in lacustrine sedimentology and reveal a wide range of information regarding depositional processes, composition, and origin of the sediments. The diverse sedimentological techniques used in high latitude studies are discussed in this section. Sediment accumulation Long- and short-term mass accumulation rates (MAR) are both important indicators of catchment processes and development. In many cases, changing MAR has been interpreted as a proxy of hydroclimatic activity, land cover stability, and vegetation cover. A key requirement for assessing MAR is a sedimentary chronology. In high latitude sites, there are considerable challenges involved in dating lacustrine sediments. These issues are discussed by Wolfe et al. (this volume) but several points are highlighted below. Common sedimentary dating methods fall into two categories: radioisotopic and stratigraphic. The former are the most commonly applied to polar lacustrine sediments and include 14C (radiocarbon) dating of sediments ranging up to approximately 40,000 years before present (BP) and 210Pb and 137Cs that are used to establish the age of sedimentary horizons for the last 150 and 50 years, respectively (Appleby 2001). Radiocarbon dating of organic-poor polar sediments is problematic due to the paucity of terrestrial organic remains. Moreover, slow organic decay rates in terrestrial soils and

Figure 10. Digital chirp sub-bottom acoustic profile from Bear Lake, Devon Island, Nunavut (75ºN, 85ºW) showing conformable (c) and filling (f) sediments deposited over bedrock (b). Dewatering structures (d) are the result of compaction of the sediment or groundwater movement from faults in the bedrock. Multiples (m) and crossed reflectors due to parallax (x) are indicated. Inset maps show the location of Bear Lake near the Devon Island Ice Cap (DIC) and the location of transect in the lake (isobath interval is 20 m).

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eventual transport of these materials to lakes introduces further chronological uncertainty to radiocarbon dates. In a comprehensive study of the radiocarbon content of various lake and soil organic constituents in a small arctic catchment, Abbott and Stafford (1996) demonstrated a wide range of potential ages depending on the material dated. These results suggest that application of radiocarbon dating to polar sediments should be used with caution (Wolfe et al., this volume). Dating recent sediments using 210Pb and 137Cs profiles is also limited by the dispersal and deposition of these isotopes in high latitude locations (Hermanson 1990). Although both approaches have been used with success (e.g., Lamoureux 1999), measurable levels of 137Cs and unsupported 210Pb are frequently too low to generate an interpretable profile for dating purposes. Moreover, sedimentation rates during the past 150 years may only constitute several centimetres of accumulation thereby limiting the number of 210 Pb samples that can be analysed (e.g., Douglas et al. 1994). Paleomagnetic records also hold the potential for chronological constraint of sedimentary records (King and Peck 2001). Work with Scandinavian lakes has demonstrated the potential of using regional paleointensity records as dating tools (Saarinen 1998; Snowball et al. 1998). Unlike varved sediments, paleointensity records may be suitable for use in lakes without clear or consistent sedimentation processes (Wolfe et al., this volume). Further, improvements in paleomagnetic sample resolution and processing should provide wider availability of the technique to the research community. Finally, when combined with varved sediments, the paleointensity records can be constrained more accurately, effectively extending the varve chronology to nonvarved lakes (Saarinen 1998). Sediment accumulation: varves Dating sediments using stratigraphic methods has been used successfully in a variety of high latitude sites. Varved sediments are especially useful for precise chronological control for sedimentary and paleoenvironmental analyses and have been found in glacial, coastal (meromictic) and freshwater high latitude lakes (Bradley et al. 1996; Retelle et al. 1996; Gajewski et al. 1997; Lamoureux 1999; Ojala and Saarnisto 1999; Hasholt et al. 2000; Hughen et al. 2000; Moore et al. 2001; Lamoureux et al. 2002). Construction of accurate, reproducible varve chronologies is discussed in detail by Lamoureux (2001). The primary requirements are establishing the processes that lead to annual sediment structures and independent verification, typically using 210Pb and 137Cs profiles of recent varves and, in some cases, radiocarbon dates of long varve sequences (Moore et al. 2001). Varves are not only a potential chronological tool, but can also be used to accurately measure changing MAR in detail. Several studies have used interannual MAR measurements from high latitude varves as proxy hydroclimatic record (Lamoureux and Bradley 1996; Hughen et al. 2000; Lamoureux 2000; Moore et al. 2001). As previously discussed, the strong seasonality of high latitude hydroclimatic processes is highly conducive to the deposition of sedimentary units associated with individual sediment-water inflows. Because the dominant inflow in most locations is associated with the spring freshet, most varves contain a deposit produced during this high-energy period (Hughen et al. 2000). The varve couplet is completed by the delayed settling of fines that remain suspended in the water column until ice cover isolates and

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stills the lake. This comparatively simple sedimentary succession (e.g., Figure 7) can be found in a wide range of high latitude environments. In some cases, biogenic deposits are also produced during the summer period (Gajewski et al. 1997; Hughen et al. 2000). Alternatively, in proglacial settings, nival melt is succeeded by glacial melt during the summer, resulting in one or more phases of sediment input and deposition (Lamoureux et al. 2002). Finally, in unglacierized catchments, where sufficient sediment is available, summer rainfall generates additional subannual sedimentary units (Lamoureux 2000). Until recently, the primary environmental information obtained from varved sediments in high latitude lakes was interannual sediment accumulation variability (e.g., Lamoureux and Bradley 1996; Hughen et al. 2000). Additionally, the chronological precision afforded by varved sediments has been used to accurately date changes to biological proxy records (Gajewski et al. 1997). Structural information obtained by detailed analysis of varved sediments permitted Lamoureux (2000) to identify rainfallinduced deposition events. However, analysis has increasingly focused on the physical characteristics of the sediments, particularly particle size characteristics. Recently, Francus et al. (2002) demonstrated how grain texture properties obtained with sedimentary image analysis could provide paleoenvironmental information. Their results revealed that grain size measured from backscatter electron micrographs using image analysis techniques could be used as a quantitative proxy for spring snow melt intensity. Additionally, the detailed sedimentary analysis typified by varve studies holds the potential for evaluating the paleoenvironmental significance of aeolian deposits in high latitude lakes (Lamoureux et al. 2002). Mineralogy The composition of sedimentary materials has been an area of considerable research focus. Composition provides information regarding the origin, transport pathways and, in some cases, the processes of deposition. The first relates to mineralogical or lithological indicators that indicate the provenance of source rocks or sediment. By contrast, authigenic minerals (for instance evaporites) provide an indication of past environmental conditions in the lake, particularly the water balance and salinity. A survey of the available literature indicates that authigenic mineral formation appears to constitute a minor component of most high latitude lacustrine sediments. While many high latitude regions are characteristically arid, evaporation is limited due to low energy availability and saline lakes are comparatively rare. Several saline and hypersaline lakes have been described from Antarctica (Adamson and Pickard 1986; Lyons and Mayewski 1993) and in the Canadian High Arctic (Ouellet et al. 1987) and Subarctic (Pienitz et al. 2000), but these investigations largely focused on the limnological aspects and did not consider the role of salinity in the composition of the sedimentary record. One of these studies (Pienitz et al. 2000), from the arid interior of the Yukon Territory, found that evaporate deposits varied substantially through the Holocene, indicating varying degrees of aridity and brine concentration in a saline lake. Although similar studies are rare in high latitudes, this study demonstrates the value of sedimentary mineralogy as a paleoenvironmental proxy where conditions warrant. Mineralogical and geochemical characterization of sedimentary materials has been widely used to establish the provenance of source rocks, particularly using marine sediments (e.g., Darby and Bischof 1996). These studies utilize electron microprobe

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analyses to identify unique trace element geochemical signatures that match reference material from particular source areas. Similar studies utilizing lacustrine sediments have been rare in polar environments (e.g., Coakley and Rust 1968), in part because many lakes investigated drained small catchments with limited mineralogical variability. In a recent study, Lamoureux et al. (2002) investigated the variation in delivery of detrital carbonate minerals in a proglacial lake on Devon Island. Bedrock in this catchment is composed of Precambrian granite and metamorphic rocks are overlain by Paleozoic carbonates. The primary source of meltwater and sediment delivered to the lake originates on the glacierized carbonate plateau; hence, the authors were able to use low frequency variations in carbonate deposition to infer changes in glacial meltwater production through the Holocene. Despite the lack of work in polar environments to date, lithological and mineralogical investigations hold promise for improving our understanding of variability in sediment source areas and delivery which may have relevance to paleoenvironmental studies. Sedimentary texture The texture and grain size properties of lacustrine sediments can provide considerable information regarding past depositional environments. Textural analysis has a long tradition in sedimentology and the reader is directed to the many comprehensive works that describe the methods for sedimentary analysis. A recent review of sedimentary texture by Last (2001) is of particular interest to researchers studying lacustrine sediments. Detailed sedimentary descriptions of arctic and antarctic lacustrine sediments have become increasingly common, particularly with the adoption of petrographic thin sections of embedded sediments. Thin sections permit investigation of the microscopic properties of the sediments using transmitted and polarized light (e.g., Pike and Kemp 1996; Zolitschka 1996; Lamoureux 1999). The thin sections and embedded slabs are also suitable for detailed textural analysis using scanning electron microscopy (SEM), particularly backscattering secondary emission imaging (BSEI) and secondary electron imaging (SEI) (Pike and Kemp 1996; Francus 1998; Ojala and Francus 2002). The SEM imagery allows close examination of the sedimentary textures and composition, as well as the structure and bedding properties of the sediment. All of these acquisition approaches are particularly suited to digital image analysis. Recent developments have provided the means to quantify sedimentary properties including grain size, shape and orientation (Francus 1998). While still not in wide use, these methods have shown their value for detailed sedimentology and paleoenvironmental analysis using high latitude sediments (Francus et al. 2002). In addition to these approaches, high quality x-radiographs obtained from embedded sediment slabs have been used to analyse secondary properties (primarily density) in sedimentary sequences (Ojala and Francus 2002). When used in conjunction with other sedimentological analyses, these approaches reveal the wide range of opportunities available for investigating the textural properties of high latitude lacustrine sediments.

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Sediment biogeochemistry-organic matter Although considerable emphasis has been placed on the subfossil composition of high latitude lacustrine sediments (see other chapters in this volume), the composition of the overall organic matter can also provide a valuable record of past aquatic and terrestrial productivity (e.g., Wolfe and Härtling 1996; Willemse and Tornqvist 1999). The most common approach to assessing organic content of sediments is using one of several measures of total organic carbon (TOC), including loss-on-ignition (LOI) (Dean 1974; Heiri et al. 2001), chemical digestion or spectrometric techniques. Although the results from some of these methods may not always be strictly interpreted as TOC (Dean 1974), they may provide an important indication of the variations in lake biogenic productivity (e.g., Willemse and Tornqvist 1999). A key limitation in these approaches is the lack of differentiation between terrestrial and aquatic organic matter. Moreover, percentage measures (LOI) can be diluted by varying quantities of inorganic sedimentation, and therefore may not accurately reflect productivity. Additional characterization of elemental composition of the organic matter can clarify the origins substantially. Materials with high carbon to nitrogen (C/N) ratios are typical of plants found in terrestrial ecosystems, compared to the lower C/N ratios (< 10) of aquatic algae (Meyers and Takemura 1997). Although further differentiation of organic matter is possible using a variety of approaches, no research from high latitude lakes is available to ascertain how useful these analyses could be. Sediment biogeochemistry- trace element composition Trace element composition of sedimentary materials has been used to infer a variety of limnological and catchment processes. Methods developed to assess the elemental composition of chemical components, or separations, are extensively used in lacustrine sediment studies and the reader is referred to Engstrom and Wright (1984) and Boyle (2001) for a comprehensive treatment of the subject. Application of trace element studies to high latitude lakes has been limited but has provided useful evidence of catchment development and lake isolation during glacioisostatic emergence. In conjunction with other proxy studies, trace element composition is especially effective (e.g., Björck et al. 1993; Verkulich et al. 2002). For example, Wolfe and Härtling (1996) noted that decreasing early Holocene minerogenic influx and burial to several small, poorly buffered lakes on Baffin Island reduced available alkalinity to the lake and resulted in natural lake acidification through the Holocene. In a separate study of three lowland lakes on Devon Island, trace metal and organic matter profiles, particularly the deposition of MoS2, was interpreted to indicate the formation of hypolimnetic anoxia associated with isolation following emergence (Young and King 1989). Although similar studies are rare, these cases do indicate the potential of applying trace element composition of lacustrine sediments for paleoenvironmental and lake ontogeny studies (Verkulich et al. 2002).

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Figure 11. A conceptual linkage between the timescales for major high latitude sedimentary processes with the depositional timing during the season. Note that many processes vary over a range of timescales during the mid-season when available energy is at a maximum and, in particular, when lake ice is receding or absent. For instance, ice rafting of shoreline materials is minimal early in the season due to extensive ice cover. Through progressive melting, the ice pan becomes more mobile, thus, the deposition becomes dependent on wind drifting and break-up at shorter timescales. By comparison, fluvial sediment inputs are highly dependent on hourly and daily snowmelt in the spring, becoming progressively more dependent on longer timescales (multi-day rainfall) as the season progresses.

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Framework for interpreting the environmental significance of physical and biogeochemical sedimentary records in high latitude lakes There are a wide range of potential sedimentary properties for use in paleolimnological studies in high latitude sites. While regional and local conditions in both the catchment and lake are critical constraints for interpreting sedimentary and geochemical proxy records, some broader inferences can be drawn about high latitude lacustrine sedimentary environments. While not comprehensive and subject to local controls, it is potentially useful to consider the temporal scales at which these primarily operate (Figure 11). Conceptually, it is important to differentiate between the timescales of external environmental (e.g., river runoff and sediment transport) from internal lake processes (e.g., duration of ice cover) and depositional processes (e.g., settling rates). It is clear that the strong seasonality of high latitude regions confines the main processes to a relatively short period of the year. However, within the active season, there is temporal convergence of the main processes during the mid-season. In part, this convergence depends on the coincidence of high energy availability for snow and ice melt together with in-lake biological production (Figure 11). Although this model emphasises the importance of ice and snow melt in high latitude environments, the role that lake ice plays as an interface between sediment delivery and deposition in the lake also becomes a significant control over sedimentary processes. Further, when several of these sedimentary components are not relevant for a given lake (e.g., glaciers) it becomes apparent that there is frequently a seasonal progression of many sedimentary inputs. For instance, in a typical nival catchment, sediment flux from the freshet should arrive before the lake ice has effectively decayed. Hence, biological activity will likely follow significant sediment influx (and will also benefit from nival nutrient influxes). Ice rafted and aeolian deposits, where present, require some decay of the lake ice, and therefore occur later in the season. While this seasonal succession may be apparent through detailed sedimentological analysis of well-preserved sediments, this model has implications for all sedimentary records. Where sediment mixing removes structural indicators, apparent cosedimentation of different sedimentary products indicates a particular set of environmental conditions through the entire active deposition season. Therefore, paleoenvironmental interpretations will benefit from a comprehensive and holistic investigation of the sedimentary sequence. Summary Physical and chemical proxies preserved in high latitude lacustrine sediments yield a potential wealth of paleoenvironmental information. The characteristics of the periglacial environment generate several sedimentary deposits unique to these regions, and the sensitivity of many watersheds to hydrometeorological and other forcings further differentiates the high latitude lacustrine record from more temperate examples. Dating and chronology construction for lacustrine sediment records remains problematic, especially in the High Arctic and Antarctica where organic material for 14C dating is rare. This situation is further complicated by long residence times for organic materials in high latitude catchments. Low levels of other useful radioisotopes (210Pb

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and 137Cs) further limit dating recent sediments, particularly where sedimentation rates are highly variable. Future work utilizing paleomagnetic methods may provide an opportunity to date lakes using regional secular variation curves produced on well-dated or varved sequences. Despite the growing number of sedimentary records from these regions, paleoenvironmental interpretations are largely inferential and are not directly supported by field process studies. The rare exceptions to this situation are several comprehensive limnological field studies that have been carried out on polar lakes (e.g., Bradley et al. 1996; Lewis et al. 2002). In some cases these detailed process observations permit developing robust, quantitative relationships between hydroclimatic variables and sedimentation rates (Hardy et al. 1996). Despite the benefit of this approach for paleoenvironmental interpretations, these studies are uncommon, in part due to logistical and financial constraints. There is also further need for researchers to link paleoenvironmental change with the geomorphic and geochemical development of the catchment. This approach is necessary to understand the role of long-term landscape development on the sedimentary record. Several studies have demonstrated that dynamic arctic environments undergo major transformations over century-to-millennial timescales. For example in glaciated regions, emergence from postglacial marine transgressions can dramatically alter sediment availability and delivery processes (e.g., Lamoureux 1999), while changing sediment availability and chemical weathering impact water quality conditions in lakes with concomitant impacts on biota (Wolfe and Härtling 1996). While coordinated, multidisciplinary proxy studies cannot be expected in all cases, the wide range of proxy information contained in the physical and chemical properties of lacustrine sediments suggests that considerable paleoenvironmental information can be obtained from the physical and chemical information contained in high latitude lacustrine sediments. Acknowledgements B. Zolitschka and the editors provided valuable comments that improved the chapter. References Abbott M.B. and Stafford Jr. T.W. 1996. Radiocarbon geochemistry of modern and ancient arctic lake systems, Baffin Island, Canada. Quat. Res. 45: 300-311. Adamson D.A. and Pickard J. 1986. Cenozoic history of the Vestfold Hills. In: Pickard J. (ed.), Antarctic Oasis: Terrestrial Environments and History of the Vestfold Hills. Academic Press, Sydney, pp. 63-93.

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Gilbert R. 1990. Rafting in glacimarine environments. In: Dowdeswell J.A. and Scourse J.D. (eds), Glacimarine Environments: Processes and Sediments. Geological Society Special Publication No. 53, London, pp. 105-120. Gilbert R. 1999. Subaquatic acoustic techniques: In: Gilbert R. (ed.), A Handbook of Geophysical Techniques for Geomorphic and Environmental Research, Geological Survey of Canada Open File 3731, pp. 103-125. Gilbert R. 2000. Environmental assessment from the sedimentary record of high latitude fiords. Geomorph. 32: 295-314. Gilbert R. and McKenna Neuman C. 1988. Occurrence and potential significance of warm weather during winter in the eastern Canadian Arctic. Arct. Alp. Res. 20: 395-403. Gilbert R., Syvitski J.P.M. and Taylor R.B. 1985. Reconnaissance study of proglacial Stewart Lakes, Baffin Island District of Franklin. Geological Survey of Canada Paper 85-1A, Ottawa, pp. 505-510. Gilbert R., Desloges J.R. and Clague J.J. 1997. The glacilacustrine sedimentary environment of Bowser Lake in the northern Coast Mountains of British Columbia, Canada. J. Paleolim. 17: 331-346. Gravenor C.P., Von Brunn V. and Dreimanis A. 1984. Nature and classification of waterlain glaciogenic sediments, exemplified by Pleistocene, Late Paleozoic and Late Precambrian deposits. Earth Sci. Rev. 20:105-166. Håkanson J. and Jansson C. 1983. Principles of Lake Sedimentology, Springer Verlag, Berlin, Heidelberg, New York, Tokyo, 316 pp. Hamblin P.F. and Carmack E.C. 1978. River-induced currents in a fjord lake: J. Geophys. Res. 83: 885-899. Hardy D.R., Bradley R.S. and Zolitschka B. 1996. The climatic signal in varved sediments from Lake C2, northern Ellesmere Island, Canada. J. Paleolim. 16: 227-238. Harris C. and Lewkowicz A.G. 1993. Form and internal structure of active-layer detachment slides, Fosheim Peninsula, Ellesmere Island, Northwest Territories, Canada. Can. J. Earth Sci. 30: 1708-1714. Harris C. and Lewkowicz A.G. 2000. An analysis of the stability of thawing slopes, Ellesmere Island, Nunavut, Canada. Can. Geotech. J. 37: 465-478. Hasholt B., Walling D.E. and Owens P.N. 2000. Sedimentation in arctic proglacial lakes: Mittivakkat Glacier, south-east Greenland. Hydrol. Proc. 14: 679-699. Hay A.E. 1984. Remote acoustic imaging of the plume from a submarine spring in an arctic fjord. Science 225: 1154-1156. Heiri O., Lotter A.F. and Lemke G. 2001. Loss on ignition as a method for estimating organic and carbonate content in sediments: reproducibility and comparability of results. J. Paleolim. 25: 101-110. Herrmanson M.H. 1990. 210Pb and 137Cs chronology of sediments from small, shallow Arctic lakes. Geoch. Cosmoch. Acta 54: 1443-1451. Hill P.R., Simard A. and Héquette A. 1999. High-resolution seismic stratigraphy of late Quaternary deposits in Manitounuk Sound, northern Québec: effects of rapid postglacial emergence. Can. J. Earth Sci. 36: 549-563. Holdsworth G. 1973. Ice calving into the proglacial Generator Lake, Baffin Island, N.W.T., Canada. J. Glaciol. 12: 235-250. Hughen K.A., Overpeck J.T. and Anderson R.F. 2000. Recent warming in a 500-year palaeotemperature record from varved sediments, Upper Soper Lake, Baffin Island, Canada. The Holocene 10: 9-19. Intergovernmental Panel on Climate Change 2001. Climate Change 2001: Impacts, Adaptation and Vulnerability. Contribution of Working Group II to the Third Assessment Report of the Intergovernmental Panel on Climate Change WMO-UNEP, Cambridge University Press. Jopling A.V. and McDonald B.C. (eds) 1975. Glaciofluvial and Glaciolacustrine Sedimentation, SEPM Special Publication No. 23, 320 pp.

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King J. and Peck J. 2001. Use of paleomagnetism in studies of lake sediments. In: Last W.M. and Smol J.P. (eds.), Developments in Paleoenvironmental Research (DPER) Volume 1-Tracking Environmental Change Using Lake Sediments: Basin Analysis, Coring and Chronological Techniques, Kluwer, Dordrecht, The Netherlands, pp. 371-390. Krishnappan B.G. 1975. Dispersion of granular material dumped in deep water. Inland Waters Directorate, Scientific Series 55, 14 pp. Lamoureux S.F. 1999. Catchment and lake controls over the formation of varves in monomictic Nicolay Lake, Cornwall Island, Nunavut. Can. J. Earth Sci. 36: 1533-1546. Lamoureux S.F. 2000. Five centuries of interannual sediment yield and rainfall-induced erosion in the Canadian High Arctic recorded in lacustrine varves. Wat. Resources Res. 36: 309-318. Lamoureux S.F. 2001. Varve chronology techniques. In: Last W.M. and Smol J.P. (eds), Developments in Paleoenvironmental Research (DPER) Volume 1-Tracking Environmental Change Using Lake Sediments: Basin Analysis, Coring and Chronological Techniques, Kluwer, Dordrecht, The Netherlands, pp. 247-260. Lamoureux S.F. and Bradley R.S. 1996. A late Holocene varved sediment record of environmental change from northern Ellesmere Island, Canada. J. Paleolim. 16: 239-255. Lamoureux S.F., Gilbert R. and Lewis T. 2002. Lacustrine sedimentary environments in High Arctic proglacial Bear Lake, Devon Island, Nunavut. Arct. Ant. Alp. Res. 34: 130-141. Larsen C.P.S., Pienitz R., Smol J.P., Moser K.A., Cumming B.F., Blais J.M., MacDonald G.M. and Hall R.I. 1998. Relations between lake morphometry and the presence of laminated lake sediments: a re-examination of Larsen and MacDonald (1993). Quat. Sci. Rev. 17: 711-717. Last W.M. 2001. Textural analysis of lake sediments. In: Last W.M. and Smol J.P. (eds), Developments in Paleoenvironmental Research (DPER) Volume 2-Tracking Environmental Change Using Lake Sediments: Physical and Chemical Techniques, Kluwer, Dordrecht, The Netherlands, pp. 41-81. Last W.M. and Smol J.P. (eds) 2001a. Developments in Paleoenvironmental Research (DPER) Volume 1-Tracking Environmental Change Using Lake Sediments: Basin Analysis, Coring and Chronological Techniques, Kluwer, Dordrecht, The Netherlands, 548 pp. Last W.M. and Smol J.P. (eds) 2001b. Developments in Paleoenvironmental Research (DPER) Volume 2-Tracking Environmental Change Using Lake Sediments: Physical and Geochemical Techniques, Kluwer, Dordrecht, The Netherlands, 504 pp. Lawson D.E. 1993. Glaciohydrologic and glaciohydraulic effects on runoff and sediment yield in glacierized basins. Cold Regions Research and Engineering Laboratory (CRREL), U.S. Army Corps of Engineers, Monograph 93-2, 108 pp. Lemmen D.S., Gilbert R., Smol J.P. and Hall R.I. 1988. Holocene sedimentation in glacial Tasikutaaq Lake. Can. J. Earth Sci. 25: 810-823. Lewis T., Gilbert R. and Lamoureux S.F. 2002. Spatial and temporal changes in sedimentary processes at high-arctic proglacial Bear Lake, Devon Island, Nunavut, Canada. Arct. Ant. Alp. Res. 34: 119-129. Luckman B. 1975. Drop stones resulting from snow-avalanche deposition on lake ice. J. Glaciol. 14: 186-188. Lyons W.B. and Mayewski P.A. 1993. The geochemical evolution of terrestrial waters in the Antarctic: the role of rock-water interactions. In: Physical and Biogeochemical Processes in Antarctic Lakes. Antarctic Research Series. American Geophysical Union, pp. 135-146. Mangerud J. and Svendsen J.I. 1990. Deglaciation chronology inferred from marine sediments in a proglacial lake basin, western Spitsbergen, Svalbard. Boreas 19: 249-272. Matula S. 1992. Bridging the gap - creating nearshore bathymetric maps from multibeam swath sonar systems and conventional hydrographic data. U.S. Geological Survey, Circular 1092, pp. 118-126. McKenna Neuman C. 1990. Observations of winter aeolian transport and niveo-aeolian deposition at Crater Lake, Pangnirtung Pass, N.W.T., Canada. Perma. Perigl. Proc. 1: 235-247.

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McKenna Neuman C. 1993. A review of aeolian transport processes in cold environments. Prog. Phys. Geog. 17: 137-155. McLaren I.A. 1967. Physical and chemical characteristics of Ogac Lake, a landlocked fiord on Baffin Island. J. Fish. Res. Board Can. 24: 981-1015. Meyers P.A. and Takemura K. 1997. Quaternary changes in delivery and accumulation of organic matter in sediments in Lake Biwa, Japan. J. Paleolim. 18: 211-218. Moore J.J., Hughen K.A., Miller G.H. and Overpeck J.T. 2001. Little Ice Age recorded in summer temperature reconstruction from varved sediments of Donard Lake, Baffin Island, Canada. J. Paleolim. 25: 503-517. Moorman B.J. and Michel F.A. 1997. Bathymetric mapping and sub-bottom profiling through lake ice with ground penetrating radar. J. Paleolim. 18: 61-73. Nedell S.S., Anderson D.W., Squyres S.W. and Love F.G. 1987. Sedimentation in ice-covered Lake Hoare, Antarctica. Sedimentology 34: 1093-1106. Noon P.E., Birks H.J.B., Jones V.J. and Ellis-Evans J.C. 2001. Quantitative models for reconstructing catchment ice-extent using physical-chemical characteristics of lake sediments. J. Paleolim. 25: 375-392. Ojala A.E.K. and Saarnisto M. 1999. Comparative varve counting and magnetic properties of the 8400-yr sequence of an annually laminated sediment in Lake Valkiajärvi, central Finland. J. Paleolim. 22: 335-348. Ojala A.E.K. and Francus P. 2002. Comparing X-ray densitometry and BSE-image analysis of thin section in varved sediments. Boreas 31: 57-64. Østrem G. and Olsen H.C. 1987. Sedimentation in a glacier lake. Geog. Ann. 69A: 123-128. Ouellet M., Bisson M., Pagé P. and Dickman M. 1987. Physiochemical limnology of meromictic saline Lake Sophia, Canadian Arctic Archipelago. Arct. Alp. Res. 19: 305-315. Overpeck J., Hughen K.A., Hardy D., Bradley R., Case R., Douglas M., Finney B., Gajewski K., Jacoby G., Jennings A., Lamoureux S., Lasca A., MacDonald G., Moore J., Retelle M., Smith S., Wolfe A. and Zielinski G. 1997. Arctic environmental change of the last four centuries. Science 278: 1251-1256. Pienitz R., Smol J.P., Last W.M., Leavitt P.R. and Cumming B.F. 2000. Multi-proxy Holocene palaeoclimatic record from a saline lake in the Canadian Subarctic. The Holocene 10: 673-686. Pike J. and Kemp A.E.S. 1996. Preparation and analysis techniques for studies of laminated sediments. Geol. Soc. Spec. Publication no. 116, London, pp. 37-48. Retelle M.J. 1986. Stratigraphy and sedimentology of coastal lacustrine basins, northeastern Ellesmere Island, N.W.T. Géogr. phys. Quat. 40: 117-128. Retelle M.J. and Child J.K. 1996. Suspended sediment transport and deposition in a high arctic meromictic lake. J. Paleolim. 16: 151-167. Retelle M.J., Lamoureux S.F., Phelps K.J., Robertson A.D., Smith S., Bradley R.S. and Zolitschka B. 1996. Contrasting styles of laminated sediment deposition in coastal meromictic lakes, Canadian Arctic Archipelago; implications for paleoclimatic reconstruction. Geological Society of America, Proceedings from the 28th annual meeting, pp. 58. Rothlisberger H. and Lang H. 1987. Glacial hydrology. In: Gurnell A.M. and Clark M.J. (eds), Glacio-fluvial Sediment Transfer, John Wiley and Sons, New York. pp. 207-284. Saarinen T. 1998. High-resolution palaeosecular variation in northern Europe during the last 3200 years. Phys. Earth Planet. Int. 106: 299-309. Saulnier-Talbot É. and Pienitz R. 2001. Isolation au postglaciaire d’un bassin côtier près de Kuujjuarapik-Whapmagoostui, en Hudsonie (Québec): une analyse biostratigraphique diatomifère. Géogr. phys. Quat. 55: 63-74. Simmons Jr. G.M., Wharton Jr. R.A., McKay C.P., Nedell S. and Clow G. 1986. Sand/ice interactions and sediment deposition in perennially ice-covered Antarctic lakes. Antarctic Journal of the United States 21: 217-220.

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Smith I.R. 2000. Dimictic sediments within high Arctic lake sediment cores: evidence for lake ice rafting along the lateral glacial margin. Sedimentology 47: 1157-1179. Smith N.D. 1978. Sedimentation processes and patterns in a glacier-fed lake with low sediment input. Can. J. Earth Sci. 15: 741-756. Smith N.D. and Syvitski J.P.M. 1982. Sedimentation in a glacier-fed lake: the role of pelletization on deposition of fine-grained suspensates. J. Sed. Petrol. 52: 503-513. Smith N.D. and Ashley G.M. 1985. Proglacial lacustrine environment. In Ashley G.M., Shaw J. and Smith N.D. (eds), Glacial Sedimentary Environments, SEPM Short Course no. 16, pp. 135-215. Smith N.D., Vendl M.A. and Kennedy S.K. 1982. Comparison of sedimentation regimes in four glacier-fed lakes of western Alberta. In: Davidson-Arnott R., Nickling W. and Fahey B.D. (eds), Glacial, Glaciofluvial, and Glaciolacustrine Systems, Geobooks, Norwich, pp. 203-238. Smith S. 1997. A record of environmental change derived from varved lakes along the margin of the Agassiz Ice Cap, Northwest Territories, Canada. M.Sc. Thesis, University of Massachusetts, Amherst, 101 pp. Smol J.P. 1988. Paleoclimate proxy data from freshwater arctic diatoms. Verh. Internat. Verein. Limnol. 23: 837-844. Snowball I., Sandgren P., Petterson G., Petrovsky E.E., Tunyi I.E. and Orlicky O.E. 1998. Secular variations and relative palaeointensity of the Earth's magnetic field between 7,000 BC and 500 AD recorded by annually laminated lake sediments in northern Sweden. Proceedings from new trends in geomagnetism paleo, rock and environmental magnetism, 6th biennial castle meeting, 49, pp. 207-208. Snyder J.A., Werner A. and Miller G.H. 2000. Holocene cirque glacier activity in western Spitsbergen, Svalbard; sediment records from proglacial Linnevatnet. The Holocene 10: 555-563. Squyres S.W., Andersen D.W., Nedell S.S. and Wharton Jr. R.A. 1991. Lake Hoare, Antarctica: sedimentation through a thick perennial ice cover. Sedimentology 38: 363-379. Svendsen J.I. and Mangerud J. 1997. Holocene glacial and climatic variations on Spitsbergen, Svalbard. The Holocene 7: 45-57. Svendsen J.I., Mangerud J. and Miller G.H. 1989. Denudation rates in the Arctic estimated from lake sediments on Spitsbergen, Svalbard. Palaeogeogr. Palaeoclim. Palaeoecol. 76: 153-168. Verkulich S.R., Melles M., Hubberten H.-W. and Pushina Z.V. 2002. Holocene environmental changes and development of Figurnoye Lake in the southern Bunger Hills, East Antarctica. J. Paleolim. 28: 253-267. Wagner B., Melles M., Hahne J., Niessen F. and Hubberten H.-W. 2000. Holocene climate history of Geographical Society Ø, East Greenland; evidence from lake sediments. Palaeogeogr. Palaeoclim. Palaeoecol. 160: 45-68. Willemse N.W. and Tornqvist T.E. 1999. Holocene century-scale temperature variability from West Greenland lake records. Geology 27: 580-584. Wolfe A.P. and Härtling J.W. 1996. The late Quaternary development of three ancient tarns on southwestern Cumberland Peninsula, Baffin Island, Arctic Canada: paleolimnological evidence from diatoms and sedimentary chemistry. J. Paleolim. 15: 1-18. Woo M.K. 1993. Northern hydrology. In: French H.M. and Slaymaker H.O. (eds), Canada's Cold Environments, McGill-Queen's University Press, Montréal, pp. 117-142. Young R.B. and King R.H. 1989. Sediment chemistry and diatom stratigraphy of two high arctic isolation lakes, Truelove Lowland, Devon Island, N.W.T., Canada. J. Paleolim. 2: 207-225. Zolitschka B. 1996. Recent sedimentation in a high arctic lake, northern Ellesmere Island, Canada. J. Paleolim. 16: 169-186.

4. PALYNOLOGY OF NORTH AMERICAN ARCTIC LAKES

KONRAD GAJEWSKI ([email protected]) Laboratory of Paleoclimatology and Climatology Department of Geography University of Ottawa Ottawa, Ontario K1N 6N5, Canada and GLEN M. MACDONALD ([email protected]) Departments of Geography and Organismic Biology, Ecology and Evolution University of California Los Angeles, California 90095-1524, USA

Key words: Arctic, Canada, Palynology, Paleoecology, Holocene, Treeline, Tundra, Picea, Lake sediments, Pollen

Introduction Recent syntheses of late Quaternary pollen diagrams from many regions of the world illustrate the significant progress made in reconstructing past vegetation and climatic change in recent decades (Wright et al. 1993; TEMPO 1996). However, the Arctic in general, and the Canadian Arctic in particular (Figure 1), remain with relatively few pollen diagrams (Gajewski et al. 1995; Tarasov et al. 1998, 2000; Edwards et al. 2000a; CAPE 2001). The region best known from the point of view of palynology is Baffin Island, where several pollen studies have reported the postglacial vegetation history (discussed below). The middle and high Arctic are particularly under-studied, while at treeline more data are available. Many of these are based upon the analysis of lake sediments, sometimes in conjunction with analysis of other biological and chemical limnological evidence (e.g., MacDonald et al. 1993; Ponader et al. 2002). As a result, a coherent picture is emerging of the postglacial history of the treeline zone. 89 R. Pienitz, M.S.V. Douglas and J.P. Smol (eds), 2004. Long-term Environmental Change in Arctic and Antarctic Lakes. Springer. Printed in the Netherlands.

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In this review we will discuss pollen analysis in arctic environments and present some of the results from the treeline and arctic regions of North America. We will discuss separately arctic and treeline studies, since the nature of the sediment, the pollen production and accumulation and, of course, the pollen assemblages are quite different. Although standard methods and approaches can be used at treeline sites as are used in temperate regions, in the Arctic pollen analysis is more time consuming and processes of production, transport, and deposition are not as well understood as they are in forested areas. Interestingly, several pollen profiles have been prepared from ice cores, listed in Bourgeois et al. (2000). We will not be explicitly reviewing these studies, although some of the conclusions are relevant to the interpretation of lacustrine records. We will be discussing primarily North American examples and literature. There has been a significant research effort in the past few years in Greenland and in the Canadian and Alaskan arctic and boreal regions, and a review of this literature is quite pertinent. A discussion of recent palynological research in arctic Eurasia appears in MacDonald et al. (this volume). Nichols (1974) summarized older literature on North American arctic and treeline paleoecology. Prior to the 1970s, only scattered studies had been completed and the postglacial history of most arctic regions was unknown. One consequence of the lack of data was that conclusions from one area were extrapolated to others, with the result that regional differences were misunderstood. Many earlier studies examined peat deposits (Nichols 1974), which are frequently short and discontinuous. Local pollen types from plants growing on the site are over-represented (Ritchie 1974). Arctic palynology “came of age” with the publication by Fredskild (1973) of a large number of pollen diagrams from around Greenland. Ritchie (1985) and Ager and Brubaker (1985) later summarized the available vegetation records from western Canada and Alaska, respectively. Recent synthetic studies (e.g., Tarasov et al. 1998, 2000; Edwards et al. 2000a; Gajewski et al. 2000b) have collated arctic pollen records from particular time-slices in order to test paleoclimate and paleovegetation models. Synthesis of methodological aspects Field methods Collecting lake sediment cores in the Arctic remains a challenge (Nichols 1974; Ritchie 1984) that usually requires air support. Although access to charter aircraft has increased in recent years, it is rare to have exclusive access to helicopter or airplane support for the length of time that is needed for coring. It may be necessary to be left near a coring site by a floatplane or Twin Otter, and the remaining work performed on foot. On occasion canoes have been used to transport coring equipment and camps along arctic rivers from landing sites to coring lakes. Setting up small camps has the advantage of permitting the collection of several cores and surface samples in a small area, but the disadvantage of the extra time lost in camping. In most of the Arctic, scientific activity typically occurs in summer. Spring fieldwork may be possible, but if it is too cold, coring becomes difficult if not impossible, as a Livingstone piston corer can freeze immediately upon leaving the water. It is far better to work in late spring or early summer, when temperatures are warmer but when there is

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still an ice platform. However, when planning the fieldwork, the chronology of snowmelt needs to be taken into account. After the snow melts from the land surface, but before the moat on the lake becomes too wide, is the ideal time to core. However, the meltwater from the land surface may accumulate as a layer of water on the lake. At this time, there is still solid ice under the water, but the several tens of centimetres of snow and water on the lake makes coring impossible. This water will soon drain into cracks in the ice; some 20-30 cm of water have been observed to drain in a couple of days on a lake in the Boothia Peninsula. As the season continues, the moat enlarges, although ice thickness will still exceed one metre. Typically the ice will be attached to some points on the shore, affording access across the moat. As a complete moat forms, the ice starts to candle. At this point, access is still possible using waders or a small boat, although extreme caution must be used, as the ice near the edge tends to be thin and easily broken under the weight of a person. Contrary to popular opinion, it is possible to drill a hole in candled, wet and floating ice, if it is thick enough. The ice continues to melt for a period of weeks, and this floating ice still affords a platform for coring, although boats must be used to gain access to the ice pack. The ideal time for coring is thus during the spring, in the window of time after the snow has melted from the landscape, but when there is still ice on the lakes. This window is difficult to find as it varies across the Arctic and between years. A strong ice auger with extensions and back-up parts is needed, as are survival suits. In all cases, caution and all available safety precautions must be taken.

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Figure 1. Location of sites mentioned in the text. Location of treeline and major vegetation zones is indicated by solid lines.

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Coring can also be conducted in summer using a platform on small boats. Three small inflatable dinghies with a platform consisting of a piece of plywood with a hole in the centre and strengthened by three wooden beams is a minimum requirement, making for a small but functional coring platform. Two dinghies support the coring platform and a third holds equipment. The platform is attached to the dinghies using nylon rope. An extra pump and patch kit are essential. In small lakes, the platform can be attached to the shore using nylon rope (available in hardware stores). This rope should be kept on a wooden spool, as the cardboard ones the rope is supplied with will disintegrate during the first use. Three fastening points on the shore are recommended. If this is not possible because the lake is too large, anchors can be made using burlap bags filled with small rocks, and easily tied and untied onto ropes. One exception to the general field strategies listed above pertains to coring near the treeline zone in the central and western Canadian Arctic. Regional centres such as the towns of Churchill, Yellowknife and Inuvik provide bases of operation that include the availability of helicopters for coring of lakes in the early spring. Collecting the cores is done in the same way as elsewhere, although a drive rig is needed in arctic sediments. Special soil sampling drivers are effective, but heavy; however, less expensive and lighter alternatives can be fabricated. A standard Livingstone corer (Wright et al. 1984) works well in small shallow lakes (< 15 m depth). A Russian sampler (Jowsey 1966) may also be used to obtain cores once a sediment depth is reached where the sediment is sufficiently compact. “Crust-freeze” corers (Renberg 1981) have been used on occasion to obtain cores from lakes with laminated sediments, but this is rare due to the difficulty of transporting dry ice into remote arctic sites. When a Livingstone corer is used, the first metre (called Drive 0) is collected using a plastic tube of the same diameter as the Livingstone corer, and taped to the same drive rods. A spare piston and wire that is used only for this core is recommended. External casing is next lowered to the sediment surface, taking care to not penetrate below the bottom of the Drive 0. The casing is standard sewer pipe and keeps the coring rods from bending or breaking in stiff sediment. A complete sequence is obtained, and a second sequence collected a couple of metres away to fill in gaps obtained from the first core. In arctic lakes, magnetic susceptibility provides an easy and cheap way to correlate the two sequences. The cores can be shipped back using black sewer pipe split lengthwise and taped together; the 5 cm size fits exactly a standard Livingstone corer. Data sets of pollen counts from lake sediment surface samples are important for interpreting paleo-records, and surface sampling is often done in conjunction with lake sediment coring. Most surface sampling studies have been done in the summer using helicopters on floats, small floatplanes or small boats. Surface sediment samples have been obtained using a number of devices, most commonly Ekman dredges (Wetzel and Likens 2000) or Glew corers (Glew 1991). Laboratory methods Methodology for pollen analysis is standardized (Faegri and Iversen 1975). However, attempts to extract pollen from arctic sediments, where the organic matter content is low and pollen concentrations are small, initially proved unsuccessful. An important

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advance occurred with the suggestion of Cwynar et al. (1979) to use disposable microsieves to concentrate the pollen. This has permitted the analysis of sediment sequences previously considered barren. Other methods may prove useful also, including heavy liquid separation of pollen grains (e.g., Nakagawa et al. 1998) or microwave digestion (Jones and Ellin 1998). Interpretation of arctic pollen Considerations in the analysis and interpretation of pollen assemblages from the Arctic were discussed in Gajewski et al. (1995). The basis for the interpretation of fossil spectra is a sufficiently dense network of pollen extracted from surface sediments, discussed below. However, as is the case with pollen records from lake cores, many areas of the Arctic are poorly represented in modern pollen deposition data sets. Pollen concentrations in arctic sediments are low, in some cases making extraction and counting of pollen assemblages time consuming, if not impossible. However, these low concentrations also seem to be related to very low vegetation density on the landscape. There is an abrupt decrease in pollen concentrations across the middle- to high-arctic boundary (Gajewski 1995), and there has been a long-term decrease in concentrations in some arctic cores (Hyvärinen 1985; Gajewski 1995). Thus, pollen concentrations, and influx if it can be reliably computed, may provide information about vegetation density. The presence of pollen transported into the Arctic from forested regions has long been noted (e.g., Ritchie and Lichti-Federovich 1967; Christie and Ritchie 1969; Ritchie 1974; Nichols et al. 1978; Barry et al. 1981; Fredskild 1973, 1984, 1985b; Hjelmroos and Franzén 1994; Gajewski et al. 1995). Pollen assemblages north of treeline in both North America and Eurasia contain significant amounts of tree pollen. The amount of conifer tree pollen can often be greater than 20% and up to 50% of the pollen rain, particularly in low arctic sites (Figure 2) but also in the High Arctic (Figure 3) (e.g., MacDonald and Ritchie 1986; Andersen et al. 1991; Gajewski 1991, 1995, 2002; Hansen et al. 1996; Gervais and MacDonald 2001). Because pollen are mostly transported in the atmosphere, they can be found in considerable quantities far from the source plant. This makes it more difficult to document the local presence of the taxon producing the pollen, but it is also a source of information about pollen transport. Nichols et al. (1978) realized that this potentially is a significant source of information about air mass movements, however, this has also been a source of controversy (Fredskild 1973, 1985b; Kelly and Funder 1974; Barry et al. 1981) and has not been systematically studied. At this point, it is difficult to make sense of the lake sediment record of long-distance transport due to the lack of sufficient sites. However, recent work on pollen in snow and ice shows a considerable coherence in time and space of pollen assemblages (Bourgeois et al. 1985, 2000, 2001; Bourgeois 1990, 2000). Two scenarios seem possible. First, the pollen are transported northwards in air masses, either during the excursions of warm air to the Arctic during the course of the summer, or in occasional exceptional air mass movements associated with cyclone passages to the north (Hjelmroos and Franzén 1994). An alternative seems to be that pollen are incorporated into the arctic circulation and essentially load the air, to fall out slowly and evenly over the Arctic. This requires further study.

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Figure 2. Modern pollen assemblages and fossil stomata from a transect of lake sediment surface samples that cross the modern treeline in central Canada (after Hansen et al. 1996).

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Pollen of Tertiary age, probably along with other microfossils, can be redeposited in modern and Holocene sediments. This occurs everywhere where these deposits are exposed, but constitutes a consideration in the Arctic due to the low amounts of local pollen that are produced. A further complication is that it is not clear if these can be, in all cases, reliably distinguished from recent pollen grains. However, these may be an index of aeolian transport or erosion (cf. Ovenden 1988). A series of sites would be needed to document if this is feasible. Recently, palynologists working along the arctic treeline zone have turned to the analysis of conifer stomata to supplement the information obtained from pollen records. Stomate remains are obtained from sediment samples during normal pollen preparation and can be easily identified and counted at the same time the pollen is analysed (Hansen 1995). Studies of treeline lake surface samples in North America and Eurasia (Clayden et al. 1996; Hansen et al. 1996; Gervais and MacDonald 2001; Pisaric et al. 2001) have shown that the stomata of boreal genera, such as Pinus, Picea and Larix, are generally present in the sediments of lakes from the forest and forest-tundra and consistently absent from tundra lakes (Figure 2). Select North American studies Modern pollen deposition A first step in any paleoenvironmental study is the analysis of modern sediment samples for use in calibration and interpretation of the fossil assemblages. Fortunately, due to 40 years of sampling by a generation of palynologists, we are in a strong position to begin quantitative comparisons of fossil spectra to the modern pollen record. From the boreal regions and low arctic tundra, there are many surface samples now available. Several major studies had as their goal the accumulation and analysis of modern pollen data from lakes in northern North America, including Lichti-Federovich and Ritchie (1968), Ritchie (1974), Davis and Webb (1975), Birks (1977), Lamb (1984), Anderson and Brubaker (1986), MacDonald and Ritchie (1986), Gajewski (1991), Hansen et al. (1996), and Oswald et al. (2003), among others. Similar studies in Europe and Asia can also lend insight into pollen representation at treeline (e.g., Prentice 1978; Clayden et al. 1996; Gervais and MacDonald 2001; Pisaric et al. 2001). Anderson et al. (1991) analysed the modern pollen and range limits of important treeline species using response surfaces, illustrating the close relation between pollen deposition and the distribution of the major tree species in the boreal zone, if transport issues are accounted for. The relation between climate and tree species distribution has been quantified by Thompson et al. (1999) for all trees of North America. Even in regional ordinations, however, the basic vegetation zonation is reflected by pollen assemblages (e.g., Lamb 1984; MacDonald and Ritchie 1986; Gajewski 1991). From the Canadian Arctic Islands there are still few surface samples available (Figure 3). Aside from core tops, modern pollen data have been published from Banks Island (Ritchie et al. 1987), Somerset Island (Gajewski 1995), Ellesmere Island (Gajewski et al. 1995), and the central Arctic (Gajewski 2002). Samples are available from around Greenland in various papers listed below. A recent synthesis (Gajewski 2002) indicates differences in the pollen assemblages among the various regions of the

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Arctic (Figure 4) that are large enough to be resolvable using multivariate methods. This was also found in a more widespread array of snow samples (Bourgeois et al. 2001). Thus, in spite of a significant portion of the pollen grains originating from outside the islands, the pollen faithfully record the different tundra zones of the Arctic, suggesting the possibility of quantitative vegetation reconstruction.

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Figure 3. Percentages of Picea and Pinus pollen in modern pollen samples of lakes from the Canadian Arctic (from Gajewski 2002).

Case study: Greenland Over 30 pollen diagrams from Greenland have been completed, with the result that the postglacial vegetation and climate history is relatively well known for this geographic region (e.g., Fredskild 1973, 1977, 1983a,b, 1985a,b, 1992; Kelly and Funder 1974; Bick 1978; Funder 1978; Björck and Persson 1981; Funder and Abrahamsen 1988; Björck et al. 1994; Eisner et al. 1995). Greenland covers a wide latitudinal range, and the vegetation varies from a rich shrub tundra in the south to polar desert in the north. There is also a steep climatic gradient from the coast inland to the ice cap, with warmer and drier conditions inland. The detailed pollen identifications permit inferences to be made about the nature of the vegetation. A species-rich herb-dominated zone, usually interpreted as pioneer vegetation, occurred immediately after local deglaciation (Fredskild 1973, 1983b, 1985a,b; Kelly and Funder 1974; Funder 1978; Björck and Persson 1981; Funder and Abrahamsen 1988). Pollen concentrations tended to be low, but exotic pollen (transported from Europe or North America) was relatively high in some sites. Various heaths (Ericaceae) increased in abundance, and Salix immigrated between 9000 and 5000 yr BP, depending on the region. Betula nana then arrived, probably from Europe, increasing rapidly in some sites, whereas B. gl andulosa arrived from North America somewhat later (Fredskild 1985b). Alnus pollen then increased in some southern sites,

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but local arrival is hard to determine due to efficient long-distance transport of the pollen grains (Kelly and Funder 1974; Fredskild 1985b). There was a deterioration of the climate in the past several thousand years, causing shrub pollen to decrease and herbaceous pollen to increase. In north Greenland, the cooling temperatures in the past several thousand years caused fjords and some lakes to become permanently frozen (Fredskild 1973; Funder and Abrahamsen 1988). -2

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Case study: the Canadian Arctic islands Pollen cores have been analysed from Baffin Island (Short et al. 1985, 1994; Jacobs et al. 1997), Banks Island (Gajewski et al. 2000a), Somerset Island (Gajewski 1995), Prince of Wales Island (Gajewski and Frappier 2001), and Ellesmere Island (Hyvärinen 1985) (Figure 1). The presence of carbonate bedrock, low organic matter and redeposited Tertiary sediments and coals frequently causes radiocarbon dates to be too old. From Banks Island in the western Archipelago, four pollen diagrams have been reported (Gajewski et al. 2000a). Only two are shown in Figure 5 (74MS11, 74MS12) as the others are similar. Although much of Banks Island was not glaciated during the Wisconsinan, none of the sequences predates the Holocene and all sequences began at a comparable time. These diagrams show warmest conditions between 7000 and 2000 yr BP (all ages are presented here as radiocarbon years Before Present, with Present taken to be AD 1950) with the timing varying depending on the site (Figure 5). The vegetation around the sites differed in some key aspects during the early and late Holocene cooler periods. For example, Dryas pollen was relatively abundant in the late Holocene but not in the early Holocene, although the reasons for this are not clear. Poaceae were more abundant during the cooler periods, and were replaced by Cyperaceae during warmer conditions in the mid-Holocene. From Baffin Island, two pollen diagrams are available from the eastern coastal region of the island (Patricia Bay, Iglutalik Lake in Figure 5; Short et al. 1985). These studies indicate regional and temporal variation in the vegetation history. A late Holocene cooling during the past 3000 years is interpreted from these records. Analyses from Robinson Lake on southern Baffin Island include interglacial as well as postglacial sediments. The pollen analysis shows a pioneer stage with high Poaceae pollen (Miller et al. 1999), followed by high shrub pollen in the mid-Holocene. A sediment core from Burwash Bay of Nettilling Lake (southern Baffin) spanned the past 4700 years (Jacobs et al. 1997). There were few changes recorded in the pollen assemblages, suggesting relatively stable vegetation in this region. One pollen sequence from Baffin Island apparently contains a long sequence of sediment extending well into the last glacial period, although with some potential hiatuses (Wolfe et al. 2000). Sedimentation was low during glacial periods, with a resultant high pollen concentration. Maximum percentages of low-arctic shrub pollen were recorded in interglacial and early Holocene sediments, while the maximum of typical middle Arctic pollen types was found in the mid-Holocene. From northern Baffin Island, a lake sediment core (Fish Lake; Figure 6) and some peat sections yielded pollen and other fossils (Short et al. 1994). The vegetation was sparse prior to 7000 yr BP, interpreted as resulting from cold conditions. There was a further warming around 3500 yr BP and a subsequent cooling since 2000 yr BP. The different proxy records show similar environmental conditions, however the dating is problematic and the pollen sums are low. A number of peat deposits of various ages have also been analysed from around Baffin Island, but are not discussed here. Only one pollen diagram, from Baird Inlet on Ellesmere Island, has reported the postglacial vegetation succession from the northeastern part of the Arctic Archipelago (Hyvärinen 1985; Figure 6). Early peaks in Oxyria and Salix are interpreted as a

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successional period lasting 1000 years. A long-term decrease in pollen concentration suggests cooling during the late Holocene. From the central islands, pollen diagrams are available from the High Arctic (Figure 7; RS29, PWWL) of Somerset and Prince of Wales islands, and the middle Arctic of Somerset Island (RS36). As was seen on Ellesmere Island, pollen concentrations generally decreased through time at these sites. Salix was abundant in the early Holocene (except on Prince of Wales Island). Poaceae pollen increased in the late Holocene in response to a large-scale cooling of the Arctic (Gajewski 1995; Gajewski and Frappier 2001).

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Figure 5. Postglacial pollen diagrams from Banks (74MS11, 74MS12; Gajewski et al. 2000a) and Baffin (Patricia Bay, Iglutalik; Short et al. 1985) Islands. Data obtained from the North American Pollen Database (http://www.ngdc.noaa.gov/paleo/paleo.html). Pollen sum includes all upland pollen and spores.

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Figure 6. Pollen diagrams from northern Baffin Island (Fish Lake; Short et al. 1994) and eastern Ellesmere Island (Baird Inlet; Hyvärinen 1985). Data source and pollen sum as in Figure 5.

ARCTIC PALYNOLOGY co nc en tra tio n Po lle n

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Figure 7. Pollen diagrams from the central Arctic. RS36 is from the middle Arctic and RS29 from the High Arctic of Somerset Island (Gajewski 1995); PWWL is from northern Prince of Wales Island (Gajewski and Frappier 2001). Data source and pollen sum as in Figure 5.

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Case study: The arctic treeline zone The arctic treeline has been defined in a number of ways by various authors, but in general the treeline zone can be taken as an ecotone between subarctic boreal forest to the south and arctic tundra to the north. The treeline zone runs eastward along the southern flanks of the Brooks Range in Alaska to the Mackenzie Delta in Canada, then arcs southeastward to the southern shores of Hudson Bay, and finally cuts across northern Québec and adjacent Labrador (Figure 1). In general, the treeline zone in North America corresponds with the mean summer position of the Arctic Front (Bryson 1966) and serves as an important bioclimatic, cultural and now political boundary (MacDonald and Gajewski 1992; MacDonald et al. 1998; MacDonald 2001). One region of intense study is northwestern Québec and Ungava, where 17 pollen diagrams and over 70 surface samples have been published from all of the vegetation zones of this transition (Richard 1981; Gajewski and Garralla 1992; Gajewski et al. 1993, 1996). In this region, the transition from treeline to the tundra, the forest-tundra, spans several degrees of latitude, and the nature and dynamics of the forest-tundra can be studied in more detail (Payette 1983). In northwestern Québec, the pollen rain is dominated by three taxa, Picea (spruce), Alnus (alder) and Betula (birch), which together can comprise over 90% of the total grains. The modern pollen are closely related to the vegetation zones (Gajewski 1991). The cores show similar sequences, with variations due to latitude (Figure 8; Gajewski and Garralla 1992; Gajewski et al. 1993, 1996). After deglaciation, which here occurred around 6000 yr BP, there was only a brief tundra period. Alnus and Picea pollen percentages soon increased; these taxa quickly reached their northern limits. In the southern portion of the transect, pollen assemblages containing high amounts of Populus, Larix and Cupressaceae pollen were deposited by a vegetation with no apparent modern analogue. In the past 3000 years, there has been a progressive deforestation of the forest-tundra, interpreted by a decrease in Picea pollen and increase of dwarf Betula and non-arboreal pollen. A high-resolution analysis of the past 1000 years of four lakes showed a decrease in spruce pollen production and transport during the Little Ice Age (Gajewski 2000). A number of lakes north of the treeline in central Canada and in the vicinity of Yellowknife, Northwest Territories (Figure 2), have been cored to determine the history of treeline in this part of Canada (e.g., Nichols 1975; Moser and McDonald 1990). In addition, a particularly detailed comparison has been made between the pollen stratigraphy and the diatom, geochemical and isotope records from one site (Queen’s Lake) in the region (Moser and MacDonald 1990; MacDonald et al. 1993; Wolfe et al. 1996; Pienitz et al. 1999). The pollen records from sites in the modern tundra typically commence shortly after deglaciation at ca. 8000 yr BP and show an initial vegetation dominated by shrub Betula tundra (Figures 9 and 10). Alnus became an important component of the vegetation, particularly around lakes and adjacent to streams between 7000 and 6500 yr BP. At ca. 5000 yr BP, there was a marked increase in Picea pollen, mainly of Picea mariana, suggesting the establishment of boreal forest-tundra or forest in the vicinity of the lakes. This forest phase appears to have lasted until around 3000 yr BP, when the modern shrub Betula and herb-dominated tundra was established. The development of a Picea forest or forest-tundra is apparent in a number of sediment

ARCTIC PALYNOLOGY BI2 Er ic P ac e oa a c e Cy ea p e O era th ce e Sp r h ae ha er gn bs um

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Figure 8. Pollen diagrams from treeline in northern Québec. GB2 is from the lichen woodland, EC1 from the tree subzone of the forest-tundra, and LB1 from the shrub subzone of the foresttundra (Gajewski et al. 1993). The three diagrams on the right-hand side are from the tundra (Gajewski and Garralla 1992; Gajewski et al. 1993). Data source and pollen sum as in Figure 5.

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records taken from lakes located some 10 to 25 km north of the mapped treeline (Figure 10). A comparison of the pollen record with paleolimnological evidence from the diatom, geochemical and isotope stratigraphies shows that the period of treeline advance between 5000 and 3000 yr BP was marked by similarly rapid and profound changes in the lake ecosystems (Figures 9 and 10). Increasing lake productivity is indicated by the loss-on-ignition record of sediment organic content and in the diatom and chrysophyte cyst records (Moser and MacDonald 1990; MacDonald et al. 1993; Wolfe et al. 1996; Pienitz et al. 1999). For example, diatom valve concentrations were up to four times greater than prior to or after the treeline advance. The diatom assemblages also suggest decreased lake water acidity in accordance with the elemental geochemistry which indicates a decrease of in-washed inorganic material and decreased acidity. The stable isotope record suggests that the development of forest or forest-tundra in the vicinity was also marked by a decrease in the rate of lake water evaporation relative to inflow. Taken together, the results suggest that warmer and moister conditions between ca. 5000 and 4000 yr BP produced relatively rapid and marked changes in the terrestrial vegetation and lake ecosystems.

Figure 9. Fossil pollen (a), diatom (b), loss-on-ignition, geochemical and isotope (c) stratigraphy of a small lake located north of the mapped treeline near Yellowknife, Northwest Territories, Canada (from MacDonald et al. 1993).

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Several palynological studies from the Mackenzie Delta region of northwestern Canada, coupled with the radiocarbon dating of subfossil tree stumps, indicate that there was an advance of boreal forest vegetation between ca. 9000 to 5000 yr BP (Ritchie 1984; Spear 1993). Forest or forest-tundra dominated by Picea appears to have expanded northward by at least several tens of kilometres at this time. Pollen records from lakes in the adjacent Mackenzie Mountains also indicate a minor increase in higher elevation forest cover during this same period with a smaller readvance centred on ca. 3000 yr BP (Szeicz et al. 1995; Szeicz and MacDonald 2001). A comparison of the pollen record with paleolimnological evidence from the diatom, geochemical and isotope stratigraphies shows that the period of treeline advance between 5000 and 3000 yr BP was marked by similarly rapid and profound changes in the lake ecosystems (Figures 9 and 10). Increasing lake productivity is indicated by the loss-on-ignition record of sediment organic content and in the diatom and chrysophyte cyst records (Moser and MacDonald 1990; MacDonald et al. 1993; Wolfe et al. 1996; Pienitz et al. 1999). For example, diatom valve concentrations were up to four times greater than prior to or after the treeline advance. The diatom assemblages also suggest decreased lake water acidity in accordance with the elemental geochemistry which indicates a decrease of in-washed inorganic material and decreased acidity. The stable isotope record suggests that the development of forest or forest-tundra in the vicinity was also marked by a decrease in the rate of lake water evaporation relative to inflow. Taken together, the results suggest that warmer and moister conditions between ca. 5000 and 4000 yr BP produced relatively rapid and marked changes in the terrestrial vegetation and lake ecosystems. Several palynological studies from the Mackenzie Delta region of northwestern Canada, coupled with the radiocarbon dating of subfossil tree stumps, indicate that there was an advance of boreal forest vegetation between ca. 9000 to 5000 yr BP (Ritchie 1984; Spear 1993). Forest or forest-tundra dominated by Picea appears to have expanded northward by at least several tens of kilometres at this time. Pollen records from lakes in the adjacent Mackenzie Mountains also indicate a minor increase in higher elevation forest cover during this same period with a smaller readvance centred on ca. 3000 yr BP (Szeicz et al. 1995; Szeicz and MacDonald 2001). The vegetation history of Alaska was recently summarized and mapped by Anderson and Brubaker (1993, 1994). During the Holocene, spruce migrated northward and westward from central Alaska. In the early Holocene there was probably a broad band of forest-tundra surrounding boreal forest, then located only in central Alaska (Anderson and Brubaker 1993, 1994). Spruce continued to expand in the mid-Holocene and reached modern limits by 4000 yr BP. Palynological records from near the northern treeline in Alaska do not indicate a significant northward extension of Picea forest during the Holocene (Edwards et al. 2000a). It may be that the elevational barrier provided by the Brooks Range restricted significant northward movement of Picea forest in that region. Other tree species such as paper birch (Betula papyrifera), aspen and poplar (Populus tremu loides, P. balsamifera) are important components of the forests and these can be found north of the spruce limit (Anderson and Brubaker 1993). Analysis of cores from northeastern Alaska (Anderson et al. 1988) showed that Populus arrived before Picea. However, the interpretation of Populus assemblages around treeline is not clear, due in part to the poor preservation of its pollen, and these may reflect parkland or gallery forests located around watercourses and lowlands.

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Figure 10. Picea pollen and loss-on-ignition (LOI) stratigraphy of four small lakes located just north of the mapped treeline near Yellowknife, Northwest Territories (from MacDonald et al. 1993).

Several pollen diagrams from the present-day tundra extend our understanding of treeline variations and the history of the tundra communities in Alaska. A long core from Joe Lake, presently in the tundra of northwestern Alaska, extends back into the glacial period (Anderson et al. 1994). The pollen diagram distinguished several different tundra communities, some with apparently no modern analogue. Pollen diagrams from

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the Brooks Range show a replacement of herb tundra by a Betula and then Alnus shrubtundra during the Holocene (Brubaker et al. 1983; Eisner and Colinvaux 1992; Oswald et al. 1999). In an earlier study, MacDonald and Gajewski (1992) assembled palynological records from treeline sites in Québec, central Canada, and western Canada to highlight the geographic pattern of climate and vegetation change along this important ecotone (Figure 11). The palynological data suggest that, while all three regions experienced periods of limited increases in forest cover in the treeline zone, the timing of these events was asynchronous from west to east. Some of the apparent delay in the development of the northern forest in central Canada may relate to the relatively late deglaciation of those regions (between 8000 and 6000 yr BP) and some of this regional difference may reflect changes in the geometry of the Arctic Front over Canada.

Figure 11 . Increase and decrease of Picea pollen reflecting the increase and decrease of forest cover at three treeline sites in Canada (from MacDonald and Gajewski 1992).

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Case study: Charcoal analysis as a complement to palynological studies Charcoal fragments are also preserved on pollen slides, and can be used to determine fire frequency. Charcoal can be analysed in several ways. The easiest way is simply to scan the same slides used for pollen and measure the charcoal grains using an appropriate graticule. Thin sections can be made, or the charcoal retained on sieves of various sizes can be enumerated. Each of these probably measures slightly different aspects, either local or regional charcoal sources, although they tend to give comparable results (MacDonald et al. 1991; Earle et al. 1996; Carcaillet et al. 2001a). Fires are of minor importance in tundra, but become important at treeline and in boreal environments (Ritchie 1987; Payette et al. 1989). Analyses of cores indicate that fire frequency is a function of the climate and vegetation type in the boreal forest (MacDonald et al. 1991; Gajewski et al. 1993; Carcaillet et al. 2001b). Case study: Sphagnum, other mosses and aquatic macrophytes The discussion above has emphasized pollen from terrestrial plants. Aquatic and wetland plants also produce pollen and spores, and these can be analysed on pollen slides. However, they are less frequently studied, and indeed, often not even presented in published reports. Recent studies have synthesized the results of these data obtained from the North American Pollen Database. Sphagnum (peat moss) produces distinctive spores. The plant is found only rarely in the Arctic islands, yet spores are abundant and widely dispersed (Gajewski et al. 2001). Halsey et al. (2000) and Gajewski et al. (2001) interpreted the distribution of Sphagnum spores across North America since the Last Glacial Maximum from maps of spore abundance. Maps comparing the modern distribution of Sphagnum spores and peatland abundance show a close correlation, suggesting mapped patterns can be used to demonstrate areas of peatland accumulation in the past. During the full- and lateGlacial, Sphagnum was abundant in Alaska, and it spread to western Canada from this source during the Holocene. In eastern North America, Sphagnum was present just south of the ice sheet, and spread northward following the retreating ice. Brubaker et al. (1998) studied the spore morphology of true mosses (Bryidae), identifying several groups and distinctive taxa. This can potentially increase the resolution of vegetation reconstructions and was illustrated by the analysis of surface samples and a core from southwestern Alaska. Aquatic macrophytes also produce pollen or spores and these are preserved in the sediments. Fredskild (1977, 1983a,b, 1985b, 1992) summarized the lake sediment records of Greenland using pollen, spore and macrofossil records. Both climate and nutrient availability seem to determine the time-space distribution of the major taxa. Edwards et al. (2000b) used pollen of aquatic macrophytes and sediment properties to reconstruct lake-level variations in Alaska. Dieffenbacher-Krall and Jacobson (2001) and Sawada et al. (2002) analysed the migration of aquatic macrophytes by mapping the pollen and spores of several taxa through the postglacial. An interesting result of these studies was the documentation of significant amounts of aquatic pollen in Alaska during the Last Glacial Maximum. This area was the refugium of the aquatic plants in western Canada.

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Case Study: Climate and analogue reconstructions A number of methods are available for producing quantitative estimates of past climates from fossil pollen. Earlier studies used multiple regression (e.g., Andrews et al. 1980), whereas today the method of modern analogues (Overpeck et al. 1985) is commonly used both to estimate past climates and to aid in the interpretation of past vegetation communities. To be successfully applied, an extensive network of modern sites is necessary, and in North America there are hundreds of modern pollen sites, although site density decreases rapidly toward the north. Many studies have presented quantitative reconstructions and maps of climates based on pollen for North America (e.g., Webb et al. 1993). Here we will discuss a couple of recent studies that focused on the arctic and subarctic regions. Anderson et al. (1989) found modern analogues for 12 cores from western Canada. Maps illustrated the spatial distribution of modern samples that were considered "good analogues" for fossil spectra. In general, many modern samples from a wide geographic range could be identified as analogues for fossil samples from the mid- and late Holocene. Non-analogue conditions, however, were identified for many early Holocene samples. Sawada et al. (1999) reconstructed August temperatures for several sites in QuébecLabrador using the method of modern analogues. Sites in northwestern Québec showed little temperature change in the past 6000 years, except for slightly cooler conditions between 5000 and 6000 yr BP. In Labrador, conditions were cooler between 10,000 and 7000 yr BP, warm around 6000 yr BP, and there has been a slight cooling in the late Holocene. The temperatures were coherent with those estimated using dinoflagellates in shallow water marine cores. Other uses of the modern analogue method are discussed in the next section. Case Study: Palynological data synthesis and comparison studies A number of recent efforts have been made to collate the existing data into mapped pollen or climate values or biome-scale vegetation communities. Some recent examples from North America include Edwards et al. (2000a), which collated pollen records to provide a history of biome development and change in the Beringian region (western Canadian Arctic and adjacent Alaska and eastern Siberia) at 18,000 yr BP and 6000 yr BP. The results indicate that at 18,000 yr BP (the Last Glacial Maximum), the entire region of Beringia that was ice-free supported tundra vegetation, while by 6000 yr BP the vegetation was much like present, except for the slight northward advance of forest in northwestern Canada and westward advance in northern Alaska. Using a similar data set, Edwards et al. (2001) derived point maps of temperature and precipitation anomalies inferred from pollen data from Alaska and eastern Siberia and compared these to similar spatial anomaly patterns seen in climatological station data from the same region. Gajewski et al. (2000b) used five different approaches to reconstruct the conditions at 6000 yr BP in North America and the adjacent oceans. Pollen analyses were used in two different ways. A subjective analysis of pollen and other data showed warmer conditions than today at the southern limit of the boreal forest, but that the forest-tundra ecotone was further south in Labrador. Maps of the climate reconstructed

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using the method of modern analogues showed a complex pattern of warming and cooling, with northern Québec drier and Keewatin and the Yukon wetter than today. The CAPE (Circum-Arctic PaleoEnvironments) project collated and interpreted pollen records from sites throughout the Arctic to derive circum-arctic paleovegetation and temperature departure maps and compared these to the estimates of past vegetation and climatic conditions produced by linked climate model – vegetation model experiments. One important result from the initial study was the conclusion that the climatic model being tested over-estimated temperatures in central Eurasia during the two time periods being examined (CAPE 2001). Conclusion – Outlook Arctic palynology has reached the stage where the basic postglacial chronology of vegetation and climate history has been determined for a number of regions. Refinements such as the ancillary use of conifer stomata or the detailed analysis of minor pollen taxa, aquatic pollen and spores has increased the scope of paleoenvironmental questions we might address using palynological records and the resolution of the answers we might derive. Numerous surface samples are now available for quantitative interpretation of pollen diagrams. The collation and application of large spatial data arrays of pollen records from the Arctic has allowed the testing of climate and vegetation results and raises new questions regarding the dynamics of climate and vegetation change. There remain regions of the North American and Eurasian Arctic that are poorly represented by fossil pollen records and many of the available sequences tend to have a relatively coarse temporal resolution, often further compromised by poor dating. However, sufficient data are now available to permit syntheses, mapping and quantitative paleoenvironmental reconstructions. Future work is needed to obtain records from under-represented regions and derive higher-resolution sequences, which can investigate interesting questions such as asynchronous changes in arctic climate and vegetation during the Holocene and the impact of different temporal modes of climatic variability on the arctic system. Summary Arctic pollen analysis provides information about many aspects of environmental change. The vegetation succession following deglaciation has been identified using pollen and spores from lake sediment cores, and pollen concentrations or influx may be related to vegetation density. The analysis of tree pollen transported from forested regions into the Arctic has provided information about past atmospheric circulation. Arctic pollen analyses identify maximum temperatures in the early to mid-Holocene depending on the region, and nearly all sites record a deterioration in the climate during the past several thousand years. The arctic treeline has varied through time in response to changes in the location of the Arctic Front, but results from one region cannot be extrapolated to other areas. Several treeline sites have shown a response to global warming of the past century, and it remains to be seen how extensive this impact has been on arctic vegetation.

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Nichols H. 1974. Arctic North American palaeoecology: the recent history of vegetation and climate deduced from pollen analysis. In: Ives J. and Barry R.G. (eds), Arctic and Alpine Environments. Methuen, London, pp. 637-667. Nichols H. 1975. Palynological and paleoclimatic study of the late Quaternary displacements of the boreal forest-tundra ecotone in Keewatin and Mackenzie, N.W.T., Canada. Institute of Arctic and Alpine Research Occasional Paper 15. Boulder, Colo., 87 pp. Nichols H., Kelly P. and Andrews J.T. 1978. Holocene palaeo-wind evidence from palynology in Baffin Island. Nature 273: 140-142. Oswald W.W., Brubaker L.B and Anderson P.M. 1999. Late Quaternary vegetation history of the Howard Pass area, northwestern Alaska. Can. J. Bot. 77: 570-581. Oswald W.W., Anderson P.M., Brubaker L.B., Hu F.S. and Engstrom D.R. 2003. Representation of tundra vegetation by pollen in lake sediments of northern Alaska. J. Biogeogr. 30: 521-535. Ovenden L. 1988. Holocene proxy-climate data from the Canadian Arctic. Geological Survey of Canada Paper 88-22: 1-11. Overpeck J.T., Prentice I.C. and Webb III T. 1985. Quantitative interpretation of fossil pollen spectra: dissimilarity coefficients and the method of modern analogs. Quat. Res. 23: 87-108. Payette S. 1983. The forest-tundra and present tree lines of the northern Québec-Labrador Peninsula. Nordicana 47: 3-24. Payette S., Morneau C., Sirois L. and Desponts M. 1989. Recent fire history of the northern Québec biomes. Ecology 70: 656-673. Pienitz R., Smol J.P. and MacDonald G.M. 1999. Paleolimnological reconstruction of Holocene climatic trends from two boreal treeline lakes, Northwest Territories, Canada. Arct. Ant. Alp. Res. 31: 82-93. Pisaric M.F.J., MacDonald G.M., Cwynar L.C. and Velichko A.A. 2001. Modern pollen and conifer stomates from north-central Siberian lake sediments: Their use in interpreting Late Quaternary fossil pollen assemblages. Arct. Ant. Alp. Res. 33: 19-27. Ponader K., Pienitz R., Vincent W. and Gajewski K. 2002. Limnological conditions in a subarctic lake (northern Québec, Canada) during the late Holocene: analyses based on fossil diatoms. J. Paleolim. 27: 353-366. Prentice I.C. 1978. Modern pollen spectra from lake sediments in Finland and Finnmark, north Norway. Boreas 7: 131-153. Renberg I. 1981. Improved methods for sampling, photographing and varve-counting of varved lake sediments. Boreas 10: 255-258. Richard P.J.H. 1981. Paléophytogéographie postglaciaire en Ungava par l’analyse pollinique. Paléo-Québec No. 13. Université de Québec à Montréal, Montréal, Québec, 153 pp. Ritchie J.C. 1974. Modern pollen assemblages near the arctic tree line, Mackenzie Delta region, Northwest Territories. Can. J. Bot. 52: 381-396. Ritchie J.C. 1984. Past and Present Vegetation of the Far Northwest of Canada. University of Toronto Press, Toronto, 251 pp. Ritchie J.C. 1985. Quaternary pollen records from the western interior and the arctic of Canada. In: Bryant Jr. V. and Holloway R. (eds), Pollen Records of Late-Quaternary North American Sediments. American Association of Stratigraphic Palynologists, Dallas, Texas, pp. 327-352. Ritchie J.C. 1987. Postglacial Vegetation of Canada. Cambridge University Press, Cambridge, 178 pp. Ritchie J.C. and Lichti-Federovich S. 1967. Pollen dispersal phenomena in arctic-subarctic Canada. Rev. Palaeobot. Palynol. 3: 255-266. Ritchie J.C., Hadden K. and Gajewski K. 1987. Modern pollen assemblages from the high Arctic of western Canada. Can. J. Bot. 68: 1605-1613. Sawada M., Gajewski K., de Vernal A. and Richard P. 1999. Comparison of marine and terrestrial Holocene climates in eastern North America. The Holocene 9: 267-278. Sawada M., Viau A. and Gajewski K. 2002. The biogeography of aquatic macrophytes in North America since the last glacial maximum. J. Biogeogr. 30: 999-1017.

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5. ALGAL INDICATORS OF ENVIRONMENTAL CHANGE IN ARCTIC AND ANTARCTIC LAKES AND PONDS

MARIANNE S.V. DOUGLAS ([email protected]) Paleoecological Assessment Laboratory Department of Geology University of Toronto Toronto, Ontario M5S 3B1, Canada PAUL B. HAMILTON ([email protected]) Canadian Museum of Nature Life Sciences Section, Research Division P.O. Box 3443, Station D Ottawa, Ontario K1P 6P4, Canada REINHARD PIENITZ ([email protected]) Paleolimnology-Paleoecology Laboratory Centre d'études nordiques Département de Géographie Université Laval Québec, Québec G1K 7P4, Canada and JOHN P. SMOL ([email protected]) Paleoecological Environmental Assessment and Research Laboratory Department of Biology Queen’s University Kingston, Ontario K7L 3N6, Canada

Key words: Algae, Diatoms, Chrysophytes, Cyanoprokaryotes/Cyanobacteria, Climate change, Environmental change, Arctic, Antarctica, Paleolimnology, Lakes, Ponds

117 R. Pienitz, M.S.V. Douglas and J.P. Smol (eds), 2004. Long-term Environmental Change in Arctic and Antarctic Lakes. Springer. Printed in the Netherlands.

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Introduction High latitude regions have been repeatedly identified as important reference areas for the study of long-term ecosystem change, as arctic and antarctic ecosystems are expected to show the first signs of environmental shifts (such as climatic warming). In addition, due to positive feedback mechanisms, circumpolar ecosystems are affected more strongly by environmental changes (both natural and anthropogenic) than temperate and tropical regions. Long-term monitoring data, however, are noticeably lacking for almost all high latitude regions, and so indirect proxy methods must be used to reconstruct past environmental trajectories. As shown by the chapters in this volume, paleolimnological approaches (i.e., the study of lake and pond histories from the physical, chemical, and biological information preserved in sediment profiles) have played a leading role in providing these missing data sets. Using these long-term records, natural and anthropogenic forcing factors can be examined, natural variability can be assessed, and episodic events can be identified and studied (Smol 2002; Cohen 2003). Such data are required to help track the causes and consequences of environmental change, to better understand the natural modes of environmental change, and to make more effective decisions regarding environmental management issues. A characteristic feature of high latitude regions is the thousands of ponds and lakes that, along with wetlands, dominate most circumpolar landscapes. The sediments that accumulate in these systems archive a diverse "library" of information that can be used by paleolimnologists to reconstruct past environments. Algal indicators play a leading role in many of these long-term assessments. Although the word “algae” has no formal taxonomic meaning, the word is commonly used to describe a diverse group consisting of primarily aquatic, photosynthetic protists and prokaryotes (i.e, cyanoprokaryotes, which are now often referred to as cyanobacteria or blue-green algae) that contain chlorophyll-a (Chl-a) as their primary photosynthetic pigment and have simple reproductive structures. As a group, they exhibit tremendous diversity in life forms, morphological features, as well as reproductive and physiological strategies. Other chapters in this volume provide details on the application of algal indicators (especially diatoms) to paleolimnological reconstructions in different high latitude regions (Chapters 9 to 15). The reader is also referred to seven recent reviews that are relevant to algal-based paleolimnological reconstructions in high latitude regions. Douglas and Smol (1999), Lotter et al. (1999) and Spaulding and McKnight (1999) recently reviewed the applications of diatoms to high arctic, subarctic, and antarctic paleolimnology, respectively. Smol and Douglas (1996) summarized the potential applicability of using changes in algal assemblages as part of arctic biomonitoring programs. Moser et al. (2000) discussed the ways that diatoms can be used to track paleohydrology in arctic regions, whereas Pienitz (2001) provided an overview of the use of diatoms in northern peatlands. In addition, Smol and Cumming (2000) reviewed the application of algae to paleoclimatic reconstructions in general, which included examples from high latitude regions. Many chapters in this book provide more recent examples of algal-based paleolimnology, and so only a brief overview is provided here of how algal indicators have been used to track long-term environmental change in arctic and antarctic regions. We begin with a summary of some of the key historical studies on the taxonomy, distribution, and ecology of algal assemblages from high

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latitudes. We then discuss the current approaches that are typically employed, and provide a survey of some of the major types of questions that can be addressed using algal remains in paleolimnological studies. Historical overview of algal research in the Arctic and Antarctic1 The first surveys The first algae studies from polar regions were published in the mid-19th century with Joseph Dalton Hooker’s (1847) records from marine antarctic waters and Christian Gottfried Ehrenberg’s (1843, 1844, 1853) observations from Alaska, Spitsbergen (Svalbard), and the Antarctic Ocean. Even during this early period of discovery, both Hooker and Ehrenberg were pondering about micro-organism habitat selectivity and species biogeographic distributions. The observations of Hooker (1847) from his Antarctic Voyage of H.M. Discovery ships Erebus and Terror, and Sutherland (1852) during the search for the missing crew of these ships from Baffin Bay and Barrow Straits (1850-1851), make specific reference to the brown and green colours of the bottom ice and pore water. Hooker more specifically says “The Water and the Ice to the South Polar ocean were alike found to abound with microscopic vegetables”, while Sutherland at the opposite pole described the algae as a “green slimy-looking substance” in the bottom ice representing “infusory animalcules” and “minute vegetable forms of exquisite beauty” (Horner 1985). Following these early observations, the studies of freshwater and marine microbes multiplied quickly from across the Arctic, with many publications documenting the diatom floras (e.g., Cleve 1867, 1873, 1883, 1896, 1898, 1900; Lagerstedt 1873; Cleve and Grunow 1880; Dickie 1880; Grunow 1884; Østrup 1895, 1897a,b), whereas more limited diatom work was done in the subantarctic and antarctic ecozones (e.g., Castracane 1886; Reinsch 1890). Cleve (1896) was the first to observe the similarity of diatom distributions across the Arctic and states “The great resemblance between diatoms found in the ice at Cape Wankarema, between Franz Josefs Land and Novaja Semlja, and at the east coast of Greenland and those observed in the Labrador-stream, tends to show that the ice-flakes are drifted from Behring strait to the north of Greenland, where one portion of them continues to drift along the east-coast of Greenland and another along the Labradorstream”. Fridtjof Nansen (1897, 1906) was probably the first arctic researcher to use a simple “multi-proxy” approach to scientifically examine the drift theory using diatoms (identified by Cleve), Ciliata and Flagellata “germs” (documented by Nansen). Through this research, which included documenting the similarity of marine and freshwater diatoms and microbes in meltwater pools from the Bering Strait and Greenland, Nansen proposed that drift ice could pass across the North Pole. Marine and freshwater floristic surveys at both poles continued during the early 20th century. In the Antarctic, the studies varied from diatoms (Holmboe 1902), general freshwater algae (e.g., Wille 1902, 1924; Gain 1911, 1912; West and West 1911; Carlson 1913), to marine taxa (Petit 1908; Van Heurck 1909; Peragallo 1921; Heiden and Kolbe 1928), whilst arctic surveys were also expanding into both freshwater and 1

Appendix 1 lists the location of historical algal collections from polar phycologists.

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marine diatom floristic studies (e.g., Østrup 1901, 1910, 1920; Gran 1904; Petersen 1924; Palibin in Melnikow 1997). Through the research publications of Cleve, Østrup, Lagerstedt and Petersen, many new taxa were identified and preliminary diatom distributions across the Arctic were documented. Arctic research (1920-1990) Two Canadian expeditions, the 1913-1918 Canadian Arctic Expedition of the Western Arctic and the 1938-1939 Canadian Arctic Expedition of the Eastern Arctic, provided a more complete biogeographical understanding of brackish and freshwater algae from the Canadian Arctic Archipelago. Lowe (1923) studied the algae of small brackish ponds in both Alaska and the Northwest Territories. Lowe’s diatom flora is brief and he does not provide illustrations or discuss ecology in any detail. A second Canadian Arctic Expedition was undertaken in 1938-1939, with a focus on the Eastern Arctic, including several high arctic sites. Diatom identifications were delayed by World War II, but completed when Ross (1947) published descriptions of 192 freshwater taxa. Wheldon (1947) documented other freshwater algae. Between 1950 and 1990, most of the algal studies were ecologically or ecosystem-based (for a review see Hamilton et al. 2001). Friedrich Hustedt (1942), using diatoms collected from Swedish Lapland, was probably the first to make a concerted effort to link environmental variables to diatom ecology. Throughout his career, Hustedt established the foundations for the first simple transfer functions relating diatom autecology to pH and general water quality. Meanwhile, the prominent Danish phycologist, Niels Foged, appreciated the significance of Hustedt’s work and proceeded, whenever possible, to link environmental conditions, especially pH and temperature, with geological regions and diatom community composition, especially in Greenland (1953, 1955, 1958, 1972, 1973, 1977), Iceland (1974), Spitsbergen (1964), and Alaska (1981). Foged (1972, 1977, 1989) was one of the first researchers to use diatoms from arctic postglacial deposits to reconstruct isostatic processes, marine-lacustrine transitions, and emerging lake patterns after the last glaciation. A milestone for North American work was the monograph by Patrick and Freese (1961), which described freshwater diatoms from northern Alaska. Meanwhile, Croasdale (1973) described and illustrated 225 freshwater algal taxa (exclusive of diatoms) from northern Ellesmere Island in the Canadian High Arctic. The first detailed paleolimnological study of a high arctic lake was completed by Smol (1983), who used diatom and chrysophyte microfossils to infer paleoclimatic trends from a small lake on eastern-central Ellesmere Island. Antarctic research (1920-1990) As summarized by Spaulding and McKnight (1999), the earliest phycological work in the antarctic regions was from expeditions to more accessible sites, such as the subantarctic Kerguelen and South Georgia islands (Reinsch 1890; Carlson 1913). Before 1960, limnological research on continental Antarctica was almost non-existent

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(< 10 publications). However, in the following two decades, over 100 taxonomic or floristic papers on algae were completed (Prescott 1979). Over 70 investigations on limnology and terrestrial ecology were also published, with a diversity of studies ranging from algal-soil interactions (primarily cyanobacteria), to salinity and nutrient dynamics in permanently frozen lakes. Diatom research during this time was not extensive, but researchers such as Fukushima (1962a,b, 1963, 1964, 1966, 1969, 1970), Fukushima et al. (1973, 1974, 1975) and Kobayashi (1962a,b, 1963a,b,c, 1965a,b) published regional floras and taxonomic treatises. These floristic studies were further complemented by the work of other researchers, including Bourrelly and Manguin (1954), Aleshinskaja and Bardin (1965), Opalinski (1972a,b, 1974), and Seaburg et al. (1979) from many other sectors across the antarctic region. One significant finding was the observation that the number of antarctic algal taxa decreased with latitude (Hirano 1965; Fukushima 1970). Similar in timing to research developments in the Arctic, Brady (1982) used diatoms from sediments to infer climatic conditions through the Late Cenozoic, and Burckle et al. (1988) examined diatoms in ice cores to interpret aspects of the glacial history of Antarctica. High latitude studies (1991- present) Over the last ca. 15 years, research using algae (and especially diatoms) from both polar regions has increased exponentially (reviewed in Douglas and Smol 1999; Lotter et al. 1999; Spaulding and McKnight 1999). As summarized by the chapters in this volume, the development of algal-based research from primarily descriptive floristic studies to their application in elucidating patterns of global environmental significance has been striking. Algal indicators Algae are often the dominant primary producers in lakes and ponds, and are found in virtually every water body where there is sufficient light for photosynthesis. Algae are important bioindicators of environmental conditions for a variety of reasons (Stevenson and Smol 2003). First, they are typically abundant and diverse, and reflect all the major habitats present in aquatic systems. Algae have rapid dispersal rates, short life cycles, and respond quickly to environmental changes. Importantly, many taxa have welldefined environmental optima and tolerances (Birks 1998) to important limnological variables, and so, as described later in this chapter, species assemblage data can be used to characterize environmental conditions. As many taxa leave reliable morphological and biogeochemical records in lake and pond sediments, they can be used by paleolimnologists to reconstruct past environments. Below, we first describe the major algal indicators that are used by paleolimnologists. We then summarize some of the definitions typically used in algal assessments of environmental change. Finally, we describe some of the main types of applications that have been undertaken using algal-based paleolimnological techniques.

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Figure 1. Light micrographs (DIC: differential interference contrast) of diatoms from Ellesmere Island, Canadian High Arctic. (a) Left to right: Cymbella botellus (Lagerstedt) A. Schmidt, Eunotia praerupta Ehrenberg, Encyonopsis sp.; (b) Diploneis sp.; (c) Complete clavate frustule in girdle view of Meridion circulare (Greville) Agardh.; (d) Cyclotella antiqua W. Smith.; (e) top to bottom: Aneumastus tusculus Ehrenberg Mann and Stickle (synonym Navicula tuscula Ehrenberg) and Cymbopleura designata Krammer (synonym Cymbella designata Krammer). (f) Raphe valve of Achnanthes coarctata (Brébisson) Grunow. Scale bar represents 10 µm.

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Diatoms (class Bacillariophyceae) The most commonly used biological indicators of past limnological conditions are diatoms, considered by many to be the most successful group of algae. Round et al. (1990) describe many aspects of the biology and taxonomy of diatoms, whereas Stoermer and Smol (1999) have provided a series of review chapters describing the many ways that diatoms can be used in environmental assessments. Diatoms have many characteristics that make them ideal biomonitors of environmental change in high latitude systems where they are especially common in benthic, shallow water habitats. Diatoms are an extremely diverse group of algae, with global estimates ranging from 10,000 to 100,000 or more taxa (Mann and Droop 1996). Many species have well defined optima to important environmental variables, such as lake water pH, nutrients, salinity, and climate-related variables, and are therefore reliable indicators of environmental conditions (Stoermer and Smol 1999). The taxonomy of diatoms is based primarily on the size, shape, and ornamentation of their siliceous cell walls (called frustules), each composed of two valves (Figure 1). As these siliceous valves are typically well preserved in lake and pond sediments, they can be used by paleolimnologists to track environmental changes. Because of their small size (many less than 30 µm in diameter or length), highresolution light microscopy, as well as transmission and scanning electron microscopy, are typically required for critical taxonomic assessments (Round et al. 1990). Diatom valves can easily be separated from the sediment matrix using standard digestion techniques, which are well established and documented (Battarbee et al. 2001). Under most circumstances, no special preparation techniques are required for arctic and antarctic assemblages. Chrysophytes (classes Chrysophyceae and Synurophyceae) Although used less commonly than diatoms, chrysophyte algae hold considerable potential in paleolimnological studies. Sandgren et al. (1995) provided a series of review chapters dealing with many aspects of chrysophyte biology, including chapters detailing ecological (Siver 1995) and paleolimnological (Smol 1995) applications.

Figure 2. Scanning electron micrographs of chrysophycean cysts from Greenland lake sediments. Cyst morphotypes numbers follow Duff et al. (1995) and Wilkinson et al. (2001). (a) Stomatocyst #11. (b) Stomatocyst #337. (c) Stomatocyst #374. Scale bar represents 2 µm. Images courtesy of Sergi Pla.

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Figure 3. Light micrographs (DIC: differential interference contrast) of disarticulated scales of the chrysophyte Mallomonas from the sediments of Kekerturnak Lake, eastern Baffin Island, Nunavut. Scales of M. allorgei- lychenensis (a-g); M. pseudocor onata (h-i). Both of these taxa were restricted to the lake’s most recent sediments. Scale bars are 5 µm. From Wolfe and Perren (2001); used with permission.

Chrysophytes often dominate the plankton of high latitude lakes and ponds (e.g., Sheath 1986), and recent work suggests that many periphytic forms are also common (Douglas and Smol 1995a; Wilkinson et al. 1997). Chrysophytes are represented in the fossil record primarily by their endogenous resting stages (Figure 2) known as stomatocysts (also called statospores in the older literature). Similar to diatoms, stomatocysts are siliceous and well preserved in sediments, and can be studied using similar techniques developed for diatoms (Zeeb and Smol 2001). Different taxa produce cysts that are believed to be species-specific. Relatively few cyst morphotypes have been linked to the taxa that produce them; however, well-defined taxonomic guidelines

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have been developed for describing cyst morphotypes using electron and light microscopy (Duff et al. 1995; Pla 2001; Wilkinson et al. 2001). Although many cysts are now simply identified as numbered morphotypes, work continues on linking cysts to the taxa that produced them. Nonetheless, from a paleoecological perspective, numbered cyst morphotypes can still be used in paleolimnological assessments, provided that the ecological characteristics of morphotypes can be determined using surface sediment calibration sets, as described below. Compared to other geographic regions (e.g., Duff et al. 1997), a moderate number of studies have been completed for chrysophyte cysts in arctic and subarctic regions (e.g., Duff and Smol 1988, 1989; Duff et al. 1992; Pienitz et al. 1992; Brown et al. 1994, 1997; Gilbert et al. 1997; Wilkinson et al. 1997). Studies on cyst morphotypes from subantarctic regions have also been initiated (e.g., Van de Vijver and Beyens 1997a,b, 2000). In addition to stomatocysts, the siliceous scales and bristles of taxa in the classes Synurophyceae and Chrysophyceae (which include important genera such as Mallomonas, Synura, Spinifermonas, Paraphysomonas, and Chrysosphaerella) are species-specific and typically well preserved in sediments (Zeeb and Smol 2001). Scaled chrysophytes are rare or absent in many high latitude settings, however they are present in some lakes and ponds (e.g., Kristiansen 2001), and have been used by paleolimnologists to track recent environmental changes (Figure 3; Wolfe and Perren 2001). Other morphological indicators Although siliceous microfossils, such as diatom valves and, to a lesser extent, chrysophyte scales and cysts are the most commonly used algal indicators in paleolimnological studies, lake and pond sediments also preserve a diverse array of other morphological fossils. Many of these have yet to be explored in high latitude environments. As summarized by van Geel (2001) and Jankovská and Komárek (2000), many green (Chlorophyta), dinoflagellate (Pyrrhophyta) and blue-green algae (Cyanoprokaryotes or often referred to as Cyanobacteria) leave some morphological fossils. Probably the most commonly reported non-siliceous remains used by paleoecologists are colonies of the green alga Pediastrum (Komárek and Jankovská 2001), which are counted by palynologists as Pediastrum colonies often survive the sediment preparation techniques used for pollen (Bennett and Willis 2001). Although some of the earliest paleolimnological work in the Arctic considered algal remains such as Pediastrum and Botryococcus colonies (e.g., Fredskild 1983), the abundances of these microfossils are still rarely reported. Given the low temperatures that characterize high latitude lakes and ponds, it is likely that many other morphological algal remains are preserved in these deposits, and represent a potentially valuable and as yet untapped source of paleoenvironmental information. For example, filamentous and sheet-forming Cyanobacteria are often very common in many polar lakes and ponds (Figure 4). Their reproductive and resting structures, such as akinetes, are well preserved in some temperate deposits (e.g., van Geel et al. 1994). Stratigraphic analyses of these morphological remains should provide important information on ecosystem development.

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Figure 4. Filamentous and sheet-forming Cyanobacteria (blue green algae) are common in many polar lakes and ponds.

Biogeochemical indicators Not all algal groups leave morphological fossils. However, a suite of biogeochemical indicators can often be used to characterize essentially all algal groups, at least in a general way. Foremost amongst these biogeochemical indicators are fossil pigments (i.e., chlorophylls and carotenoids, the latter represented by carotenes and xanthophylls). As described by Leavitt and Hodgson (2001), fossil pigment analyses can provide important information on past algal and cyanobacterial populations, as many groups have specific pigment compositions. For example, alloxanthin is specific to cryptophytes, whereas oscillaxanthin is only produced by some Cyanobacteria (see Table I in Leavitt and Hodgson 2001). There is clearly considerable potential for using fossil pigment analyses in high latitude paleolimnology, although little work has been completed thus far. Paleo-pigment analyses also hold considerable potential for tracking past ultra-violet (UV) light penetration (e.g., Leavitt et al. 1997, 2003a,b; Hodgson et al., this volume), which is a major concern in high latitude regions (Vincent et al. 1998). Furthermore, blue-green algae are often important primary producers in many high latitude lakes and ponds (Vincent 2000). Although akinetes and other morphological fossils are occasionally recovered from sediments (van Geel 2001), fossil pigments may provide more reliable estimates of past populations in arctic and antarctic ecosystems. Given that high latitude systems are cold and dark for long periods of time, potential problems of pigment diagenesis may be lessened in these environments. Clearly, considerable potential remains for applying these techniques to future studies.

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The field of organic biogeochemistry is developing rapidly, with many new potential indicators and applications. Some examples of these include biomarker hydrocarbons (e.g., alkanes, lignins, phenols), ratios of various elements, such as carbon to nitrogen, as well as many other biomarkers which may potentially be related, at least in an indirect way, to past algal populations. Meyers and Teranes (2001) summarized many of these relatively new approaches and applications; however, most freshwater applications are still from temperate regions. Biogenic silica Estimates of past population sizes are possible by calculating absolute abundances of microfossils (either as concentration data or correcting these data for changing sedimentation rates, to provide accumulation data). However, biogeochemical techniques can also be used to estimate past shifts in total population sizes of past diatom (and other siliceous algal) assemblages. For example, biogenic silica analysis (Conley and Schelske 2001) uses a timed chemical digestion technique to estimate the total amount of silica (SiO2) in sediments that is derived from biogenic sources. As climatic conditions, such as changes in ice and snow cover on lakes (Smol 1983, 1988), may pose important limitations on past overall primary production, biogenic silica analyses can provide a proxy of production of siliceous algae, and therefore also climatic conditions. For example, Willemse and Törnqvist (1999) used biogenic silica levels, as well as other paleolimnological data, to track past primary production in western Greenland lake sediment records, and correlated these changes to climatic shifts, as inferred from isotope data gathered from the Greenland Ice Core Project. Similarly, Wolfe (2003) used biogenic silica to explore the relationship between past siliceous algal production and climate in the Canadian Arctic. Ecological classifications of algae Habitat classification Different algal taxa often characterize specific aquatic habitats. As the availability of these habitats is often closely related to climatic and other environmental conditions in polar regions, documenting the nature and degree of algal habitat specificity has important paleolimnological implications (Smol 1988; Douglas and Smol 1999; Pienitz 2001). For example, certain taxa are characteristic of open water assemblages, while other species are adapted to attached (periphytic) shallow water habitats. The term phytoplankton refers to a diverse group of algae that characterize the open water regions of lakes and ponds. Phytoplanktonic organisms have little control over their position in the water column, and are largely dependent on currents. However, some taxa (such as cryptophytes and many chrysophytes) are flagellated and therefore have some mobility. For example, during periods of low light penetration (such as under ice cover) these algae may move to the surface of the water column and take advantage of any light for photosynthesis that may penetrate the snow and ice cover.

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The second major functional group of algae is the periphyton, which includes taxa that characterize a suite of shallow, littoral zone habitats. Periphytic algae, which include many different phyla (using the most recent guidelines from the Botanical Code, algae are now classified into phyla, whereas previously the taxonomic category of division was used), are further differentiated by the type of substrates that they exploit. For example, the epilithon are algae attached to rocks, the epiphyton attached to plants, the epipsammon attached to sand grains, and the epipelon living on the sediment surface. Given that many high latitude lakes have extended snow and ice covers (thus precluding extensive phytoplankton growth), benthic algae often dominate these systems. Some taxa are found in both planktonic and periphytic habitats due to entrainment into the water column by wind and water current action. These are often referred to as the tychoplankton. Environmental optima and tolerances For many decades, limnologists have been attempting to classify aquatic ecosystems based on the algal assemblages characterizing these environments (for a historical review, see Stevenson and Smol 2003). For example, Hustedt (1937-1939) divided the diatom taxa he identified in Sumatra, Bali and Java based on pH categories, namely acidobiontic (occurring at pH values < 7, optimum distribution at pH = 5.5 and under), acidophilous (occurring at pH of about 7, and most common at pH < 7), indifferent (equal occurrences on both sides of pH 7), and alkaliphilous (occurring at pH values of about 7, and most common at pH > 7). Subsequent researchers added the category circumneutral to designate taxa that are most commonly found near pH 7. Although these terms are still used today, data derived from surface sediment calibration sets (discussed below) have largely supplanted these categories. Surface sediment calibration sets or training sets Over the last ca. 20 years, paleolimnologists have been cataloguing the ecological optima and, to some extent, tolerances of algal (especially diatom) assemblages employing more quantifiable approaches by using surface sediment calibration sets or training sets and appropriate multivariate statistics. The overall approach is straightforward, and has been reviewed in Charles and Smol (1995) and Smol (2002), with a synopsis of the major statistical techniques summarized in Birks (1998). Briefly, a paleolimnologist chooses an appropriate set of calibration lakes and ponds (typically about 50 in number) that encompasses the major environmental gradients characterizing the study region. Limnological and other environmental data are gathered for the lake set. It is desirable to have several years of limnological data for a training set, but in high latitude regions, where logistical constraints often override scientific aspirations, environmental data from a single sampling date are often the only available measurements. However, given that the open water period is usually very short for many circumpolar lakes, and if an appropriate and consistent time window is used to collect the environmental data, ecologically relevant data can still be collected from a

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single sampling date. The limnological data collected as part of a calibration set typically includes chemical variables such as pH, specific conductivity, nutrient concentrations, dissolved organic carbon (DOC), as well as physical variables such as morphometric data and temperature, and occasionally biological data, such as grazers and the distributions and types of plant substrates (e.g., mosses). This information comprises the first data matrix used in the calibration set, namely the measured environmental variables characterizing the lake set. The second matrix of data includes the indicators (typically expressed as relative abundance data) collected for the calibration lakes and ponds. In order to provide a temporally and spatially integrated sample of indicators, paleolimnologists collect the surface (e.g., top 0.5 cm or top 1 cm) of sediment from each calibration site. This surface sediment sample contains the preserved remains of diatoms, chrysophytes, and other indicators that lived in the lake over the last few years, and so provides the “response variables” for the calibration set. Using a variety of statistical approaches (Birks 1998), the ecological optima and tolerances of taxa can then be estimated, and transfer functions can be constructed that allow environmental conditions to be inferred from fossil assemblage data. Although surface sediments are routinely used for training sets, arctic paleolimnologists have employed similar calibration approaches to describe the environmental optima and tolerances of periphytic diatoms, such as those attached to mosses or rocks (e.g., Douglas and Smol 1995b; Vinocur and Pizarro 2000; Michelutti et al. 2003a). As habitat specificity is an important ecological variable in paleolimnological reconstructions, these biological data can be used to reconstruct past climatic and other environmental conditions. Paleolimnological reconstructions Algal assemblages as indicators of past climatic change Although long-term monitoring data are rarely available, fortunately many algal taxa provide direct or indirect proxy data of climatically relevant variables. Given that many of the traditional paleoclimate indicators, such as pollen grains or tree ring sequences, are more difficult or impossible to use in polar regions (e.g., absence of trees), and that lakes and ponds are often characteristic features of these landscapes, fossil algal assemblages are now being used routinely in many paleoclimatic studies. Lake ice and snow cover, and related limnological variables An overriding factor influencing the physical, chemical, and biological characteristics of high latitude aquatic systems is the extent and duration of ice and snow cover, which itself is closely related to prevailing climatic conditions. In the first detailed high arctic paleolimnological study using diatoms and chrysophytes, Smol (1983) proposed that the extent of snow and ice cover on an Ellesmere Island lake was a dominant factor influencing algal assemblages. As elucidated further in subsequent publications (Smol 1988; Douglas and Smol 1999), changing snow and ice cover may have an overriding influence on many physical, chemical, and biological characteristics of high latitude

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lakes, and so it is not surprising that algal indicators would track, at least indirectly, past climate (Figure 5). Ice and snow cover on lakes and ponds is influenced by several climatic factors. Certainly air temperature is a critical variable, but other factors, such as wind, precipitation, and local morphometric and topographic characteristics are also important (Douglas and Smol 1999). In most arctic regions, lakes and ponds are totally frozen by September and only begin to thaw in June or July. In some high arctic and antarctic lakes, persistent ice cover during the summer is common (e.g., an ice thickness of over 5 m has been recorded in August on Ellesmere Island; Blake 1989), with only a shallow moat of ice-free water in the lake’s littoral zone. Although some algae still persist under an ice cover, the concentration and community composition are markedly affected (Smol 1983, 1988; Douglas and Smol 1999). During warmer years, ice and snow cover is reduced, and progressively larger portions of the lake system are more amenable to the growth of different algal communities (Figure 5). As summarized by Douglas and Smol (1999), changes in ice cover have a cascading effect throughout the aquatic ecosystem, affecting the length of the growing season, the percentage of plankton, the development of substrates such as mosses, diversity, pH, nutrient levels, and other physical, chemical and biological factors. As diatoms and other indicators track these limnological changes, shifts in past algal assemblages can be used to estimate past climatic and other environmental changes. Smaller water bodies, particularly in the High Arctic, may be expected to be even more responsive to environmental changes (Rouse et al. 1997). In polar regions, ponds are often defined as water bodies that freeze completely in winter. These shallow sites, many of them less than 0.5 m in depth, often represent the dominant type of surface waters in high latitude ecosystems. Given their small volumes, ponds should be especially sensitive bellwethers of environmental shifts related to climate change (Douglas and Smol 1994). The first paleolimnological studies from shallow arctic ponds revealed striking successional changes beginning in the 19th century that appeared to be related to climatic warming (Douglas et al. 1994). As summarized by numerous examples in this volume, many other researchers have used similar approaches to track climatic changes in a diverse range of lakes and ponds (see the chapter by Wolfe and Smith for a thorough review). Although ice cover is an important limnological variable in all high latitude ecosystems, related environmental variables will also influence algal assemblages. Lakes in high polar regions are rarely thermally stratified, but as one moves farther from the poles, shifts in thermal stratification may also affect algal assemblages. For example, Sorvari and Korhola (1998) and Sorvari et al. (2002) recorded marked diatom changes in sediment cores from several Finish Lapland lakes. In deeper lakes, they reported synchronous increases in small planktonic Cyclotella diatoms in the more recent sediments, with concurrent relative decreases in benthic taxa. They interpreted these stratigraphic changes as evidence of longer ice-free periods and enhanced thermal stratification. Similarly, Rühland et al. (2003) recorded increases in a planktonic Cyclotella taxon in the 19th century sediments of a lake in subarctic Canada, which they suspected may have been at least indirectly related to enhanced thermal stratification with climate warming.

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Figure 5. Although a variety of climate-related factors influence the limnological characteristics (and hence the biota) of high latitude lakes, changes in ice and snow cover are often overriding variables. This schematic diagram shows ice and snow conditions on a polar lake during relatively cold (A), moderate (B), and warm (C) conditions. A permanent float of ice and snow may persist throughout the summer during colder years (A), precluding the development of large populations of planktonic algae, and restricting much of the primary production to the shallow, open water moat. Many other physical, chemical and biological changes occur in lakes that are either directly or indirectly affected by snow and ice cover (Douglas and Smol 1999). From Smol (1988); used with permission.

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Undertaking detailed paleolimnological studies of sediment cores is time-consuming, and so it is often difficult, due to practical and logistical reasons, to attain regional assessments of environmental change. Detailed, contiguous sampling of lake sediment cores certainly provides much information on the timing and trajectories of environmental changes. However, if a large number of lakes and ponds need to be studied, and if the overall research question can be restricted to simply asking “Are recent algal assemblages similar to those present before the period of major human impacts, such as from the early 1800s?”, then paleolimnologists can still provide ‘snapshots’ of past environmental change by using the so-called “top-bottom” or “before and after” paleolimnological approach (Smol 2002). The top-bottom approach is fairly simple. Paleolimnologists remove sediment cores spanning the last ca. 200 years from the study lakes, as they would in a typical stratigraphic study. However, instead of analysing many sediment sections from each core, they only analyse indicators, such as diatoms or chrysophytes, from the surface centimetre of sediment (representing recent algal assemblages, the “top” sediment sample) and from a sediment slice that represents the time period known to pre-date significant anthropogenic activities (the “bottom” sediment sample). Comparing algal assemblages from these two time periods allows the paleolimnologist to estimate if change has occurred in an ecosystem. This approach is an approximation, and certainly does not provide the same answers attainable from a detailed paleolimnological investigation. However, it is a practical and reasonable approach where regional questions are posed. For example, Rühland et al. (2003) used the top-bottom approach to track overall patterns in diatom assemblage changes from 50 lakes across the Canadian arctic treeline ecosystem. Similar to the studies noted earlier, they showed that Cyclotella taxa percentages increased in almost every study lake, at the expense of benthic diatom taxa, with the greatest change occurring in the deeper lakes. They interpreted this regional assemblage shift as indicative of climate warming, resulting in a shorter duration of ice cover, a longer growing season, and/or stronger thermal stratification. Similarly, Betts-Piper et al. (2004) used a top-bottom approach to track marked changes in chrysophyte cyst assemblages in a suite of Svalbard lakes. Perhaps some of the most striking evidence of environmental change was first documented from shallow ponds on eastern-central Ellesmere Island, Canada. Paleolimnological analyses from several sites at Cape Herschel showed an unprecedented shift in diatom assemblages beginning ca. AD 1850 (Douglas et al. 1994). At this time there was an almost complete shift from an assemblage consisting mainly of benthic Fragilaria c onstruens-pinnata taxa (now positioned in the genera Staurosira and Staurosirella, respectively) to one of much greater diversity. The authors concluded that a lengthening in the growing season, due to climate warming, had been responsible for this marked change in the paleolimnological record. Further work by Overpeck et al. (1997) supported these findings. In their study, a high resolution 400-year climate record was examined using various proxy records (i.e., ice cores, dendrochronology etc.) demonstrating that, starting ca. 1845, climate forcing factors such as volcanism, solar radiation, and other variables had combined to influence warming at high latitudes. As described in subsequent chapters, many other studies have since reported similar findings throughout the Arctic. The timing of the shifts in algal assemblage varies, reflecting the heterogeneity of climate across the Arctic, as well as the sensitivity of small shallow ponds versus larger lakes, which have a greater thermal inertia and hence respond more slowly (Michelutti et al. 2003b). Comparing studies

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between regions emphasizes the diverse limnological responses occurring at different latitudes. For example, in high arctic regions, shallow sites typically record the most pronounced changes, whereas in subarctic regions (where thermal stratification may be a factor), deeper lakes generally track the most striking changes. Moreover, regions that were not expected to have warmed (e.g., parts of northern Québec and Labrador; Laing et al. 2002; Ponader et al. 2002; Paterson et al. 2003) have recorded relatively complacent sedimentary algal profiles. Dissolved organic carbon (DOC) reconstructions at northern treeline Lakes located at or near the treeline boundary in the vast boreal forest region of the northern hemisphere are receiving considerable attention because of their sensitivity to climatic change. Reconstructing the past position of the arctic treeline boundary has important implications, as it corresponds fairly closely to the mean summer position of the Arctic Front and, in particular, to the 10°C July isotherm (Bryson 1966; Larsen 1989; Finney et al., this volume; Pienitz et al., this volume). In addition, lakes and ponds located along latitudinal and ecotonal transects that include steep climate and vegetation gradients show profound changes in water chemistry and physical properties. These changes are reflected in the highly variable algal community compositions across northern treeline (e.g., Pienitz and Smol 1993; Pienitz et al. 1995a,b; Korsman and Birks 1996; Korsman and Segerström 1998; Fallu and Pienitz 1999; Gregory-Eaves et al. 1999; Lotter et al. 1999; Fallu et al. 2002; Rühland and Smol 2002). Two major gradients emerge from algal (mostly diatom) distributional studies across the treeline ecotone (based on lake surface sediment studies). First, the concentration of lake water dissolved organic carbon (DOC, which is primarily controlled by external inputs of humic and fulvic acids from the catchment vegetation and the position of treeline) and related variables (water colour and transparency, which affect the underwater optical conditions in the water column). Second, temperature (surface water and air temperature), which affects many other lake properties, such as the intensity and duration of water column stratification. As discussed later in this chapter, as well as elsewhere in this volume, transfer functions have been constructed to infer both lake water DOC concentrations and temperature (e.g., Pienitz et al. 1999; Seppä and Weckström 1999; Korhola et al. 2000; Rosén et al. 2000; Saulnier-Talbot et al. 2003; Chapters 9-13, this volume). Tracking past changes in ultraviolet (UV) radiation Paleolimnological reconstructions of past irradiance regimes can contribute to a better understanding of the scales, causes and consequences of temporal variability in ultraviolet radiation (UVR). Using microfossil remains, such as diatoms (e.g., Pienitz et al. 1999), retrospective analyses permit reconstructions of concentrations of UVRabsorbing dissolved organic carbon (DOC) or its correlate chromophoric dissolved organic matter (CDOM), which can act as natural sunscreens. Important optical properties of water bodies include spectral irradiance regimes (UVR/photosynthetic active radiation (PAR) ratios) and water transparency (depth of UVR penetration), as well as algal responses to variable UVR flux in freshwater ecosystems. Diatom-based paleolimnological approaches have shown that climatically-induced changes in the

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export of terrestrial (allochthonous) DOC to lakes near treeline is up to 100-fold more effective than modern stratospheric ozone depletion at altering biological exposure to UVR (Pienitz and Vincent 2000). Furthermore, paleolimnological studies have also shown that historical variability in UVR exposure has been high during the Holocene (Pienitz and Vincent 2000; Saulnier-Talbot et al. 2003). To address the potential impact of long-term climate change relative to that of ozone depletion, Pienitz and Vincent (2000) combined paleolimnological analyses of diatom assemblages with bio-optical models based on present-day conditions in northern Canadian lakes. This new paleo-optical approach allowed them to estimate past underwater light conditions from DOC concentrations that were inferred from fossil diatom assemblages preserved in Holocene sedimentary deposits from a lake near the northern treeline (Queen’s Lake; 64°07’N, 110°34’W) in the central Northwest Territories, Canada. Analysis of fossil pollen records indicated that regional vegetation cover was sparse and tundra-like following deglaciation ca. 8000 yr BP and persisted until trees colonized the catchment ca. 5000 years ago (MacDonald et al. 1993). Diatom community structure and inferred DOC levels (Pienitz et al. 1999) showed three distinct and abrupt changes during the history of Queen’s Lake. Both diatom biomass and inferred DOC concentrations were low (< 2 mg DOC l-1) during the initial lake phase, with particularly few fossils recovered from sediments older than 5000 yr BP. This initial period of lake development was followed by major and rapid shifts in species composition and inferred chemical conditions ca. 5000 yr BP, with increased ratios of periphytic/planktonic taxa to > 70% of the total diatom assemblage. This second period of forest advance also corresponded to a major increase in algal production, recorded as the sediment mass-specific concentration of diatom valves, as well as a three-fold increase in inferred DOC levels. Based on fossil pollen analyses, Pienitz and Vincent (2000) argued that changes in lake chemistry and production resulted from climatic warming that stimulated treeline advance and increased forest density for about 2000 years. Finally, diatom-based reconstructions indicated that DOC concentrations declined more than 85% after 3000 yr BP, concomitant with the onset of Neoglacial climatic cooling and a southward retreat of treeline (MacDonald et al. 1993; Pienitz et al. 1999). The large and rapid changes in DOC suggested that Queen’s Lake experienced major shifts in the underwater optical environment over the last 6000 years as a consequence of climate-induced variation in forest development (Pienitz et al. 1999). Consistent with this hypothesis, application of bio-optical models derived from measurements in highlatitude waters (Vincent et al. 1998) showed that the inferred DOC shifts were equivalent to a decrease in exposure of two orders of magnitude of biologicallyeffective UVR during the mid-Holocene vegetation maximum between ca. 5000 and 3000 yr BP. In contrast, the most recent 3000 years were characterized by a > 50-fold increase in levels of damaging UVR, with recent inferences agreeing closely with present-day estimates of UVR exposure. Overall, changes in DOC concentrations arising from climatic variability increased exposure to photosynthetically damaging UVR several orders of magnitude more than did the relatively moderate (30%) decline in stratospheric ozone levels (Pienitz and Vincent 2000). Saulnier-Talbot et al. (2003) used a diatom-based paleo-optical approach to estimate past depths of UV penetration in coastal Lake Kachishayoot (northwestern Québec, Canada; 55°20’N, 77°37’W; 102 m a.s.l.) following its isolation from the marine waters due to glacio-isostatic rebound of the Hudson Bay lowlands. Their multi-proxy

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investigation revealed abrupt increases in diatom-inferred DOC concentrations and water colour that coincided with the retreat of postglacial marine waters and the arrival of spruce trees to the lake’s catchment. They also tracked large changes in the underwater irradiance environment over the course of the postglacial period, from extremely high UV exposure following the initial formation of the lake and its isolation from the sea, to an order-of-magnitude lower exposure associated with the establishment of spruce forests in the lake’s catchment. The use of additional macrofossil markers revealed that UVR penetration remained low even following forest retreat due to the development of alternate DOC sources, such as Sphagnum mats. In another investigation of a lake chronoseries from Glacier Bay, Alaska, Engstrom et al. (2000) demonstrated that DOC concentrations were low for at least the first century following lake formation and that this initially high UVR exposure can structure biotic communities (Fritz and Engstrom 1995; Williamson et al. 2001). Finally, the observations that diatoms are particularly sensitive to changes in UVR exposure (e.g., Vinebrooke and Leavitt 1998; Watkins et al. 2001), and that these organisms are rare in early postglacial sediments (Pienitz et al. 1999), provide evidence that extreme UVR transparency is a mechanism directing the early development of glacial lake ecosystems in high latitude regions. In addition to the type of studies described above, fossil pigment analyses also provide important insights on past trajectories of UV penetration in lakes and ponds (e.g., Leavitt et al. 1997, 2003a,b). Hodgson et al. (this volume) summarize some of the ongoing studies in antarctic regions. Although further research is required to validate fossil interpretations, paleoecological analyses of lake sedimentary records can provide valuable insights into the history of UVR exposure and its potential impacts on freshwater ecosystems. For example, when used in combination with long-term environmental monitoring (e.g., Schindler et al. 1996), historical reconstructions may be valuable at identifying the importance of UVR relative to other stressors in regulating lake structure and function. Climate-related changes in lake water pH A number of within-lake and catchment processes may link lake water pH to a variety of climate-related variables, such as temperature (Psenner and Schmidt 1992). Using these relationships, paleolimnological studies from soft-water alpine lakes have suggested diatom-inferred pH changes may be linked to climate (e.g., SommarugaWögrath et al. 1997; Koinig et al. 1998). Similar processes may be occurring in polar lakes. For example, Wolfe (2002) used diatom-inferred pH to estimate past climaterelated changes in two Baffin Island (Nunavut) lakes over the last ca. 5000 years. Recently, Antoniades (2004) applied similar approaches to two high arctic lakes. Temperature inferences As noted throughout this review, most attempts at inferring past climatic changes use algae to track some limnological variable that is indirectly related to temperature (e.g., pH, salinity, habitat availability, etc.). Using fossil algal assemblages, such as diatoms, to directly track temperature is neither a straightforward subject, nor one without controversy. As discussed by Smol and Cumming (2000), from a paleoclimatic

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perspective, it does not matter if the relationships between algal indicators and temperature are direct or indirect, only how reliably the variable can be reconstructed and whether the relationship between climate and indirect factor(s) remains linear over time. For example, Pienitz et al. (1995a) found that water depth and summer temperature were strongly related to diatom species assemblage data from northern Canadian (Yukon and Northwest Territories) lakes. The authors acknowledged that the relationship may not be a simple one, and that lake water temperature may be the measured environmental variable that simply captured most of the variance. A number of temperature transfer functions have since been developed and applied for other arctic regions (e.g., Weckström et al. 1997; Rosén et al. 2000; Joynt and Wolfe 2001; Bigler and Hall 2002; Wolfe 2003). Athalassic lakes are systems known to occur in the semi-arid polar and subpolar regions of both hemispheres, yet remain relatively untapped sources of paleoclimatic information. The sediments of these saline lakes potentially contain high-resolution records of past climates and hydrological regimes (Pienitz et al. 1992; Fritz et al. 1999), as brine concentrations and lake-levels are governed, to a large extent, by the balance between precipitation and evaporation (or effective moisture). Because lake water salinity concentration and composition have a strong influence on algal distributions that can be quantified (e.g., Wilson et al. 1994; Zeeb and Smol 1995), past changes in climatic and hydrologic conditions can be reconstructed from the algal microfossils preserved in the paleolimnological records of these northern lakes. Using this approach, Pienitz et al. (2000) provided sub-centennial scale evidence for ca. 1500-year cyclic changes in paleosalinity and paleoproductivity from a climate-sensitive saline basin in the central Yukon, Canada (62°45’N, 136°38’W). This study demonstrated the potential of these high latitude saline sites as recorders of paleoclimatic shifts. Paleolimnological work has recently been initiated on saline lake systems in Greenland (Anderson et al., this volume), and similar types of studies can be undertaken with antarctic lakes (e.g., Roberts et al. 2001). For example, a number of researchers working on continental antarctic lakes have recently combined their data sets and developed diatom-based transfer functions to infer lake water salinity and water depth, from which past changes in moisture balance can potentially be deduced (Verleyen et al. 2003). Tracking past river flow Paleolimnological studies of river systems may be challenging, as rivers are high-energy systems and typically have the opposite characteristics that paleolimnologists rely upon to attain reliable archives of environmental change (Smol 2002). Nonetheless, under certain conditions, diatoms and other paleolimnological information pertaining to rivers (e.g., past river inflow, which may be an important climatic indicator) can be archived in arctic lake sediments. For example, Ludlam et al. (1996) and Antoniades and Douglas (2002) noted that certain diatoms (e.g., Meridion circular e and Hannaea arcus) were commonly associated with fluvial habitats in northern Ellesmere Island and Cornwallis Island rivers, respectively, than with nearby lake environments. Using these relationships, Ludlam et al. (1996) proposed a Lotic Index, which was simply calculated as the ratio of the relative abundances of Hannaea and Meridion diatoms, divided by the total number of pennate diatoms. Using this index, they reconstructed past river

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discharge using diatoms preserved in the varved sediments of a downstream lake on northern Ellesmere Island. They interpreted the diatom changes over the last two centuries as indicating a decline in runoff beginning ca. AD 1800 until the late 1800s, followed by increased runoff until the mid-20th century. These types of studies hold considerable promise for other high latitude regions. For example, Potapova (1996) noted that flow regime was the measured environmental variable that best described the species composition of stream and river epilithic diatom communities from northeastern Siberia. Diatoms are also important components of antarctic rivers (e.g., Kawecka and Olech 1993), and so these approaches may well be applicable in the southern hemisphere as well. River diatoms can also be used to study river flow in much more complex arctic river systems, although the approaches are somewhat different. As it is generally recognized that the current state of the Arctic Ocean and its influence on global climate is at least partly dependent on the freshwater input from the discharge of large circumpolar rivers (currently estimated at about 10% of the global freshwater runoff), there is considerable interest in tracking past river inflow. Perhaps the most promising approach to estimate past river inflow from these large complex deltas is to track diatoms and other paleolimnological indicators in sediment cores collected from the delta floodplain lakes. The Mackenzie River Delta (Northwest Territories, Canada) is the second largest river delta system in the Arctic (surpassed only by the Lena River Delta in Siberia). Delta lakes are closely linked to their parent river systems, with the degree of river attachment being a prime variable determining limnological characteristics (e.g., Lesack et al. 1998). For example, delta lakes can typically be divided into three major categories along a gradient of connectivity to the parent river, namely: (1) no closure lakes, which have a continuous connection to the river; (2) low closure lakes, which flood every spring, but lose connection with the river during summer; and (3) high closure lakes, which flood only during extreme events, such as major ice jams that dam the river. Hay et al. (1997, 2000) examined the diatoms preserved in the surface sediments of 77 Mackenzie River delta lakes. Their calibration data suggested that diatoms can be used to estimate trends in past river connectivity of delta lakes. As limnological variables are strongly influenced by the amount of river input a delta lake receives (Lesack et al. 1998), these conclusions were not surprising. Michelutti et al. (2001a) subsequently used these calibration data to infer patterns of past river flow in eight Mackenzie River delta lakes. Marine/lacustrine transgressions and sea-levels Diatoms and other algae are excellent indicators of lake water salinity, with most taxa relatively easily differentiated between freshwater, marine, and brackish forms. Diatom analysis has been widely used at the interface of fresh and saline environments to identify lake isolation from the sea in areas of land uplift in formerly glaciated regions of the Arctic (e.g., Young and King 1989; Pienitz et al. 1991; Douglas et al. 1996; Saulnier-Talbot and Pienitz 2001) and Antarctica (e.g., Roberts and McMinn 1999) to indicate marine and brackish water transgressions. Traditional approaches mostly relied on the identification and analysis of transgressive and regressive overlaps by analysing stratigraphic boundaries between terrestrial freshwater sediments and marine littoral

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facies. In these studies, diatom analysis was used to validate sea-level index points by considering changes in the composition of diatom groups of different salinity preferences. However, the potential for using diatom assemblages to quantitatively reconstruct past sea-level changes has rarely been explored, mainly because the environmental optima of individual taxa with respect to water depth are poorly documented. By determining the relationship between water depth and diatom species distributions and abundances in 74 modern sedimentary environments along the southern shores of the Beaufort Sea coast (Arctic Ocean, Canada), Campeau et al. (1998, 1999, 2000) have been able to overcome the limitations of traditional approaches and quantitatively inferred past relative sea-level changes using a diatom-based waterdepth model that was applied to fossil assemblages in several long sediment cores. Acidification and other contaminants Algal-based paleolimnological studies have played leading roles in assessing many water quality problems (Smol 2002), and foremost amongst these has been research related to lake acidification (Battarbee et al. 1999). Based on available data, however, it appears that few high latitude regions have been acidified as a result of atmospheric deposition of strong acids, although certainly many high latitude systems are subjected to various sources of pollution (Muir and Rose, this volume). For example, the metal smelting facilities at Noril’sk (Taymyr Peninsula, Siberia, Russia) are located approximately 300 km north of the Arctic Circle, and presently represent the largest point source of sulfur dioxide emissions in the world. However, despite very high loadings of acids, Duff et al. (1999) showed that Noril’sk area lakes are still typically alkaline due to local bedrock geology and overlying glacial deposits. Not surprisingly, diatom-based paleolimnological studies in the area could not detect any acidification trends (Michelutti et al. 2001b). Similarly, Weckström et al. (2003) used diatoms in the sediments of 32 Kola Peninsula (Russia) lakes to reconstruct recent changes in lake water pH related to emissions from local industries. Signs of recent acidification could only be recorded in seven of the lakes, and they did not find evidence for any large-scale acidification in the region. Meanwhile, Moiseenko et al. (1997) noted evidence of recent lake acidification from lakes in the poorly buffered regions of Kola North, Russia. The sensitivity of algae to other contaminants is less well documented. For example, organochlorides, such as DDT and PCBs, are transported via atmospheric pathways or by other vectors (e.g., military installations) to high latitude regions (Muir and Rose, this volume). Paterson et al. (2003) tracked diatom and chrysophyte changes in two lakes that had been subjected to high concentrations of PCB contamination from local military facilities in Labrador, Canada. They found no noticeable influence on the algal assemblages. Recently, Outridge et al. (2004) have documented positive correlations of mercury (Hg) concentrations with diatoms in two Canadian high arctic lakes over the last 50 to 100 years. This correlation (when diatoms were present in high numbers) indicated a close association between biology and geochemistry, and may even suggest biological feedback mechanisms that enhance some attributes of geochemical processes in polar lakes.

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Local effects of eutrophication Diatoms are widely recognized as reliable indicators of lake water nutrient levels, and so have become the mainstay of many paleolimnological investigations of eutrophication (Hall and Smol 1999). However, cultural eutrophication is not generally considered to be a major water quality issue in arctic and antarctic regions as human settlements are spaced widely apart and have low populations. There are, however, some exceptions. Much of what we know about high arctic eutrophication has come from work that was initiated in the late 1960s and early 1970s as part of the International Biological Programme (IBP). Two Canadian high arctic lakes in the hamlet of Resolute Bay (Cornwallis Island, Nunavut) were chosen for detailed studies of freshwater production: ultra-oligotrophic Char Lake and eutrophic Meretta Lake (Rigler 1974). Meretta Lake (72°41.75’N, 94°59.58’W) received sewage from a small number of buildings from 1949 to 1998. The Department of Transport base (referred to locally as “the North Base”), associated with the local airport, discharged its grey water and sewage through a central collecting pipe and then via an utilidor onto the land (Figure 6). Utilidors are systems of elevated horizontal, insulated pipes that transport a variety of liquids above ground to prevent them from freezing and to prevent melting of the permafrost. The utilidor outfall discharged sewage into a series of streams that ultimately drained into Meretta Lake. For almost half a century, this nutrient-rich effluent fertilized the lake. Schindler et al. (1974) summarized the limnological characteristics of Meretta Lake during the short time window of the IBP programme but, as pre-eutrophication conditions were not known, it was not possible to place current limnological conditions into a historical context. As the director of this IBP programme, Frank Rigler (1974) noted in the final report on this IBP project: “The conclusions we can draw are limited because the original condition of Meretta Lake is unknown”. The absence of long-term monitoring data for Meretta Lake prompted Douglas and Smol (2000) to undertake a diatom-based paleolimnological study of this important reference site. Using a high-resolution sediment record, they showed that eutrophication from the “North Base” had significantly affected diatom assemblages; however the species changes were markedly different from those recorded in nutrient-enriched temperate lakes. Despite the increase in nutrients, periphytic diatoms continued to overwhelmingly dominate the assemblage, further confirming that extended ice covers, and not nutrients, were precluding the development of large concentrations of planktonic diatoms. In 1993, when the sediment core was collected, the North Base still had a moderate number of users. However, during the 1990s, the number of people using the North Base declined steadily, and in 1999 the entire utilidor system was dismantled and sewage was thereafter trucked to a designated area of the hamlet’s dump. Douglas and Smol (2000) measured summer nutrient levels since 1991, and tracked the decline in eutrophication throughout the 1990s. Comparing the diatom assemblages preserved in the surface sediments of a core collected in 2000 (Michelutti et al. 2002a) with those at the surface of the 1993 core (Douglas and Smol 2000) revealed recent diatom species shifts that were consistent with nutrient reductions. These recent assemblage shifts also confirmed that there were no significant lags between environmental changes and the deposition of diatoms in the profundal zone in these largely ice-covered lakes. In addition, changes in periphytic assemblages collected

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Figure 6. This utilidor, which was recently dismantled, is releasing sewage from the Resolute Bay (Cornwallis Island), North Base into a series of streams and ponds. The effluent ultimately reaches Meretta Lake, about 1.6 km downstream.

each summer since 1991 from Meretta Lake’s littoral zone also tracked the recovery from eutrophication as taxa characteristic of more oligotrophic environments increased relative to more eutrophic diatoms (Michelutti et al. 2002b).As northern communities increase rapidly in size (Canadian Inuit birth rates are among the highest in the world), concerns over issues of waste disposal and other environmental problems will only escalate in the future. Paleolimnological approaches will likely become more important in water quality assessments in these remote regions, as they have been in many other parts of the world. Reconstructing changes in nutrient levels have other applications besides tracking the effects of cultural eutrophication. For example, nutrient levels in high latitude lakes and ponds are also indirectly linked to climatic conditions (Douglas and Smol 1999; Moser et al. 2002), and diatom-based transfer functions to infer lake water nutrient and related trophic variables are being developed for polar regions (e.g., Jones and Juggins 1995; Lim et al. 2001). In addition, wildlife, such as musk ox, caribou, and other grazers, as well as birds, marine mammals, and fish can affect nutrient concentrations, and diatoms can play important roles in tracking these sources (see sections below).

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Tracking the influences of past fish and wildlife populations on freshwater systems Recent research has suggested ways that algal assemblage data can be used, in conjunction with other paleolimnological techniques, to track past changes in ecologically and economically important animal populations. Below, we summarize three recent examples demonstrating how diatoms have been used to track past changes in sockeye salmon, caribou, and fur seal populations, although many other applications are likely possible. Many important fish stocks are decreasing as a result of over-fishing and other human interventions (such as habitat destruction, dam construction, competition from hatchery fish, etc.). Sockeye salmon (Oncorhyncus nerka) is the most important commercial fish species on the Pacific west coast of North America. However, many stocks are believed to be nearing extirpation or are at least threatened, and so there is considerable interest in tracking long-term changes in salmon abundance. Paleolimnological approaches have been developed that take advantage of the anadromous life strategy exploited by sockeye salmon. Fry of the sockeye salmon hatch in a nursery lake and then, after about one to three years, they leave the freshwater environment and migrate downstream to the Pacific Ocean. After living a further one to three years in the mid-Pacific Ocean where they accumulate over 95% of their biomass, they return with very high fidelity to their natal stream and nursery lakes, where they spawn and then die. For some lake systems, the carcasses of spawned salmon may contribute over 50% of the lake water nitrogen and phosphorus levels, and thus are an important subsidy of marine-derived nutrients for nursery lakes. As diatoms are sensitive indicators of lake water nutrient levels, stable isotope analyses (e.g., į15N) can also be used to track past salmon returns. Combined paleolimnological approaches have been used to reconstruct past sockeye salmon returns for a variety of Alaskan nursery lakes (Finney et al. 2000, 2002; Gregory-Eaves et al. 2003, 2004; Holtham et al. 2004). Research is now in progress using similar approaches on nursery lakes in other arctic and subarctic regions, such as the Yukon Territory and northern British Columbia in Canada (D. Selbie and J.P. Smol, unpublished data). The effects of land-based mammals can also be examined using paleophycological approaches. For example, the world’s largest caribou herd in the Rivière George region of northern Québec was known to have undergone a large population increase since the 1950s, which was accompanied by pronounced impacts on vegetation cover, soil erosion, and other terrestrial disturbances. Laing et al. (2002) used diatom-based paleolimnological techniques to assess the effects of wildlife fluctuations on several northern lake ecosystems. They found no evidence of marked changes in lake conditions as a result of the fluctuating herd populations. The results of the Laing et al. (2002) study confirm that, in order to be registered in the diatom record, a perturbation must be sustained and significantly outside baseline natural variability for a lake. While this is often evident for anthropogenic impacts on water quality, evaluating the relative importance of natural factors as influences on limnological conditions is more complicated. In the southern hemisphere, diatoms have been used to assist conservation managers determine the effects of increasing fur seal (Arctocephalus gazella ) populations on coastal antarctic lakes. Although seals were almost extirpated by over-hunting in the 19th century, the cessation of sealing in 1912, coupled with the intensification of the

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South Georgian whaling industry between 1902 and 1965 (baleen whales being the principal competitors of seals for food), has resulted in a striking increase in fur seal populations on some of the South Orkney Islands. Using a variety of paleoindicators, including preserved seal hairs, diatom assemblages, and other data, paleolimnologists were able to ascertain that historical sealing and whaling activities have affected fur seal populations, which in turn have impacted the ecology of coastal lakes (Hodgson and Johnston 1997; Hodgson et al. 1998). Higher seal populations, and therefore higher loading of nutrients from seal excrements (which can be tracked using diatom assemblages; Jones and Juggins 1995), have led to lake eutrophication, which in turn has raised conservation and management concerns. Complementing archeological investigations It is now well documented that high latitude ecosystems have been affected by longrange transport of atmospheric pollutants (Muir and Rose, this volume). However, it is generally assumed that high arctic lakes and ponds have been unaffected by direct local human activities before the arrival of Europeans, as most native peoples were primarily nomadic with relatively low population densities, and had un-intrusive hunting and gathering technologies and customs. Using similar techniques to those developed for the previously described salmon research, Douglas et al. (2004) combined archeological and paleolimnological approaches to challenge some of these assumptions. The Thule Inuit migrated to the Canadian Arctic islands about 1000 years ago from Alaska and brought with them an effective whaling technology (e.g., seal-skin or sometimes walrus-skin boats, harpoons with toggled points, seal skin bladder floats). The Thule Inuit hunted a variety of marine mammals, but their most important prey were large bowhead whales, which provided a food staple as well as material for house construction and tools. Although nomadic during the short summer, the Thule Inuit maintained a number of winter settlements consisting of houses constructed from the bones of bowhead whales (Figure 7). The Thule Inuit abandoned most of the high arctic islands about 400 years ago. Nonetheless, the bones and other debris still present near and in the ponds adjacent to the abandoned over-wintering settlements continue to influence water chemistry (i.e., higher nutrient levels) and diatom assemblages (Douglas et al. 2004). The highest concentrations of abandoned Thule over-wintering sites are present on south-eastern Somerset Island, Nunavut. Douglas et al. (2004) used stable isotope and diatom analyses for one such site (PaJs-13; 72’08.66°N, 94’01.50°W) to show that 13th century whaling activities markedly influenced pond ecology. Site PaJs-13 consists of 11 large semi-subterranean whalebone houses, a number of smaller, shallower sod dwellings, and a ring of 10 bowhead crania of probable ceremonial function. In addition, the adjacent beach area is strewn with several thousand bowhead bones, resulting from the flensing and caching of whales (Habu and Savelle 1994). Diatom assemblages from the pond’s early history were similar to those recorded in other high arctic sites. However, at approximately AD 1200, the paleolimnological record indicated a marked shift in diatom assemblages to moss epiphytes (e.g., Pinnularia balfouriana), coupled with an elevated į15N ratio. This zone identified the period of nutrient enrichment from decaying whale carcasses and other refuse. Diatom

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Figure 7. A Thule over-wintering site, as reconstructed by the Canadian Museum of Civilization. This whale bone house is situated at Resolute Bay, Cornwallis Island, Nunavut, which is north of the Somerset Island site described in Douglas et al. (2004). As no wood is present in the High Arctic (except possibly driftwood from more southern locations), the Thule Inuit used whale bones as the main foundations for their over-wintering houses. This structure would be covered by seal and other animal skins, and then sod for further insulation.

assemblages similarly tracked the abandonment of the site about four centuries ago, although, as noted above, the pond still has elevated nutrient levels and a somewhat distinctive diatom assemblage. The Douglas et al. (2004) Somerset Island study site possibly represents the oldest record of changes in aquatic ecology associated with human activities yet reported for any high arctic ecosystem, as well as the earliest such changes documented for a water body in Canada or the USA. Diatoms have been used to assist archeologists in a number of applications in temperate regions (reviewed in Juggins and Cameron 1999), but have only recently been applied to polar research. Archeological sites are common throughout many arctic regions, and the patterns of settlement and abandonment of native peoples remain an area of controversy. Working collaboratively with archeologists, diatom-based paleolimnolgical techniques can potentially be used to decipher the influences of native cultures on local environments and help track the spread of different peoples throughout the Arctic.

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Other applications Paleoenvironmental information preserved in lake and pond sediments is the focus of this volume, however algae can also provide proxy data from other freshwater habitats in high latitude regions. For example, many high latitude regions (provided they are not too cold or dry) accumulate peat deposits. Incorporated in many peat profiles is a large suite of microscopic indicators, including diatoms, chrysophytes (e.g., Gilbert et al. 1997; Pienitz 2001), and other algal remains. For example, Brown et al. (1994) and Rühland et al. (2000) used diatom and chrysophyte remains in peat cores from Ellesmere Island and Siberia, respectively, to infer past climatic conditions. Diatoms have also been recovered from polar ice cores (e.g., Burckle et al. 1988; Gayley et al. 1989), where they have been used to supplement paleoenvironmental interpretations. Diatoms and other algal indicators have been used in a variety of forensic studies around the world (Peabody 1999). One arctic example includes Foged’s (1982) examination of the eight so-called “Greenland mummies”, who died ca. AD 1460 and whose mummified remains were discovered in 1972. By examining lung tissue from the bodies for diatom remains, Foged concluded that the cause of death was not drowning. Stoermer et al. (1988) noted similarities between the diatom assemblages they found associated with the remains of a mastodon skeleton in Michigan (USA) to those found in the littoral zones of arctic lakes and ponds (Smol 1988). Using this forensic approach, Stoermer et al. (1988) concluded that the mastodon died in a pond, similar to those currently found in arctic or periglacial environments. Summary Algal indicators have been used in a broad spectrum of paleolimnological applications from high latitude regions, even though most studies have only been completed over the last decade. Research has focused primarily on attempts to reconstruct climatic trends, which have included a variety of approaches, ranging from using habitat preferences of algal taxa to reconstruct past snow and ice cover on lakes, to inferring changes in lake water chemistry (e.g., DOC, pH, salinity) that are indirectly related to climatic changes. Efforts have also been made to infer temperature and past underwater light regimes from diatom assemblages, as well as distinguish riverine from lacustrine diatoms to track past river inflows. In addition, paleophycological approaches have been used to track the effects of local human disturbances (e.g., sewage inflows) and other sources of pollution, the influences of animal populations (such as anadromous fish and terrestrial mammals) on lake systems, as well as tracking marine to lacustrine transgressions in coastal systems. Collaborative work with other groups of researchers, such as archeologists, has also proved fruitful. Given that most studies have been published over the past decade, we believe that considerable advancements have been made in a relatively short time. However, much work remains to be done. Although a variety of morphological and biogeochemical indicators are potentially available, diatom assemblage data have overwhelmingly dominated most paleolimnological studies. More reliable inferences of past environmental conditions will be possible by employing a larger spectrum of

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paleoindicators. For example, chrysophytes and blue-green algae are especially common in many high latitude lakes and ponds. It is also becoming increasingly clear that we have only a rudimentary understanding of the limnological characteristics of high latitude systems. As the distributions and abundances of algal taxa are closely linked to environmental conditions, it is important to better understand present-day limnological processes. Fortunately, transportation and northern logistics, which have limited earlier research, are improving in many high latitude regions, and so some scientific progress is being made in these areas. Despite these shortcomings, it is very encouraging to see the advancements made in the diversity and geographic coverage of algal-based paleolimnological studies, and the important data that this research has generated. Over a relatively short period of time, considerable progress has been made in describing, at least in a broad sense, the environmental requirements of many taxa, as well as information on the taxonomy and biogeography of indicators. As illustrated by many examples presented in subsequent chapters of this volume, paleophycological approaches have already made substantial contributions to elucidating the environmental histories of high latitude systems, and have tremendous potential for future applications. Acknowledgements Polar research in our laboratories is primarily funded by the Natural Sciences and Engineering Research Council of Canada (NSERC), the Canadian Museum of Nature and the Polar Continental Shelf Project (PCSP). Many helpful comments were provided by K. Rühland, I. Gregory-Eaves, D. Antoniades, B. Keatley, N. Michelutti, A. Wolfe and W. Last. References Aleshinskaja A.V. and Bardin V.I. 1965. Diatomaceous flora of the Schirmacher Ponds. Soviet Antarc. Expedition Bull. 5: 432-433. Antoniades D.A. 2004. The limnology, diatom autecology, and paleolimnology of lakes and ponds from Alert, Ellesmere Island, Isachsen, Ellef Ringnes Island, and Mould Bay, Prince Patrick Island, Canadian High Arctic. Ph.D. Thesis, University of Toronto, Dept. Geology, Toronto, 569 pp. Antoniades D.A. and Douglas M.S.V. 2002. Characterization of high arctic stream diatom assemblages from Cornwallis Island, Nunavut, Canada. Can. J. Bot. 80: 50-58. Battarbee R., Charles D.F., Dixit S.S. and Renberg I. 1999. Diatoms as indicators of surface water acidity. In: Stoermer E.F. and Smol J.P. (eds), The Diatoms: Applications for the Environmental and Earth Sciences. Cambridge University Press, Cambridge, pp. 185-127. Battarbee R.W., Carvalho L., Jones V.J., Flower R.J., Cameron N.G., Bennion H. and Juggins S. 2001. Diatoms. In: Smol J.P., Birks H.J.B. and Last W.M. (eds), Tracking Environmental Change Using Lake Sediments. Volume 3: Terrestrial, Algal, and Siliceous Indicators. Kluwer Academic Publishers, Dordrecht, The Netherlands, pp. 155-210. Bennett K.D. and Willis K.J. 2001. Pollen. In: Smol J.P., Birks H.J.B. and Last W.M. (eds), Tracking Environmental Change Using Lake Sediments. Volume 3: Terrestrial, Algal, and Siliceous Indicators. Kluwer Academic Publishers, Dordrecht, The Netherlands, pp. 5-32.

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Ross R. 1947. Freshwater Diatomae (Bacillariophyta). In: Polunin N. (ed.), Botany of the Canadian Eastern Arctic. Part II: Thallophyta and Bryophyta. Ottawa. National Museum of Canada, Bull. 97. (Biol. Ser. 26), pp. 178-233. Round F.E., Crawford R.M. and Mann D.G. 1990. The Diatoms: Biology and Morphology of the Genera. Cambridge University Press, Cambridge, 757 pp. Rouse W., Douglas M., Hecky R., Kling G., Lesack L., Marsh P., McDonald M., Nicholson B., Roulet N. and Smol J. 1997. Effects of climate change on fresh waters of Region 2: Arctic and Sub-Arctic North America. Hydrologic Proc. 11: 873-902. Rühland K. and Smol J.P. 2002. Freshwater diatoms from the Canadian arctic treeline and development of paleolimnological inference models. J. Phycology 38: 249-254. Rühland K., Smol J.P., Jasinski P. and Warner B. 2000. Response of diatoms and other siliceous indicators to the developmental history of a peatland in the Tiksi Forest, Siberia. Arct. Ant. Alp. Res. 32: 167-178. Rühland K., Priesnitz A. and Smol J.P. 2003. Evidence for recent environmental changes in 50 lakes across the Canadian arctic treeline. Arct. Ant. Alp. Res. 35: 110-123. Sandgren C., Smol J.P. and Kristiansen J. (eds) 1995. Chrysophyte Algae: Ecology, Phylogeny and Development. Cambridge University Press, Cambridge, 384 pp. Saulnier-Talbot É. and Pienitz R. 2001. Isolation au post-glaciaire d’un bassin côtier près de Kuujjuarapik-Whapmagoostui, en Hudsonie (Québec): une analyse biostratigraphique diatomifère. Géogr. phys. Quat. 55: 63-74. Saulnier-Talbot É., Pienitz R. and Vincent W.F. 2003. Holocene lake succession and palaeooptics of a subarctic lake, northern Québec, Canada. The Holocene 13: 517-526. Schindler D.W., Kalff J., Welch H.E., Brunskill G.J., Kling H. and Kritsch N. 1974. Eutrophication in the high arctic Meretta Lake, Cornwallis Island (75° N lat.). J. Fish. Res. Bd. Can. 31: 647-662. Schindler D.W., Curtis P.J., Parker B.R. and Stainton M.P. 1996. Consequences of climatic warming and lake acidification for UV-B penetration in North American boreal lakes. Nature 379: 705-708. Seaburg K.C., Parker B.C., Prescott G.W. and Whitford L.A. 1979. The algae of southern Victoria Land, Antarctica. A taxonomic and distribution study. J. Cramer, Berlin, Bibliotheca Phycologica 46. Seppä H. and Weckström J. 1999. Holocene vegetational and limnological changes in the Fennoscandian tree-line area as documented by pollen and diatom records from Lake Tsuolbmajävri, Finland. Ecoscience 6: 621-635. Sheath R.G. 1986. Seasonality of phytoplankton in northern tundra ponds. Hydrobiologia 138: 75-83. Siver P.A. 1995. The distribution of chrysophytes along environmental gradients: their use as biological indicators. In: Sandgren C., Smol J.P. and Kristiansen J. (eds), Chrysophyte Algae: Ecology, Phylogeny and Development. Cambridge University Press, Cambridge, pp. 303-329. Smol J.P. 1983. Paleophycology of a high arctic lake near Cape Herschel, Ellesmere Island. Can. J. Bot. 61: 2195-2204. Smol J.P. 1988. Paleoclimate proxy data from freshwater arctic diatoms. Verh. Internat. Verein. Limnol. 23: 837-844. Smol J.P. 1995. Application of chrysophytes to problems in paleoecology. In: Sandgren C., Smol J.P. and Kristiansen J. (eds), Chrysophyte Algae: Ecology, Phylogeny and Development. Cambridge University Press, Cambridge, pp. 303-329. Smol J.P. 2002. Pollution of Lakes and Rivers: A Paleoenvironmental Perspective. Arnold Publishers, London, 280 pp. Smol J.P. and Douglas M.S.V. 1996. Long-term environmental monitoring in arctic lakes and ponds using diatoms and other biological indicators. Geosci. Canada 23: 225-230. Smol J.P. and Cumming B.F. 2000. Tracking long-term changes in climate using algal indicators in lake sediments. J. Phycology 36: 986-1011.

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Sommaruga-Wögrath S., Koinig K., Schmidt R., Sommaruga R., Tessadri R. and Psenner R. 1997. Temperature effects on the acidity of remote alpine lakes. Nature 387: 64-67. Sorvari S. and Korhola A. 1998. Recent diatom assemblage changes in subarctic Lake Saanajärvi, NW Finnish Lapland and their paleoenvironmental implications. J. Paleolim. 20: 205-215. Sorvari S., Korhola A. and Thompson R. 2002. Lake diatom response to recent Arctic warming in Finnish Lapland. Global Change Biology 8: 153-163. Spaulding S.A. and McKnight D.M. 1999. Diatoms as indicators of environmental change in Antarctic freshwaters. In: Stoermer E.F. and Smol J.P. (eds), The Diatoms: Applications for the Environmental and Earth Sciences. Cambridge University Press, Cambridge, pp. 245-263. Stevenson R.J. and Smol J.P. 2003. Use of algae in environmental assessments. In: Wehr J. and Sheath R. (eds), Freshwater Algae of North America: Classification and Ecology. Academic Press, San Diego, pp. 725-804. Stoermer E.F. and Smol J.P. (eds) 1999. The Diatoms: Applications for the Environmental and Earth Sciences. Cambridge University Press, Cambridge, 484 pp. Stoermer E.F., Kociolek P., Shoshani J. and Frisch C. 1988. Diatoms from the Shelton Mastodon site. J. Paleolim. 1: 193-199. Sutherland P.C. 1852. Journey of a voyage in Baffin Bay and Barrow Straits, in the years 18501851, Performed by H.M. ships “Lady Franklin” and “Sophia”, under the command of Mr. William Penny, in search of the missing crews of H.M. ships “Erebus” and “Terror”. Longman, Brown & Green, London. 2 vols, 232 pp. Van de Vijver B. and Beyens L. 1997a. The subfossil chrysophyte cyst flora of some peat samples from Kerguelen Islands. Arch. Protistenkd. 148: 491-503. Van de Vijver B. and Beyens L. 1997b. The subfossil chrysophyte cyst flora of the moss vegetation from Strömness Bay area, South Georgia. Arch. Protistenkd. 148: 505-520. Van de Vijver B. and Beyens L. 2000. Chrysophycean stomatocysts from freshwater habitats of Strömness Bay area, South Georgia, Antarctica. Can. J. Bot. 78: 88-97. van Geel B. 2001. Non-pollen palynomorphs. In: Smol J.P., Birks H.J.B. and Last W.M. (eds), Tracking Environmental Change Using Lake Sediments. Volume 3: Terrestrial, Algal, and Siliceous Indicators. Kluwer Academic Publishers, Dordrecht, The Netherlands, pp. 99-119. van Geel B., Mur L.R., Ralska-Jasiewiczowa M. and Goslar T. 1994. Fossil akinetes of Aphanizomenon and Anabaena as indicators for medieval phosphate-eutrophication of Lake Gosciaz (Central Poland). Rev. Palaeobot. Palyn. 83: 97-105. Van Heurck H. 1909. Diatomées. In: Expédition Antarctique Belge. Résultats du Voyage du S.Y. Belgica 1897-1859. Botanique 6: 1-126. Verleyen E., Hodgson D.A., Vyverman W., Roberts D., McMinn A., Vanhoutte K. and Sabbe K. 2003. Modelling diatom responses to climate induced fluctuations in moisture balance in continental Antarctic lakes. J. Paleolim. 30: 195-215. Vincent W.F. 2000. Cyanobacterial dominance in the polar regions. In: Whitton B.A. and Potts M. (eds), The Ecology of Cyanobacteria: Their Diversity in Time and Space. Kluwer Academic Publishers, Dordrecht, The Netherlands, pp. 321-340. Vincent W.F., Laurion I. and Pienitz R. 1998. Arctic and Antarctic lakes as optical indicators of global change. Ann. Glaciol. 27: 691-696. Vinebrooke R.D. and Leavitt P.R. 1998. Direct and interactive effects of allochthonous dissolved organic matter, inorganic nutrients, and ultraviolet radiation on an alpine littoral food web. Limnol. Oceanogr. 43: 1065-1081. Vinocur A. and Pizarro H. 2000. Microbial mats of twenty-six lakes from Potter Peninsula, King George Island, Antarctica. Hydrobiologia 437: 171-185. Watkins E.M., Schindler D.W., Turner M.A. and Finlay D. 2001. Effects of solar ultraviolet radiation on epilithic metabolism, and nutrient and community composition in a clear-water boreal lake. Can. J. Fish. Aquat. Sci. 58: 2059-2070. Weckström J., Korhola A. and Blom T. 1997. Diatoms as quantitative indicators of pH and water temperature in subarctic Fennoscandian lakes. Hydrobiologia 347: 171-184.

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Weckström J., Snyder J.A., Korhola A., Laing T.E. and MacDonald G.M. 2003. Diatom inferred acidity history of 32 lakes in the Kola Peninsula, Russia. Wat. Air Soil Pollut. 149: 339-361. West W. and West G.S. 1911. Freshwater Algae. In: Murray J. (ed.), Biology, Vol. 1, Reports on the Scientific Investigations, British Antarctic Expedition 1907-9. Heinemann, London, pp. 263-287. Wheldon R.M. 1947. Algae. In: Polunin N. (ed.), Botany of the Canadian Eastern Arctic. Part II: Thallophyta and Bryophyta. Ottawa. National Museum of Canada, Bull. 97. (Biol. Ser. 26), pp. 13-137. Wilkinson A.N., Zeeb B.A., Smol J.P. and Douglas M.S.V. 1997. Chrysophyte stomatocyst assemblages associated with periphytic, high arctic pond environments. Nord. J. Bot. 17: 95-112. Wilkinson A.N., Zeeb B.A. and Smol J.P. 2001. Atlas of Chrysophycean Cysts, Volume II. Kluwer Academic Publishers, Dordrecht, The Netherlands, 180 pp. Wille N. 1902. Antarktische Algen. Mitteilungen über einige von Borchgrevnik auf den antarktischen Pflanzen. III. Nyt Magazin for Naturvidenskaberna 40: 209-222. Wille N. 1924. Süsswasseralgen von der Deutschen Südpolar-Expedition auf dem Schiff “Gauss” I. Teil. Süsswasseralgen vom antarktischen Festlande. Deutsche Südpolar-Expedition 19011903. Reichsministerium des Innern, E.V. Drygalski (ed.), 8. Botanik: 373-407. Willemse N.W. and Törnqvist T.E. 1999. Holocene century-scale temperature variability from West Greenland lake records. Geology 27: 580-584. Williamson C.E., Olson O.G., Lott S.E., Walker N.D., Engstrom D.R. and Hargraves B.R. 2001. ultraviolet radiation and zooplankton community structure following deglaciation in Glacier Bay, Alaska. Ecology: 82: 1748-1760. Wilson S.E., Cumming B.F. and Smol J.P. 1994. Diatom-salinity relationships in 111 lakes from the Interior Plateau of British Columbia, Canada: Development of diatom-based models for paleosalinity and paleoclimatic reconstructions. J. Paleolim. 12: 197-221. Wolfe A.P. 2002. Climate modulates the acidity of Arctic lakes on millennial time scales. Geology 30: 215-218. Wolfe A.P. 2003. Diatom community responses to late Holocene climatic variability, Baffin Island, Canada: a comparison of numerical approaches. The Holocene 13: 29-37. Wolfe A.P. and Perren B.B. 2001. Chrysophyte microfossils record marked responses to recent environmental changes in high- and mid-arctic lakes. Can. J. Bot. 79: 747-752. Young R.B. and King R.H. 1989. Sediment chemistry and diatom stratigraphy of two high arctic isolation lakes, Truelove Lowland, Devon Island, N.W.T., Canada. J. Paleolim. 2: 207-225. Zeeb B.A. and Smol J.P. 1995. A weighted-averaging regression and calibration model for inferring lakewater salinity using chrysophycean stomatocysts from western Canadian lakes. Int. J. Salt Lake Res. 4: 1-23. Zeeb B.A. and Smol J.P. 2001. Chrysophyte scales and cysts. In: Smol J.P., Birks H.J.B. and Last W.M. (eds), Tracking Environmental Change Using Lake Sediments. Volume 3: Terrestrial, Algal, and Siliceous Indicators. Kluwer Academic Publishers, Dordrecht, The Netherlands, pp. 203-223.

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Appendix 1. Some polar phycologists and the locations of their algal collections of historical significance. International herbarium codes are in brackets. Phycologist

Location of Historical Collections

Bourrelly P. Castracane F.

Paris Museum, Paris, France (PC) Dr. Henri Van Heurck Museum, Antwerpen, Belgium (AWH); British Museum, London, UK (BM) Friedrich-Hustedt-Arbeitsplatz für Diatomeenkunde, Bremerhaven, Germany (BRM); Paris Museum, Paris, France (PC); Philadelphia Academy of Natural Sciences, Philadelphia, USA (PH) Canadian Museum of Nature, Ottawa, Canada (CANA) British Museum, London, UK (BM) Museum für Naturkunde der Humboldt University, Berlin, Germany (BHU) Botanical Museum, Copenhagen, Denmark (C) Museo de La Plata, Buenos Aires, Argentina (LP) National Science Museum, Tokyo, Japan (TNS) Oslo, Norway (not listed in Index Herbariorum, IMBB) Dr. Henri Van Heurck Museum, Antwerpen, Belgium (AWH); British Museum, London, UK (BM); Friedrich-Hustedt-Arbeitsplatz für Diatomeenkunde, Bremerhaven, Germany (BRM); Naturhistorisches Museum Wien, Austria (W) Friedrich-Hustedt-Arbeitsplatz für Diatomeenkunde, Bremerhaven, Germany (BRM) National Science Museum, Tokyo, Japan (TNS) Swedish Museum of Natural History, Stockholm, Sweden (S-PA-k) British Museum, London, UK (BM) Friedrich-Hustedt-Arbeitsplatz für Diatomeenkunde, Bremerhaven, Germany (BRM) Philadelphia Academy of Natural Sciences, Philadelphia, USA (PH) Paris Museum, Paris, France (PC); Philadelphia Academy of Natural Sciences, Philadelphia, USA (PH) Botanical Museum, Copenhagen, Denmark (C) Philadelphia Academy of Natural Sciences, Philadelphia, USA (PH) Paris Museum, Paris, France (PC); Philadelphia Academy of Natural Sciences, Philadelphia, USA (PH) Botanical Museum, Copenhagen, Denmark (C) Dr. Henri Van Heurck Museum, Antwerpen, Belgium (AWH); Paris Museum, Paris, France (PC); Philadelphia Academy of Natural Sciences, Philadelphia, USA (PH); Swedish Museum of Natural History , Stockholm, Sweden (S) Farlow Herbarium, Cambridge, USA (FH); Field Museum of Natural History, Chicago, USA (F); New York Botanical Garden, New York, USA (NY) British Museum, London, UK (BM); Farlow Herbarium, Cambridge, USA (FH); Canadian Museum of Nature, Ottawa, Canada (CANA) Dr. Henri Van Heurck Museum, Antwerpen, Belgium (AWH); British Museum, London, UK (BM); Philadelphia Academy of Natural Sciences, Philadelphia, USA (PH); Naturhistorisches Museum Wien, Austria (W) British Museum, London, UK (BM)

Cleve P.T. Croasdale H. Dickie G. Ehrenberg C.G. Foged N. Frenguelli J. Fukushima H. Gran H.H. Grunow A.

Heiden H. Kobayashi T. Kolbe R.W. Hooker J.D. Hustedt F. Lagerstedt N.G.W. Manguin E. Østrup E. Patrick R. Peragallo M. Petersen J.B. Petit P.

Prescott G.W. Ross R. Van Heurck H.

West G.S.

6. AQUATIC INVERTEBRATES AND HIGH LATITUDE PALEOLIMNOLOGY

OLE BENNIKE ([email protected]) Geological Survey of Denmark and Greenland Øster Voldgade 10 1350 Copenhagen K Denmark KLAUS P. BRODERSEN ([email protected]) Freshwater Biological Laboratory University of Copenhagen 51 Helsingørsgade 3400 Hillerød Denmark ERIK JEPPESEN ([email protected]) Department of Freshwater Ecology National Environmental Research Institute Vejlsøvej 25 8600 Silkeborg Denmark and Department of Plant Biology University of Aarhus Nordlandsvej 68 8240 Risskov Denmark and IAN R. WALKER ([email protected]) Departments of Biology and Earth and Environmental Sciences Okanagan University College 3333 College Way Kelowna, British Columbia V1V 1V7, Canada

Key words : Invertebrate remains, Paleolimnology, Paleoecology, Climatic change, Zoogeography, Arctic, Antarctic, Canada, Greenland 159 R. Pienitz, M.S.V. Douglas and J.P. Smol (eds), 2004. Long-term Environmental Change in Arctic and Antarctic Lakes. Springer. Printed in the Netherlands.

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Introduction Almost all major groups of invertebrate animals may leave remains in sediments. An excellent review of animal remains (mainly invertebrates in Quaternary lake and bog deposits) was published by Frey (1964). Since that time, studies of invertebrate remains in lake sediments have undergone a tremendous development (Walker 1993). Much of this work has focussed on changes in trophic status, salinity and lake type. However, much work has also been devoted to paleoclimatic studies. Johansen (1904) was the first to use invertebrate remains for paleoclimatic reconstruction, but his work was much criticized and debated (Johansen et al. 1906). History repeated itself when invertebrate remains were reintroduced to paleoclimatology in the 1980s (Walker and Mathewes 1987; Warner and Hann 1987; Walker et al. 1991). However, just as most terrestrial plants and animals have northern geographical range limits, so do most limnic plants and animals. As these limits are sensitive to temperature, shifts in their northern limits can be used in paleoclimatic reconstructions, although other factors may also be involved. David Frey’s comprehensive review from 1964 did not mention invertebrate remains from the Arctic or the Antarctic, and there are still few studies from these vast regions devoted to invertebrate remains in Quaternary deposits. However, in studies of the icefree parts of Greenland, there has been a tradition that some invertebrate remains have been included in the more detailed paleobotanical studies, and some information on invertebrate remains frequently appears in Greenland's Quaternary literature. Lakes in the Arctic (and Antarctic) hold an enormous potential for investigations of invertebrate remains. Such studies would no doubt greatly enhance our understanding of lake development, and also shed light on the history of the animals, including especially long-term shifts in their former geographical distribution. In this chapter we restrict the discussion mainly to areas north of the arctic treeline and south of the antarctic treeline. We focus on the Holocene (the last ca. 11,500 cal. yr BP) but include some notes on older Quaternary and Neogene sediments. Radiocarbon dates are calibrated into calendar years, and uncalibrated dates are not discussed. Lake sediments contain abundant remains of land plants, such as pollen, leaves, seeds, moss remains, etc. Remains of land animals may also be preserved in lake sediments, and remains of terrestrial species of Coleoptera, Hemiptera, Hymenoptera, Lepidoptera and Arachnida have been recovered from Holocene lake sediments (e.g., Bennike 2000), but such remains are not reviewed here. Lakes are common over large parts of the Arctic, and the sediments of the lakes normally contain remains of aquatic invertebrates. However, pre-Holocene Quaternary lake sediments are rare in the Arctic. On Baffin Island, interglacial organic lake gyttja has been reported from seven lakes (Miller et al. 2002). In one sequence, seven invertebrate taxa have been identified (Miller et al. 1999). In South Greenland, late-glacial lake sediments have been found in three lakes, and seven invertebrate taxa were reported by Bennike and Björck (2000). Limnic invertebrates have been found in various other sediment types of interglacial and Neogene age and provide some insight into lake biota over longer time spans. Many different kinds of invertebrate remains are well preserved and often very abundant in sediments from arctic freshwater lakes. Many invertebrates are excellent paleoindicators because they are: (1) widely distributed (ubiquitous) in the arctic and subarctic region; (2) identifiable to subfamily, genus or species (-group) level, using

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available identification literature; and (3) a diverse group adapted to particular ecological conditions, thus closely reflecting the environmental (lake-ecological) conditions at the time of sedimentation. A short life cycle and a high mobility allow invertebrates to respond rapidly to environmental change (e.g., climate), resulting in distributions that are probably in near equilibrium with environmental conditions. Notes on different zoological indicators present in lake sediments Protozoa Shells of testate amoebae and some other protozoans may be present in huge numbers in peat, but lake deposits may also contain abundant tests of these rhizopods. The use of these small, unicellular protists has been reviewed by Beyens and Meisterfeld (2001) and by Douglas and Smol (2001). Few subfossil records are available from the Arctic and Antarctic, and only a single paleolimnological study devoted to testate amoebae has been published (Dallimore et al. 2000). Ten species were recorded from three samples in the upper part of a sediment core. Distinct changes were observed between the samples and these may be related to geomorphological and hydrological processes. A few taxa of limnic agglutinated testate amoebae, identified by L. Beyens, were observed in two Holocene lake sequences in West Greenland (Eisner et al. 1995; Bennike 2000), but no systematic analyses were conducted. Tests of the rhizopod Assulina survive normal treatment of sediment samples for pollen analysis, and have been reported from western Greenland (Fredskild 1973, 1983, 1985). According to Fredskild (1973), the species present is probably A. muscorum. Assulina spp. live on the surface of bogs, and their tests were presumably washed out into the lakes. In addition, four types of protozoan plates tentatively referred to four genera, were recorded from a Holocene peat deposit in northwestern Greenland (Brown et al. 1994). In some samples, protozoan plates constituted over 80% of the siliceous microfossils found in the peat. Another Holocene peat deposit, from Edgeøya, Svalbard, provided evidence for a local hydrological change dated to 4500-4300 yr BP (Beyens and Chardez 1987). Rühland et al. (2000) studied siliceous protozoan plates along with other microfossils in a peat core from north-central Siberia from just south of the treeline. The core penetrated deposits from an open water environment. The deposits were devoid of protozoan plates, whereas un-differentiated protozoan plates dominated the siliceous microfossil assemblages in the peat. This study showed the usefulness of protozoans in analyses of surface wetness in peatlands, at least in the Subarctic. Protozoan plates have been analysed from surface sediment samples along a transect from southern Finland to northern Norway, but no noticeable geographical trends were found in the assemblages (Pienitz et al. 1995). Porifera There are few records of freshwater sponges from the Arctic (Frost 2001). To what degree this reflects the dearth of paleolimnological research versus a limited distribution or abundance of Porifera at high latitudes is unclear. Surface sediment samples from

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West Greenland commonly contain sponge gemmules (O. Bennike, unpublished data), and several species are present in western and southern Greenland (O. Tendal, pers. comm. 1999). In western Greenland, gemmules of Spongilla lacustris were recorded from a lake sequence (Eisner et al. 1995). Fredskild (1983) reported sponge spicules from Johannes Iversen Sø, but during routine pollen analyses of lake sediments, spicules are dissolved by hydrofluoric acid treatment. Sponge gemmules were reported from Holocene peat from the Northwest Territories, Canada, by Vardy et al. (1997). Turbellaria Cocoons of Rhabdocoela flatworms (Figure 1) may be present in high concentrations in lake sediments. A key for such fossils (called Neorhabdocoela oocytes) was published by Haas (1996). One (large) type has been photographed in Greenland sediments (Bennike 1995). This type has also been reported from other parts of Greenland and from Canada (e.g., Björck et al. 1994a; Miller et al. 1999; Bennike 2000; Bennike et al. 2002). Neorhabdocoela egg capsules were found in Holocene peat from the Northwest Territories (Vardy et al. 1997). Rotifera Published records of rotifer remains from Holocene lake sediments are scarce and thus far restricted to western Greenland, where Filinia was reported (Fredskild 1983, 1985). However, in recent studies of surface sediments, rotifer remains were frequently recorded from lakes in western and northeastern Greenland. Resting eggs of the genera Filinia and Brachionus were found (E. Jeppesen et al., unpublished data). Oligochaeta Egg cocoons of Oligochaeta were recorded in Greenland lake sediments by Bennike (1995), who suggested that they might belong to Tubifex sp. Mollusca Only a few species of freshwater bivalves and gastropods occur in the Arctic (Miller and Tevesz 2001). In Greenland, three species were found, and fossil remains have been recovered from a few lakes in western Greenland (Bennike 2000). The known temporal range of the small bivalve Pisidium steenbuc hii extends back to the early Holocene, whereas the gastropods Lymnaea vahlii and Gyraulus arcticus extend back to the midHolocene. The species are thought to be endemic to Greenland, but their taxonomic relation to other species needs to be established.

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D

B

C

E

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Figure 1. Scanning electron microscope photographs of invertebrate remains. (A) Cocoon of Rhabdocoela indet. from a Holocene lake deposit in West Greenland (from Bennike 1995). (B) Shell of Ilyocypris bradyi (Crustacea: Ostracoda) from a mid-Holocene lake deposit in West Greenland (from Bennike 2000). (C) Shell of Potamocypris parva (Crustacea: Ostracoda) from a Holocene lake deposit in West Greenland. (D) Statoblast of Plumatella repens (Bryozoa: Phylactolaemata) from a Holocene lake deposit in West Greenland (from Bennike 1995). (E-F) Statoblasts of Cristatella muced o (Bryozoa: Phylactolaemata) from Pliocene deposits in North Greenland. Scale bars: 200 Pm.

Arthropoda Branchiopoda The use of Branchiopoda and Cladocera in paleolimnological studies has been reviewed by Korhola and Rautio (2001). Skeletal remains of Lepidurus arcticus are commonly found in lake sediments from the Arctic. The mandibles are rather heavily sclerotised, but other parts of the skeleton may also be preserved. In addition, eggs are occasionally found. L. arcticus has been recorded from many parts of Greenland, and its fossil range extends back to the earliest Holocene (Fredskild et al. 1975; Fredskild 1983, 1985; Kelly and Bennike 1992; Bennike and Funder 1997; Bennike et al. 1999). In Svalbard, it has also been reported from the earliest Holocene (Bennike and Hedenäs 1995; Wohlfarth et al. 1995). Lepidurus has also been found from a number of interglacial and Pliocene sites in Greenland and Canada (Meldgaard and Bennike 1989; Bennike 1990; Bennike and Böcher 1992, 1994; Miller et al. 1999). Without supra-anal plates, the

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species identification is uncertain, and, in these instances, the remains have been referred to as Lepidurus sp. or Lepidurus cf. arcticus. In several lakes in western Greenland, a distinctive succession has been recorded. L. arcticus mandibles show an early peak shortly after the onset of organic sedimentation in the lake. Lake formation may result from deglaciation of the area, or from isolation of marine bays from the sea following relative land emergence due to crustal rebound. The early peak is followed by much lower abundances, and in some cases L. arcticus apparently disappeared from the lake, or even from the whole region. This development has been attributed to rising water temperatures (Fredskild 1983). However, L. arcticus is an important prey for fish, and may suffer a severe decline or extermination once fish colonise a lake. The observed decline may therefore alternatively reflect fish colonisation of the lake (Bennike 2000). The role of fish is evident from surface sediment studies of 41 lakes in northeastern Greenland (Jeppesen et al. 2001a). L. arcticus remains were found in only one of the lakes with fish, and here only in very low numbers. The genus Branchinecta is widely distributed and occurs both in the Arctic and the Antarctic. We are only aware of fossil Branchinecta records from Antarctica. Fossil eggs from mid- to late Holocene lake deposits on James Ross Island were identified as B. gaini (Björck et al. 1996). This crustacean does not occur on the island at present, presumably because of a too short ice-free season. Eggs of Branchinecta were also recorded from Signy Island by Jones et al. (2000). In the Arctic, Branchinecta is sensitive to fish predation and is most common in fishless ephemeral freshwater ponds. Yet, sediments from such environments are rarely analysed for invertebrate remains, which may explain their apparent absence in the arctic record. Thus, in a recent study of surface sediments from 27 lakes in western Greenland, remains of Branchinecta were found in 10 lakes, all without fish but with a maximum depth ranging from 1 to 30 m (E. Jeppesen et al., unpublished data). Cladocera Remains of cladocerans may constitute a major part of lake sediments, and are often present in large quantities. Most common are shells, head shields (Figure 2), postabdomens and other parts of the exoskeleton. Ephippia may also occur in large numbers. Exoskeletal remains can often be identified to the species level. In past decades, many studies of Cladocera remains have been conducted as part of paleolimnological investigations of Holocene lake sequences, but so far few studies have been published from the Arctic. However, in several of the studies of pollen and macrofossils from Greenland lake and peat sediments, remains of some cladoceran species have been recorded. A pioneering study was conducted by Fredskild et al. (1975), who reported seven species, mainly from southernmost Greenland. A total of 65 samples, sieved through a 0.4 mm mesh, were analysed, and the frequency of the remains was recorded simply as "rare", "frequent" or "common". Based on data from analyses of lake deposits from western Greenland, Fredskild (1983, 1985) published a series of pollen and macrofossil diagrams that included some information on cladocerans. Since then a few studies have included quantitative records of cladocerans from the entire Holocene or the top sediment (e.g., Björck et al. 1994a; Bennike 2000; Jeppesen et al. 2001a).

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It is also known, from contemporary data, that species richness is generally poor in northern and eastern Greenland (Blake et al. 1992; Jeppesen et al. 2001a,b and unpublished data) as well as in Iceland (Einarsson 1982), while it is somewhat higher in western Greenland, which is located closer to the North American continent (Lauridsen et al. 2001; E. Jeppesen et al., unpublished data). Thus, in a survey of cladoceran remains in surface sediments from 15 lakes, species richness averaged 9.1 in western Greenland, 3.0 in northeastern Greenland, 8.4 in Iceland, 11.8 in the Faeroe Islands and 21 in sediments from Denmark, southern Sweden, southern Finland and Estonia (E. Jeppesen et al., unpublished data). In accordance with these results, Bennike and Funder (1997) recorded three species in Holocene sediments from eastern Greenland. One of these, Acroperus harpae , is a relatively thermophilous species. Its fossil range here extends back to ca. 7000 cal. years BP, but in southern Greenland its range extends back to 13,500 cal. yr BP (Bennike and Björck 2000). Likewise, Einarsson (1982) recorded nine species in a detailed study of a 7 m core from eutrophic Lake Mývatn, Iceland. From northern Greenland, only two lake sequences have been subjected to a detailed study of invertebrates (Fredskild 1985; Blake et al. 1992). Three Cladocera taxa were present, but one of these (Bosmina sp.) was only recorded at trace levels in a single sample. Some data on Cladocera from peat and pond deposits are also available. Alona guttata (Figure 2) has been reported from two early or mid-Holocene sites in northern Greenland (Bennike 1983; Frey 1991); these may represent extralimital occurrences (occurrences outside the present geographical range), since there are no other records of this species so far north (Røen 1962, 1992). A few records of cladocerans exist from before the Holocene. In southernmost Greenland, records from the late-Glacial extend 14,000 years back in time. Five Cladocera taxa were reported from these late-glacial sediments (Bennike and Björck 2000). Two taxa occurred before and after the Younger Dryas cooling. Studies of last interglacial deposits from Greenland have so far only recorded ephippia of Daphnia and Simocephalus vetulus (Bennike et al. 2000; Hedenäs and Bennike 2003). The latter represents a southern extralimital faunal element of which many others have already been reported from this time period (Bennike and Böcher 1992, 1994). A single sample from the Late Pliocene Kap København Formation in northern Greenland contained various remains of cladocerans, and 12 taxa were identified by Røen (1988). The most surprising record was of Pleuroxus, which has not been recorded previously in contemporary samples from arctic lakes, but seven of the other taxa also represent extralimital species. Cladocerans are useful paleoindicators as different species dominate in different environments (Frey 1986; Hofmann 1987; Hann 1989; Jeppesen et al. 2001b; Korhola and Rautio 2001). Some are mainly found in the pelagic zone, some are sediment dwellers and others are associated with macrophytes. Moreover, the cladoceran species composition is influenced by changes in environmental variables such as salinity, pH and temperature. Cladoceran remains in the sediment have been used as indicators of changes in lake productivity, the relative importance of benthic and pelagic production, lake depth, salinity, macrophyte abundance, fish predation pressure, temperature and microevolution (Frey 1986; Lotter et al. 1997; Jeppesen et al. 2001b, 2003a). Increased accumulation of cladocerans may indicate increased productivity, whereas changes in the ratio of pelagic and benthic forms have been attributed to fluctuating

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lake-levels (Whiteside and Swindoll 1988). Recently, Korhola et al. (2000) developed a transfer function between cladocerans and water depth from surface sediments of 53 subarctic Fennoscandian lakes, but, as also discussed by the authors, the relation must be used with caution as depth and temperature in the lakes were highly correlated. To this should be added that a number of other factors might influence the ratio of benthic to pelagic forms, such as acidification, eutrophication and variation in predation pressure from fish (Jeppesen et al. 2001b). Since many arctic lakes are nutrient-poor and clear, ample possibilities of developing transfer functions relating plant-associated cladocerans to macrophyte abundance exist. To date, cladoceran fossils have only been used to describe qualitatively the changes in abundance of plants in high latitude lakes (Einarsson 1982; Duigan and Birks 2000). A nice example is the study of a sediment core from the highly dynamic eutrophic Lake Mývatn, Iceland, where changes in the chydorid assemblage have been used to identify major changes in abundance of plants including Cladophora (Einarsson 1982; Á. Einarsson et al., unpublished data). Work on cladoceran-plant abundance transfer functions for arctic and subarctic lakes is presently in progress (E. Jeppesen et al., unpublished data). Cladocerans are useful indicators of climate change as well (Battarbee 2000). Based on sediment data from 68 lakes along an altitudinal gradient from 300 to 2350 m in Switzerland, Lotter et al. (1997) developed a transfer function for mean summer air temperature using benthic and pelagic cladocerans. This function was subsequently adjusted by Duigan and Birks (2000) to describe temperature changes during the lateGlacial and early Holocene at Kråkenes, Norway. Likewise, Korhola (1999) developed a temperature transfer function for subarctic Fennoscandian lakes. Cladoceran remains from long Holocene sequences in Finnish Lapland were also studied by SarmajaKorjonen (1999) and Sarmaja-Korjonen and Hyvärinen (1999), the latter study providing evidence for lake-level changes. Changes in the life cycle of cladocerans may also be used as a climate indicator. In a comparative study of lakes from the Arctic to the temperate zone, Jeppesen et al. (2003a) showed that the ephippia to carapace ratio of Bosmina in the surface sediment was highly negatively related to summer mean air temperature and changed 3-4 orders of magnitude from the Arctic to the temperate zone. The primary explanation is that the number of parthenogenetic generations increases with temperature as the effect of other potential influential factors, such as food and predation, was comparatively low. The ephippia to carapace ratio has also been used as an environmental signal in the family Chydoridae (Sarmaja-Korjonen 2003). In a study of late Holocene sediments from Karluk Lake in southern Alaska, Bosmina concentration showed concordance with diatom, G15N and fish records (Finney et al. 2000). The cladoceran fossil record may also provide information on changes in fish predation. Fish feed selectively on large-bodied cladocerans, and remains of cladocerans are therefore indicative of fish abundance (Kitchell and Kitchell 1980; Jeppesen et al. 2001a). For lakes in Greenland, Jeppesen et al. (2003a) found a positive relation between the abundance of fish and the contribution of Bosmina resting eggs to the total sum of Daphnia and Bosmina resting eggs in the surface sediment. Generally, remains of large-bodied cladocerans were dominant in the Greenland lakes without fish while small-bodied forms dominated in lakes with fish, and remains of Daphnia were only occasionally found in lakes with fish (Jeppesen et al. 2003a), which is confirmed by contemporary data (Jeppesen et al. 2003b).

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A

B

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C

Figure 2 . Drawings of invertebrate remains. (A) Head shield of Daphnia pulex (Crustacea: Cladocera: Daphniidae) from a Holocene pond deposit in northern Greenland (from Frey 1991). Scale bar = 500 Pm. (B) Postabdomen and postabdominal claw of Alona guttata (Crustacea: Cladocera: Chydoridae) from a Holocene pond deposit in northern Greenland (from Frey 1991). Scale bar = 30 Pm. (C) Egg sac of Diaptomus castor (Crustacea: Copepoda) from a lake in Denmark (from Røen 1957). Scale bar = 500 Pm.

There is no doubt that cladocerans will be used extensively in the future to reconstruct past changes in environmental variables and trophic structure in cold lakes. Because of the relatively simple pelagic structure and the low species richness (in arctic lakes) and the diverse but species-specific habitat coverage (pelagic, benthic, plant-associated species), cladocerans function as ”on-off” indicator organisms at many ecological levels. Ostracoda Carbonate shells of ostracods are usually only found in carbonate-rich sediments, such as lake marls or calcareous gyttja (Holmes 2001). Even if the carbonate shells are dissolved, the chitinous linings may still be preserved. In Greenland, eight different species of non-marine ostracods have been reported from Holocene lake deposits (Bennike et al. 2000). Two of these, Ilyocypris br adyi (Figure 1) and Sarscypridopsis aculeata, have not been recorded from the extant Greenland fauna. They may represent southern extralimital species that have disappeared from the subcontinent as a consequence of cooling during the past millennia. A third species, Potamocypris parva (Figure 1b), is considered endemic to Greenland. The oldest remains found so far have an age of around 6000 years (Bennike 2000), but the fossil record can probably be extended backwards by analysing sediment cores from other lakes. This species must have evolved in less than 10,000 years, since the lakes it inhabits were completely glaciated before then.

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Copepoda Copepods constitute a major part of the zooplankton of arctic lakes, but their exoskeleton does not normally preserve in lake sediments. There is a single record of egg sacs from Holocene sediments in a small, shallow lake in western Greenland (Bennike 1998). The egg sacs were probably produced by Diaptomus castor , whose resting eggs have an unusually thick egg membrane (Figure 2c). This species is most common in ponds, although there are also a few records from lakes. Copepod egg sacs (called egg shells) have also been reported from alpine lakes (Knapp et al. 2001). The dearth of copepod egg sac records is probably attributable to the fact that up to now sediment has only been analysed from a limited number of small lakes. Eggs of Boeckella were recorded from Holocene lake sediments from Signy Island (Jones et al. 2000).

Colymbetes dolabratus Holocene

Pleistocene

Pleistocene and Holocene

Hydroporus morio

Gyrinus opacus

Modern range

Figure 3. Modern geographical ranges and fossil records of three species of water beetles in Greenland. The modern ranges are based on Böcher (1988).

Coleoptera Many species of beetles, such as predaceous diving beetles and whirligig beetles, live in lakes and ponds for most of their life cycle (Elias 2001). A few of these have been recorded from Neogene deposits in the north, and many remains have been identified to the species level (e.g., Böcher 1995; Matthews and Fyles 2000). Also, Pleistocene interglacial sites have produced remains of water beetles (e.g., Blake and Matthews 1979; Matthews et al. 1986; Morgan et al. 1993; Bennike and Böcher 1994). The two diving beetles Colymbetes dolobratus and Hydroporus morio have been recorded from a rather large number of Holocene deposits in Greenland (Figure 3), whereas the

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whirligig beetle Gyrinus opacus has only been recorded from two lake sequences (Bennike et al. 2000). On Svalbard, the diving beetle Agabus bioust ulatus has been found in early Holocene deposits (Wohlfarth et al. 1995). Diptera Non-biting midges (Chironomidae) are amongst the best adapted and most successful insect groups in the Arctic (Oliver 1968). Some species live close to their limit of existence in the cold arctic region, while others have adapted to a life at low temperatures, a short growth season, a temporally restricted emergence period, and low food supply. Most genera complete their larval stage in freshwater habitats or in humic soils, often well protected from extremely low temperatures. Chironomid life cycles can extend up to several years. The adults, however, are typically short-lived and emerge after a certain number of degree-days have accumulated. Some species are parthenogenetic and do not invest expensive time and energy in sexual reproduction. Chironomids play a significant role in degradation and recycling of organic matter in arctic lakes and streams. Larvae, pupae and winged adults are all important food items for predaceous invertebrates, fish and birds associated with freshwater habitats and their catchments. In addition to the chironomids, remains of some other Diptera (e.g., phantom midges - Chaoboridae) have occasionally been reported from arctic lake sediments (Frey 1964; Walker 2001). In the last decade, extensive surface sampling programmes in the Arctic, and along transects spanning the arctic treeline have been initiated in Canada, Greenland, northern Europe, and Siberia. These sampling programmes are providing extensive databases for the interpretation of chironomid paleoecological studies, and especially for the development of transfer functions facilitating quantitative reconstructions of air or water temperature, or other environmental variables. In the beginning of the 1990s, Walker et al. (1991) published the first quantitative chironomid-temperature transfer function from Labrador in eastern Canada. This surface sample transect extended from the boreal forest in southern Labrador, north into the low arctic tundra. The transect has since been extended to Devon Island in the Canadian arctic archipelago (Walker et al. 1997), and D.R. Francis and I.R. Walker (unpublished data) have completed analyses on an extensive surface sample set from Baffin Island. In western Canada, Walker and MacDonald (1995) examined the distributions of midges across treeline near Yellowknife, Northwest Territories. Similar analyses are in progress across treeline in Alaska, Yukon and the Mackenzie River delta (Walker et al. 2003; E. Barley et al., unpublished data). A number of quantitative studies of chironomid assemblages spanning the treeline in the western Palaearctic have now been published. Thus Olander et al. (1997) presented a chironomid-surface water temperature transfer function based on 30 lakes in northern Finnish Lapland. The transect encompasses four different vegetation zones, and has since been expanded with 23 additional lakes, now including a total of 53 lakes (Olander et al. 1999). More recently, Larocque et al. (2001) have published a transfer function based on a 100-lake data set from the boreal and alpine zones of northern Swedish Lapland. This model indicated that chironomids can provide useful estimates of past changes in mean July air temperature with an estimated error of prediction of ± 1.13°C. Walker et al. (1997), however, note that the prediction errors of such models

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are likely to be under-estimated when based on training sets spanning such short climatic gradients. In northeastern Siberia, Porinchu and Cwynar (2000) analysed chironomid remains from the surficial sediments of 31 lakes. Using canonical variate analysis (discriminant functions) they were able to differentiate among lakes in three vegetation zones, suggesting that subfossil chironomids may be used to infer past changes in treeline position. Chironomid distribution in southwestern Greenland was studied by Brodersen and Anderson (2002). They investigated 47 lakes along a gradient extending from the coast to the continental climatic region, adjacent to the margin of the inland ice. They also obtained high-resolution water temperature data using automatic data-loggers (Brodersen and Anderson 2000). Trophic parameters such as total nitrogen and total phosphorus came out as the strongest variables explaining the chironomid distribution patterns, closely followed by the third strongest variable, mean July surface water temperature. Because of the strong trophic gradient, it was recommended to use, for example, a discriminant function to match subfossil samples into one of eight different lake types, along with interpretation of chironomid-inferred paleotemperatures from the regional training set. Chironomid exuviae (the cast pupal skins) were collected from a subset of the western Greenland lakes to verify the head capsule identifications (Brodersen et al. 2001). Presence-absence data derived from the exuviae were sufficient to distinguish among four different lake types: shallow ponds, oligosaline nutrient-rich lakes, dilute nutrient-rich lakes, and dilute nutrient-poor lakes. Among published chironomid training sets, the estimated temperature optima are in very good agreement. Lotter et al. (1999) compared a North American with a European surface sediment data set and found that the North American and European temperature optima were highly correlated, and that the models revealed similar patterns of lateglacial temperature conditions. The rank-order of temperature optima for four subarctic data sets is shown in Figure 4a. Genera and taxa such as Heterotrissocladius, Micropsectra, Mesocricotopus, Hydrobaenus and Corynocera oliveri are considered to be indicators of cold environmental conditions, while taxa such as Microtendipes, Dicrotendipes, Polypedilum, Chironomus and Pseudochironomus indicate warmer temperatures. A general bias in most training sets, however, is the strong correlation among environmental variables also influencing the chironomid distribution (Brodersen and Anderson 2002). The overall pattern is that cold lakes are also deep and unproductive, whilst warm lakes are shallow and (in the Arctic moderately) highly productive. In both the Swedish and the Finnish data sets, sediment organic content was the strongest explanatory environmental variable, and in the Finnish data set lake depth (Figure 4b) was such a strong variable that a chironomid-depth transfer function was developed (Korhola et al. 2000). Evidently, the chironomid response to climate is complex and indirect (Walker and Mathewes 1989; Battarbee 2000). Interpretation of chironomid-inferred paleotemperature in the arctic and subarctic regions should, thus, be weighted against the knowledge of other ecological optima such as depth, trophic conditions and oxygen availability (Figure 4). The quality of the environmental data used in quantitative transfer functions in high-latitude and remote areas is highly dependent on logistics and resources (Korhola et al. 2001; Seppälä 2001). Olander et al. (1999) noted that Finnish chironomid distributions were better

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correlated to air than water temperature. However, the air temperature data used in these studies is generally of much higher quality and easier to obtain than the water temperature data; thus, the strength of the correlations is probably a reflection of temperature data quality rather than underlying mechanisms regulating midge distributions. In contrast to the extensive surface sampling programs, few downcore studies have yet been published from arctic regions (Walker 2001). This situation is changing as cores from many parts of the Arctic are currently being analysed. Apart from the few midge records provided in Bennike et al. (2000) and Smith (2000), more detailed stratigraphic studies have already been published from Norway, Iceland, Svalbard and Russia. Investigations are also underway in Canada (Yukon, northern Québec and Baffin Island) and Greenland. Alm and Willassen (1993) have analysed a sediment core from Nedre Æråsvatn, on Andøya in northern Norway. Porinchu and Cwynar (2002) describe the late Quaternary history of midge communities at a tundra site near the lower Lena River in northeast Siberia, and long-term changes in Lake Mývatn (Iceland) are reported by Einarsson and co-workers (Einarsson 1981, 1982; Einarsson and Haflidason 1988; Einarsson et al. 1988; Gardarsson et al. 1988). From the island of Spitsbergen (Svalbard), Brooks and Birks (2004) took sediment cores from three lakes to detect ecosystem responses to natural small-amplitude climatic changes over the last 700 years. They interpret the chironomid stratigraphies as indicating a lake-ecological climate signal corresponding to the onset and end of the "Little Ice Age". The signal from one lake also indicates a response to human settlement, beginning AD 1930 and ending AD 1988. Brodersen and colleagues are currently undertaking multi-proxy studies of sediment cores covering the last 7000 to 9000 years from western Greenland. Preliminary results reveal a strikingly good concordance between chemical and biological paleo-proxies. The abrupt cold event 8200 yr BP (Alley et al. 1997) is clearly evident in the chironomid stratigraphy (and other proxies), with a distinct shift in dominance from Psectrocladius to Micropsectra, and back again to Psectrocladius dominance. In an analysis of an interglacial sample from eastern Greenland for chironomids and other invertebrates, Simuliidae and Diamesa, which are restricted to streams, were found (Bennike et al. 2000). The sample also contained one head capsule of Brillia/Euryhapsis that represents a southern extralimital element. Interglacial midge remains have also been reported from northwestern Greenland (Brodersen and Bennike 2003). Francis et al. (unpublished data) have completed analyses on two lake sediment cores from eastern Baffin Island. Sediments from the last interglacial are represented in the core base and include remains of Chaoborus (Chaoboridae), and other warm water midges. These fossils are consistent with other evidence that the height of the last interglacial was warmer than at any time during the Holocene. From the ca. 2.4 million years old Kap København Formation in northern Greenland, 16 taxa of Diptera have been reported (Böcher 1995). The fauna includes head capsules of Simuliidae, which indicate fast running water, and 12 taxa of Chironomidae. The latter comprise Axarus sp. and Corynocera ambigua that represent extralimital occurrences.

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W. Greenland

28 26 24

N. Finland NE. Canada

22 20

TN optima - rank number

Temperature optima - rank number

N. Sweden

18 16 14 12 10 8 6 4 2

A

VWHO optima - rank number

30

Depth optima - rank number

32

B

C

D

Pseudodiamesa sp. Abiskomyia sp. Paracladopelma sp. Eukiefferiella/Tvetenia Mesocricotopus sp. Protanypus sp. Corynocera oliveri Micropsectra spp. Hydrobaenus sp. Heterotrissocladius sp. Arctopelopia sp. Orthocladius consobrinus Stictochironomus Corynoneura sp. Paratanytarsus sp. Tanytarsus chinyensis Cricotopus sp. Corynocera ambigua Stempellinella/Zavrelia Heterotanytarsus Procladius sp. Tanytarsina Tanytarsus lugens Sergentia Pentaneurini Stempellina Psectrocladius spp. (avg) Tanytarsus gracillentus Zalutschia spp. (avg) Allopsectrocladius-gr Ablabesmyia sp. Pagastiella Parakiefferiella bathophila Microtendipes Dicrotendipes spp. Cryptochironomus Cladopelma spp. Cladotanytarsus sp. Chironomus spp. Polypedilum sp. Monopsectrocladius-gr Lauterborniella/Zavreliella Pseudochironomus

0

Figure 4. (A) Diagram showing the ranked temperature optima for selected and common (shared) chironomid taxa from published arctic/subarctic studies. Northeastern Canada: Walker et al. (1997); northern Finland: Olander et al. (1997, 1999); northern Sweden: Larocque et al. (2001); western Greenland: Brodersen and Anderson (2002). (B) The ranked lake depth optima from the northern Finland data set (Korhola et al. (2000). (C) The ranked trophic optima (total nitrogen) from the western Greenland data set. D: The ranked optima to end-of-summer volume weighted hypolimnetic oxygen concentration (VWHO; Quinlan et al. 2001). B, C, D have the same horizontal axis as A.

In three million year old sediments from Meighen Island in the Canadian High Arctic, Walker and Matthews (unpublished data) have recorded several chironomid taxa, including Corynocera ambigua and Abiskomyia. Abiskomyia is a common arctic midge, which does not occur south of treeline. Corynocera ambigua is often abundant at low arctic sites, and is also known from scattered temperate locations in Europe, including both Denmark and Germany (Brodersen and Lindegaard 1999). In the Canadian Arctic, the present distribution of C. am bigua does not appear to extend farther north than southernmost Baffin Island. The Meighen Island assemblage, therefore, suggests a warmer Pliocene climate, similar to that presently found in the low Arctic, or close to present treeline, where the distributions of these two taxa now overlap. This inference

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corresponds well with studies of beetle and plant remains from the same site (Matthews 1974; Matthews and Ovenden 1990). Chironomid stratigraphies have recently been published for two sites separated by about 25 km in northernmost Finland, Tsuolbmajarvi and Toskaljarvi (Korhola et al. 2002; Seppä et al. 2002). The Toskaljarvi chironomid-temperature reconstruction is based on Olander et al.’s (1999) WA-PLS (weighted-averaging partial least squares) model, and indicates that the early and mid-Holocene were generally warmer than the late Holocene. Peak postglacial temperatures were recorded ca. 4.0 to 4.5 and 7.0 to 8.0 ka BP. Korhola et al. (2002) present two reconstructions for Tsuolbmajarvi. The WAPLS reconstruction indicates a similar postglacial trend towards cooler late Holocene temperatures, but with a slight warming after 2.0 ka BP. In contrast, a reconstruction for the same data set, using an alternative Bayesian modelling technique, suggests that temperatures were stable throughout the Holocene, with the exception of a prolonged cool interval between 4.0 and 6.0 ka BP. Given the apparent contradiction, additional work will be needed to better characterize Holocene climate in this region. In addition to these conventional fossil studies, isotope analyses of midge head capsules are also intriguing. Comparisons of AMS dates from chironomid head capsules with bulk sediment dates suggest that the latter may yield better chronologies for arctic lakes (Fallu et al. 2004). Furthermore, oxygen isotopes preserved in the chitin of chironomid head capsules offer a new paleotemperature proxy (Wooller et al. 2004). Trichoptera The larvae of caddisflies live in freshwater. Cases (from case building species), anal sieve plates and sclerites from the frontal parts of the larvae may preserve (Elias 2001). Larval sclerites of a few species were reported from Pliocene deposits in northern Greenland and northern Canada (Böcher 1995; Matthews and Fyles 2000). From Holocene lake deposits in Greenland and Svalbard, cases and sclerites have also been reported (Fredskild et al. 1975; Fredskild 1983, 1985; Wohlfarth et al. 1995; Bennike 2000). Larval cases from Holocene deposits from the northern part of Greenland probably belong to Apatania zonella , the only trichopteran species presently found in northernmost Greenland. Araneae Oribatida, or soil mites, are frequently encountered in lake deposits (Solhøy 2001). Most oribatids live in soil, but a few species are truly aquatic, and the majority of mite remains in lake deposits apparently come from these species. Certainly Hydrozetes lacustris is common, but little is known about its present distribution in the Arctic. During a study of oribatid remains from late-glacial and early Holocene sediments from Kråkenes in western Norway, four aquatic species and 34 non-aquatic species were recorded (Solhøy and Solhøy 2000). Fredskild et al. (1975) and Fredskild (1995) mentioned the occurrence of Hydrachnidae, but unfortunately these unexpected finds have not been documented. We are not aware of any other records of Hydrachnidae mite remains from arctic or antarctic lake sediments.

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Present day northern tree-line

Holocene

Last Interglacial

Neogene

Figure 5. Circumpolar map showing the position of the present-day arctic treeline (according to Hare and Ritchie 1972) and fossil records of the freshwater bryozoan Cristatella mucedo.

Bryozoa The statoblasts of freshwater bryozoans are composed of chitin and preserve well, and both ungerminated statoblasts and single valves are found (Figure 1; Francis 2001). Statoblasts of Cristatella mucedo are common in Neogene deposits from Canada and Greenland (e.g., Kuc 1973; Matthews et al. 1986, 1990; Bennike 1990; Fyles et al. 1994; Matthews and Fyles 2000; Figure 5), and have also been found in deposits from the last interglacial stage in Greenland (Bennike and Böcher 1992, 1994). In southern and southwestern Greenland, C. mucedo statoblasts have been recovered from Holocene lake deposits from five lakes (Fredskild et al. 1975; Fredskild 1983). In one of the lakes, statoblasts occurred in a surface sediment sample. Kuc (1973) and Vardy et al. (1997) mentioned finds from Holocene deposits in arctic Canada. Many of these finds may represent extralimital occurrences, because C. mucedo is said to have a northern range limit that coincides with the arctic treeline (Lacourt 1968). However, the presence of statoblasts in a Greenland lake indicates that it still lives there, although there are no extant Greenland records. Statoblasts of another freshwater bryozoan, Plumatella repens (Figure 1), have been recovered from seven Holocene lake sequences in western and eastern Greenland (Bennike 2000). Its statoblasts are rather inconspicuous, and they are probably more

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widespread than indicated by the limited number of finds. There is only one record of an extant colony of this animal from southernmost Greenland, but whether the fossil finds represent extralimital occurrences is unknown. Discussion Nutrient and pH status Sediment records have shown that many lakes in the north have experienced long-term oligotrophication and acidification after they were formed, either as a result of deglaciation or land uplift (Fredskild 1983). This may be especially pronounced in areas with acid bedrock, such as gneisses and granites. Aquatic macrophytes that prefer high lake water conductivity were abundant in the pioneer phase, and the sediments may also be extremely rich in colonies of the planktonic green algae Pediastrum bory anum, perhaps indicating higher primary productivity (Figure 6). This initial phase is followed by lower primary production and many of the macrophytes disappeared or became rare. The decrease in trophic state and primary production is also reflected in the invertebrate community, with most invertebrate taxa showing decreasing concentrations in the sediment with time and also changes in community composition reflecting different habitat preferences of the invertebrates (e.g., Fredskild 1983). Climate change Because of their great abundance in arctic lake sediments, invertebrate remains are potentially powerful indicators for paleolimnological studies. Many species have northern range limits in the Arctic, and past changes in these limits can be used as a climate proxy. However, this requires that the range limits of the species are well known, and also well defined. Another more promising approach is to use calibration sets based on analyses of invertebrate remains from surface sediments from a range of lakes. Such work is now in progress. Neogene lakes at high latitudes supported a much more diverse fauna than at present, which shows that pronounced shifts in range limits have taken place (e.g., Matthews et al. 1986, 1990; Morgan et al. 1993; Böcher 1995; Matthews and Fyles 2000). This reflects that Neogene temperatures in the far north were much higher than at present. A similar, though less extreme situation, pertains to the last interglacial (e.g., Bennike and Böcher 1992, 1994). Range shifts also occurred during the early to mid-Holocene thermal optimum (Figure 3). Chironomid-inferred salinity shifts have been reported in a saline lake from the Yukon (Pienitz et al. 1992). These salinity changes reflect changes in effective moisture and thus changes in climate. Zoogeography The study of invertebrate remains in lake deposits or other sediments is the only way to reconstruct shifts in geographical distributions of species backwards through time. With

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Spongilla sp.

Eurycercus glacialis

Cristatella mucedo

Daphnia pulex

Lepidurus arcticus

Pediastrum kawraiskyi

Pediastrum duplex

Invertebrates

Pediastrum angulosum

Pediastrum integrum/muticum

Pediastrum boryanum

Isoetes echinospora

Potamogeton pusillus

Hippuris vulgaris

Myriophyllum spicatum

Chara and Nitella

Green algae

Vascular plants

Potamogeton filiformis

0

Acidophilous diatom species (%)

Age (cal. years BP)

respect to the formerly glaciated regions at high latitudes, it is obvious that species had to immigrate to these regions. However, non-glaciated areas may have been more or less devoid of aquatic invertebrates because of the extreme temperatures. Modelling of present-day temperatures in boreholes on the Greenland Ice Sheet indicates a temperature about 25°C colder than at present during the last glacial maximum (DahlJensen et al. 1998). It appears that the numerous lakes, which formed at high latitudes after the last deglaciation, were rapidly invaded by some invertebrates. Some species could spread as aerial plankton, but many species must have relied on transport by birds. The problem of colonising newly formed lakes is particularly acute for lakes on oceanic islands, but even lakes in Greenland appear to have been rapidly colonised by some species (Bennike 1999; Figure 7).

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Figure 6. Model for the long-term development of lakes in the Godthåbsfjord area, southwestern Greenland, according to data by Fredskild (1983) and Foged (1977). Based on records of diatoms, pollen, seeds, fruits, leaf fragments, oospores, coenobia, mandibles, ephippia, statoblasts, spicules and shells. The plant species indicative of the highest conductivity and nutrient-rich water are placed to the left in the diagram.

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Chydorus arcticus Chironimidae Acroperus harpae Alona sp. Daphnia pulex Simocephalus vetulus Lepidurus arcticus Alona guttata Tonnacypris glacialis Plumatella repens Colymbetes dolabratus Apatania zonella Hydroporus morio Bosmina sp. Heterocypris incongruens Cristatella mucedo Eurycercus glacialis Pisidium steenbuchii Limnocythere sanctipatricii Ilyocypris bradyi Gyraulus arcticus Potamocypris parva Gyrinus opacus Lymnaea vahlii

Age (cal. ka BP)

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2 4 6 8 10 12 14 Figure 7 . Late-glacial and Holocene temporal range of different invertebrates in Greenland. Sources: Bennike 2000; Bennike and Björck 2000; Bennike et al. 2002; O. Bennike, unpublished data. Sediments from ca. 15 lakes were surveyed for this summary diagram.

Isolation and transgression of lakes Remains of invertebrates and aquatic macrophytes have been used to pinpoint isolation contacts in lakes situated below the marine limit (Björck et al. 1994a,b; Bennike 1995; Bennike et al. 2002). Such lakes became isolated from the sea as a consequence of relative land emergence due to isostatic rebound. Just after isolation from the sea, the lake ecosystems underwent rapid and drastic changes. Often a very pronounced peak in Cladocera abundance can be observed, indicating an initial short-lived productive stage, just after isolation from the sea. In southern Greenland, analyses have been performed on seven isolation basins that span the interval from the Allerød interstadial, through the

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Plumatella repens

4

Chironomidae indet.

Lepidurus arcticus

100

Cladocera indet.

Calitriche hermaphroditica

Isoetes lacustris

LAKE

Hippuris vulgaris

Macoma balthica

Pectinaria sp.

Foraminifera indet. Tricladida indet.

Depth (cm)

SEA

Batrachium confervoides

Younger Dryas stadial to early Holocene deposits (Bennike et al. 2002). Although the Younger Dryas in southern Greenland appears to have been anomalously mild (Björck et al. 2002), a number of thermophilous species did not appear until the early Holocene (Figure 7). Submarine Holocene lake deposits appear to be widespread in southern and western Greenland (Long et al. 1999; O. Bennike et al., unpublished data; Figure 8). Invertebrate remains can also be used to pinpoint transgression levels in sequences from such basins.

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Figure 8. Isolation of a lake basin following emergence of land and following marine transgression led to a sequence of lake deposits under- and overlain by marine sediments. The isolation and transgression contacts can be identified by analysing remains of invertebrates and plants (O. Bennike, unpublished data, from southern Greenland). Scale = number per 30 ml of wet sediment.

Summary Little research has been devoted to the study of aquatic invertebrate remains in lake sediments from the Arctic, and only a single record from the Antarctic is known to us. To some degree this reflects our limited knowledge of the taxonomy, autecology and geographical ranges of many taxa, and especially the many small and inconspicuous

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ones. However, more attention is given to invertebrates with the increasing interest in multi-disciplinary paleo-studies. This trend will probably accelerate as regional calibrations and transfer functions are developed, allowing for detailed, quantitative assessments of paleoenvironmental conditions from the remains of testate amoebae, chironomids, cladocerans and other invertebrates from high latitude lake and pond sediments. Acknowledgements We thank referee Steve Brooks, London, for his comments on the manuscript. References Alley R.B., Mayewski P.A., Sowers T., Stuiver M., Taylor K.C. and Clark P.U. 1997. Holocene climate instability: A prominent, widespread event 8200 yr ago. Geology 25: 483-486. Alm T. and Willassen E. 1993. Late Weichselian Chironomidae (Diptera) stratigraphy of Lake Nedre Æråsvatn, Andøya, northern Norway. Hydrobiol. 264: 21-32. Battarbee R.W. 2000. Palaeolimnological approaches to climate change, with special regard to the biological record. Quat. Sci. Rev. 19: 107-124. Bennike O. 1983. Palaeoecological investigations of a Holocene peat deposit from Vølvedal, Peary Land, North Greenland. Rapp. Grønlands Geol. Unders. 115: 15-20. Bennike O. 1990. The Kap København Formation: stratigraphy and palaeobotany of a PlioPleistocene sequence in Peary Land, North Greenland. Medd. Grønland, Geosci. 23: 85 pp. Bennike O. 1995. Palaeoecology of two lake basins from Disko, West Greenland. J. Quat. Sci. 10: 149-155. Bennike O. 1998. Fossil egg sacs of Diaptomus (Crustaceae: Copepoda) in Late Quaternary lake sediments. J. Paleolim. 19: 77-79. Bennike O. 1999. Colonisation of Greenland by plants and animals after the last ice age: a review. Polar Record 35: 323-336. Bennike O. 2000. Palaeoecological studies of Holocene lake sediments from West Greenland. Palaeogeogr. Palaeoclim. Palaeoecol. 155: 285-304. Bennike O. and Böcher J. 1992. Early Weichselian interstadial land biotas at Thule, Northwest Greenland. Boreas 21: 111-117. Bennike O. and Böcher J. 1994. Land biotas of the last interglacial/glacial cycle on Jameson Land, East Greenland. Boreas 23: 479-487. Bennike O. and Hedenas L. 1995. Early Holocene land floras and faunas from Edgeøya, eastern Svalbard. Polar Res. 14: 204-214. Bennike O. and Funder S. 1997. Macrofossil studies of Holocene lake sediments from Jameson Land, East Greenland. Geol. Greenland Surv. Bull. 176: 80-83. Bennike O. and Björck S. 2000. Lake sediment coring in South Greenland in 1999. Geol. Greenland Surv. Bull. 186: 60-64. Bennike O., Björck S., Böcher J., Hansen L., Heinemeier J. and Wohlfarth B. 1999. Early Holocene plant and animal remains from North-east Greenland. J. Biogeogr. 26: 667-677. Bennike O., Björck S., Böcher J. and Walker I.R. 2000. Quaternary arthropods from Greenland: a review with new data. Bull. Geol. Soc. Denmark 47: 111-134. Bennike O., Björck S. and Lambeck K. 2002. Estimates of South Greenland late-glacial ice limits from a new relative sea level curve. Earth Planet. Sci. Letters. 197: 171-186.

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Morgan A.V., Kuc M. and Andrews J.T. 1993. Palaeoecology and age of the Flitaway and Isortoq interglacial deposits, north-central Baffin Island, Northwest Territories, Canada. Can. J. Earth Sci. 30: 954-974. Olander H., Korhola A. and Blom T. 1997. Surface sediment Chironomidae (Insecta: Diptera) distributions along an ecotonal transect in subarctic Fennoscandia: developing a tool for palaeotemperature reconstructions. J. Paleolim. 18: 45-59. Olander H., Birks H.J.B., Korhola A. and Blom T. 1999. An expanded calibration model for inferring lakewater and air temperatures from fossil chironomid assemblages in northern Fennoscandia. The Holocene 9: 279-294. Oliver D.R. 1968. Adaptations of arctic Chironomidae. Ann. zool. fenn. 5: 111-118. Pienitz R., Walker I.R., Zeeb B.A. and Smol J.P. 1992. Biomonitoring past salinity changes in an athalassic sub arctic lake. Intern. J. Salt Lake Res. 12: 91-123. Pienitz R., Douglas M.S.V., Smol J.P., Huttunen P. and Meriläinen J. 1995. Diatom, chrysophyte and protozoan distributions along a latitudinal transect in Fennoscandia. Ecography 18: 429-439. Porinchu D.F. and Cwynar L.C. 2000. The distribution of freshwater chironomidae (Insecta: Diptera) across treeline near the lower Lena River, Northeast Siberia, Russia. Arct. Ant. Alp. Res. 32: 429–437. Porinchu D. F. and Cwynar L.C. 2002. Late-Quaternary history of midge communities and climate from a tundra site near the lower Lena River, Northeast Siberia. J. Paleolim. 27: 59-69. Quinlan R. and Smol J.P. 2001. Chironomid-based inference models for estimating end-ofsummer hypolimnetic oxygen from south-central Ontario shield lakes. Freshw. Biol. 46: 1529-1551. Røen U. 1957. Contributions to the biology of some Danish free living freshwater copepods. Biol. Skr. Kgl. Dan. Vid. Selsk. 9 (2): 101 pp. Røen U. 1962. Salt- og brakvandsdyr i grønlandske ferskvande. Grønland 62: 335-341. Røen U. 1988. Remains of early glacial Cladocera from North Greenland. Verh. Int. Ver. Limnol. 23: 845-847. Røen U. 1992. Review of Greenlandic species of Alona Baird 1850, with descriptions of three new species (Cladocera: Chydoridae: Aloninae). Steenstrupia 18: 101-109. Rühland K., Smol J.P., Jasinski J.P. and Warner B.G. 2000. Response of diatoms and other siliceous indicators to the developmental history of a peatland in the Tiksi Forest, Siberia. Arct. Ant. Alp. Res. 32: 167-178. Sarmaja-Korjonen K. 1999. Headshields of ephippial Chydorus piger Sars (Cladocera, Chydoridea) females from northern Finnish Lapland: a long period of gamogenesis? Hydrobiologia 390: 11-18. Sarmaja-Korjonen K. 2003. Chydorid ephippia as indicators of environmental change biostratigraphical evidence from two lakes in southern Finland. The Holocene 13: 691-700. Sarmaja-Korjonen K. and Hyvärinen H. 1999. Cladoceran and diatom stratigraphy of calcareous lake sediments from Kuusamo, NE Finland. Indications of Holocene lake-level changes. Fennia 177: 55-70. Seppä H., Nyman M., Korhola A. and Weckström J. 2002. Changes of treeline and alpine vegetation in relation to post-glacial climate dynamics in northern Fennoscandia based on pollen and chironomid records. J. Quat. Sci. 17: 287-301. Seppälä M. 2001. Unsatisfactory field data in a calibration model for inferring past temperatures from chironomid assemblages in northern Fennoscandia: a comment on Olander, Birks, Korhola and Blom. The Holocene 11: 613-615. Smith I. R. 2000. Diamictic sediments within high arctic lake sediment cores: evidence for lake ice rafting along the lateral glacial margin. Sedimentology 47: 1157-1179.

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Solhøy T. 2001. Oribatid mites. In: Smol J.P., Birks H.J.B. and Last W.M. (eds), Tracking Environmental Change Using Lake Sediments. Volume 4: Zoological Indicators. Kluwer Academic Publishers, Dordrecht, The Netherlands, pp. 81-104. Solhøy I.W. and Solhøy T. 2000. The fossil oribatid mite fauna (Acari: Oribetida) in late-glacial and early-Holocene sediments in Kråkenes Lake, western Norway. J. Paleolim. 23: 35-47. Vardy S.R., Warner B.G. and Aravena R. 1997. Holocene climate effects on the development of a peatland on the Tuktoyaktuk Peninsula, Northwest Territories. Quat. Res. 47: 90-104. Walker I. R. 1993. Paleolimnological biomonitoring using freshwater benthic macroinvertebrates. In: Rosenberg D.M. and Resh V.H. (eds), Freshwater Biomonitoring and Benthic Macroinvertebrates. Chapman and Hall, New York, pp. 306-343. Walker I.R. 2001. Chironomidae and related Diptera. In: Smol J.P., Birks H.J.B. and Last W.M. (eds), Tracking Environmental Change Using Lake Sediments. Volume 4: Zoological Indicators. Kluwer Academic Publishers, Dordrecht, The Netherlands, pp. 43-66. Walker I.R. and Mathewes R.W. 1987. Chironomids, lake trophic status, and climate. Quat. Res. 28: 431-437. Walker I.R. and Mathewes R.W. 1989. Chironomidae (Diptera) remains in surficial lake sediments from the Canadian Cordillera: analysis of the fauna across an altitudinal gradient. J. Paleolim. 2: 61-80. Walker I.R. and MacDonald G.M. 1995. Distributions of Chironomidae (Insecta: Diptera) and other freshwater midges with respect to treeline, Northwest Territories, Canada. Arct. Alp. Res. 27: 258-263. Walker I.R., Smol J.P., Engstrom D.R. and Birks H.J.B. 1991. An assessment of Chironomidae as quantitative indicators of past climatic change. Can. J. Fish. Aquat. Sci. 48: 975-987. Walker I.R., Levesque A.J., Cwynar L.C. and Lotter A.F. 1997. An expanded surface-water palaeotemperature inference model for use with fossil midges from eastern Canada. J. Paleolim. 18: 165-178. Walker I.R., Levesque A.J., Pienitz R. and Smol J.P. 2003. Freshwater midges of the Yukon and adjacent Northwest Territories, Canada: a new tool for reconstructing Beringian paleoenvironments? J. N. Am. Benthol. Soc. 22: 323-337. Warner B.G. and Hann B.J. 1987. Aquatic invertebrates as paleoclimate indicators? Quat. Res. 28: 427-430. Whiteside M.C. and Swindoll M.R. 1988. Guidelines and limitations to cladoceran paleoecological interpretations. Palaeogeogr. Palaeoclim. Palaeoecol. 62: 405-412. Wohlfarth B., Lemdahl G., Olsson S., Persson T., Snowball I., Ising J. and Jones S. 1995. Early Holocene environment on Bjørnøya (Svalbard) inferred from multidisciplinary lake sediment studies. Polar Res. 14: 253-275. Wooller M.J., Francis D., Fogel M.L., Miller G.H., Walker J.R. and Wolfe A.P. 2004. Quantitative paleotemperature estimates from G18O of chironomid head capsules preserved in arctic lake sediments. J. Paleolim. 31: 267-274.

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Appendix: Authorities for species names Acroperus har pae (Baird); Alona gu ttata Sars; Apatania zonella Zetterstedt; Assulina muscorum Greeff; Branchinecta gaini (Daday); Chydorus arcticus Røen; Chydorus sphaericus (O.F. Müller); Colymbetes dolabr atus (Paykull); Corynocera amb igua Zetterstedt; Corynocera oliveri Lindeberg; Cristatella mucedo Cuvier; Daphnia p ulex (De Geer); Diaptomus cast or (Jurine); Eurycercus gl acialis Liljeborg; Eurycercus lamellatus (O.F. Müller); Gyraulus arct icus (Møller); Gyrinus opacus Sahlberg; Heterocypris in congruens (Ramdohr); Hippuris vulgaris L.; Hydroporus mori o Aubé; Hydrozetes lacustris (Michael); Ilyocypris bradyi Sars; Isoetes echinos pora L.; Lepidurus arc ticus (Pallas); Limnocythere sanctipatric ii (Brady and Robertson); Lymnaea vahliii Møller; Myriophyllum sp icatum L.; Orthocladius consobri nus Holmgren; Parakiefferiella bat hophila (Kieffer); Pediastrum angul osum (Ehrenberg) Meneghini; Pediastrum bor yanum (Turpin) Meneghini; Pediastrum duplex Meyen; Pediastrum integr um Naegeli; Pediastrum kawraisky Schmidle; Pediastrum muticum Kützing; Pisidium steenbuc hii (Møller); Plumatella re pens (Linnaeus); Potamocypris parva Schmidt; Potamogeton filiformis Pers.; Potamogeton pusillus L.; Sarscypridopsis aculeata (Costa); Simocephalus vetul us (O.F. Müller); Spongilla la custris (Linnaeus); Tanytarsus chinyensis Kieffer; Tanytarsus gracillentus (Holmgren); Tanytarsus lugens Kieffer; Tonnacypris glacialis (Sars); Warnstorfia exannulata (B.S. and G.) Loeske.

7. USE OF WATER ISOTOPE TRACERS IN HIGH LATITUDE HYDROLOGY AND PALEOHYDROLOGY

THOMAS W.D. EDWARDS ([email protected]) Department of Earth Sciences University of Waterloo Waterloo, Ontario N2L 3G1, Canada BRENT B. WOLFE ([email protected]) Department of Geography and Environmental Studies Wilfrid Laurier University Waterloo, Ontario N2L 3C5, Canada JOHN J. GIBSON ([email protected]) Water and Climate Impacts Research Centre Department of Geography University of Victoria Victoria, British Columbia V8W 3P5, Canada and DAN HAMMARLUND ([email protected]) GeoBiosphere Science Centre Quaternary Sciences Lund University Sölvegatan 12 SE-223 62 Lund Sweden

Key words: Stable isotopes, G18O, G2H, Meteoric Water Line, Local Evaporation Line, Hydroclimatology, Hydrology, Lake water balance, Water isotope archives, Paleohydrology, Paleoclimatology

Introduction The stable isotopes 2H and 18O are highly valuable tracers for investigating present and past hydrology and hydroclimatology because of the existence of robust physically187 R. Pienitz, M.S.V. Douglas and J.P. Smol (eds), 2004. Long-term Environmental Change in Arctic and Antarctic Lakes. Springer. Printed in the Netherlands.

188 T.W.D. EDWARDS, B.B. WOLFE, J.J. GIBSON AND D. HAMMARLUND based understanding of isotopic partitioning in the global water cycle (e.g., Rozanski et al. 1993; Araguás-Araguás et al. 2000). Although applicable in any geographic location, water isotope data are particularly effective for characterizing hydrology and hydroclimatology in remote regions, especially at high latitudes, providing information about runoff generation processes and surface water balance that may be difficult or impossible to acquire using other methods (e.g., Burgman et al. 1987; Gibson et al. 1993, 2002; Wolfe and Edwards 1997; Gibson 2001, 2002; Gibson and Edwards 2002; Leng and Anderson 2003; Maric 2003). Stratigraphic records of water isotope data in lake sediments can also be used to reconstruct or constrain past hydrology and hydroclimatology at high latitudes, relying on the same mechanistic principles and benefiting from the framework provided by knowledge and inference about global paleo-isotope fields (e.g., MacDonald et al. 1993, this volume; Edwards et al. 1996; Wolfe et al. 1996, 2000, 2003; Anderson et al. 2001; Sauer et al. 2001a; Shemesh et al. 2001; Hammarlund et al. 2002; Hu and Shemesh 2003). This chapter briefly reviews key aspects of the use of stable water isotope tracers in hydrological and paleohydrological investigations applicable to the study of long-term environmental change in arctic and antarctic lakes, illustrated by selected examples from recent investigations in northern Canada and Sweden (see Figure 1 for locations of sites mentioned in the text). The aim is to provide the reader with basic working knowledge of the major controls on the isotopic partitioning of water in high latitude catchments and the nature of the information that can be gleaned from records of the isotopic composition of paleowaters obtained from lake sediment archives in such environments. Isotopic labelling in the hydrological cycle Background Isotopic partitioning (fractionation) of water in the hydrological cycle arises because of differences in the behaviour of water molecules containing various combinations of the naturally occurring stable isotopes of hydrogen (1H, 2H) and oxygen (16O, 17O, 18O). Most isotope hydrology studies are based on variations in the relative abundances of the two rare heavy isotopic species (isotopomers) containing a single 2H or 18O atom (1H2H16O and 1H1H18O) with respect to common light water (1H1H16O). Isotopic compositions are expressed conventionally as G-values, representing deviation in per mil (‰) from the isotopic composition of a specified standard, such that G2H or G18O = 1000((Rsample/Rstandard) – 1), where R refers to the 2H/1H or 18O/16O ratios in sample and standard, respectively. The most widely used standard in hydrological applications is Vienna Standard Mean Ocean Water (VSMOW), which approximates the bulk isotopic composition of the present-day global ocean reservoir, and hence has G2H and G18O values both defined to be exactly 0 ‰. This is a logical datum for hydrological studies since evaporation from the oceans is the fundamental source of global atmospheric moisture, which provides the precipitation input for continental water cycling, and the isotopic composition of the oceans is essentially invariant at human timescales. Studies using limnic carbonates (e.g., sedimentary calcite, mollusc shells, ostracode valves) as oxygen-isotope archives also make use of the Vienna Pee Dee Belemnite (VPDB)

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standard, which approximates the isotopic composition of modern marine carbonate. G18O values expressed relative to the VSMOW and VPDB standards are related such that (1000 + G18Osample/VSMOW)/(1000 + G18Osample/VPDB) = 1.03086 (Fritz and Fontes (1980) or Clark and Fritz (1997) are recommended for further information about isotopic standards and nomenclature).

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190 T.W.D. EDWARDS, B.B. WOLFE, J.J. GIBSON AND D. HAMMARLUND Isotopic labelling of precipitation – Global and Local Meteoric Water Lines The distribution of isotopes in precipitation and surface waters is characterized by the existence of strongly linear relations between G2H and G18O values over a broad range of spatial and temporal scales, reflecting systematic mass-dependent partitioning of water molecules in the hydrologic cycle. The most striking expression of the coupling between hydrogen- and oxygen-isotope labelling of water is the Global Meteoric Water Line (GMWL) (Craig 1961), described by G2H = 8G18O + 10, which closely approximates the observed relation between mean annual amount-weighted precipitation G2H and G18O world-wide (Figure 2a). The existence of this fundamental baseline is consistent with the notion that global atmospheric moisture arises primarily from a well-mixed source (i.e., the subtropic ocean surface) and undergoes progressive rain-out of mass and heavy isotopes during subsequent poleward atmospheric transport. The linearity and slope of the GMWL reflect the pervasive influence of temperature-dependent equilibrium partitioning of the two heavy isotopic species between atmospheric vapour and

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Figure 2 . (a) Isotopic labelling of weighted-mean annual precipitation (GP) collected at selected high latitude stations, based on data compiled by Birks et al. (2002) from the Global Network for Isotopes in Precipitation (GNIP), operated jointly by the International Atomic Energy Agency and the World Meteorological Organization (IAEA/WMO 2001). Modern mean annual composition of precipitation for ice core sites from Greenland and Antarctica are also indicated, with arrows showing the approximate ranges of variability between average glacial and interglacial conditions (ice core data archived at the World Data Centre for Paleoclimatology, Boulder, Colorado, USA: http://www.ngdc.noaa.gov/paleo/data.html). Vienna Standard Mean Ocean Water (VSMOW) is also shown, offset slightly below the Global Meteoric Water Line (GMWL). (b) Isotopic labelling of mean composite monthly precipitation at Alert (Canada) and Naimakka (Sweden), based on data compiled by Birks et al. (2002). Data from both stations are strongly localized along the GMWL (dotted line), defining individual Local Meteoric Water Lines (LMWLs) having slopes slightly less than 8 and large ranges of monthly variability. Maximum and minimum mean composite monthly G2H and G18O values at both stations occur, respectively, in July and January. Note that the GMWL commonly provides an excellent proxy for the LMWL in the absence of direct isotopic characterization of local precipitation, especially at high latitudes.

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condensing precipitation. This can be modelled conceptually as a multi-step Rayleightype open-system distillation process, leading to the combination of low amount and depleted heavy-isotope content (low G values) of precipitation at high latitudes. This important coupling of mass and isotope depletion is well-illustrated by comparison of the global fields of precipitable moisture and precipitation G values (e.g., see Figure 2 of Birks et al. 2002). It is also reflected indirectly in the spatial relations that are observed between precipitation G values and air temperatures at mid- to high latitudes, often used as the basis for paleotemperature reconstruction from the isotopic composition of paleoprecipitation stored in glacier ice or groundwater (e.g., Rozanski et al. 1993, 1997; Schotterer et al. 1996). Amount-weighted G2H and G18O values of precipitation received over the year at individual sites also commonly plot in strongly linear clusters in G2H-G18O space close to the GMWL, and best-fit LMWLs drawn through these clusters can provide highly useful characterization of local isotopic input functions for hydrological studies (Figure 2b). Substantial seasonal variability is typical along LMWLs at mid- to high latitudes, especially in cold regions, with winter precipitation generally strongly depleted and more variable in heavy-isotope content compared to that received during the summer season. Additional isotopic characterization of local precipitation is provided by the derived "deuterium excess" parameter (d-excess = G2H – 8G18O), which gives an indication of location in G2H-G18O space relative to the GMWL reference (d-excessGMWL = +10‰). This characteristic feature of the distribution of isotopes in precipitation reflects the combined influence of equilibrium and kinetic isotopic partitioning as atmospheric moisture is generated by evaporation from the sea surface, which tends to be conserved during subsequent transport and distillation. Local precipitation typically has a seasonal cycle marked by lower d-excess values in summer, often attributable to slight evaporative isotopic enrichment of falling raindrops, and higher d-excess in winter due to kinetic effects associated with ice formation in clouds (Merlivat and Jouzel 1979; Jouzel and Merlivat 1984). Isotopic enrichment of surface waters - Local Evaporation Lines Strong coupling of hydrogen- and oxygen-isotope variations is also evident from the development of characteristic heavy-isotope build-up in surface waters undergoing evaporation. The isotopic signatures of neighbouring water bodies receiving input of similar isotopic composition typically lie along more-or-less well-defined linear arrays in G2H-G18O space, termed Local Evaporation Lines (LELs; Figure 3), which deviate from the LMWL along slopes that usually range between 4 and 6, depending on local atmospheric conditions during the thaw season (primarily relative humidity, temperature, and the isotopic composition of ambient moisture). Intersection of the LEL with the LMWL often provides an excellent approximation of the weighted-mean isotopic composition of input waters to a catchment (GI), while displacement of a given lake water (GL) along the LEL provides an index of water balance, which can be

192 T.W.D. EDWARDS, B.B. WOLFE, J.J. GIBSON AND D. HAMMARLUND quantified in terms of evaporation/inflow ratio (E/I) via isotope-mass balance considerations, as described for hydrogen or oxygen isotope data by: E/I = (GI – GL)/(GE – GL)

(1)

where GE represents the isotopic composition of the evaporative flux. GE can be evaluated using the well-validated linear resistance model of Craig and Gordon (1965), which accounts for the differing volatilities of the water isotopomers as a combination of mass-dependent variations in equilibrium vapour pressures ("equilibrium effects") plus variations in molecular diffusivities arising from the combination of differing absolute mass and its distribution within water molecules ("kinetic effects"). Assuming negligible resistance to liquid-phase mixing, GE can be calculated by: GE = ((1 + 10-3H*) GL – hGA – H)/(1 – h + 10-3HK)

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where H* and HK represent the respective equilibrium and kinetic effects, expressed as per mil (‰) separations between the liquid and vapour phases, H = H* + HK, and h is the atmospheric relative humidity normalized to the saturation vapour pressure at the temperature of the air-water interface. H* values for both hydrogen and oxygen can be estimated from the empirical relations derived by Horita and Wesolowski (1994). For G2H, H* = 1158.8(T3/109) – 1620.1(T2/106) + 794.84(T/103) – 161.04 + 2.9992(109/T3) and for G18O, H* = –7.685 + 6.7123(103/T) – 1.6664(106/T2) + 0.3504(109/T3), where T is the interface temperature (in K) and HK values for typical natural conditions can be accurately approximated as a function of relative humidity deficit (Gonfiantini 1986; Araguás-Araguás et al. 2000). For G2H, HK = 12.5(1 – h) and for G18O, HK = 14.2(1 – h). Equations (1) and (2) can be applied readily to assess instantaneous lake water balances in systems approximating hydrologic steady-state (e.g., Gibson 2001; Gibson and Edwards 2002) and a similar approach has been used for reconstructing or constraining past hydrologic variability from isotopic paleorecords (e.g., Edwards and Fritz 1988; Edwards and McAndrews 1989; MacDonald et al. 1993; Edwards et al. 1996; Wolfe et al. 2001a; Hammarlund et al. 2003). Substitution and rearrangement of equations (1) and (2) also yields: E/I = [(1 – h + 10-3HK)/(h – 10-3H)][(GL – GI)/(G* – GL)]

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where G is the limiting isotopic enrichment attainable when a water body evaporates to near-zero volume, and has hence progressed well beyond hydrologic steady-state (see Figure 3). Modelling of the non-steady-state isotopic enrichment leading to G* is also possible, affording information about mass and isotope fluxes under transient conditions (e.g., Gibson 2002).

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G O (‰) Figure 3. Isotope labelling of water balance components, consistent with a steady-state E/I ratio of about 0.8 for a hypothetical high latitude lake, as described by equations (1)-(4) in the text (based on mean thaw season relative humidity and air temperature, respectively, of about 0.7 and 10°C). The isotopic composition of the lake water at the time of sampling (GL) is offset from the input water composition (GI) along a local evaporation line (LEL) having a slope of about 5. Although impossible to sample directly, mass conservation dictates that the evaporative flux from the lake (GE) lies on the extension of the LEL to the left of the LMWL. Evaporation and exchange occur in association with ambient atmospheric moisture during the summer evaporation season (GAs), weighted for evaporation flux, which is in approximate isotopic equilibrium (İ*) with the mean isotopic composition of summer precipitation (GPs). Also shown is the limiting isotopic composition (G*), which is the maximum enrichment attainable under local atmospheric conditions as a water body approaches dryness, providing an indication of potential extreme nonsteady-state behaviour. See Gibson and Edwards (2002) for additional discussion.

Isotope hydrology at high latitudes Runoff generation and evaporation Hydrological processes at high latitudes are almost universally characterized by intense climatic and isotopic seasonality. Winter is typically marked by significant reduction or even total suspension of runoff generation because of deep frost penetration into soils

194 T.W.D. EDWARDS, B.B. WOLFE, J.J. GIBSON AND D. HAMMARLUND and the storage of highly 2H- and 18O-depleted seasonal precipitation in the frozen state within the overlying snow pack. In spite of dry atmospheric conditions, open-water evaporation is strongly suppressed by development of ice cover on lakes and all but the most rapidly flowing streams, though significant quantities of moisture may be recycled to the atmosphere via sublimation directly from the snow. Warming in spring results in rapid melting of the snow pack, introducing pulses of isotopically-depleted runoff into streams and lakes, while still-frozen soils inhibit infiltration. Lake ice begins to thin, with early break-up on small, shallow water bodies leading to immediate onset of seasonal evaporation, while persistence of ice cover for days or weeks on large, deep lakes may suppress this process well into the thaw season. The isotopic expression of these two primary hydrological processes, snowmelt dilution

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Figure 4 . (a) Changing isotopic composition of a small upland tundra lake in the West Exeter Lake catchment (64º07’N, 111º01’W), a headwater tributary of the Coppermine River, Canada, sampled on three occasions, immediately prior to freeze-up (late August 2000 and 2001) and during spring melt (mid-June 2001), reported by Maric (2003). The plot clearly shows the seasonal cycle produced by the opposing effects of snowmelt dilution and evaporative enrichment. Pronounced enrichment attained during summer 2000, conserved under winter ice cover, was strongly offset by isotopically depleted snowmelt runoff in spring 2001. Subsequent accumulated evaporative enrichment during summer 2001 almost exactly countered this effect, returning the lake to a composition very similar to that of the previous year. A best-fit line through these data points provides a good approximation of the LEL and the weighted-mean isotopic composition of input waters (GI). (b) Isotopic data from 30 small tundra lakes in the West Exeter Lake catchment (including the lake in Figure 4a), sampled on the same three occasions, showing pronounced clustering along the inferred LEL and the equivalent signals of snowmelt dilution and evaporative enrichment at catchment scale. The existence of limited scatter about the LEL reflects the occurrence of slightly differing evaporation lines for each lake, arising because of minor variations in snowmelt isotopic composition and varying contributions from rainfall and groundwater runoff, but generally also confirms that the LEL-GMWL intersection is a reasonable proxy for the isotopic composition of local weighted-mean annual precipitation (i.e., GI § GP). Also shown for comparison is the isotopic composition of outflow from Yamba Lake, a large lake downstream, which exhibited no measurable variation between the three sampling episodes, reflecting the damping of seasonal fluctuations because of its large volume (and hence long hydraulic residence time).

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and evaporative enrichment, is often strongly evident at both lake and catchment scales, as shown for data from a selection of small upland lakes in the continental Low Arctic of northern Canada (Figure 4). The isotopic evolution of a given lake during any particular thaw season is highly variable, depending on the fluxes and isotopic compositions of inputs (surface and subsurface runoff, inflow from upstream water bodies, precipitation falling directly upon the lake) and outputs (surface and subsurface outflow, evaporative flux), as well as the storage capacity within the lake itself, which acts to buffer seasonal variation (note data for Yamba Lake in Figure 4b). Annual thawing of the active layer in permafrost terrain also serves to suppress seasonal heavy-isotope build-up in small lakes by sustaining input of unevaporated precipitation recharged late in the previous thaw season (e.g., Bursey et al. 1991). Although most upland lakes exhibit systematic seasonal cycles of snowmelt depletion and progressive evaporative enrichment, much less predictable seasonal isotopic behaviour may occur in lakes in fluvial and deltaic settings, as in the myriad water bodies along the Mackenzie, Kolyma, Lena and other major northern river systems. As illustrated in Figure 5, episodic high-water events (often induced by high spring runoff, local ice-jam flooding or excess precipitation in upstream reaches of the river) can significantly influence the seasonal isotopic cycling in a lake in such environments by abruptly introducing water of strongly differing composition originating outside of the lake’s catchment.

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Figure 5. Strongly differing isotopic evolution of two lakes in the Peace-Athabasca Delta (PAD), Canada (55º08’N, 111º05’W), sampled immediately prior to freeze-up in October 2000 and again in early June and late August 2001. The relative depletion of PAD 5 between October 2000 and June 2001 reflects the overriding influence of snowmelt dilution, likely offset somewhat by evaporative enrichment. In contrast, evaporation completely dominated the early-season evolution of PAD 42, at that time a much shallower, desiccating water body, as shown by extreme nonsteady-state heavy-isotope enrichment. The subsequent evolution of PAD 5 between June and August 2001 was dominated by evaporative enrichment, yet PAD 42 became prominently depleted as a result of an influx of Athabasca River water during a summer high-water event (after Wolfe et al. 2002).

196 T.W.D. EDWARDS, B.B. WOLFE, J.J. GIBSON AND D. HAMMARLUND Various other factors can also potentially affect the regularity of the seasonal cycle, including variation in the timing and abundance of precipitation, which may force a lake across key hydrologic thresholds. Particularly strong shifts accompany the development of seasonal or perennial closure of a lake, for example, because of the rapid onset of non-steady-state evaporative enrichment (e.g., Wolfe et al. 1996). Seasonal variations in water-levels can also influence important processes like snowmelt bypassing, which can occur when high water-levels lead to rapid flushing of snowmelt under or around ice cover with minimal mixing, potentially leading to under-representation of isotopicallydepleted winter precipitation in a lake’s water budget (e.g., Edwards and McAndrews 1989). Perhaps most fundamentally, climate variability and change can be expected to affect the isotopic baseline provided by the LMWL through shifts in the range and weighted-mean isotopic composition of annual precipitation, possibly in company with temperature change and/or shifting seasonality of precipitation, as well as through alteration of the LEL due to changes in relative humidity and the isotopic composition of atmospheric moisture. Water isotope tracers in paleolimnology Development of water isotope records from lake sediments is typically undertaken with the goal of reconstructing the isotopic history of the overlying lake water, as the basis for inferences about paleoclimate and paleohydrology. The majority of such investigations have focused on the use of specific carbonate or organic sediment constituents as oxygen-isotope archives, building upon pioneering studies by Stuiver (1968, 1970) on lacustrine carbonates and Edwards and McAndrews (1989) on algal cellulose. Promising results have also been achieved using other substrates, including biogenic silica as an alternative oxygen-isotope archive (Shemesh et al. 2001) and kerogen and lipids as hydrogen-isotope archives (Sternberg 1988; Krishnamurthy et al. 1995; Sauer et al. 2001b; Huang et al. 2002). Although advances in the development of hydrogen-isotope archives may eventually afford the opportunity for simultaneous reconstruction of both G18O and G2H histories, only single-isotope records have yet been produced. As a result, the interpretation of lake water isotopic history is fundamentally dependent on the ability to partition shifts in lake water G18O (or G2H) into MWL- and LEL-parallel components, as shown schematically in Figure 6. Various strategies can be employed to deconvolute lake water G18O histories, including use of independently derived records of paleo-precipitation G18O to account for variations in input water composition, thus allowing direct resolution of lake-specific changes in water balance (e.g., Edwards and Fritz 1988; Wolfe et al. 2000, 2001a). The latter can also be achieved to varying degrees of confidence by comparing inferred lake water G18O records from multiple lakes in the same region spanning a range of hydrologic sensitivities, on the assumption that they share a common input water G18O history controlled by the isotopic composition of regional precipitation (e.g., Edwards et al. 1996). Clues from other indicators in lake sediment records (e.g., G13C or G15N as nutrient isotope tracers or other biological proxies like diatoms) can also inform or constrain hydrologic interpretations (e.g., Wolfe et al. 1996; Hammarlund et al. 1997).

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time

Comprehensive discussions of archive-specific methods and interpretation of lake sediment oxygen isotope records have been presented by Edwards and McAndrews (1989), Talbot (1990), Talbot and Kelts (1990), Edwards (1993), Holmes (1996), Li and Ku (1997), Ito (2001), Shemesh et al. (2001), Wolfe et al. (2001b) and others. Lake sediment oxygen-isotope studies at high latitudes have mainly been conducted as part of multidisciplinary investigations aimed at documenting regional late-Glacial and Holocene climatic and hydrological history (e.g., Anderson et al. 2001; Shemesh et al. 2001; Hammarlund et al. 2002; Hu and Shemesh 2003), and paleohydrological conditions associated with the advance and retreat of the circumpolar boreal treeline (e.g. Wolfe et al. 1996; MacDonald et al., this volume). Below we highlight two examples from recent investigations demonstrating, respectively, the derivation and interpretation of millennial-scale MWL-parallel fluctuations in lake water G18O inherited by sedimentary carbonates in a groundwater-fed lake in northern Sweden, and the analysis of decadal-scale LEL-parallel changes in the oxygen-isotope composition of a lake in northern Canada recorded by algal cellulose.

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Figure 6. Schematic diagram showing that changes in lake water G18O over time are controlled by a combination of changes in the isotopic composition of input waters, which should occur parallel to the GMWL (or LMWL) in G18O-G2H space, modified by changes in the extent of secondary evaporative enrichment, which occur along a shallower slope parallel to the LEL. The equivalent hydrogen-isotope input and evaporative enrichment records would obviously display similar trends, although the MWL-LEL geometry would result in a differing lake water G2H history.

198 T.W.D. EDWARDS, B.B. WOLFE, J.J. GIBSON AND D. HAMMARLUND Lake Tibetanus, Sweden Abisko National Park, northern Sweden, is situated east of the main drainage divide of the Scandes mountain range, close to the Norwegian border. Following the establishment of the Abisko Scientific Research Station 100 years ago, numerous ecological research projects have been carried out in this subarctic area, many of which have focused on the distinct climatic and vegetational gradients, particularly the altitudinal forest-tundra ecotone. Holocene vegetation dynamics are relatively well known through paleoecological investigations of peat sequences and lake sediments (Sonesson 1968, 1974; Barnekow 1999, 2000), and the climatic forcing of past environmental change in the area has received increased attention as part of multi-disciplinary research based at the Climate Impacts Research Centre at Abisko since the late 1990s (e.g., Bigler et al. 2002). -20

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Figure 7. Plot of G18O versus G2H values for samples of lake water collected at intervals over a three-year period (1995-98) from Lake Tibetanus, northern Sweden (68º20’N, 18º42’E), reported by Hammarlund et al. (2002). The LMWL was estimated from isotopic data obtained on samples of local precipitation and groundwater and the LEL was calculated based on local climate normals and data from the Global Network for Isotopes in Precipitation (IAEA/WMO 2001). Owing to the short water residence time, the lake water isotopic composition is strongly controlled by that of inflowing groundwater with negligible evaporative enrichment, even during the summer months. Independent evidence suggests that rapid flushing has characterized the lake throughout its history and hence that reconstructed lake water G18O should be a good proxy for the G18O of local paleoprecipitation.

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Multi-proxy investigations of the sediments underlying Lake Tibetanus, a small hardwater lake within the upper mountain-birch zone close to Abisko, have yielded particularly crucial information, including the development of a paleo-precipitation G18O history, derived through coupling of paleo-isotope data obtained from sedimentary carbonate with pollen-inferred paleotemperature estimates (Hammarlund et al. 2002). As revealed in Figure 7, the modern isotope hydrology of Lake Tibetanus is characterized by complete absence of evaporative enrichment in the lake water, attributable to rapid throughflow of local groundwater. Independent evidence suggests that this hydrologic situation has persisted throughout the lake's existence (Hammarlund et al. 1997) and thus the lake water G18O history should be characterized by purely MWL-parallel variability. G18O records were obtained on three different carbonate components from a sediment sequence spanning the last 10,000 years (Figure 8). The longest record was developed from fine-grained sedimentary calcite, originating primarily from photosynthesis in submerged Chara algae during early summer, while slightly shorter records were gained from analyses of aragonitic shells of Pisidium sp. molluscs and calcitic valves of adult Candona candida ostracodes precipitating, respectively, throughout the summer and in late autumn. Consistent oxygen-isotope offsets between the three records can be Carbonate G18O (‰ VPDB) -14

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Figure 8. Oxygen-isotope records for a core from Lake Tibetanus plotted against calibrated radiocarbon age, along with pollen-inferred precipitation and temperature, based on data from Hammarlund et al. (2002). Samples with poor modern pollen analogues are circled. Thick lines represent five-point running averages. Modern temperature and precipitation values for Abisko are shown by vertical lines (dashed line represents calculated modern mean July air-temperature at Lake Tibetanus). Sed. = fine-grained sedimentary carbonate (primarily Chara calcite); Pis. = Pisidium sp. mollusc aragonite; Can. = adult Candona candida ostracode calcite. The heavy dashed line represents the inferred G18O of local paleo-precipitation (G18OP) after accounting for temperature-dependent fractionation between sedimentary carbonate and lake water using polleninferred mean July temperature.

200 T.W.D. EDWARDS, B.B. WOLFE, J.J. GIBSON AND D. HAMMARLUND attributed to a combination of mineralogical, biological and temperature-dependent effects (the latter related to the differing annual periods of calcification). Based on the assumption of stable open-drainage conditions resembling the modern hydrological situation throughout the lake's history, long-term changes in carbonate G18O were interpreted as reflecting variations in lake water G18O inherited from precipitation G18O, modified slightly due to temperature-dependent variations in carbonate-water equilibrium isotope fractionation. The pollen-based reconstruction of mean July air temperature was used to account for the effect of changing water temperature on the sedimentary carbonate G18O record, assuming temperature sensitivity of -0.25 ‰/K (Friedman and O'Neil 1977). Since inferred temperatures were similar to present at the beginning and end of the record, but warmer during the intervening period, this temperature correction resulted in a modest increase in inferred lake water G18O throughout the mid-part of the record (see Figure 8). Comparison of the derived paleo-precipitation G18O record and the pollen-based temperature reconstruction from Lake Tibetanus reveals clear evidence for changing temporal isotope-temperature relations over the Holocene at this site. As addressed in detail by Hammarlund et al. (2002), use of the traditional "isotope paleothermometer", based on the observed modern spatial relation between G18O of local amount-weighted mean annual precipitation and mean annual temperature of about +0.7‰/K (Dansgaard 1964), is only applicable over the past ca. 6000 years. Evidently, significantly different and variable G18O-MAT relations prevailed 10,000 to 6000 years ago, presumably reflecting substantially different atmospheric and oceanic circulation patterns than those of the present. As suggested by Hammarlund et al. (2002), these differences likely included enhanced westerly air-mass circulation leading to relatively moist maritime conditions in response to high summer insolation, elevated sea-surface temperatures in the Norwegian Sea and steep meridional pressure gradients. The subsequent transition to drier conditions, which is clearly reflected in the pollen-based reconstruction of mean annual precipitation shown in Figure 8, was accompanied by major changes in forest extent and composition, recorded unequivocally by pollen and plant macrofossil data (Barnekow 1999). Jemis Lake, Canada The Peace-Athabasca Delta (PAD), situated at the confluence of the Peace, Athabasca and Fond-du-Lac (Lake Athabasca) drainages in northern Alberta, Canada, is an extensive high latitude fluvio-deltaic complex of considerable cultural, historical and ecological significance (e.g., Peace-Athabasca Delta Project Group 1973; Prowse et al. 1996). The PAD is also a key node in the Mackenzie River basin, the major North American source of freshwater to the Arctic Ocean. Concerns about the combined influence of climate change, regulation of Peace River flow due to hydro-electric power generation, increasing consumptive use of Athabasca River water associated with tar sands development, and changes in land use upstream have led to numerous environmental studies in the PAD over the past 30 years. Many of these investigations have focussed on the frequency and severity of periodic ice-jam floods on the Peace and Athabasca rivers, which are believed to provide crucial fluxes of water and nutrients to lakes in the

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delta that are perched beyond the reach of normal seasonal high-water flooding from these two rivers (e.g., see Prowse and Conly 2000). Increasing attention is also being paid to the implications downstream in the Mackenzie River system. Jemis Lake, a small shallow water body located centrally in the PAD, has been targeted by a number of investigators as being potentially representative of broader conditions in the delta over time, as well as being relatively well-known to the residents of Fort Chipewyan, the major local settlement. Seasonal monitoring of the isotopic composition of Jemis Lake as part of longer term isotope hydroclimatology studies in the PAD suggests that the water balance of this lake (Figure 9), like many others in the delta, is strongly controlled by the opposing effects of evaporation, leading to heavy-isotope enrichment, and periodic inflow of isotopically depleted river water and precipitation (see also Figure 5). Because river water inputs to Jemis Lake and local precipitation have similar isotopic compositions, fluctuations in the lake’s water balance are expressed as LEL-parallel shifts in G18O-G2H space, which can thus be translated directly into shifting

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202 T.W.D. EDWARDS, B.B. WOLFE, J.J. GIBSON AND D. HAMMARLUND E/I ratio using equation (3). Indeed, consideration of reconstructed lake water G18O obtained from algal cellulose in a short core of sediment from Jemis Lake suggests that changes in the isotopic composition of lake water throughout at least the past 150 years (and perhaps much longer; Wolfe et al. 2002) have also been strongly controlled by LELparallel fluctuations, as revealed by strong inverse correspondence between inferred E/I and an independently derived flood history (Figure 10). Such insight and context are critical to understanding the hydro-ecological evolution of the Peace-Athabasca Delta, both because of its substantial natural heritage value and because of overarching concerns about the responsible stewardship of water resources within the Mackenzie River system as a whole.

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Figure 10 . (a) Speculative isotope-inferred history of evaporation/inflow ratio (E/I) in Jemis Lake AD 1850-2000, based on lake water G18O reconstructed from algal cellulose in a short sediment core (after Wolfe et al. 2002). E/I was calculated using equation (3), assuming input of either river water or precipitation approximating the G18O of modern local mean annual precipitation (see Figure 9). The chronology was developed from 210Pb and 14C data. The solid line on the E/I plot is a three-point running mean. (b) Peace River spring ice-jam flood frequency record in the Peace-Athabasca Delta derived from anecdotal and historical sources by Timoney et al. (1997), expressed as a 10-year mean. The water balance of Jemis Lake has apparently fluctuated in concert with flood frequency throughout the past 150 years, as shown clearly by the inverse relation between isotope-inferred E/I and estimated floods/year, with particularly well-marked depiction of the prolonged period of wetter conditions between about 1910 and 1955.

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Summary and future perspectives This chapter has outlined the major features of water-isotope labelling in the hydrological cycle, highlighting the characteristic signals in high latitude lakes that may provide clues to interpret stratigraphic variations in sediment cores. Although lakes respond individualistically, the strongest signals at seasonal time scales in most water bodies having short hydraulic residence times are likely to derive from snowmelt influx and evaporative enrichment. Samples from lake sediment cores tend to naturally average and integrate seasonal signals (as do the waters in lakes having long residence times), thus often affording the opportunity to resolve longer term fluctuations in the isotopic composition of lake water into two respective components, one reflecting shifting input water composition (and hence expected to vary parallel to the Global or Local Meteoric Water Line in G18O-G2H space) and the other reflecting varying evaporative enrichment along trajectories parallel to the Local Evaporation Line. Deconvolution of MWL-parallel signals and associated estimates of the isotopic composition of paleo-precipitation may provide direct links to global paleo-precipitation isotope fields, analogous to those afforded by high latitude ice cores (e.g., Rozanski et al. 1997), whereas LEL-parallel signals provide the basis for reconstructing local hydroclimate (e.g., MacDonald et al. 1993). Future advances in the application of water-isotope tracers in high latitude lakes are likely to be realized from the development of coupled water-isotope records, using new archive materials to reconstruct lake water isotopic history in G18O-G2H space, thus permitting fuller and more sophisticated paleohydrologic and paleoclimatic interpretation. Acknowledgements We wish to acknowledge support received from a number of agencies in our studies of high latitude isotope hydrology and paleohydrology, including the Natural Sciences and Engineering Research Council of Canada, Environment Canada, Indian and Northern Affairs Canada (and especially the Northern Scientific Training Program), the British Columbia Hydro and Power Authority, and the Swedish Research Council, as well as the residents of the various communities in which we have worked. The comments of the editors and an anonymous reviewer led to considerable improvements in our contribution. References Anderson L., Abbott M.B. and Finney B.P. 2001. Holocene climate inferred from oxygen isotope ratios in lake sediments, central Brooks Range, Alaska. Quat. Res. 53: 313-321. Araguás-Araguás L., Froehlich K. and Rozanski K. 2000. Deuterium and oxygen-18 isotope composition of precipitation and atmospheric moisture. Hydrol. Proc. 14: 1341-1355. Barnekow L. 1999. Holocene tree-line dynamics and inferred climatic changes in the Abisko area, northern Sweden, based on macrofossil and pollen records. The Holocene 9: 253-265. Barnekow L. 2000. Holocene regional and local vegetation history and lake-level changes in the Torneträsk area, northern Sweden. J. Paleolim. 23: 399-420.

204 T.W.D. EDWARDS, B.B. WOLFE, J.J. GIBSON AND D. HAMMARLUND Bigler C., Larocque I., Peglar M., Birks H.J.B. and Hall R.I. 2002. Quantitative multiproxy assessment of long-term patterns of Holocene environmental change from a small lake near Abisko, northern Sweden. The Holocene 12: 481-96. Birks S.J., Gibson J.J., Gourcy L., Aggarwal P.K. and Edwards T.W.D. 2002. Maps and animations offer new opportunities for studying the global water cycle. EOS Trans. AGU 83: 406 plus electronic supplement at: http://www.agu.org/eos_elec/eeshome.html Burgman J.O., Calles B. and Westman F. 1987. Conclusions from a ten year study of oxygen-18 in precipitation and runoff in Sweden. Symposium on Isotope Techniques in Water Resources Development. International Atomic Energy Agency, Vienna, 30 March-3 April 1987, IAEASM-299/107. Bursey G.G., Edwards T.W.D. and Frape S.K. 1991. Water balance and geochemistry studies in a tundra watershed, District of Keewatin, N.W.T. In: Prowse T.D. and Ommanney C.S.L. (eds), Northern Hydrology: Selected Perspectives. National Hydrology Research Institute Symposium No. 6, Environment Canada, pp. 17-31. Clark I.D. and Fritz P. 1997. Environmental Isotopes in Hydrogeology. CRC Press/Lewis Publishers, Boca Raton, Florida, USA, 328 pp. Craig H. 1961. Isotopic variations in meteoric waters. Science 133: 1702-1703. Craig H. and Gordon L.I. 1965. Deuterium and oxygen-18 variations in the ocean and marine atmosphere. In: E. Tongiorgi (ed.), Stable Isotopes in Oceanographic Studies and Paleotemperatures. Cons. Naz. Rich. Lab. Geol. Nucl. Pisa, Italy, pp. 9-130. Dansgaard W. 1964. Stable isotopes in precipitation. Tellus 16: 436-468. Edwards T.W.D. 1993. Interpreting past climate from stable isotopes in continental organic matter. In: Swart P.K., Lohmann K.C., McKenzie J. and Savin S. (eds), Climate Change in Continental Isotopic Records. Geophysical Monograph 78, American Geophysical Union, Washington, pp. 333-341. Edwards T.W.D. and Fritz P. 1988. Stable isotope paleoclimate records for southern Ontario: comparison of results from marl and wood. Can. J. Earth Sci. 25: 1397-1406. Edwards T.W.D. and McAndrews J.H. 1989. Paleohydrology of a Canadian Shield lake inferred from 18O in sediment cellulose. Can. J. Earth Sci. 26: 1850-1859. Edwards T.W.D., Wolfe B.B. and MacDonald G.M. 1996. Influence of changing atmospheric circulation on precipitation G18O-temperature relations in Canada during the Holocene. Quat. Res. 46: 211-218. Friedman I. and O'Neil J.R. 1977. Compilation of stable isotope fractionation factors of geochemical interest. In: Fleischer M. (ed.), Data of Geochemistry. U.S. Geol. Surv. Prof. Paper, 440-KK, pp. 1-12. Fritz P. and Fontes J.C. 1980. Introduction. In: Fritz P. and Fontes J.C. (eds), Handbook of Environmental Isotope Geochemistry. Volume 1. The Terrestrial Environment, A. Elsevier Scientific Publishers, Amsterdam, pp. 1-19. Gibson J.J. 2001. Forest-tundra water balance signals traced by isotopic enrichment in lakes. J. Hydrol. 251: 1-13. Gibson J.J. 2002. Short-term evaporation and water budget comparisons in shallow arctic lakes using non-steady isotope mass balance. J. Hydrol. 264: 247-266. Gibson J.J. and Edwards T.W.D. 2002. Regional surface water balance and evaporationtranspiration partitioning from a stable isotope survey of lakes in northern Canada. Glob. Biogeochem. Cyc. 16: 10.1029/2001GB001839. Gibson J.J., Edwards T.W.D., Bursey G.G. and Prowse T.D. 1993. Estimating evaporation using stable isotopes: quantitative results and sensitivity analysis for two catchments in northern Canada. Nordic Hydrol. 24: 79-94. Gibson J.J., Prepas E.E. and McEachern P. 2002. Quantitative comparison of lake throughflow, residency, and catchment runoff using stable isotopes: modelling and results from a regional survey of Boreal lakes. J. Hydrol. 262: 128-144.

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206 T.W.D. EDWARDS, B.B. WOLFE, J.J. GIBSON AND D. HAMMARLUND in Continental Isotopic Records. Geophysical Monograph 78, American Geophysical Union, Washington, pp. 1-36. Rozanski K., Johnsen S.J., Schotterer U. and Thompson L.G. 1997. Reconstruction of past climates from stable isotope records of palaeo-precipitation preserved in continental archives. Hydrol. Sci. 142: 725-745. Sauer P.E., Miller G.H. and Overpeck J.T. 2001a. Oxygen isotope ratios of organic matter in arctic lakes as a paleoclimate proxy: field and laboratory investigations. J. Paleolim. 25: 43-64. Sauer P.E., Eglinton T.I., Hayes J.M., Schimmelmann A. and Sessions A.L. 2001b. Compoundspecific D/H ratios of lipid biomarkers from sediments as a proxy for environmental and climatic conditions. Geochim. Cosmochim. Acta 65: 213-222. Schotterer U., Oldfield F. and Fröhlich K. 1996. GNIP - Global Network for Isotopes in Precipitation. PAGES, Bern, 48 pp. Shemesh A., Rosqvist G., Rietti-Shati M., Rubensdotter L., Bigler C., Yam R. and Karlen W. 2001. Holocene climatic change in Swedish Lapland inferred from an oxygen-isotope record of lacustrine biogenic silica. The Holocene 11: 447-454. Sonesson M. 1968. Pollen zones at Abisko, Torne Lappmark, Sweden. Botaniska Notiser 121: 491-500. Sonesson M. 1974. Late Quaternary development of the Torneträsk area, North Sweden: 2. Pollen analytical evidence. Oikos 25: 288-307. Sternberg L.S.L. 1988. D/H ratios of environmental water recorded by D/H ratios of plant lipids. Nature 333: 59-61. Stuiver M. 1968. Oxygen-18 content of atmospheric precipitation during the last 11,000 years in the Great Lakes region. Science 162: 994-997. Stuiver M. 1970. Oxygen and carbon isotope ratios of fresh-water carbonates as climatic indicators. J. Geophys. Res. 75: 5247-5257. Talbot M.R. 1990. A review of the palaeohydrological interpretation of carbon and oxygen isotopic ratios in primary lacustrine carbonates. Chem. Geol. (Iso. Geosci. Sect.) 80: 261-279. Talbot M.R. and Kelts K. 1990. Paleolimnological signatures from carbon and oxygen isotopic ratios in carbonates from organic carbon-rich lacustrine sediments. In: Katz B.J. (ed.), Lacustrine Basin Exploration: Case Studies and Modern Analogs. The American Association of Petroleum Geologists, Tulsa, pp. 99-112. Timoney K., Peterson G., Fargey P., Peterson M., McCanny S. and Wein R. 1997. Spring ice-jam flooding of the Peace-Athabasca Delta: Evidence of a climatic oscillation. Clim. Change 35: 463-483. Wolfe B.B. and Edwards T.W.D. 1997. Hydrologic control on the oxygen-isotope relation between sediment cellulose and lake water, Taimyr Peninsula, Russia: Implications for the use of surface-sediment calibrations in paleolimnology. J. Paleolim. 18: 283-291. Wolfe B.B., Edwards T.W.D., Aravena R. and MacDonald G.M. 1996. Rapid Holocene hydrologic change along boreal treeline revealed by G13C and G18O in organic lake sediments, Northwest Territories, Canada. J. Paleolim. 15: 171-181. Wolfe B.B., Edwards T.W.D., Aravena R., Forman S.L., Warner B.G., Velichko A.A. and MacDonald G.M. 2000. Holocene paleohydrology and paleoclimate at treeline, north-central Russia, inferred from oxygen isotope records in lake sediment cellulose. Quat. Res. 53: 319-329. Wolfe B.B., Aravena R., Abbott M.B., Seltzer G.O. and Gibson J.J. 2001a. Reconstruction of paleohydrology and paleohumidity from oxygen isotope records in the Bolivian Andes. Palaeogeogr. Palaeoclim. Palaeoecol. 176: 177-192. Wolfe B.B., Edwards, T.W.D., Beuning K.R.M. and Elgood R.J. 2001b. Carbon and oxygen isotope analysis of lake sediment cellulose: methods and applications. In: Last W.M. and Smol J.P. (eds), Tracking Environmental Change Using Lake Sediments: Volume 2: Physical

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and Chemical Techniques. Kluwer Academic Publishers, Dordrecht, The Netherlands, pp. 373-400. Wolfe B.B., Edwards T.W.D. and Hall R.I. 2002. Past and present ecohydrology of the PeaceAthabasca Delta, northern Alberta, Canada: water isotope tracers lead the way. PAGES News 10: 16-17. Wolfe B.B., Edwards T.W.D., Jiang H., MacDonald G.M., Gervais B.R. and Snyder J.A. 2003. Effect of varying oceanicity on early to mid-Holocene palaeohydrology, Kola Peninsula, Russia: isotopic evidence from treeline lakes. The Holocene 13: 153-160.

8. LAKE SEDIMENTS AS RECORDS OF ARCTIC AND ANTARCTIC POLLUTION

DEREK C.G. MUIR ([email protected]) National Water Research Institute Environment Canada Burlington, Ontario L7R 4A6, Canada and NEIL L. ROSE ([email protected]) Environmental Change Research Centre University College London 26 Bedford Way London, WC1H 0AP United Kingdom

Key words: Arctic, Sediments, Heavy metals, Mercury, Lead, Persistent organic pollutants, PCBs, PAHs, Flyash, Particulates

Introduction The use of lake sediments to infer current fluxes and depositional histories of heavy metals and persistent organic pollutants (POPs) is well established and well documented for many temperate environments including the Great Lakes (Goldberg et al. 1981; Czuczwa and Hites 1984; Swackhamer and Armstrong 1986; Eisenreich et al. 1989) and lakes in Scandinavia (Johansson 1985; Verta et al. 1989; Rognerud and Fjeld 1993) as well as in European alpine and arctic lakes (Fernández et al. 2000; Grimalt et al. 2001). The use of carbonaceous anthropogenic particles, to examine pollution impacts in temperate lakes is also well established (Rose 2001; Cameron et al. 2002). Since the early 1980s, studies of cores from temperate lakes have demonstrated that sediments preserve historical records for a wide range of organic chemical pollutants, heavy metals and anthropogenic particles. Sediment cores have become valuable tools for the assessment of large scale spatial trends of pollutant deposition, as well as for development of retrospective information on emissions and use of various chemicals, and for estimating current fluxes from specific source regions. Use of dated sediment cores from arctic environments is more recent, with most studies conducted from the early 1990s onwards. This work has been driven to a large extent by the need to understand the spatial and temporal trends of contaminants in the Arctic and the path209 R. Pienitz, M.S.V. Douglas and J.P. Smol (eds), 2004. Long-term Environmental Change in Arctic and Antarctic Lakes. Springer. Printed in the Netherlands.

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ways of exposure to wildlife and humans. This is especially the case for mercury (Hg) which is elevated in predatory fish in lakes in Scandinavia, Greenland, and Canada (Dietz et al. 1998; Rognerud et al. 2002). While sediment records are increasingly being used to examine POPs, heavy metals, and anthropogenic particles at high latitudes, published reports on these pollutants in antarctic lakes are limited to analysis of surface sediments (Sarkar et al. 1994) and lake waters (Gasparon and Burgess 2000). Factors unique to polar environments make the interpretation of sediment records of anthropogenic chemical inputs particularly challenging. Because of the lack of sediment cores analysed for chemical pollutants from Antarctica the focus will be on arctic sediments. Our definition of the Arctic is that defined by the Arctic Monitoring and Assessment Program (Figure 1). In this review we will critically examine the studies on heavy metals, POPs, and anthropogenic particles in arctic sediment cores. Our goal is to provide guidance for the design and interpretation of results of future sediment core studies in high latitude regions.

Figure 1 . Map showing boundaries of the Arctic as defined by the Arctic Monitoring and Assessment Program and the 10°C July isotherm.

Challenges in the study of high latitude lake sediment cores Logistical and limnological challenges Logistical considerations Remote arctic lakes present logistical challenges for investigators interested in obtaining samples for determination of POPs, metals and anthropogenic particles. In all high

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latitude areas except for parts of northern Norway, Sweden, Finland and northwestern Russia, access has to be via aircraft, by boat or by overland/ice travel from the nearest commercial airport. Yet it can be argued that lack of road and boat access is a prerequisite for helping to ensure background, i.e., atmospheric deposition or natural geological sources, rather than primary sources of anthropogenic inputs. This requires special consideration of size and weight of sampling equipment and limits the number of cores that can be taken at one time. Investigators must consider possible contamination from various transport vehicles used for the entire trip. Contamination can also be introduced from ice augers used to drill holes for coring equipment. In most cases there is only limited bathymetric information for the candidate lakes. This problem can be overcome with combined global positioning/bathymetry equipment. Nevertheless, the lack of information on bathymetry, basic water and sediment chemistry makes lake selection difficult ahead of time. In this regard, surveys of lakes for water chemistry and paleolimnological cores often provide a guide for appropriate sites, and discussions with the scientists responsible for those studies are an essential part of project planning. Limnological considerations The unique limnological features of arctic lakes also present a challenge for investigators. Within the Arctic and Subarctic (Figure 1) lakes range from typical oligotrophic freshwater systems near the treeline (Pienitz et al. 1997; Fallu et al. 2002) to ultra-oligotrophic systems that are ice-covered year round in high arctic environments (Hamilton et al. 2001; Michelutti et al. 2002). Low sedimentation rates are rather typical of most lakes in the High Arctic (Rigler 1975). For example, in northern Québec, Laing et al. (2002) found sedimentation rates ranging from 39 to 120 g m-2 yr-1 in seven lakes in the Rivière George region south of Ungava Bay. Lockhart et al. (1998) reported sedimentation rates of 70 to 498 g m-2 yr-1 for five arctic lakes while nine more southerly and glacier-influenced lakes had high rates (264 to 2050 g m-2 yr-1). In the same study, a glacier-influenced lake in the High Arctic had a high sedimentation rate as well (2930 g m-2 yr-1). Furthermore, primary production in high arctic lakes is often low (< 10 g C m-2 yr-1) (Hobbie 1984). This limits inputs to bottom sediments via sorption to sedimenting particles and makes them less significant reservoirs for hydrophobic organics than temperate lakes. In small arctic lakes with long ice cover, benthic primary production from moss and algal beds is a significant source of organic matter (Welch et al. 1989). Allochthonous (external) inputs of organic and inorganic particles to these systems are also generally low due to sparsely developed catchment vegetation and soils. Rapid spring melting may result in limited mixing with lake waters owing to water density differences such that pollutants that have been deposited in snow are moved through the lake rather than deposited in sediments (Bergman and Welch 1985; Macdonald et al. 2000; Helm et al. 2002). Thus the extent of retention of the snow deposited contaminants will depend on watershed topography, lake morphometry, and surface area. The low sedimentation rates in many systems, due to low water column productivity and allochthonous inputs, limit temporal resolution of polar lake sediments. This can be overcome by high resolution sampling, e.g., at 0.2 or 0.5 cm intervals. However, the low

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sediment mass in such small slices presents an analytical challenge if concentrations of contaminants are low as well. This is particularly a problem for POPs and PCBs, as background contamination from the laboratory can be a significant problem (Alcock et al. 1994; Wallace et al. 1996). Smearing along the sides of the sampling tube is potentially a problem in cores with low sedimentation rates where there are fewer slices with which to distinguish historical profiles. However, KB (Kajak-Brinkhurst) type corers have been shown to provide unbiased estimates of contaminant inventory in lake sediments (Stephenson et al. 1996). Other studies have shown that artefacts associated with the hydraulic “bow wave” (associated with lowering of the corer) and thinning of sediment layers due to frictional resistance as the corer moves downward through the sediment, are not a problem for open barrel gravity corers (Cumming et al. 1993; Glew et al. 2001). Uncertainties in the dating of sediment cores In addition to low sedimentation rates in many polar lakes, 210Pb deposition declines exponentially with latitude and, in North America, is about 5-fold lower at 60°N than at 30°N because of lower precipitation at higher latitudes (Preiss et al. 1996). They may also reflect the influence of permafrost in reducing radon gas (the parent of 210Pb) emissions from soil and extended periods of ice cover which prevent atmospheric 210Pb from reaching lake sediments (Hermanson 1990). This can be overcome to some extent with longer counting times when using gamma-counting of 210Pb and 226Ra, but it limits use of alpha-counting where dried sediment samples are acid-digested prior to analysis. Use of 137Cs, and also short-lived 134Cs, can aid 210Pb dating of lake sediments, particularly in the Arctic, where deposition from both nuclear weapons testing (peak activity around 1963) and the Chernobyl accident (1986) are readily discerned in most lake sediments (Lockhart et al. 1998; Johnson-Pyrtle et al. 2000). Annually laminated or varved sediments are ideal for studies of historical profiles of POPs and heavy metals (Simola 2000). However, while more lakes with these sediments are gradually being found in the Arctic they are nevertheless rare, which limits their use in large regional studies of pollutant deposition. Gottgens et al. (1999) have discussed errors associated with paleoecological studies of Hg and many of their comments apply generally to any work with dated sediment cores. They noted that careful mixing of wet sediments was necessary prior to subsampling for Hg or bulk density analysis and moisture analysis. They found losses of Hg upon ovendrying of sediments, especially in low porosity top slices of sediment cores. However, there were also errors in the measurement of moisture content for converting wet weights to a dry weight basis. In general, analysis of wet sediments or freeze-dried sediments appears to be the approach used in most studies of arctic contaminants. Post-depositional changes in the contaminant record Diagenesis (alteration of metal profiles) in sediments is a concern for important contaminants such as Hg with respect to marine sediment cores, and possibly also for lake sediments. It has been argued that processes such as adsorption of Hg complexes (to iron and manganese oxides and hydroxides) may cause Hg accumulation near the surface in the redox boundary layer (Rasmussen 1994). However, Jackson (1997) and

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Fitzgerald et al. (1998) concluded that, based on the weight of available evidence, postdepositional remobilization of Hg was insignificant in temperate lakes and that diagenetic processes could not explain substantial increases in Hg fluxes during the 20th century. Lockhart et al. (2000) demonstrated convincingly that Hg inputs to sediments of three lakes (two related to mining activities and one to chlor-alkali plant discharges) were well preserved over a 30-50 yr period and corresponded with the known history of emissions. These lakes all had relatively high sedimentation rates and anoxic sediments. Mercury was shown to be present as insoluble HgS in two of the lakes which reduced its mobility. Low sedimentation rates and low sulfide concentrations, however, could lead to some diffusion of Hg. In Arctic Ocean profundal sediments, Gobeil et al. (1999) found that Hg concentrations decreased from the sediment surface to a depth of 10 cm and concluded that the Hg profiles were produced by its redistribution during diagenesis. They observed a relationship between profiles of Hg and reactive Fe which implies that a portion of the total Hg deposited was recycled along with Fe during redox changes. Muir et al. (2003) found significant correlations between Hg concentrations, in dated sediments from subarctic and arctic lakes, and “reactive” Fe and Mn, in 7 of 24 study lakes. These correlations were generally in lakes with very low sedimentation rates, i.e. < 50 g m-2 yr-1, and they imply possible remobilization of Hg with Fe and Mn. In general, the sedimentation rates in arctic lakes (ca. 0.01 to ca. 0.2 cm/yr) compared to deep Arctic Ocean basins (ca. 0.5 cm/1000 yr), combined with low rates of biomixing, suggest that of Hg would be relatively unimportant in most systems. There are fewer concerns about post-depositional mobility of lead (Pb) owing to its strong sorption to fine-grained sediment particles at neutral pH. Furthermore, at least in subarctic lakes, the anthropogenic signal of 206Pb/207Pb helps to delineate the recent deposition profile and source region (Bindler et al. 2001a; Outridge et al. 2002). Lake acidification can potentially mobilize Pb, Cd, Zn and Cu and redistribute them in the core via diffusion in interstitial water (Schindler et al. 1980; Nriagu et al. 1982; Rognerud and Fjeld 1993) although sorption to organic carbon appears to limit this (Carignan and Nriagu 1985). However, acidification is not a major consideration for most arctic lakes at present except those near smelters in the Kola Peninsula and near Noril’sk (Dietz et al. 1998). POPs can also undergo post-depositional movement due to diffusion and biomixing as well as biodegradation. Eisenreich et al. (1989) estimated effective diffusion rates of 0.03, 0.09 and 0.11 cm2/yr for mirex, PCBs and HCB, respectively, in Lake Ontario sediments. These results illustrate that organic compounds of higher water solubility will be more prone to diffusion in sediment pore waters because of higher molecular diffusion coefficients and lower sediment water partition coefficients. Dechlorination is a significant transformation pathway for chlorinated organics in anoxic sediment. Rawn et al. (2001) observed a slow conversion of DDT to DDD in subsurface sediments of Lake Laberge and Watson Lake in the Yukon, reflecting slow dechlorination of DDT (Oliver et al. 1989). Toxaphene (a mixture of chlorobornanes and camphenes) is an example of a chemical that has lower sediment sorption than PCBs and which appears to slowly degrade by dechlorination in all sediments (Howdeswell and Hites 1996; Rose et al. 2001). Rose et al. (2001) concluded that diffusive penetration for heptachlorobornanes and nonachlorobornanes was in the 1-2 cm range, which would help to explain the toxaphene profile in a core from remote Lochnagar in Scotland. On the other hand, there is no evidence for degradation of PCBs in Lake Ontario sediments

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where total PCBs were in the range of 100-500 ng g-1 dw (Wong et al. 1995) and little evidence in arctic lake cores analysed to date (Muir et al. 1996; Rawn et al. 2001). However, PCBs have been observed to undergo biotransformation at relatively low concentrations (0.05-154 Pg g-1) in river bed sediments (Wong et al. 2001). Perylene is also observed to increase with depth in lake sediments as a result of post-depositional transformation of hydrocarbon precursors under anoxic conditions (Wakeham et al. 1980). Advantages of arctic lake sediments Despite the potential problems and challenges described above, use of sediments from high latitude lakes offers unique advantages over other types of environmental samples such as cores from peat bogs, soil or ice for determining POPs and heavy metals. There are thousands of undisturbed lakes in all countries of the circumpolar Arctic thus theoretically permitting wide geographical coverage. While there are far fewer lakes to choose from in the Antarctic, there are a number of well characterized lake systems there which have been analysed for pesticides (Sarkar et al. 1994; Gasparon and Burgess 2000), and are candidates for determination of historical deposition profiles. Lake sediments are available for sampling where ice cores may not be readily available owing to low altitude or where peat bogs are uncommon due to historically low plant growth rates. Lake sediments are more accessible and present fewer analytical challenges than snow or ice cores. Historical trends and fluxes can therefore be determined for a wide range of polar environments. Furthermore, sediments can be resampled if necessary within a few metres of the original sampling location. This is also true of glacial ice, but re-sampling of surface snow on a glacier can be problematic, because the surface properties of snow crystals change during aging after deposition (Franz et al. 1997). As well, sediments also have typically higher temporal resolution than soils owing to lower turnover of net depositional material and to particle focusing. Spatial and temporal trends of metals, persistent organic pollutants and anthropogenic particles Trends of mercury and lead inferred from arctic sediment cores Background There have been a large number of studies of heavy metals, especially of Hg and Pb, in sediment cores from arctic lakes. Landers et al. (1998) and the Arctic Monitoring and Assessment Program (Dietz et al. 1998) have reviewed the work done to the mid-1990s. Therefore, it is mainly the more recent work that is reviewed here. Our objective is to assess the importance of anthropogenic inputs from the circumpolar Arctic, and examining the design of each study. To examine the importance of anthropogenic inputs of Hg and Pb, enrichment factors (EFs) reported by the authors were examined on a circumpolar basis. The EFs were based either on ratios of concentrations in recent (near surface) sediments to pre-industrial concentrations or on the recent versus historical flux (ng m-2 yr-1) ratios.

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Mercury Bindler et al. (2001b) studied Hg deposition in dated sediment cores from three mountain lakes in northernmost Sweden and reported an EF for Hg of about 2. A study of 100 lakes in Sweden, including 12 arctic lakes, examined the deposition pattern for Hg (Swedish Monitoring Program 2002). Metal analyses were confined to three levels in the sediment cores, 0-2, 2-4 and 30-32 cm rather than individual slices. The 12 most northerly lakes in the study had mean EFs for Hg of 1.12 based on concentrations in the 0-4 cm depth interval relative to the concentrations at 30-32 cm. By comparison, Hg in lake sediments from central/northern Sweden had EF values from 2.2 to 3.5. Both studies observed a slight decline of Hg in the most recent sediment layers reflecting lower deposition of Hg in the Swedish Arctic in the 1990s. Previous studies of Hg deposition in the 1980s in Sweden had found maximum Hg concentrations in the surface 0-1 cm interval (Johansson 1985, 1989). Thus, the selection of sampling depth of the surface slices was an important consideration for estimating the EF and has implications for comparisons between EF studies undertaken at different times. Bindler et al. (2001b) also found significant correlations between Hg and spheroidal carbonaceous particle (SCP) concentrations in six cores from Swedish lakes; however, the Hg maximum did not always coincide with maximum SCP. The measurement of SCP along with Hg served to highlight the differences between deposition via particles with short atmospheric residence times versus a species with long atmospheric lifetime that is both on particles as a cation (Hg2+) and in the gaseous phase as Hg0. Bindler et al. (2001c) investigated Hg in sediments of 21 lakes in southwestern Greenland (Kangerlussuaq region, 67°N) along a 150 km transect from the coast to the ice sheet margin. Sampling design was similar to the Swedish lake survey – surface slices versus deep slices were analysed from 18 of 21 lakes whilst full historical profiles were examined in three lakes. Mercury EFs ranged from 0.6 to 6.0 (mean 2.1) with the highest values in the lakes closest to the ice margin. Hg was found to be strongly correlated with % organic carbon in sediment cores and therefore the authors calculated EFs based on organic carbon normalized concentrations. Although a similar concentration gradient is observed for SCP (Bindler et al. 2001c; N.L. Rose, unpublished data), the opposite trend was observed for Pb pollution which was found to be higher in lakes near the coast (Bindler et al. 2001a). The historical profile in dated sediment cores from the region indicated that the Hg pollution started in the late 19th century, possibly as early as the 17th century (Bindler et al. 2001c). Rognerud et al. (1998) examined Hg and other heavy metals (see below) as well as Fe, Al and organic carbon (as loss-on-ignition) in surface and pre-industrial sediments in cores collected from 66 lakes in the Norwegian and Russian Arctic. Samples were collected at the deepest point in each lake with the help of a portable echosounder. This study used analysis of covariance to adjust the data for important covariates, organic matter, Al and Fe. Other variables such as lake size, catchment area, pH, and total organic carbon in water were not considered significant. This large spatial survey revealed significantly higher concentrations of Hg in surface sediments than in preindustrial sediments. The differences decreased with increasing latitude and east longitude, i.e., away from regions of major industrial/urban emissions. The most distant area sampled, including Wrangel Island off the northeastern Siberian coast, exhibited 2-fold increases of Hg over pre-industrial sediments possibly reflecting Chinese

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sources. On Svalbard and along the arctic coast of Russia, only Hg was elevated in surface sediments among the heavy metals that were analysed. Blais et al. (1998) found that Hg did not show a significant decline in concentrations in sediment cores with distance from the smelters of Noril’sk, reflecting its long atmospheric half-life compared to most other heavy metals, which are particle associated in the atmosphere. EFs for Hg in lakes around Noril’sk averaged 2.6, similar to other arctic lakes (Figure 2) and less than in mid-latitude lakes in North America (Swain et al. 1992; Lorey and Driscoll 1999) and Scandinavia (Johansson 1989) sampled in the 1980s. Another study showed that a lake sediment core collected from the eastern Taymyr Peninsula in northern Russia in the mid-1990s had no recent enrichment by Hg or other atmospherically transported heavy metals (Landers et al. 1998).

Figure 2. Circumpolar view of Hg enrichment factors (EFs) in sediment cores. Results from Landers et al. (1998) for Alaska, Taymyr Peninsula and Canada; Lockhart et al. (1998), Cheam et al. (2001) and Muir et al. (2003) for other Canadian sites; Bindler et al. (2001c) for southwestern Greenland; Rognerud et al. (1998) for northern Norway, Svalbard and Wrangel Island; Blais et al. (1998) for Noril’sk area; Bindler et al. (2001b) for northern Sweden; Mannio et al. (1997) for northern Finland. Results with open-crosshatched bars are averages of cores from lakes in the region of the bar.

Hg concentrations have been determined in about 36 dated sediment cores collected from lakes from across the Canadian Arctic and Subarctic during the 1990s (Lockhart et al. 1995, 1998, 2000b; Hermanson 1998; Cheam et al. 2001; Muir et al. 2003). EFs of Hg in these arctic sediment cores ranged from 0.8 to 7.1 (mean 1.9). Even lower EFs for Hg were found in seven lakes in central and northern Alaska, ranging from 0.7 to 1.3

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(mean 1.1) (Gubala et al. 1995; Landers et al. 1998). These Hg EFs for the North American Arctic were generally lower than those found in small, isolated mid-latitude lakes in boreal forest regions of southern Canada (1.8-2.9; Lockhart et al. 1995, 1998; Lucotte et al. 1995) and mid-latitude/northern USA (3 to 5; Swain et al. 1992; Lorey and Driscoll 1999). Much higher Hg EFs have been found in lakes impacted by nearby urban areas such as Lake Tahoe (EF = 24; Heyvaert et al. 2000). The EFs for Hg in the North American Arctic are also generally lower than those of southwestern Greenland (Bindler et al. 2001c) and in the European Arctic (Rognerud et al. 1998, 2001). The pattern of EFs for Hg in the Arctic, as defined by the sediment core studies, points to a region of low anthropogenic input in Alaska and the Canadian High Arctic. In addition to low EFs in Alaska, Lake Hazen on Ellesmere Island, and Yaya Lake in the Mackenzie Delta in the western Canadian Arctic showed no recent Hg increases (Landers et al. 1998). EFs can be difficult to compare among lakes because different time intervals may be used. For example, Lockhart et al. (1998) used EFs based on average flux ratios for slices dated post-1950 that were compared with pre-industrial sediment layers, rather than surface slices only. Landers et al. (1998) selected various intervals to try to obtain average recent fluxes. In so doing they may have understated the EFs compared to other studies in Scandinavia, Russia and northern USA where a “top-bottom” ratio has generally been used (e.g., Johansson 1989; Verta et al. 1989; Rognerud and Gjeld 1993, 1998; Bindler et al. 2001b). Sediment core studies in the Canadian Arctic and Alaska, based on cores collected in the late 1980s and early ‘90s, did not show the decline in Hg deposition that has been observed in Sweden (Bindler et al. 2001b) and in some lakes near sources in midlatitude North America (Engstrom and Swain 1997; Lockhart et al. 1998; Lorey and Driscoll 1999) as well as in mid-latitude glaciers (Schuster et al. 2002). However, more recent collections (1998-2001) have shown sub-surface maxima for Hg and declining fluxes during the 1990s in cores with high sedimentation rates (Muir et al. 2003). This relatively slow response to reduced Hg emissions in North America and Europe may reflect the fact that remote arctic lakes are influenced more by global background concentrations and unique atmospheric processing of Hg in the Arctic rather than primary sources. Inputs of Hg2+ to lake surfaces and catchments may occur as a result of annual mercury depletion events (MDE) in spring time owing to interaction of sunlight and bromine oxide (BrO) formed from bromine emitted from the Arctic Ocean (Schroeder et al. 1998; Lu et al. 2001). The MDEs appear to be unique to regions within or adjacent to the Arctic Ocean and antarctic marine waters (Ebinghaus et al. 2002). They occur between polar sunrise and complete snowmelt, and appear to be closely coupled with stratospheric ozone depletion events (Lu et al. 2001; Ebinghaus et al. 2002). Total Hg concentrations measured in the sea ice snow pack north of Alaska were about one order of magnitude higher downwind of the major leads and polynyas off Point Barrow (Alaska) than in snow from sea ice locations predominantly upwind of the open water areas surrounding Point Barrow (Garbarino et al. 2002). High total Hg concentrations in snow are well correlated to areas of high atmospheric BrO concentrations within the Canadian Archipelago as well (Welch et al. 1999; Lu et al. 2001). A Hg deposition gradient away from coastal areas is implied by these results and indeed Landers et al. (1995) showed Hg levels in Alaskan tundra vegetation were inversely related to distance from the nearest coast. Snyder-Conn et al. (1997) also

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showed total Hg levels in Alaskan tundra snow pack were inversely related to distance from the nearest coast. However, the pattern seen in sediment cores (Figure 2) in the North American Arctic and Greenland does not conform to this trend. EFs tend to be higher in the more southerly and inland lakes, especially those in the eastern/central Canadian Arctic (Muir et al. 2003), and were higher in lake sediments near the ice cap rather than near the coast in southwestern Greenland (Bindler et al. 2001c). Similarly, Hg EFs are higher in southern Norway and Sweden compared to the ocean-influenced environment of Svalbard. This confirms the conclusions of Landers et al. (1998) who worked with a slightly older data set, that anthropogenic Hg deposition is the dominant regional deposition process in some northern regions near sources but of lesser importance in areas remote from urban/industrial Hg emissions. Another factor that might contribute to the observed lack of decline of Hg in arctic sediments is the warming trend which has become apparent especially in the North American Arctic (Macdonald et al. 2003). Longer open water seasons and higher temperatures observed during the 1990s may be leading to increased phytoplankton productivity thereby increasing the rate of Hg scavenging via sedimenting particles from the water column into sediments (Outridge et al. 2001; Lockhart 2003). The profile of Hg in a varved sediment core from a lake on Devon Island corresponded well to the increase in deposition of some diatom species observed in a separate core from the same lake. Furthermore, the chemical conditions in the sediment (especially low organic content and lack of available sulfides for Hg binding) could have promoted Hg diffusion up-core (Outridge et al. 2001). These factors remain hypotheses at this stage but could be examined in more detail by a combined geochemical and paleolimnological investigation of high arctic lake sediments. Roos-Barraclough et al. (2002) have demonstrated, using a dated peat core from Switzerland, that pre-industrial Hg deposition was climate-controlled, with higher deposition during cold periods, and also influenced by volcanic eruptions. Mercury deposition in peat was also correlated with bromine in historic (pre-1300 AD) sediments implying a link of the geochemical cycles of these two elements. Lead and other metals In addition to Hg, large scale spatial trends in the deposition of Pb and other heavy metals have been conducted in Scandinavia and on the Kola Peninsula. In Sweden, a monitoring study of 100 lakes, that included 12 arctic lakes, examined the anthropogenic deposition pattern for Pb and Cadmium (Cd) (Swedish Monitoring Program 2002). Lead and Cd EFs averaged 4.6 and 1.9, respectively, in the 12 lakes illustrating the continuing importance of anthropogenic sources for Pb. Rognerud et al. (1998) examined the heavy metals Pb, copper (Cu), nickel (Ni), zinc (Zn), as well as selenium (Se), iron (Fe), aluminium (Al) and organic carbon (as losson-ignition) in surface and pre-industrial sediments in cores collected from 66 lakes in the Norwegian and Russian Arctic. This large spatial survey found significantly higher concentrations of Pb closer to regions of major industrial/urban emissions. No other elements (Cu, Ni, Zn, Se) showed increases. There was a shift in Pb association from inorganic matter (as measured by Al) in the reference sediments to organic matter in the surface sediments, which was interpreted as an historic change from geologic sources to atmospheric deposition. The study did not detect the regional deposition of heavy

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metals particularly Ni and Cu, around smelters on the Kola Peninsula and near Noril’sk in the Russian Arctic. This was probably due to sampling in small lakes along the Russian coastline and in northern Norway, rather than those in close proximity to the smelters, and it illustrates the relatively small heavily impacted area associated with the smelters (Rognerud et al. 1998). Another limitation of the study was that the cores were undated and thus the time period represented by the top layer and the age of the reference layer were unknown. However, the authors were careful to use only the top 0-1 cm layer, and a very deep layer for the reference or background metal concentration, and they had general information on average deposition rates in small lakes in their study regions. A survey of Norwegian lakes in 1996-1997 based on 210 sediment cores showed that the elements with highest EFs were antimony (Sb), bismuth (Bi), arsenic (As), Cd, and Pb (Rognerud and Fjeld 2001). Higher concentrations of Sb, Bi and Pb were found in southern Norway, generally decreasing with latitude and altitude. The exception was the northeastern part of Norway, close to the Russian border, where elevated concentrations, especially of Ni and Cu, were found due to the influence of smelters at Nikel and Zapolyarny in northwestern Russia. The regional distribution pattern of these elements was most likely due to long-range anthropogenic atmospheric deposition, with some influences from local smelters on the levels of primarily Cu and Ni. This is the first report of anthropogenic enrichment of Sb and Bi in remote lake sediments and it illustrates the need to consider a wide range of metals to evaluate sources and longrange transport. Dated sediment cores collected closer to the large base metal smelters on the Kola Peninsula in Russia have confirmed the importance of local inputs of cobalt (Co), Cu, Ni, and Pb to recent sediments (Norton et al. 1996) and demonstrate transport in a northeasterly direction. However, deposition declines rapidly with distance from these sources, reflecting the association of the metals with the coarse aerosol fraction which declines exponentially from the smelters (Sivertsen et al. 1992; Shevchenko et al. 2003). Blais et al. (1998) found that concentrations of Cu and Ni were greatly enriched in sediment cores from lakes near the smelters of Noril’sk. Concentrations of Cu, Ni, barium (Ba) and Co in surface (0.5 cm) samples declined exponentially with distance up to 100 km from the smelters indicating a point source influence. A lake sediment core from the eastern Taymyr Peninsula in northern Russia indicated no recent enrichment by atmospherically transported elements (Landers et al. 1998). Inputs of recent anthropogenic and natural Pb to arctic lake sediments have been studied by examining stable Pb isotope ratios in recent and pre-industrial sediments. In southwestern Greenland, Pb concentrations in dated lake sediment cores showed that significant concentration increases occurred during the 18th and 19th centuries, with a maximum in the 20th century around 1970 (Bindler et al. 2001a). The enrichment factors were 2.5, with slightly higher values closest to the coast. There was a decline in the 206 Pb/207Pb ratio in the recent sediments as compared to deeper sediment layers, indicating an increased influence of airborne anthropogenic Pb. By calculating an excess Pb isotope ratio using an isotope mixing model, it was concluded that the Pb isotope ratios of the lake sediments in southwestern Greenland (1.14-1.15) showed a strong influence of western European emission sources (206Pb/207Pb = 1.14). These findings differed from results for atmospheric aerosols in the Canadian Arctic at Alert and Mould Bay (Sturges and Barrie 1989) and at Barrow, Alaska (Sturges et al. 1993) in

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the mid-1980s which had shown a strong contribution from Russian sources (206Pb/207Pb = 1.15-1.16). However, no previous study had examined the deposition of the aerosolassociated Pb isotopes in arctic freshwater environments. Outridge et al. (2002) conducted a similar study in dated sediment cores from three lakes in the Canadian Arctic Archipelago in the eastern High Arctic and from Yaya Lake in the Mackenzie Delta area of the western Canadian Arctic, as well as two lakes near Hudson Bay. Pb concentrations and anthropogenic Pb accounted for no more than 19% of acid-leachable Pb in recent sediments from the High Arctic and western lakes. In contrast, two lakes situated in the western and southeastern Hudson Bay lowlands showed significant anthropogenic Pb sources which accounted for 70-90% of acid-leachable Pb by the 1980s or 1990s. Isotopic trends through time indicated that Eurasian sources contributed most of the anthropogenic Pb to northwestern Hudson Bay, with possibly a minor Canadian contribution at the southeastern Hudson Bay site. Similar latitudinal differences were reported by Cheam et al. (2001), with cores from four high latitude lakes exhibiting no recent Pb concentration increases while one lake in northern Québec (latitude 54°N) showed evidence of anthropogenic Pb. Outridge et al. (2002) concluded that reduced atmospheric Pb deposition at higher latitudes was associated with lower precipitation rates, while over Hudson Bay the Polar Front, an area of meteorological disturbance, may play a role in increasing both precipitation and Pb deposition. Bindler et al. (2001a) also noted the importance of the precipitation gradient from the coast to the ice sheet margin in southwestern Greenland (Kangerlussuaq region) for explaining the higher fluxes near the coast. Trends of POPs inferred from arctic lake sediments Background There are relatively few studies of POPs in dated arctic lake sediment cores compared with heavy metals. Circumpolar geographic coverage is very incomplete with almost no results from Scandinavia, Russia, or Greenland. In the European Arctic the emphasis has been on analysis of surface sediments, particularly from the marine environment (de March et al. 1998). While the reasons for this information gap are not clear, they probably have to do with the high cost of analyses and the analytical and sampling challenges. Concentrations of POPs such as PCBs and DDT are typically very low (low ng g-1 dry wt) in most arctic lake sediments compared to heavy metals (de March et al. 1998; Macdonald et al. 2000) and large samples and “clean” techniques are needed for reliable results. Investigations to the mid-1990s have been reviewed by de March et al. (1998) and Macdonald et al. (2000), thus mainly the more recent work is reviewed here with the objective of assessing the importance of historical and spatial trends in inputs. Organochlorines Dated sediment cores from approximately 27 lakes in the Canadian Arctic, Alaska and Svalbard were analysed for PCBs, DDT and other persistent organochlorines during the 1990s (Gubala et al. 1995; Muir et al. 1995, 1996; Cleverly et al. 1996; Evans et al. 1996; Lockhart 1997; Rawn et al. 2001; Muir et al. 2002; Rose et al. 2004). Recent (1990s) fluxes (ng m-2 yr-1) of 6PCB and 6DDT in these lakes, corrected for particle

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focusing, are plotted in Figure 3. Lakes in the southern Yukon Territory (Laberge, Watson, and Little Atlin; Rawn et al. 2001) and Lake Tenndammen near coal mining towns in Svalbard were omitted because of the influence of local sources of contamination. No latitudinal trend of 6PCB fluxes was observed. However, 6PCB fluxes in the western Canadian arctic lakes (Fox, Kusawa, Lindeman, Hanson, Great Slave, and Yaya) were similar and higher than fluxes in lakes from the Canadian High Arctic, Alaska and Svalbard. Highest 6PCB fluxes were found in a core from the western basin of Great Slave Lake (Evans et al. 1996), which receives inputs from northern Alberta and British Columbia via the Peace and Athabasca rivers. In the case of 6DDT, higher fluxes were found in the subarctic lakes than in high arctic and Alaskan lakes. This is due to a combination of long-range transport/deposition from southern sources and from the influence of past use of DDT which was particularly evident in Yukon lake sediments (Rawn et al. 2001).

Figure 3 . Fluxes (ng m-2 yr-1) of 6PCB and 6DDT in dated lake sediment cores (averaged for slices dated to the 1990s) from lakes in the North American Arctic and Svalbard. While fluxes of both organochlorines are generally higher in subarctic lakes, no significant geographical trend with increasing north latitude was found. Results for the North American Arctic are compiled from Gubala et al. (1995), Muir et al. (1995, 1996), Cleverly et al. (1996), Evans et al. (1996), Lockhart (1997), Rawn et al. (2001), and Muir et al. (2002), and for Svalbard from Rose et al. (2004).

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A latitudinal study of organochlorine contamination in bottom sediments of 100 Swedish lakes, many of which were part of the National Swedish Environmental Monitoring program (Swedish Monitoring Program 2002), included two lakes from the Swedish Arctic mountains and four other high latitude lakes (Söderström et al. 2002). This study compared concentrations of PCBs (seven congeners), three DDT-related compounds, HCB and three HCH-isomers in surface (0-2 cm) samples of each lake. The six Swedish arctic lakes had significantly lower mean concentrations of 6PCB (0.9 ± 0.5 ng g-1 dw), 6DDT (1.0 ± 1.0 ng g-1 dw) and 6HCH (0.1 ± 0.05 ng g-1 dw) than southern lakes (9.2 ± 6.3 ng g-1 dw, 14.6 ± 12.1 ng g-1 dw, and 1.3 ± 1.5 ng g-1 dw, respectively). For HCB the six northern lakes had about 3-fold lower concentrations than the 60 southern lakes in the study reflecting the longer atmospheric half-life of HCB than the other compounds and lack of major past local uses in the south. Skotvold and Savinov (2003) examined the regional distribution of PCBs in surface slices (0-2 cm) of sediment cores from 26 lakes in northern Norway, Svalbard, Bear Island and tundra ponds along the Russian arctic coastline. Concentrations of sum of 10 PCB congeners 610PCB) were generally low ranging from geometric means of 0.34 ng g-1 dw in tundra ponds along the Laptev Sea, 11.2 ng g-1 dw in four lakes in central/northern Norway (Nordland) and 6.0 ng g-1 dw in a lake on Svalbard. Sediment from Lake Ellasjøen on Bear Island (74ºN, 19ºE) in the Barents Sea had much higher concentrations (35 ng g-1 dw) than all other locations. In Ellasjøen, guano from seabirds was also identified as a potential source of contaminants. Indeed, high concentrations of recalcitrant PCB congener 153 point to a non-atmospheric source for PCBs in this lake. These data are in good agreement with the sum of 10 PCB congener concentrations for surface sediments from five lakes along the west coast of Svalbard sampled in 1997 (Rose et al. 2004) where the 610PCB ranged between 2.5 and 13.5 ng g-1 dw. This latter study involved the analysis of surface (0-1 cm) and pre-industrial sediments from dated cores. Highest concentrations and fluxes were recorded in the surface slice of a core from Lake Tenndammen located near the coal mining towns of Barentsburg and Longyearbyen. This lake also showed elevated SCP and PAHs in surface slices compared to the other lakes in Svalbard although levels of contamination were low compared to European sites (Rose et al. 2004). These spatial trend studies in Sweden and the Norwegian and Russian Arctic illustrate how a large array of lakes can be used to assess broad geographical trends in deposition of POPs. Similar approaches have been used in North America (e.g., Muir et al. 1995, 1996) but with far fewer and less well characterized lakes than in Sweden or Norway. A limitation of this approach is that the sediment cores were not dated, thus sedimentation rates and the degree of particle focussing were not known. Given generally lower sedimentation rates especially in the arctic mountain lakes and in other lakes above the treeline, it is reasonable to assume that a much longer period of deposition was represented by the 2 cm layer in the high latitude lakes than in those from southern Sweden or northern/central Norway. This complicates the geographical comparison since only recent deposition may be measured in high sedimentation sites versus longer term deposition at low sedimentation sites. The sediment records of POPs in arctic lakes have also provided information on temporal trends of deposition of these hydrophobic contaminants in the Arctic. The AMAP assessment of POPs included results for PCBs and polychlorinated dibenzo-pdioxins and dibenzofurans (PCDD/Fs) in sediment cores from Alaska, Canada, and

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Finland collected and analysed in the early/mid-1990s (de March et al. 1998). The historical profiles of total PCBs and trichlorobiphenyls, a more volatile homolog group, are shown in Figure 4 for 12 lakes in the Canadian Arctic ranging from 60 to 81 degrees

Figure 4. Historical profiles of PCBs in dated sediment cores from lakes in the Canadian Arctic (redrawn from Muir et al. 1996, 2002; Rawn et al. 2001). Black portion of the bar represents trichlorobiphenyls which have been found to generally increase in proportion to6PCB with northern latitude. Entire bar represents 6PCB concentration in each slice. Dotted bell-shaped curve represents the "source function" for global PCB production, i.e., starting in the 1930s with a maximum in the early 1970s (from Brevik et al. 2002) adjusted for the dating of each core.

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northern latitude. Most of the high latitude lake cores show a later onset of PCB deposition than the historical global PCB production curve depicted in each graph, as well as maximum inputs near or at the surface indicating continuing inputs from atmospheric deposition. In general, lakes nearer sources, for example the Great Lakes (Eisenreich et al. 1989) and lakes in northwestern Ontario (Muir et al. 1996), tend to show PCB deposition patterns resembling the PCB production curve in which maximum production occurred in the 1960s and early 1970s prior to phase-out in the late 1970s (Breivik et al. 2002). Exceptions among the 12 lakes in Figure 4 are Romulus Lake, Lake Laberge and Char Lake. The latter two can be explained by proximity to pollution sources. Lake Laberge is downstream of the city of Whitehorse, Yukon, and has received inputs of DDT and other chemicals (Rawn et al. 2001). Char Lake is near the airport in the village of Qausuittuq (Resolute) and could have received small inputs during the airport development in the 1950s. The pattern in Romulus Lake is more difficult to explain because it is an isolated, meromictic lake on Ellesmere Island. The historical profile in Romulus fits trends in global PCB production rather well but shows a sharp decline in the late 1960s. Further analysis of the data from Canadian lakes shows that lower chlorinated PCBs generally predominate in cores from the High Arctic while sediments from temperate lakes have higher proportions of hexa-, hepta-, and octachlorobiphenyls (Muir et al. 1996, 2002) which have shorter atmospheric travel distances (Beyer et al. 2000). Polyaromatic hydrocarbons While historical records of deposition of persistent organochlorines are very limited from the European Arctic, deposition of PAHs in lake sediments has been studied in greater detail. Fernández et al. (1999, 2000) calculated fluxes of 23 PAHs for Lake Arresjøen on northwestern Svalbard (79º40’N) and Øvre Neådalsvatn (62º46’N) in the Caledonian mountain range of central Norway. Major PAHs in Lake Arresjøen were primarily pyrolytic in origin (i.e., resulting from combustion of coal and hydrocarbon fuel burning) such as phenanthrene, fluoranthene, pyrene, chrysene/triphenylene, indeno[1,2,3-cd]pyrene and benzo[ghi]perylene. Fernández et al. (1999) concluded that the PAH deposition pattern in high altitude mountain lakes, including Arresjøen, paralleled sulfate deposition, pointing to combustion particles as the main input pathway. Rose et al. (2004) determined 15 PAHs in surface and pre-historical slices of sediment cores from four lakes in Svalbard. Highest fluxes were found in Lake Tenndammen (360 µg m-2 yr-1; N.L. Rose, unpublished data) a lake within 20 km of the coal mining towns of Barentsburg and Longyearbyen. PAH fluxes in the four Svalbard lakes appeared to decline with distance from the Barentsburg/Longyearbyen area. Fluxes of total PAH (same 23 PAHs as analysed by Fernández et al. (1999) excluding perylene and retene) have also been reported in a series of studies of dated sediment cores from the Canadian Arctic (Lockhart 1994, 1996, 1997; Muir and Lockhart 1994; Lockhart et al. 1995, 1997). Total PAH fluxes in southern Yukon lakes ranged from 9.1 µg m-2 yr-1 in remote Kusawa Lake to 174 µg m-2 yr-1 in Little Atlin Lake (Lockhart et al. 1997). Two remote lakes in the Mackenzie River basin, Lac Ste. Therese and Yaya Lake, had much higher PAH fluxes (68 and 140 µg m-2 yr-1, respectively) (Lockhart 1997). The source of PAHs for the core from Yaya Lake was thought to be annual inputs from the Mackenzie River during spring floods. High PAH values have also been

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found in Mackenzie River suspended particulates and they are thought to be mainly of petrogenic origin (Yunker and Macdonald 1995). The recent (late 1980-1990s) PAH fluxes for the European and Canadian Arctic are compared in Figure 5. In general, higher fluxes of substituted PAHs have been found in the Mackenzie River valley and Yukon lakes compared with the European arctic lakes. Pyrolytic PAH inputs to Arresjøen showed a steady increase from the 19th century with a maximum in top slices (Fernández et al. 2000). Lake Arresjøen had the lowest PAH fluxes of any of the 10 European mountain lakes studied. The more southerly Øvre Neådalsvatn had maximum inputs in the 1930-50s and recent inputs about 4-fold higher than in the mid-19th century. This is in agreement with results for PAHs in dated sediment cores from North American subarctic and arctic lakes.

Figure 5. Recent fluxes (Pg m-2 yr-1) of total PAH (23 PAHs minus perylene and retene) in dated lake sediment cores from the Canadian and Norwegian Arctic. Results compiled from Fernández et al. (1999) and Rose et al. (2004) for Norway/Svalbard, and for the Canadian Arctic from Lockhart (1994, 1996, 1997), Muir and Lockhart (1994), and Lockhart et al. (1995, 1997).

In the historical profiles of PAHs reported for eight lakes (Lockhart 1994, 1996, 1997; Muir and Lockhart 1994; Lockhart et al. 1995, 1997), total PAHs, especially the threeand four-ring pyrolytic PAHs, e.g., fluoranthene, pyrene, retene, benzo(a)pyrene increased in concentration from the mid-19th century. Hawk Lake, in the central Canadian Arctic (Figure 5), had a distinctive sub-surface maximum for total PAHs dated to the 1950s, similar to mid-latitude lakes in northwestern Ontario (Lockhart et al. 1995). Hawk Lake also had elevated levels of retene, a tracer of wood smoke. The distinctive mid-20th century peak of pyrolytic PAHs was not observed in Lac Ste.

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Therese, which is further west in the watershed of Great Bear Lake but at the same latitude as Hawk Lake. This lake also had low EFs for Hg suggesting that the basin was little influenced by combustion-related sources (Lockhart 1997). Yaya Lake, in the Mackenzie River Delta (Figure 5), did show a sub-surface maximum of total PAHs from the period 1940 to 1970 but returned to pre-1940 levels in the 1980s (Lockhart 1997). The sources of PAH to this lake could include annual spring floods as well as atmospheric deposition, and the relative importance of these pathways is difficult to evaluate. The historical profile of PAHs in Lake Laberge, which is part of the Yukon River system, showed very similar fluxes for the period 1900-1990 and higher fluxes in slices dated to 1880-1900. These latter slices also had higher concentrations of retene corresponding to the period of the Klondike gold rush and this may reflect substantial wood burning during that time period (Muir and Lockhart 1994). Anthropogenic particles in arctic lake sediments Background It is now almost 50 years since Mitchell (1956) coined the term ‘Arctic Haze’ observing that it comprised particles, but not ice crystals, d 2 Pm in diameter. Although it took some time for the implications of this study and its significance for the widespread contamination of this ‘pristine’ area to be realized, a considerable volume of work has since been published on the presence, composition and transport of aerosols to the Arctic (e.g., Barrie 1986; Heintzenberg 1989; Anderson et al. 1992; Virkkula et al. 1995; Lowenthal et al. 1997; Xie et al. 1999; Hung et al. 2002). By comparison, the Antarctic appears to have remained less affected (e.g., Hara et al. 1996; Correia et al. 1998; Legrand et al. 1999; Arimoto et al. 2001). In contrast to the number of studies dealing with atmospherically sampled aerosols in polar regions, the number relating to paleolimnological studies of anthropogenicallyderived particulates is rather limited. There are probably a number of reasons for this. Anthropogenic particles, detectable in the lake sediment record, are of necessity quite large (typically > 2 Pm) and hence will be detectable in very low concentrations in areas remote from sources. Therefore, the limited long-range transport of these particles (in contrast to gases) and the detection limits of available techniques conspire to confound the polar paleolimnologist. For these reasons, particle types which are unambiguous, and hence identifiable to anthropogenic sources, are invaluable for studies in minimally impacted regions and, as a consequence, polar paleolimnological studies on anthropogenic particles are almost exclusively based on fly-ash. Indeed, to our knowledge, the only paleolimnological studies not solely relating to fly-ash particles in the Arctic or Antarctic are those of Doubleday et al. (1995) and Doubleday (1999). Despite these limitations, there is evidence for long-range transport of fly-ash particles over thousands of kilometres. Industrial spherules > 5 Pm in diameter were recorded in all atmospheric samples taken on a transect across the Atlantic Ocean (Folger 1970), constituting 5% of the total number in mid-ocean samples and more than 60% near land (Parkin et al. 1970). Spherules have also been found in high latitude ice deposits in both Greenland (Hodge et al. 1964) and the Antarctic (Fredriksson and Martin 1963; Hodge et al. 1967). A mechanism that would explain the transport of observed continental soot to Antarctica has been proposed (Murphey and Hogan 1992)

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although the inorganic spheres may have been confused with spheres of volcanic or meteoritic origin (Hodge and Wright 1964). However, in some regions of the Arctic it is not necessary to invoke long-range transport, as industrial sources occur within the polar circle. These include the well known smelters in Nikel on the Kola Peninsula and Noril’sk in Siberia previously discussed as sources of heavy metal pollution. Smaller sources include coal-fired power stations on Svalbard and potentially other municipal waste burning sites throughout the circumpolar Arctic. Paleolimnological data may therefore provide useful information on the spatial and temporal distribution of particulates from these sources and the relative roles of local versus long-range sources over the broader region. Fly-ash The definition of fly-ash in relation to past paleolimnological studies has been dealt with more fully in Rose (2001), but briefly, these are particles emitted with the fluegases after the high temperature combustion of fossil fuels such as coal or oil. For the paleolimnologist there are two types of fly-ash particles, SCPs produced by the incomplete combustion of the fuel and composed mainly of elemental carbon, and inorganic ash spheres (IASs) formed by the fusing of the inorganic component within the fuel (hence mainly from coal-series fuels) and composed primarily of aluminosilicates with varying quantities of other elements, such as iron. Due to their composition, SCPs are easier than IASs to extract from lake sediments by chemical means. Further, they are easier to unequivocally identify and hence most paleolimnological studies that deal with fly-ash use SCPs. This appears to be especially true for the polar regions despite the fact that IASs tend to be smaller and may travel to these remote areas more readily than their larger carbonaceous counterparts. The geographical distribution of arctic paleolimnological particle studies is similarly skewed with the majority having been undertaken in Europe with smaller scale studies in Greenland, Canada and Siberia. In Europe this has been a combination of both spatial (i.e., a contemporary ‘snap-shot’ of SCP distribution) and temporal (historical trends in SCP deposition) studies. Spatial studies have been undertaken in Sweden (as part of a national survey; Wik and Renberg 1991), Finland (N.L. Rose, unpublished data) and Svalbard (Rose et al. 2004). The Swedish survey included 11 sites within the Arctic Circle and SCP concentrations varied considerably within this area from < 500 per gram dry mass of sediment (gDM-1) to 4000-8000 gDM-1. This latter concentration was a single point resulting from the impact of local sources in the Kiruna and Gallivare/Malmberget area superimposed on elevated concentrations from a large smelter further south on the northeastern coast (I. Renberg, pers. comm.). The spatial SCP data also showed good agreement with the geographical distributions of both Pb in moss and excess sulphate in wet deposition. The Svalbard study (Rose et al. 2004) involved the analysis of surface sediments taken from 21 lakes on the west coast of the archipelago between Bellsund (77q33’N) and Amsterdamøya (79q46’N) and included sites within a few kilometres of the small coal-fired power station at Longyearbyen (also at Pyramiden and Barentsburg) in Isfjord. In contrast to the Swedish study, the SCP concentrations in these surface sediments were quite similar with all values falling in the range 200-1500 gDM-1, and although the higher concentrations were generally seen to occur around the Isfjord area,

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it was not until SCP inventories were calculated that the influence of the local coal-fired power station could be fully appreciated. Inventories at the sites closest to Isfjord were seen to be an order of magnitude higher than those at the most remote locations. It was concluded that the local sources were having a measurable but spatially limited (within ca. 50 km) influence superimposed on a more diffuse long-range pattern. This study highlights the caution required when comparing surface sediment concentration data and shows that without information on sediment accumulation rates any conclusions are at best speculative. The Finnish survey undertaken in 2000-2001 analysed SCP concentrations in the surface sediments of 34 lakes in northeastern and northwestern Finland (N.L. Rose, unpublished data). All concentrations were within the range < 200-1000 gDM-1. These lakes were selected to fulfil strict criteria and it is therefore unlikely that sediment accumulation rates vary greatly between sites. Therefore these concentrations, being remote from sources, probably reflect patterns of deposition from long-range transport to the area. Indeed, SCP concentrations of d ca. 1000 gDM-1 in contemporary sediments appear to be typical of many remote areas around the northern hemisphere. Temporal studies have also been conducted within the Arctic Circle in both Finland (Korhola et al. 2002) and Svalbard (Rose et al. 2004) whilst studies in remote northern areas (but just south of the Arctic Circle) have been undertaken in Sweden (Wik and Renberg 1996) and Iceland (N.L. Rose, unpublished data). In general, the patterns in the sediment cores from all these areas are similar. The SCP profiles are short, due to a combination of slow accumulation rates and concentrations below analytical detection limits, and either show increases in concentration to maxima at the surface or are highly variable. Of these, the former profile type indicates both the low temporal resolution resulting from slow accumulation rates and the shift from SCP absence to presence as concentrations increase above the detection limit through time, whilst the latter result from the level-to-level variability in very low particle numbers. In summary, the characteristic SCP profiles seen in lake sites throughout Europe are often absent and hence the use of SCPs as sediment dating tools (e.g., Rose et al. 1995) in these systems is usually limited. A good example of this is seen in the Svalbard data (Rose et al. 2004) where six SCP profiles were 210Pb-dated. The longest temporal record was observed at the sites closest to the local sources whereas in the remainder the SCP record began either in the 1950s or the 1980s, periods with significant expansions in local industrial coal-use thus representing the shift from below to above analytical detection limit at each site. In these cases interpretations are probably limited to the positive identification of the presence of pollutants from atmospheric sources and to discussion of the direction of change in impact status. A notable exception to the above is the sediment core taken in 1995 from Saanajärvi in northern Finland (69q05’N, 20q52’E) (Korhola et al. 2002). The study yielded a SCP profile showing all the typical features. The start of the record in the mid- to late 19th century, the rapid increase in SCP concentration in the mid-20th century and the subsurface concentration peak (here, in the early 1980s) and subsequent decline to the sediment surface are all present, suggesting that similar profiles from lakes in this area could be used for dating in the future. This may partly be due to the concentration peak in the Saanajärvi core being slightly higher (1850 r 600 gDM-1) than the ‘typical’ arctic values suggested above, although the reason for this is not certain. Apart from this core,

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all the reported concentration maxima in European lakes are in the range 400-1500 gDM-1. In contrast, sediment cores taken from within contaminated areas in the Arctic show full profiles and, as might be expected, higher SCP concentrations. Rose (1995) produced a SCP profile for a sediment core taken from Stepanovichjärvi (69q28’N, 30q40’E), a lake within 20 km of Nikel on the Kola Peninsula. This profile, although unfortunately not independently dated, shows all the typical SCP features but with a much longer record and a peak SCP concentration of 12,000 gDM-1. Comparison of the full SCP inventory of this core with the other arctic sites reveals the extent of the contamination in this area. The Stepanovichjärvi inventory is between one and two orders of magnitude higher than most of the Svalbard sites and three times higher than the most contaminated site, Tenndammen, which is located close to the coal-fired power station at Longyearbyen (Rose et al. 2004). Beyond Europe, paleolimnological studies involving fly-ash particles are remarkably scarce. Doubleday et al. (1995) produced a profile of “spherical black carbon particles” for Lower Dumbell Lake (82q29’N, 62q29’W) on northern Ellesmere Island, Nunavut, reporting concentrations of up to 30,000 gDM-1. However, this particle classification included diesel particles from a local military base as well as the SCPs described above (N. Doubleday, pers. comm.) and hence the concentrations are not directly comparable with other arctic SCP data and cannot be considered typical of the area. In contrast, Rose (unpublished data) produced an undated SCP profile for Amituk Lake on Cornwallis Island (75q03’N, 93q48’W), where peak concentrations of 350 gDM-1 were determined in the surface sediments. Bindler et al. (2001c) reported SCP concentrations of ca. 200 gDM-1 for recent sediments in Nunatak Lake, western Greenland (67q57’N, 49q48’W). A SCP presence was recorded in sediments back to the early 20th century, whilst the inventory was seen to be similar to the more remote sites on Svalbard. Rose (unpublished data) also analysed 13 surface sediments from the Søndre Strømfjord area (to the west of Nunatak Lake, western Greenland) and a full core from ‘Lake 53’ (66º30’N, 53º32’W) on the coast. Where concentrations were above the analytical detection limit, all were < 100 gDM-1. Highest SCP concentrations were thus furthest inland at Nunatak Lake but declining towards the coast where concentrations were below the limit of detection. This is, therefore, consistent with the conclusions of Bindler et al. (2001c) of an east to west declining gradient of Hg pollution in the area. Finally, a core taken from Schuchie Lake in northern Siberia (68q45’N, 161q15’E) was analysed for SCPs by Rose (unpublished data). Again, low concentrations (maximum 600 gDM-1) were found and the profile was rather variable. In the southern hemisphere, little work on fly-ash has been conducted. However, the presence of SCPs was detected in the recent sediments from lakes on both Signy Island, South Orkneys (60q43’S, 45q36’W) and in the Larsemann Hills area of eastern Antarctica (69q23’S, 76q20’E), an area where no previous evidence for contamination from global air circulation has been recorded (Gasparon and Burgess 2000). In both cases SCP concentrations were ca. 100 gDM-1 (N.L. Rose, unpublished data).

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Other anthropogenic particles Doubleday (1999) produced an atlas of particles derived from combustion sources in the Arctic and related these to paleolimnological studies along a transect of lake sites from Alert and Cape Herschel on Ellesmere Island to lakes on the east and west coasts of Hudson Bay. Sources included not only fly-ash from coal and oil combustion but also particles from wood combustion, incineration of garbage, stationary sources such as diesel generators providing power for small communities in the Arctic, as well as domestic oil stoves and mobile sources such as truck, tractor and helicopter exhausts. Particle data throughout this study were expressed as percentage data and hence are difficult to compare with data from other arctic areas. However, it is interesting to compare the trends within the transect. SCPs and IASs contributed more than 90% of combustion particles at the most northerly sites around Alert (e.g., Lower Dumbell Lake; discussed above), whilst wood charcoal particles comprised ca. 60% of combustion particles at Hawk Lake (63q38’N, 90q40’W) on the west coast of Hudson Bay and 20 to 45% in lakes on the Belcher Islands off the east coast (56q34’N, 79q16’W). At all sites, however, a range of combustion particle types was identified suggesting the influence of contamination from a number of different sources varying both spatially and temporally. Summary Dated sediment cores have been used effectively to discern historical records and current fluxes of anthropogenic pollutants in arctic lakes. A similar database does not yet exist for antarctic lakes possibly owing to challenges in obtaining samples and to lack of a strong incentive to examine pathways for human exposure that has driven a lot of the investigations in the Arctic. The use of sediment cores for assessment of longrange transport and deposition of contaminants is particularly challenging in polar environments because of the generally low concentrations of the target chemicals and the potential for contamination during sampling and analysis. Low sedimentation rates typical of most arctic lakes limit the time resolution and thus the usefulness of the historical record. On the other hand, the hundreds of thousands of lakes available within the circumpolar Arctic permits wide geographical coverage compared to glacier or peat cores. National programs within Norway, Sweden and Finland (with extensions into western Russia) have successfully used the wide availability of lake sediments to examine detailed geographical trends of deposition of Hg, Pb and other heavy metals. Combined with results from a relatively smaller number of cores from Alaska, Canada, and Greenland, it is now possible to evaluate Hg fluxes and anthropogenic enrichment factors on a circumpolar basis. One interesting result of this work is that EFs for Hg do not correspond well with recent observations of higher Hg deposition in snow near the Arctic Ocean, and instead tend to be higher in more southerly and inland lakes. This confirms the conclusions of Landers et al. (1998) that anthropogenic Hg deposition is the dominant regional deposition process in some northern regions near sources but of lesser importance in areas remote from urban/industrial Hg emissions. Lead deposition also shows distinctive geographical trends with higher concentrations in sediments of more southerly lakes in Scandinavia and Canada corresponding to the short atmospheric

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residence time of particle-associated metals. Use of Pb isotope ratios has also enabled source regions (i.e., Eurasia versus North America) to be discerned. A similar (large lake number; top slice only) standardized approach has also been used to examine spatial trends of PCBs in Sweden, northern Norway and the Russian Arctic. Fluxes of PCBs and other persistent organochlorines have been examined in about 27 sediment cores from the Canadian Arctic and northern Alaska. These cores show generally higher fluxes nearer sources and later onset of inputs compared with the known global PCB production. These studies clearly demonstrate the value of sediment cores for discerning anthropogenic inputs; however, only limited temporal trend information has been compiled so far for persistent organic pollutants in lakes of the European Arctic due to the emphasis on “top/bottom” sampling and lack of core dating. PAH deposition is an exception because the historical record in high latitude lakes in Norway has been compared with that of high altitude lakes in central Europe. The large range of PAH fluxes in Canadian arctic lake sediment cores demonstrates that sources of PAHs are complex and could be due to petrogenic sources, while at the same time showing some characteristics of pyrolytic inputs such as increasing concentrations during the early to mid-20th century. The strength of the paleolimnological approach has been further emphasized in studies of anthropogenic particles. Although there have been relatively few studies, mainly in the European Arctic, spatial comparisons have been possible across a region whilst sediment core analyses at individual sites have allowed both rate and direction of depositional change to be determined. Away from local sources within the Arctic, SCP surface concentrations are observed to be similar throughout the arctic region (typically 200-1000 gDM-1), whilst the few unpublished data for the Antarctic suggest still lower values. Although there have been few studies to date, the increasing interest in particles as a transport medium for other pollutants to remote regions suggests that there is value in expanding this work and linking it to parallel measurements of particle-borne substances such as high molecular weight POPs and heavy metals. Further, there is a wealth of ‘unexplored’ information, particularly in the area of source apportionment studies, in the use of a multi-proxy pollutant approach by comparing the sediment records of a number of different pollutant types. Acknowledgements We thank Nancy Doubleday, Lyle Lockhart, Peter Outridge, Ingemar Renberg and John Smol for discussing their data and interpretations, and Tog Jackson and an anonymous reviewer for a critical review of the manuscript.

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Outridge P.M., Percival J., Stern G. and Lockhart W.L. 2001. Unpublished data. Presented at the Northern Contaminants Program Results Workshop, Calgary AB Sept 2001. Natural Resources Canada, Ottawa, Ontario. Outridge P.M., Hermanson M.H. and Lockhart W.L. 2002. Regional variations in atmospheric deposition and sources of anthropogenic lead in lake sediments across the Canadian Arctic. Geochim. Cosmochim. Acta 66: 3521-3531. Parkin D.W., Phillips D.R., Sullivan R.A.L. and Johnson L. 1970. Airborne dust collections over the North Atlantic. J. Geophys. Res. 75: 1782-1793. Pienitz R., Smol J.P. and Lean D.R.S. 1997. Physical and chemical limnology of 24 lakes located between Yellowknife and Contwoyto Lake, Northwest Territories (Canada). Can. J. Fish. Aquat. Sci. 54: 347-358. Preiss N., Mélières M.-A. and Pourchet M. 1996. A compilation of data on lead-210 concentration in surface air and fluxes at the air-surface water and water-sediment interfaces. J. Geophys. Res. 101D: 28847-28862. Rasmussen P.E. 1994. Current methods of estimating atmospheric mercury fluxes in remote areas. Environ. Sci. Technol. 28: 2233-2241. Rawn D.F.K., Lockhart W.L., Wilkinson P., Savoie D.A., Rosenberg B. and Muir D.C.G. 2001. Historical contamination of Yukon Sediments by PCBs and organochlorine pesticides: Influence of local sources and watershed characteristics. Sci. Total Environ. 280: 17-37. Rigler E.H. 1975. The Char Lake Project. In: Cameron T.W.M. and Billingsley L.W. (eds), Energy Flow - Its Biological Dimensions: A Summary of the IBP in Canada, 1964-1974. Royal Society of Canada, Ottawa, pp. 171-198. Rognerud S. and Fjeld E. 1993. Regional survey of heavy metals in lake sediments in Norway. Ambio 22: 206-212. Rognerud S. and Fjeld E. 2001. Trace element contamination of Norwegian lake sediments. Ambio 30: 11-19. Rognerud S., Skotvold T., Fjeld E., Norton S.A. and Hobaek A. 1998. Concentrations of trace elements in recent and pre-industrial sediments from Norwegian and Russian Arctic lakes. Can. J. Fish. Aquat. Sci. 55: 1512-1523. Rognerud S., Grimalt J.O., Rosseland B.O., Fernández P., Hofer R., Lackner R., Lauritzen B., Lien L., Massabuau J.C. and Ribes A. 2002. Mercury and organochlorine contamination in Brown Trout (Salmo trutta) and Arctic Charr (Salvelinus alpinus) from high mountain lakes in Europe and the Svalbard Archipelago. Wat. Air Soil Pollut. Focus 2: 209-232 Roos-Barraclough F., Martinez-Cortizas A., García-Rodeja E. and Shotyk W. 2002. A 14,500 year record of the accumulation of atmospheric mercury in peat: volcanic signals, anthropogenic influences and a correlation to bromine accumulation. Earth Planet. Sci. Let. 202: 435-451. Rose N.L. 1995. Carbonaceous particle record in lake sediments from the Arctic and other remote areas of the northern hemisphere. Sci. Total Environ. 160/161: 487-496. Rose N.L. 2001. Fly-ash particles. In: Last W.M. and Smol J.P. (eds), Tracking Environmental Change Using Lake Sediments. Volume 2: Physical and Geochemical Methods. Kluwer Academic Publishers, Dordrecht, The Netherlands, pp. 319-349. Rose N.L., Harlock S., Appleby P.G. and Battarbee R.W. 1995. The dating of recent lake sediments in the United Kingdom and Ireland using spheroidal carbonaceous particle concentration profiles. The Holocene 5: 328-335. Rose N.L., Backus S., Karlsson H. and Muir D.C.G. 2001. An historical record of toxaphene and its congeners in a remote lake in western Europe. Environ. Sci. Technol. 35: 1312-1319. Rose N.L., Rose C.L., Boyle J.F. and Appleby P.G. 2004. Lake sediment evidence for local and remote sources of atmospherically deposited pollutants on Svalbard. J. Paleolim. 31: 499-513. Sarkar A., Singbal S.Y.S. and Fondekar S.P. 1994. Pesticide residues in the sediments from the lakes of Schirmacher Oasis, Antarctica. Polar Record 30: 33-38.

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9. PALEOLIMNOLOGY OF THE MIDDLE AND HIGH CANADIAN ARCTIC

ALEXANDER P. WOLFE ([email protected]) Department of Earth and Atmospheric Sciences University of Alberta Edmonton, Alberta T6G 2E3, Canada and I. ROD SMITH ([email protected]) Geological Survey of Canada Terrain Sciences Division 3303 - 33rd St. NW Calgary, Alberta T2L 2A7, Canada

Key words: Arctic, Diatoms, Neoglacial, Paleoclimate, Paleolimnology, Quaternary, Baffin Island, Ellesmere Island, Canada

Introduction The objective of this chapter is to complement, rather than duplicate, previous reviews of arctic paleolimnology (Smol and Douglas 1996; Douglas and Smol 1999), paleoclimatology (Ovenden 1988; Bradley 1990), and developments in the application of paleolimnological techniques to the reconstruction of climate change (Smol and Cumming 2000). Our approach is to first provide a detailed environmental overview of this vast region, then a brief history of limnological and paleolimnological research, before assessing regional paleolimnological records from the oldest to youngest deposits. Methodological considerations are limited, since these are covered in separate contributions to this volume. Emphasis is placed on climatic influences on lake development and, consequently, on demonstrating the utility and limitations of paleolimnology in the generation of proxy paleoclimate data. In many ways, this is the most pressing agenda for paleoenvironmental research in the high latitudes, given (a) the relatively poor spatial coverage of available data by comparison to temperate regions, and (b) the acknowledged sensitivity of the Arctic to climate change, borne out of its intimate linkage to ocean circulation, atmospheric feedback mechanisms involving snow and ice cover, and biospheric modulation of greenhouse gas concentrations. 241 R. Pienitz, M.S.V. Douglas and J.P. Smol (eds), 2004. Long-term Environmental Change in Arctic and Antarctic Lakes. Springer. Printed in the Netherlands.

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Environmental background Physiography and bioclimatology The middle and high Canadian Arctic, as formally defined by ecotonal boundaries (Edlund 1984), incorporates over 1.4 million km2 of the Arctic Archipelago, Boothia and Melville peninsulas, and the northernmost coasts of Québec and Labrador bordering Hudson Strait (Figure 1). The Arctic Archipelago contains six of the world’s 30 largest islands, including Baffin Island, the fifth largest (507,450 km2). Physiographically, the western and central arctic islands, as well as Foxe Basin, are of low relief, comprised of variably folded and faulted, platform and basinal, Paleozoic to Cenozoic carbonaceous sedimentary rocks (Trettin 1989). The eastern axis of the Arctic Archipelago is more mountainous, with peaks up to 2665 m above sea-level (asl) on northern Ellesmere Island (Barbeau Peak), and 2156 m asl on Baffin Island (Tête des Cirques). Precambrian granitic basement rocks outcrop from east-central Ellesmere Island south through eastern Devon Island, extending across most of eastern Baffin Island and Boothia and Melville peninsulas. Glaciers and larger ice caps (up to 800 m thick) occupy much of the mountainous regions of Ellesmere and Axel Heiberg islands, while to the west, only Meighen and Melville islands retain small ice caps (Figure 1). Glaciation levels and equilibrium line altitudes largely mirror topography, but decline (< 400 m asl) along northern coasts and westward across the Arctic Archipelago (400 to 600 m asl), reflecting a gradient of decreasing melting degree day totals (Miller et al. 1975; Edlund and Alt 1989). The Devon Island Ice Cap, and ice caps and glaciers on southeastern Ellesmere Island and northern and easternmost Baffin Island, exhibit a strong northwest-oriented precipitation gradient with increasing distance from moisture sources in Baffin Bay (Koerner 1979). Vegetation in the Canadian Arctic is strongly influenced by both summer climate patterns and geologically-mediated edaphic factors (Edlund and Alt 1989). In the High Arctic, adjacent to the Arctic Ocean, only herbaceous and non-vascular plants survive, with < 10% ground cover. Southeastward towards Baffin Bay, plant diversity increases, sedges and herbaceous plants predominate, and ground cover increases to ” 25%. Vegetation is often concentrated along lower and/or south-facing slopes, and in lowlying areas, reflecting localized thermal enclaves with abundant moisture in a region of otherwise strong summer moisture deficits. On Baffin Island, and across the middle Arctic, vegetation cover remains sporadic on upland and exposed surfaces, but elsewhere there is increased abundance of ericaceous plants, sedges, grasses, and some areas of peat accumulation, resulting in an overall ground cover of 25 to 50% (Edlund 1986). Regional climatic patterns in the middle and high Canadian Arctic are governed by (1) the influence of anticyclonic activity centered on the Arctic Ocean, (2) the frequency of cyclonic systems originating from Davis Strait and Baffin Bay, (3) the extent and persistence of sea ice within intervening channels, and (4) local topoclimatic variability (Maxwell 1981). Annual precipitation varies from a high of 600 mm along the easternmost coast of Baffin Island to < 100 mm in the polar desert of northern Ellesmere Island. Mean daily July temperatures are less variable, averaging between +3 and +8°C, while daily January temperatures range from -25 to less than -40°C (Maxwell 1981). Intermontane areas surrounding Fosheim Peninsula and Lake Hazen (Figure 1)

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Figure 1. Map of the Arctic Archipelago, indicating the numbered locations of sites listed in Table 1, bioclimatic zones, and other locations discussed in the text. Grey shading depicts contemporary glacier cover.

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are climatically anomalous with minimal cloud cover, relatively warm summer temperatures, and prolonged melt seasons (Edlund and Alt 1989). Elsewhere, steep topoclimatic gradients exist between coastal sites and those inland (Jacobs and Leung 1981). The large lake systems of south-central Baffin Island (Nettilling and Amadjuak) and adjacent regions of Cumberland Sound and Frobisher Bay also form an anomalously warm subregion that supports the growth of dwarf birch (Betula glandulosa and B. nana) and other low arctic plant species (Jacobs et al. 1997). Wisconsinan glacial history Regional glacial history is important in that it constrains the length of the paleolimnological record that can be expected from lakes of any given region of the Arctic Archipelago. During Late Wisconsinan glaciation, the high arctic area north of Parry Channel (Figure 1) was inundated by the Innuitian Ice Sheet, which coalesced with the Laurentide Ice Sheet to the south and the Greenland Ice Sheet to the east (Dyke et al. 2002). Ice retreat commenced around 11 ka BP, and fiord heads along Axel Heiberg and western Ellesmere islands became deglaciated by 8 ka BP (Hodgson 1989; Ó Cofaigh et al. 2000). However, on east-central Ellesmere Island, fiord-based ice continued to occupy the outer coast until at least 8 ka BP, retreating to fiord heads as late as 5.5 ka BP (England et al. 2000). Interior regions of Ellesmere Island appear to have retained extensive Innuitian and/or plateau ice cover until the mid-Holocene (Bell 1996; Smith 1999), after which ice margins retreated to positions near or behind those of the present. Modern ice margins largely reflect pronounced post-3 ka BP Neoglacial advances (Blake 1989a), coupled to recent volume losses in many areas (Dyurgerov and Meier 2000). Most of Banks Island remained unglaciated during the Late Wisconsinan (Vincent 1989; Dyke et al. 2002). Retreat of the Laurentide Ice Sheet from the western and central middle Arctic began by 13 ka BP and was completed by 8.5 ka BP (Dyke and Prest 1987). Laurentide ice inundated all of northwestern and southern Baffin Island, while eastern Baffin Island preserved isolated ice-free areas on inter-fiord uplands north of Cumberland Sound, with outlet glaciers terminating near the mouths of fiords and sounds (Dyke et al. 2002; Miller et al. 2002). Deglaciation of northwestern Baffin Island began by 10 ka BP, yet the Laurentide Ice Sheet continued to occupy Foxe Basin until ca. 6.7 ka BP. By 6 ka BP, Laurentide ice became centered along the axis of Baffin Island, with final retreat focused around three sectors, the Barnes and Penny ice caps (Figure 1) and the region west of Frobisher Bay. The limnological legacy Because glaciation is a pervasive feature of most of the Arctic Archipelago, lake basins typically occur in glacially-scoured bedrock basins, over-deepened structural features, meltwater channels, or lowland depressions within glacigenic sediment. In coastal regions, a large number of lakes were formed as a result of glacio-isostatic uplift and the isolation of marine embayments, partially or entirely dammed by raised beach deposits (Retelle et al. 1989; Young and King 1989; Williams 1990). The range of geological

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materials surrounding and underlying individual lake basins results in a broad spectrum of water chemistries, ranging from hypersaline meromictic brines (Ouellet et al. 1987), to extremely dilute surface waters on crystalline terrain (Church 1974). Lakes having developed within raised marine environments typically have higher solute concentrations than counterparts above the limit of marine submergence (Rust and Coakley 1970; Hamilton et al. 2000). This reflects the aggradation of permafrost into uplifted marine sediments and the exsolution of solutes into lakes, a process that may continue to influence lake chemistry long after initial isolation from the sea. The unifying characteristic of lakes across the middle and high Canadian Arctic is the severity of lake ice environment in terms of both thickness and duration. Spring ice thickness in excess of 2 m is commonly encountered. Lake ice typically persists until mid-summer (June-July), only to reform in late August or September. Examples of typical summer conditions, with and without ice cover, are shown on Figure 2. Solidly frozen lakes occur in certain extreme cases (Blake 1989b). Perennially ice-covered lakes are rare (Doubleday et al. 1995; Belzile et al. 2001), and restricted to the north coast of Ellesmere Island, or to the highest non-glaciated environments. Certain lakes retain ice cover in cold summers only, resulting in a warming feedback to underlying waters, which reduces the likelihood of consecutive summers of persistent lake ice (Doran et al. 1996). The limnological consequences of extensive lake ice are numerous. Integrated water column temperatures are generally cold (< 4°C), although surface heating can occur rapidly during summer ice-free conditions when inputs of solar radiation may be continuous. Lake ice also exerts profound influences on circulation patterns (Welch and Bergmann 1985). The duration of ice cover, coupled to a seasonal (versus diurnal) photoperiod, combine to reduce biological production in arctic lakes, especially in the plankton. Detailed studies of arctic lake metabolism are few, but catchment inflow and vegetation appear to play key regulatory roles (Welch 1974). This is consistent with extensive water chemical data that suggest the greater importance of allochthonous nutrient sources relative to autochthonous cycling (Hamilton et al. 2001). Therefore, because soil development is minimal and vegetation is sparse, oligotrophy is ubiquitous, except in rare instances of direct anthropogenic nutrient loading (Douglas and Smol 2000). In addition to lakes, ponds are a common limnological feature across much of the Arctic Archipelago. Ponds are shallower (< 2 m) water bodies that freeze solid in winter but may attain much higher maximum summer temperatures (> 10°C), suggesting heightened sensitivity to climatic variability (Douglas and Smol 1994). Ponds also have higher levels of biological activity and hence nutrient cycling (Hamilton et al. 2001). To paleolimnologists, the most important consequence of prolonged ice cover and low productivity is that sediment accumulation rates in non-glacial arctic lakes and ponds are very low. The entire Holocene may be represented by < 100 cm of sediment (e.g., Wolfe and Härtling 1996). Such low sedimentation rates present significant challenges in terms of attaining high temporal resolution from the sediment record. On the other hand, lake ice serves to buffer arctic lakes from the influences of low-frequency, high intensity climatic events for up to 90% of the year, thereby protecting the stratigraphic integrity of the sediment record.

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Figure 2. Photographs of two of the Baffin Island lakes discussed in this chapter ((A) Fog Lake; (B) Tulugak Lake), providing typical examples of the appearance of non-glacial arctic lakes in (A) late and (B) early summer. Relatively subtle environmental changes, such as slight summer cooling or increased snow cover, are likely sufficient to sustain the conditions seen in (B) throughout the summer.

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Historical development Early research on lakes in the Canadian Arctic was primarily in the form of multi-year limnological investigations: Defense Research Board’s Operation Hazen (Oliver and Corbet 1966; Figure 1, Lake Hazen region), McGill University’s Axel Heiberg Island Project (Caflisch 1972; Figure 1, site 8), and International Biological Program initiatives at Char Lake (Schindler et al. 1974; Rigler 1978; Figure 1, site 13) and Truelove Lowland (Minns 1977; Figure 1, site 17). Despite the early establishment of pollen (Iversen 1953) and diatom (Foged 1972) studies on sediment cores from lakes and ponds on western Greenland, paleolimnological studies were relatively slow to come to the Canadian Arctic. Pollen analyses of lake sediments on Baffin Island were conducted in the 1970s and 1980s (Short et al. 1985; Gajewski and MacDonald, this volume), while coring of lakes and ponds along east-central Ellesmere Island was initiated around the same time (Blake 1978, 1981). Rock Basin Lake (Figure 1, site 12), flanked by the Ekblaw Glacier near Baird Inlet, was cored at this time and subsequently analysed by Smol (1983) for diatoms and chrysophytes, and by Hyvärinen (1985) for pollen. For all practical purposes, Smol (1983) represents the first modern paleolimnological study from the Arctic Archipelago. With recognition of the sensitivity of polar regions to past and future global climate change (Overpeck et al. 1997), and the heightened vulnerability of high latitude aquatic ecosystems to these forcings (Rouse et al. 1997), the variety and extent of paleolimnological research in the middle and high Canadian Arctic has blossomed in the last decade. Table 1 provides a geographically referenced (Figure 1) listing of various paleolimnological studies that have been undertaken in the Arctic Archipelago. While not intended to be exhaustive, it provides a sense of the overall distribution of published research on lake sediments, and illustrates well both the paucity in spatial coverage (particularly in the western Arctic Archipelago), as well as strong bias towards coastal localities. For example, Ellesmere Island is roughly the size of Great Britain, yet has fewer than 10 locations where detailed paleolimnological research has been conducted (e.g., Taconite Inlet, Cape Herschel (Figure 1, sites 1 and 11, respectively); Table 1). From the studies listed in Table 1, we have chosen to highlight a cross-section of studies that reflect the temporal breadth of the arctic lake sediment record. This survey progresses from the oldest to the youngest sediments, in each case highlighting salient results and their implications, as well as remaining questions and future research needs. Pre-Holocene lake sediment records Although uncommon, pre-Holocene paleolimnological records from the middle and high Canadian Arctic are important in terms of (a) constraining the geometry of former ice sheets, and (b) providing analogs for equable interglacial climates in the absence of any anthropogenic impacts. Lichti-Federovich (in Blake 1974) first documented lacustrine diatoms from organic sediment beneath diamicton in a section on Bathurst Island. Similar assemblages, dominated by small colonial, benthic Fragilaria diatoms, have been subsequently recovered from beneath marine transgressive and postglacial regressive facies on eastern Devon Island (Wolfe and King 1999). Although the dating of these deposits remains problematic, they necessarily relate to ice-free intervals, and

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may thus delimit the extent of Late Wisconsinan glaciation, especially when considered alongside other, more reliably AMS 14C-dated terrestrial biota from the Arctic Archipelago older than 10 ka BP (Dyke et al. 2001). Stratified pre-Holocene sediments have also been documented from eastern Baffin Island, specifically beneath Late Wisconsinan till on outer Hall Peninsula (Miller et al. 1999), and in lakes situated above flights of moraines on southern (Wolfe and Härtling 1996) and northern (Steig et al. 1998) Cumberland Peninsula. Although the possibility of episodic sedimentation during the Last Glacial Maximum (LGM) has been suggested for some sites (Wolfe et al. 2000), AMS 14C dates on sediment humic acid extracts are now thought to be in error, due to admixtures of ancient (14C-depleted) and contemporaneous dissolved organic carbon pools. At Fog Lake (Figure 1, site 24), this has been confirmed by early Holocene dates obtained on moss and insect remains from the same sediments that yield humic acid dates in the 14-16 ka BP range (Wolfe et al., this volume). Despite the persistent geochronological difficulties surrounding these records, several lakes preserve a distinctive facies of highly compacted, organic-rich gyttja which is clearly beyond the range of 14C, regardless of the dating target considered (Miller et al. 1999; Wolfe et al. 2000). On the basis of optically- and thermally-stimulated luminescence ages, which have been calibrated against lithologically similar Holocene gyttjas, these sediments are securely dated to the latter portion of Marine Oxygen Isotope Stage (MIS) 5, that is ca. 75 to 90 ka BP. The emerging picture is one in which these unglaciated lakes accumulated organic sediments during the last interglacial (sensu lato, MIS 5a), but were thereafter largely geomorphically and biologically inert during stages 4-2, prior to the onset of Holocene sedimentation. This scenario matches well with the record of summer insolation, in which MIS 5a represents the highest values of the last 100 ka (Figure 3). Frigid conditions during MIS 4-2 may have induced perennial lake ice cover or burial of lakes under non-erosive, cold-based ice caps. The possibility of such conditions is exemplified by the -20°C temperature depression suggested for the LGM at comparable latitudes in Greenland (Johnsen et al. 2001). Present mean annual air temperature on the northern coast of Cumberland Peninsula is about -10°C, implying that a shift to perennially-frozen conditions is physically plausible under these envisaged full glacial conditions. Paleoecologically, diatom and pollen assemblages in MIS 5a sediments from Baffin Island are extremely rich (Miller et al. 1999; Wolfe et al. 2000). Shrubs including birch (Betula) and alder (Alnus) are represented in much higher relative and absolute abundances than in any Holocene counterparts, indicating the local presence of these plants, more proximal pollen source areas, enhanced southerly flow, or some combination of these factors. The green alga Pediastrum was also several times more abundant than in Holocene sediments. On these grounds, it is suggested that MIS 5a summer conditions on eastern Baffin Island were the warmest of the last 100 ka. To test this hypothesis, the weighted-averaging (WA) summer water temperature transfer function of Joynt and Wolfe (2001) has been applied to diatom assemblages corresponding to MIS 5a from Fog Lake (Figure 4). The results show maximum inferred summer water temperatures early in the interglacial, several degrees warmer than anytime in the mid- to late Holocene. This result is in agreement with the pollen record, and provides a useful application of the WA transfer function.

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Figure 3. Schematic diagram of the paleoenvironmental record preserved in several upland Baffin Island lakes on Cumberland Peninsula, shown in relation to summer insolation for the past 150 ka (after Berger and Loutre 1991).

Figure 4. Selected paleoecological data from the basal gyttja of last interglacial age (sensu lato) from Fog Lake, an unglaciated site on northern Cumberland Peninsula, Baffin Island (Figure 1, site 24). The age of 90 ± 8 ka is an arithmetric mean of available optical and thermal luminescence dates. Original data are from Wolfe et al. 2000. The summer water temperature transfer function (Joynt and Wolfe 2001) has a root mean squared error of prediction (RMSE) of 1.9°C (bootstrapped RMSE = 2.8°C).

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In summary, it is now clear that stratigraphically coherent pre-Holocene paleolimnological records exist in specific locations of the Canadian Arctic. Chronological uncertainties are gradually being resolved (Wolfe et al., this volume), and a wealth of new information is being gleaned that pertains directly to (a) the reconstruction of regional glacial history (Miller et al. 2002), (b) the apparent ‘shutdown’ of limnological systems for tens of millennia during Wisconsinan time (MIS 4-2), and (c) the characterization of arctic ecosystems having evolved under warmerthan-present climatic conditions during previous interglacials. Holocene climatic evolution and paleolimnology Holocene paleoclimatology As a useful framework in which to consider the paleolimnological record, the Holocene climatic evolution of the Arctic Archipelago is summarized in Figure 5. Peak highlatitude summer insolation (Milankovitch forcing) occurred during the earliest Holocene (between 10 and 9 ka cal. BP), with a maximum radiation anomaly approximately 8% greater than present (Berger and Loutre 1991). Although much of the middle and high Canadian Arctic remained glaciated, warm summers are clearly registered by enhanced summer melting of the Agassiz Ice Cap (Koerner and Fisher 1990). At the same time, warming sea surface temperatures in Baffin Bay enhanced precipitation on Baffin Island (Miller and de Vernal 1992), leading to a widespread early Holocene glacial advance along the east coast: the Cockburn Substage, ca. 9.5 ka cal. BP (Andrews and Ives 1978). The reduction of early Holocene sea ice allowed bowhead whales and walrus to extend well beyond their current ranges (Dyke et al. 1996a, 1999) and facilitated driftwood delivery to northern coastal areas (Dyke et al. 1997). Warming is also suggested by the northward expansion of boreal molluscs (Dyke et al. 1996b). The mid-Holocene in the Canadian Arctic also appears to have been relatively warm, although records differ spatially, temporally, and by the extent to which they suggest comparable or even greater summer warmth relative to the early Holocene insolation maximum (Bradley 1990; Hardy and Bradley 1996). Restricted marine mammal distributions imply more extensive summer sea ice between 8.5 to 6 ka cal. BP (8 to 5 14C ka BP), and hence cooler conditions (Dyke et al. 1996a, 1999). In sharp contrast, warmer than present marine conditions at 6 ka BP are suggested by Gajewski et al.’s (2000) marine micropaleontological analyses, which extend from the High Arctic to the Labrador Sea, via Baffin Bay. On Baffin Island, enhanced southerly flow is suggested by peak abundances of shrub pollen at several sites around 6 ka BP (Mode 1996). A multi-proxy summary of marine and terrestrial evidence from the Baffin sector (Williams et al. 1995) suggests that warming began around 8 ka BP, intensified at 6 ka BP, and had markedly deteriorated by 3 ka BP. Between 5.5 and 3.5 ka cal. BP (5 to 3 14C ka BP), bowhead and walrus distributions suggest reduced summer sea ice (Figure 5). The least negative į18O values from the Devon Ice Cap are recorded in this interval, implying maximum winter and/or mean annual warmth (Koerner 1989). The isotopic record subsequently indicates sustained cooling (Figure 5). This trend, widely recognized as the Neoglacial, led to glacier advances (Blake 1989a) and the establishment of ice shelves (Stewart and England

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1983) along the north coast of Ellesmere Island, as well as the progressive reduction in marine mammal summer ranges (Dyke et al. 1996a, 1999). Neoglacial advances of local glaciers on Baffin Island were widespread (Davis 1985), and in some areas sufficiently extensive to override Late Wisconsinan moraines (Miller 1976). The absolute amount of cooling reflected by the Neoglacial ice core į18O decline is difficult to assess because the temperature signal is compounded by the influence of sea ice variability on the proximity to precipitation sources. Koerner and Fisher (1990) have suggested an amplitude of 2°C for Holocene temperature variability on the Agassiz Ice Cap. Others, working further south, and with different indicators, have suggested that Neoglacial cooling alone was in the order of 3°C or more (Andrews et al. 1981; Johnsen et al. 2001). Therefore, even if the broad pattern of Holocene climatic evolution (Figure 5) is assumed to be coherent across the Arctic Archipelago, the available data suggest that the amplitudes of temperature shifts varied regionally. To paleolimnologists working in these regions, the main questions become: how faithfully does the lake sediment record reflect the subtleties of regional climatic variability, and how can it contribute insights beyond what is gleaned from other terrestrial and marine proxies?

Figure 5. Holocene paleoclimatic evolution of the northern and central Arctic Archipelago shown in relation to summer insolation anomaly (Berger and Loutre 1991), Devon Ice Cap į18O (Koerner 1979), percent melt from the Agassiz Ice Cap (Koerner and Fisher 1990), and the frequency of bowhead whale radiocarbon dates from the central channels, including Lancaster Sound, Gulf of Boothia, and Prince Regent Inlet (Dyke et al. 1996a). Radiocarbon dates have been calibrated to calendar years to facilitate comparisons.

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Holocene lake sediment records: the High Arctic As noted by Bradley (1990), few early Holocene paleolimnological records exist in the High Arctic because glacier cover remained extensive. The longest paleolimnological record from Ellesmere Island is that from Rock Basin Lake (Smol 1983), which captures the entire Holocene (Figure 6a). Here, the diatom record preserves an initial colonizing assemblage (Fragilaria c onstruens var. venter and Cymbella mi nuta) overlying a basal barren zone of sandy-silt. This assemblage is analogous to those recorded in recently deglaciated terrains outside the Arctic, and presumably implies prolonged summer lake ice cover, low production, and high base cation status from catchment inputs. Diatom communities shifted rapidly towards dominance by large benthic periphytic taxa such as Pinnularia biceps, suggesting enhanced littoral habitat availability and seasonally open water until about 4 ka cal. BP. Elevated early to midHolocene algal production is suggested by trends in diatom valve and chrysophyte cyst concentrations (Figure 6a), apparently in response to regional paleoclimatic conditions (Figure 5). From 4 ka cal. BP onwards, the Rock Basin Lake record is characterized by decreased diatom absolute abundances but increased species richness, the dominance of shallow water and aerophilic forms, and high relative frequencies of chrysophyte cysts to diatoms. Each of these features likely represents a portion of the limnological response to Neoglacial cooling. The Rock Basin Lake diatom record is important for several reasons. Foremost, it is from this site that evolved the model that diatom assemblages in arctic lakes are shaped by the influences of lake ice and snow cover on habitat availability (Smol 1983, 1988). This model has been applied widely by paleolimnologists, largely because the extent and duration of ice cover is highly sensitive to climate. It has been demonstrated elegantly in alpine lakes that ice cover indeed plays a significant role in regulating diatom assemblage composition (Lotter and Bigler 2000). Furthermore, stratigraphic changes in the Rock Basin Lake record provided the first indications that algal communities in arctic lakes respond sensitively to Holocene climatic variability. This has stimulated much diatom research aimed at paleoclimatic reconstruction in arctic regions (e.g., Table 1). But at the same time, it remains perplexing how diatom assemblages from many high arctic lakes do not record pronounced species shifts in response to millennial-scale Holocene climatic variability, instead remaining dominated by relatively invariant floras of small colonial Fragilaria diatoms (e.g., Young and King 1989; Douglas et al. 1994; Smith 2002). Smol (1983) has suggested that in such instances, the severity of local climatic conditions may arrest lake development, so that these lakes are in effect extant analogs for early postglacial environments in the glaciated mid-latitudes. Given the proximity of Rock Basin Lake to a large glacier and encircling icefield, it would perhaps appear unusual for it to record such marked ecological shifts. However, local site characteristics, namely its location within the basin of a parabolic bedrock catchment, result in uncharacteristic summer warmth as demonstrated by the presence of relatively lush vegetation including Empetrum nigrum and Vaccinium uliginosum (Blake 1981). It is also suggested that edaphic factors may override climatic influences on diatom assemblages in certain lake types, such that poorly buffered lakes may exhibit increased environmental sensitivity in comparison to highly alkaline systems. In this sense, lakes in small non-carbonate, bedrock-dominated catchments, above the limit of local marine submergence, and characterized by dilute

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Figure 6. Representative paleolimnological records from Ellesmere Island: (A) Holocene diatom and chrysophyte stratigraphy from Rock Basin Lake, Baird Inlet (modified from Smol 1983), and (B) diatom concentrations and inferred lake ice variability from four lakes on Hazen Plateau (Smith 2002). All ages are calibrated radiocarbon years BP.

water chemistries, may ultimately be the most sensitive sites for paleolimnological research using diatoms. Lakes on Hazen Plateau (Figure 1, site 4), the continental interior of north-central Ellesmere Island, were deglaciated after 8.5 ka cal. BP, but diatom populations did not

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become established until ca. 5.5 ka cal. BP, when glaciers finally retreated from the Lake Hazen trough (Smith 2002). Thus, Hazen lakes show no limnological response to "warm" early Holocene conditions recorded in the Agassiz melt record (ca. 100 km to the south), nor do they reflect meltwater availability and warming as early as 7.5 ka cal. BP, as reconstructed from a peat deposit 40 km to the northwest (LaFarge-England et al. 1991). It is therefore clear that strong local topoclimatic differences can be modulated by the proximity of large glaciers (Koerner 1980; Bradley and Serreze 1987), a concept referred to as negative thermal inertia. This likely explains why many arctic paleolimnological records do not correlate directly with independent climate proxies in the early to mid-Holocene interval. Diatom assemblages from four lakes on Hazen Plateau are dominated by small colonial Fragilaria throughout the mid- to late Holocene records (Smith 2002). In the absence of pronounced taxonomic shifts, absolute abundances provide a measure of diatom variability and are considered to represent a proxy for lake ice cover (Figure 6b); high diatom concentrations are associated with increased summer open water and a prolonged growing season. Although diatom concentrations can be influenced by changes in sedimentation rate, this problem is obviated by correlating several sites. Near-synchronous rises in diatom concentrations occur in all four Hazen lakes beginning ca. 5.5 ka cal. BP, reaching their Holocene maxima in two lakes between 4.6 and 3.2 ka cal. BP. Diatom numbers then decline in the uppermost lake, but maintain high levels in a lower basin until ca. 1.8 ka cal. BP, after which diatom concentrations decline sharply (Figure 6b). Changes in diatom concentrations and differences between the lake basins are thought to reflect a regional lowering of the summer snowline and increasing Neoglacial ice cover (Smith 2002). Decreased sedimentation in varve records from Lake C2, Taconite Inlet (Figure 1, site 1), between 3 to 2.3 ka cal. BP, are also interpreted to record cooling on northern Ellesmere Island (Lamoureux and Bradley 1996). A subsequent increase in Lake C2 sedimentation between 2.3 and 1.2 ka cal. BP is considered to reflect, in part, an increase in southerly-derived moisture and hence increased snowfall. Evidence in support of an increased southerly moisture source at this time is found in detrended ice core į18O records (Lamoureux and Bradley 1996) and the Agassiz Ice Cap pollen record (Bourgeois et al. 2000). Therefore, independently of temperature changes, it must be recognized that increased snowfall will delay summer melting (higher albedo, insulation), resulting in reduced light transmission through lake ice, prolonged summer ice cover, and attendant biological impacts. Baffin Island lakes during the early Holocene On eastern Baffin Island, the interpretation of early Holocene paleolimnological records is challenged by the unusual characteristics of sediment deposited during this interval in the majority of lakes investigated to date. Most (but not all) upland sites on both the northern and southern coasts of Cumberland Peninsula preserve a distinctive early Holocene facies of black, organic-rich gyttja, often with irregularly-spaced laminations of siderite (FeCO3). This sediment contains little allochthonous clastic material, few planktonic diatoms, and is enriched in sulfur and trace elements (Wolfe and Härtling 1997). Examples of this sediment from two lakes near Pangnirtung (Figure 1, site 26) are illustrated in Figure 7. Because of the regional expression of this sediment type,

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which is generally restricted to the 10-8 ka BP interval, its origin provides important clues to early Holocene conditions in the eastern middle Arctic. The absence of similar lithologies in younger sediments (typically olive-brown, massive, silty gyttja) implies that early Holocene climatic conditions engendered a fundamentally different limnological regime than that of subsequent millennia. The presence of stable siderite coupled to the preservation of delicate sedimentary structures indicate that these sediments are chemically reduced. The question then becomes how anoxia was induced and sustained at the mud-water interface of these moderately deep (9-40 m), ultra-oligotrophic lakes. By analogy to the development of winter oxygen deficits in temperate and alpine lakes (e.g., Ohlendorf et al. 2000), one possible explanation is that prolonged ice cover restricted lake circulation, leading to hypolimnetic anoxia. Diatoms were likely restricted to shallow littoral habitats in summer moats, then transported by lateral currents. Despite pervasive lake ice, organic matter production was relatively high, and its preservation was excellent. Although the early Holocene was a period of warming in the Baffin Bay region (Figure 5), enhanced snowfall resulted in extensive regional glacier advances (Andrews 1980; Miller and de Vernal 1992). The insulation of lake ice, and indeed of lake waters (Doran et al. 1996) by snow, combined to the negative thermal inertia of nearby glaciers, are therefore consistent with the paleolimnological observations. This example demonstrates the complex and indirect influences of climate on limnological conditions and hence on the characteristics of the sediment record. It also illustrates how arctic lakes may have more than one stable state, but that the thresholds between them may be exceeded by relatively small forcing, well within the range of natural Holocene variability.

Figure 7 . (A) Reduced early Holocene gyttja with siderite laminations from Ukalik Lake near Pangnirtung, Baffin Island, and (B) stratigraphic changes associated with the same sediment facies (shaded interval) in nearby Tulugak Lake. Data are redrawn from Wolfe and Härtling (1996, 1997). Ages are all based on calibrated radiocarbon dates.

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Paleolimnological records of Neoglacial cooling Throughout the Arctic Archipelago, there is general consensus that the late Holocene was an interval of progressive cooling, culminating in the Little Ice Age (LIA), and followed by pronounced 20th century warming. Two diatom records from Baffin Island are used to explore the response of diatom communities to this climatic gradient, and to assess the downcore performance of WA transfer functions (Joynt and Wolfe 2001). The two lakes have very different Quaternary histories, Fog Lake being at least 90 ka old, whereas Kekerturnak Lake (Figure 1, site 23) was glaciated in the early Holocene. Furthermore, the two lakes support very different diatom floras. Nonetheless, both lakes record progressive decreases in diatom absolute abundances during the last five millennia in concert with Neoglacial cooling (Figure 8). The WA summer surface water reconstruction from Fog Lake indicates a 4°C amplitude of late Holocene variability. The most pronounced decline of lake water temperatures occurred between 2500 and 1500 cal. BP. This coincides with the cooling of nearshore currents along the east coast of Baffin Island, inferred from southward migration of molluscs with arctic ecological

Figure 8. Mid- to late Holocene trends in diatom-inferred summer water temperatures, valve concentrations, and reconstructed pH from Fog and Kekerturnak lakes on northern Cumberland Peninsula, Baffin Island. The records indicate cooling, decreased diatom production, and pH lowering in association with Neoglacial conditions (redrawn from Joynt and Wolfe 2001; Wolfe 2002). The root mean squared errors of prediction (RMSE) for the water temperature and pH transfer functions are 1.94°C and 0.38, respectively. Bootstrapped RMSE for these models are 2.79°C and 0.43 pH unit.

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affinities (Dyke et al. 1996b). Although few samples are available to assess the limnological response to post-LIA warming (Figure 8), changes in diatom assemblages in the upper 1.5 cm of the Fog Lake sediment core translate to a 1.5°C warming in approximately the last 150 years. Additional analyses at higher stratigraphic resolution (i.e., < 0.5 cm continuous sampling) are needed to examine this interval in greater detail. Interestingly, the late Holocene cooling trend reconstructed from Fog Lake is closely associated with expansions of acidophilous diatoms, such as Eunotia rh omboidea, E. vanheurckii, and Peronia fibul a (Joynt and Wolfe 2001). In Kekerturnak Lake, Aulacoseira distans replaced Cyclotella rossii as the dominant diatom, again suggesting a late Holocene lowering of lake water pH, in the order of 0.5 units (Figure 8). These observations have been presented as confirmation of previous evidence from other regions (Koinig et al. 1998) that lake water pH is at least partially regulated by climate in these dilute lakes (Wolfe 2002). Because pH is among the most readily reconstructed lake chemical variables from diatom assemblages, this relationship may be applied in an exploratory way in areas that lack training sets for paleoclimatic inferences, as long as pH is not overwhelmingly controlled by edaphic processes such as the leaching of carbonate lithologies. In support of this general relationship, it is noted that the highest diatom-inferred pH values from Fog Lake, slightly above 7.0, occurred during the last interglacial (Wolfe et al. 2000). The latest Holocene: a time of unprecedented change In 1994, Douglas et al. published a seminal paper demonstrating that recent (post-1850) diatom assemblages have become reorganized to degrees unparalleled in the previous several millennia. This work addressed cores from several ponds on Cape Herschel (Figure 1, site 11), each revealing shifts towards more diverse floras including greater representations by larger periphytic forms, indicating longer growing seasons and/or greater habitat availability. Although atmospheric contamination and changes in ultraviolet radiation were considered by Douglas et al. (1994), their final conclusion is that climate change is the most probable causative agent for these ecological changes. Since this time, comparable trends have been observed in diatom records from northern (Doubleday et al. 1995) and west-central (Wolfe 2000; Perren et al. 2003) Ellesmere Island, northern Devon Island (Gajewski et al. 1997), Cornwallis Island (Michelutti et al. 2003) as well as the Fennoscandian Arctic (Sorvari et al. 2002). Recent changes in diatom assemblages are further corroborated by shifts in siliceous chrysophyte microfossils from Ellesmere and Baffin islands (Wolfe and Perren 2001; Perren et al. 2003). Several of these records are reproduced here (Figure 9). The sobering reality drawn from these independent analyses is that many arctic lakes seem to be rapidly veering towards ecological states for which no prior analogs exist. It is evident that arctic climate has warmed. In this regard, the detailed multi-proxy analysis of Overpeck et al. (1997) demonstrates the uniqueness of the last ca. 70 years in the context of the last four centuries, and the high probability that anthropogenic (greenhouse gas) forcing has compounded the naturally-initiated warming that terminated the LIA. The paleolimnological component of this reconstruction is especially compelling in that it contains both biological (diatoms: Douglas et al. 1994;

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Figure 9. Summary of recent (post-LIA) paleolimnological changes recorded in a variety of lakes of the middle and high Canadian Arctic. Original sources are Douglas et al. 1994 (A) and (B), Wolfe 2000 (C), Doubleday et al. 1995 (D), Gajewski et al. 1997 (E), and Wolfe and Perren 2001 (F).

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Gajewski et al. 1997; Wolfe 2000; Perren et al. 2003) and physical sedimentological (varves: Hughen et al. 2000; Lamoureux 2000; Moore et al. 2001) indications of recent warming. However, warming of the Arctic is unlikely to be spatially homogenous over the large geographic area for which paleolimnological records register abrupt recent changes (Serreze et al. 2000; Laing et al. 2002). Furthermore, it seems improbable that present conditions, warm as they appear, exceed the thermal maxima attained during the Holocene optimum (Figure 5) or the last interglacial (Figure 4). Yet detailed analyses of the latter intervals have not revealed comparable stratigraphic changes to those found in sediments deposited in the post-LIA interval. For example, hundreds of siliceous microfossil preparations from Baffin Island lake sediments of various age have been scanned and counted, but it is only in the surface sediments of Kekerturnak Lake on Cumberland Peninsula that scaled mallomonadacean chrysophytes occur (Wolfe and Perren 2001). The data from other parts of the Arctic Archipelago manifest similar examples of 20th century ecological uniqueness, which are yet to be fully resolved. This discussion leads to the possibility that climate alone may not be solely accountable for the recent shifts observed in algal populations throughout the middle and high Canadian Arctic. As a working hypothesis, it seems entirely possible that the unquestionably dramatic recent ecological changes arise from a synergy between climate warming and some as of yet unspecified anthropogenic atmospheric input. For example, increased fluxes of fixed nitrogen to the Arctic are well chronicled in polar ice (Mayewski et al. 1990; Goto-Azuma and Koerner 2001), and may have the potential to enrich nutrient-impoverished aquatic ecosystems. Such possibilities, which remain to be adequately explored, satisfy the criterion of representing truly unprecedented environmental conditions in the context of the late Quaternary, in keeping with the very character of the primary paleoecological observations. Problems, recommendations and conclusions Although our understanding of paleolimnology in the middle and high Canadian Arctic is improving, it is clear that the density of sites with well-dated and high resolution analyses is still low given the size of the area under consideration. This is particularly evident in the islands of the western Arctic where, for example, Banks Island remains largely unexplored despite significant potential for temporally long and continuous nonglacial sediment records. The number of paleolimnological records that encompass the Holocene is still insufficient to map regional variability in the timing of maximum summer warmth, or the inception of Neoglacial cooling. The understanding of such patterns is critical in order to assess whether regional contrasts exist with respect to sensitivity towards present and future climate warming. The analysis of new sites should be complemented by additional detailed, integrative studies on carefully selected sites (cf. Bradley 1996). Additional efforts are needed to link physical, biological and chemical processes in lakes to sediment characteristics. The transport of littoral biocoenoses to central coring locations under lake ice is one example of a poorly understood process that is critical to the shaping of the sediment record. The models to be followed are those of the detailed multi-proxy paleolimnological campaigns conducted on intensely-studied lakes in the Alps (Hagelseewli: Lotter et al. 2000) and Lapland (Saanajärvi: Korhola et al. 2002). As an illustration of the paucity of current

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data, at present there are no published sediment records of fossil pigments, chironomids, cladocera or ostracodes from any site in the Arctic Archipelago. Lake sediment geochemical investigations are also in their infancy (Young and King 1989; Smith 1991; Wolfe and Härtling 1996, 1997; Sauer et al. 2001). On the other hand, advances are being made relatively rapidly in the fields of diatom taxonomy, autecology and quantification of diatom-environment relationships (Douglas and Smol 1993; Hamilton et al. 1994; Joynt and Wolfe 2001; Lim et al. 2001; Antoniades and Douglas 2002). These tools are only beginning to be applied downcore, as illustrated by several examples reported herein. Clearly, diatoms are powerful proxies for establishing linkages between limnological conditions and climate, and for identifying recent deviations from baseline limnological conditions. Varve studies have also begun to extend well beyond the purely descriptive stage, now providing specific insights into processes such as summer rainfall intensity (Lamoureux et al. 2001) and the transport of riverine biota (Ludlam et al. 1996). Thus, the last decade’s results have convincingly demonstrated that paleolimnology plays an important role in assessing the sensitivity of the Arctic towards environmental changes, whether mediated by natural or anthropogenic processes. Summary Paleolimnological research in the middle and high Canadian Arctic has a relatively brief history of about 20 years. Although paleolimnological studies in this region are currently proliferating, there remains a conspicuous paucity in the actual numbers and types of sedimentary records studied given the size of the geographical area and the high degree of geological, ecological, and climatological heterogeneity. Yet these same elements of spatial diversity are conducive to a rich array of lake and pond sediment records that, within existing studies, have provided important data on the rates and magnitude of late Quaternary environmental change. This is exemplified by studies that span a full spectrum ranging from pre-Holocene interglacial intervals to ecosystem changes of the late 20th century. Detailed investigations, primarily from localized networks of sites on Ellesmere and Baffin islands, provide useful baselines for future investigations. These and other relevant studies are highlighted primarily with respect to the climate signal contained within diatom-based lake sediment records. Acknowledgements Our paleolimnological research has been supported by the Natural Sciences and Engineering Research Council of Canada and the United States National Science Foundation. Thanks especially to John England, Robert Gilbert, Gifford Miller and John Smol for their ongoing support and encouragement, and to this volume’s editors for inviting our review. Logistical support by Polar Continental Shelf Project and the Nunavut Research Institute is also gratefully acknowledged. Paul Hamilton is thanked for his formal review of this manuscript.

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10. PALEOLIMNOLOGY OF THE NORTH AMERICAN SUBARCTIC

BRUCE P. FINNEY ([email protected]) Institute of Marine Science University of Alaska Fairbanks Fairbanks, Alaska 99775, USA KATHLEEN RÜHLAND ([email protected]) Paleoecological Environmental Assessment and Research Laboratory Department of Biology Queen’s University Kingston, Ontario K7L 3N6, Canada JOHN P. SMOL ([email protected]) Paleoecological Environmental Assessment and Research Laboratory Department of Biology Queen’s University Kingston, Ontario K7L 3N6, Canada and MARIE-ANDRÉE FALLU ([email protected]) Département de Chimie-Biologie Université du Québec à Trois-Rivières C.P. 500 Trois-Rivières, Québec G9A 5H7, Canada

Key words: Subarctic, North America, Paleolimnology, Treeline, Climatic change, Holocene, Late Quaternary, Lake systems

Introduction The North American Subarctic, the broad zone between the temperate mid-latitudes and the High Arctic, is an area that is highly sensitive to climatic changes (Schindler et al. 1990; Prentice et al. 1991; Barber et al. 2000; Pienitz and Vincent 2000; Serreze et al.

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2000). The two dominant, climate-sensitive features that often have an overriding influence on subarctic ecosystems are vegetation (e.g., the position of the treeline ecotone, as well as the composition and density of vegetation) and permafrost development, both of which, in turn, may affect hydrological and geochemical processes, substrate stability, water quality, and aquatic and terrestrial biological community composition (Rouse et al. 1997). Moreover, it is now generally recognized that any changes occurring in subarctic regions may also influence the global climate system (Bonan et al. 1995; Chapin et al. 2000). Modeling studies often emphasize the sensitivity of the global climate system to shifts in vegetation-related albedo and carbon storage (e.g., Foley et al. 1994; Bonan et al. 1995; Chapin et al. 2000), as feedback mechanisms may cause additional changes in regional and global climate (Rupp et al. 2000). For example, climatic warming may reduce albedo in tundra regions, further promoting a warming trend (Chapin et al. 2000), but may also result in an increase in albedo in the boreal forest if warming results in denser stands of deciduous trees (Eugster et al. 2000; Rupp et al. 2000). The impacts of climate change will clearly vary from region to region depending on numerous factors, including topography, degree of soil and permafrost development, vegetation type, and atmospheric circulation patterns (Chapin et al. 2000; Rupp et al. 2000). It is critical to better understand these responses and feedback mechanisms to climatic and environmental change that have occurred in the past in order to assess the impacts of future global change to this region and neighbouring areas. However, similar to other high latitude regions, long-term monitoring data are rarely available, and so paleoenvironmental approaches must be used to assess the nature and magnitude of past ecosystem changes. This chapter provides an overview of paleolimnological studies completed in the North American Subarctic. A major theme for most of these studies relates to climatic change, which is not surprising as one of the most distinctive characteristics of this vast region is the strong climatic and vegetation gradient from boreal forest (taiga) to tundra. Moreover, parts of the Subarctic are believed to be especially sensitive to the effects of greenhouse-induced climate warming. Furthermore, compared to other subarctic regions in parts of Europe and Russia, subarctic North America is relatively isolated from anthropogenic activities, and it can therefore generally be assumed that any changes recorded in the paleolimnological record are at least indirectly related to past climatic changes. Paleolimnological studies from the North American Subarctic are highlighted in this chapter with an emphasis on tracking past changes in climate, vegetation and biogeochemical processes. We focus on lake systems, but attempt to compare this research with other paleoenvironmental studies based on glaciers, terrestrial records, and pollen analysis of lake and peat cores. Several recent reviews of late Quaternary, pollen-based vegetation histories for the North American Subarctic are available (Ritchie 1987; Lamb and Edwards 1988; Wright et al. 1993; Edwards et al. 2000a, Bigelow et al. 2003; Gajewski and MacDonald, this volume). In addition, other terrestrial-based methods, such as fossil beetle analysis (Elias 2000), have been conducted in regions such as Alaska. Our chapter, however, focuses on limnological records, with the realization that multi-disciplinary studies are often required for holistic reconstructions of past ecosystem change.

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In this chapter, we subdivided the North American Subarctic into three subregions based primarily on their paleoenvironmental histories: eastern Beringia, central Canadian Subarctic, and northern Québec and Labrador (Figure 1). Most of the lake sediment chronologies presented here are based on radiocarbon (14C) dating. Although it is recognized that calibration of the radiocarbon timescale is necessary to correct radiocarbon dates to calendar years (e.g., Stuiver et al. 2000), we report dates as presented in the original publications. All radiocarbon ages are uncalibrated radiocarbon years before present (BP), except when indicated by cal. yr BP (calibrated 14C age BP). Ages, when determined by varve, 210Pb, 137Cs and other methods, are noted in the text.

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Figure 1. Map depicting the three subarctic regions of North America and the northern extent of treeline, as discussed in the text. 1 = Arolik Lake; 2 = Grandfather Lake; 3 = Birch Lake; 4 = Farewell Lake; 5 = Wolverine Lake; 6 = Lake U60; 7 = Tangled Up Lake; 8 = Meli Lake; 9 = Karluk Lake; 10 = Queen’s Lake; 11 = Toronto Lake; 12 = Lake TK-20; 13 = Slipper Lake; 14 = RLA Lake; 15 = McLeod Bay (Great Slave Lake); 16 = Beaufort Sea Coast and Tuktoyaktuk Peninsula; 17 = Leech Lake; 18 = Square Lake; 19 = Hendry and Tasirlaq-sud lakes; 20 = Isiurqutuuq Lake; 21 = Lake Kachishayoot and Lake Kaapumticuumac; 22 = Lake Karinbou and Lake B2; 23 = Lake Saglek; 24 = Lake K2 and Lake PC5; 25 = Lakes PC4, LM and LR; 26 = Lakes PC1, PC2 and PC3; 27 = Lake Oksana. AK = Alaska; YT = Yukon Territory; NWT = Northwest Territories; NU = Nunavut; QUE = Québec; LAB = Labrador; BC = British-Columbia; AB = Alberta; SK = Saskatchewan; MB = Manitoba; ON = Ontario.

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Description of study region Physiography The subarctic regions of North America cover a large geographic area including Alaska, the Yukon Territory, central parts of the Northwest Territories (NWT) and Nunavut, northern Alberta, northern Manitoba, northern Québec’s Ungava Peninsula, and Labrador (Figure 1). This landscape includes a wide range of vegetation, geomorphologic, and ecoclimatic gradients. For example, eastern Beringia contains relatively steep, high-elevation mountains with complex geological structures and variable bedrock, including sedimentary, metamorphic, volcanic, and igneous rocks. To the east, in the central Canadian Subarctic, the landscape is of relatively low relief underlain by Precambrian Shield granitic-gneissic bedrock that extends into northern Québec and Labrador. Three broad ecozones can be distinguished along a latitudinal gradient across the North American Subarctic. As the conditions favourable for tree growth decline northward, there are gradually changing zones of vegetation from dense boreal forest, to the forest-tundra transitional zone, to the treeless arctic tundra in the north (Figure 2). In the mountainous parts of the Yukon Territory and Alaska, an alpine tundra ecozone also exists, whereas a maritime-influenced ecozone consisting of a western hemlocksitka spruce forest and a spruce-fir forest-tundra is present in coastal parts of southern Alaska and southeastern Labrador, respectively. These boreal ecozones are generally characterized by the close relationship between climate and vegetation, where plant taxa meet their ecological limits, making subarctic treeline regions particularly well-suited for studies of environmental and climatic change. The steep climatic and environmental gradients in subarctic North America are largely a reflection of warm, moist Pacific air masses of the south converging with cold, dry, arctic air from the north known as the Arctic Front (Bryson 1966). Marked gradients in both precipitation and temperature are linked with notable differences in albedo between forested boreal regions and treeless arctic tundra regions (Bonan et al. 1995; Pielke and Vidale 1995). For example, average January and July temperatures range from -23 to 17°C, respectively, in interior Alaska, and -29 to 8°C in northern Alaska, respectively. Annual precipitation decreases from about 32 to 10 cm along this climatic/vegetation gradient. The position of the treeline (and, hence, albedo) is influenced by these climatic influences, although edaphic factors, including the degree of soil and permafrost development, bedrock lithology, and topography, also play an important role (Ritchie 1993; Timoney et al. 1993). The degree of soil development generally decreases from west to east corresponding to the timing of regional deglaciation (i.e., the longer the region has been ice-free the more time for soils to develop) and is influenced by the presence of underlying permafrost (i.e., a higher degree of permafrost development will inhibit soil development). Due to its restricting effects on vegetation, the southern limit of continuous permafrost generally follows the position of arctic treeline (MacDonald and Gajewski 1992).

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Figure 2. Photographs of lakes in different subarctic ecozones of North America. (A) Unnamed lake in tundra zone, northern Québec, Canada. (B) Unnamed lake in the forest-tundra transition zone, Northwest Territories, Canada.

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Figure 2 (continued). Photographs of lakes in different subarctic ecozones of North America. (C) Unnamed lake in the boreal forest zone, Northwest Territories, Canada. (D) Brooks Lake (58.5ºN, 155.9ºW), a large sockeye salmon nursery lake in the forest-tundra transition zone, Katmai National Park, Alaska.

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Glacial and climatic history The North American Subarctic can be subdivided into regions of different physiography and spatial patterns of glaciation and deglaciation. Much of Alaska, the Yukon Territory and parts of the Mackenzie River District of the western Northwest Territories were icefree during the Last Glacial Maximum (LGM; ca. 18,000 yr BP), whereas the rest of continental arctic Canada was covered by the Laurentide Ice Sheet (Dyke and Prest 1987; Hamilton 1994). Differences in glacial history played an important role in the climatic, vegetation, and paleolimnological history of the region. The process of glacial retreat followed a northeastern pattern over the Holocene resulting in asynchronous lake development and climatic change (e.g., the Holocene Thermal Maximum; Kaufman et al. 2004), with the eastern arctic region experiencing limnological and climatic changes much later than the western regions. The timing of maximum Late-Wisconsinan ice extent in Alaska and the Yukon occurred between ca. 24,000 to 14,000 yr BP, with deglaciation largely completed by 10,000 yr BP (Hamilton 1994; Mann and Hamilton 1995). In the central Canadian Subarctic, deglaciation lagged the western Canadian Subarctic by a few thousand years and likely occurred between 8000 to 7000 yr BP (Fulton and Andrews 1987). The persistence of ice may have contributed to a general delay in climatic warming in the central Canadian Subarctic compared to western regions, which may be related to changes in the geometry of the Arctic Front (MacDonald and Gajewski 1992). For example, Moser and MacDonald (1990) proposed that the asymmetric behaviour of long-wave westerly winds may have affected the summer position of the Arctic Front. A northward displacement of the Arctic Front in northwestern North America would have resulted in the southward displacement of the Front to the east (Namais 1970), which Moser and MacDonald (1990) contended may explain the asynchronies in Holocene warming and treeline shifts between northwestern and central Canada. Therefore, a northward shift of the arctic frontal zone ca. 5000 yr BP would have resulted in warmer and moister conditions associated with southern Pacific air masses. After 3000 to 3500 yr BP, climate cooling would have caused a southward shift in this front with an injection of cool, dry arctic air masses, resulting in regional treeline retreat (MacDonald et al. 1993; Pienitz et al. 1999). Most of the eastern Canadian Subarctic was still covered by the Laurentide Ice Sheet at 10,000 yr BP and deglaciation was a relatively slow process (Williams et al. 1995), principally occurring between 7000 and 6500 yr BP (Gray et al. 1993). The pattern of deglaciation was generally believed to have involved a residual ice cap that persisted over the Ungava and Labrador regions until approximately 6000 yr BP (Dyke and Prest 1987; Williams et al. 1995). However, two new concepts have been proposed for the regional retreat of the Laurentide Ice Sheet. The first postulates that the ice sheet fragmented into several individual ice caps and glaciers that may have persisted until ca. 5000 yr BP (Clark et al. 2000). The second asserts that the ice sheet retreated northwards and that ice persisted over southern Ungava Bay and the adjacent shores until ca. 6000 yr BP (Jansson 2003). Regardless of which scenario actually occurred, it is generally acknowledged that delayed deglaciation maintained a cool climate in northern Québec and Labrador for a much longer period than elsewhere in the North American Subarctic (Richard 1995).

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Aquatic ecosystems in subarctic North America Lentic systems in the North American Subarctic show tremendous diversity. For example, they vary in size from small (often < 0.5 ha), shallow (usually < 1 m deep) ponds to large (e.g., Great Bear Lake with an area of ca. 3130 ha), deep (some > 200 m deep) lakes. Lakes in the region originate from many processes, including volcanic, glacial, fluvial, and thermokarst activity. Ice cover may last for more than six months, and ice thicknesses more than 1.5 m are common, although coastal lakes in southern Alaska may occasionally remain ice-free during warm winters. In interior continental regions characterized by relatively warm summers, lake surface temperatures may exceed 20°C and thermal stratification is common. The decreasing trend in lake temperatures from south to north, combined with the progressive loss of wind barriers in the form of trees, results in a decreasing tendency for thermal stratification in tundra regions. Throughout the North American Subarctic, a variety of lake stratification types occur, including dimictic, monomictic, meromictic, and polymictic systems, which are influenced by climatic conditions, topographic and vegetation barriers, and lake and basin morphometry. The trophic states of these lakes also vary widely from oligotrophic to eutrophic, depending on a range of factors including watershed characteristics, lake bathymetry, and climatic setting. Generally, more productive lakes are located in the boreal forest region (e.g., Pienitz et al. 1997a,b; Fallu and Pienitz 1999; Gregory-Eaves et al. 2000; Rühland et al. 2003a). The steep climatic and vegetation gradients across the treeline ecotone result in substantial differences in catchment characteristics and, consequently, the limnological conditions of lakes on either side of this border. For example, colder climates and continuous permafrost in the arctic tundra impede hydrological processes to a greater degree than in the boreal forest, where precipitation and runoff is higher (Schindler et al. 1996; Henriksen et al. 1998). These differences in catchment characteristics across the North American subarctic treeline affect lake water alkalinity and dissolved inorganic carbon (DIC), with decreases in ionic concentrations in lakes from the boreal forest biome to the arctic tundra (Pienitz et al. 1997a,b; Rühland and Smol 1998; Fallu and Pienitz 1999; Gregory-Eaves et al. 2000; Fallu et al. 2002; Rühland et al. 2003a). Coniferous leaf litter in the catchments of boreal forest lakes provides a major source of allochthonous coloured humic substances, often characterized in limnological surveys by elevated concentrations of dissolved organic carbon (DOC) (Forsberg 1992; Meili 1992; Hongve 1999). Not surprisingly, pronounced differences in DOC are often measured by limnological sampling programs that cross from forest to tundra zones in the North American Subarctic (Pienitz et al. 1997a,b; Fallu and Pienitz 1999; GregoryEaves et al. 2000; Fallu et al. 2002; Rühland et al. 2003a). In some sections of the North American Subarctic, precipitation is relatively low, but summer temperatures are relatively warm (e.g., interior Alaska, average annual precipitation < 275 mm, summer temperatures range from ca. 13ºC to ca. 23ºC). In such regions, watershed evapotranspiration and lake surface evaporation are relatively high, with potential evapotranspiration equal to annual precipitation (Barber and Finney 2000). Thus, similar to some temperate localities, the water balance of subarctic lakes may be tightly linked to climate changes (Barber and Finney 2000; MacDonald et al. 2000), potentially resulting in shifts in lake-levels (Abbott et al. 2000; Barber and

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Finney 2000), as well as the chemical and isotopic composition of lake water (Hu et al. 1998; Pienitz et al. 2000; Anderson et al. 2001; Edwards et al., this volume). Paleoindicators of the North American Subarctic The location of the northern treeline is strongly influenced by the mean July position of the Arctic Front, which itself is closely related to the summer position of the 10°C isotherm (Bryson 1966; Pielke and Vidale 1995). Boreal ecosystems in North America (and in other subarctic regions) are geographically vast and regionally variable (e.g., Korhola and Weckström, this volume; MacDonald et al., this volume). Within the vast North American Subarctic, marked local differences underscore the need for numerous regional calibration data sets to more accurately document the ecology of aquatic organisms from different geographic regions. Understanding local processes and environmental controls is also key for more accurate inferences from isotopic and geochemical indicators. These data are the foundation for building robust and meaningful inference models. Below, we provide a brief overview of how some of the main paleolimnological indicators are used in the North American Subarctic. More detailed discussions, however, are provided in chapters two to eight of this volume. Diatoms Diatoms (class Bacillariophyceae) are one of the most extensively used biological indicators in paleolimnological investigations, as they are abundant, ecologically diverse, preserve well in lake sediments, and their short life-spans enable them to respond rapidly to environmental changes (Stoermer and Smol 1999). Many diatoms have definable optima to alkalinity (or alkalinity-related variables, such as DIC, specific conductivity, and major ions), DOC, silica, and nutrients (such as phosphorus and nitrogen). In subarctic North America, limnological surveys have consistently noted that these variables were strongly related to vegetation/climatic gradients from boreal forest to open tundra (Pienitz and Smol 1993; Pienitz et al. 1995a, 1997a,b; Rühland and Smol 1998; Fallu and Pienitz 1999; Gregory-Eaves et al. 1999; Fallu et al. 2002; Rühland et al. 2003a), which could also be linked to diatom assemblages characterizing different ecozones (e.g., tundra, forest-tundra transition, boreal forest lakes; Rühland et al. 2003b) within each sub-region. However, in some areas, such as to the south of the coniferous treeline border within the boreal forest in Wood Buffalo National Park (northern Alberta and the Northwest Territories), differences in underlying lithology are the primary drivers of lake water chemistry, and therefore become overriding factors in determining diatom distributions (Moser et al. 2004). Trends in measured limnological variables across treeline are known to be indirectly related to other climatic factors. For example, compared to the arctic tundra, the boreal forest is warmer, the degree of permafrost development is less, soils are thicker and more highly developed, precipitation is higher, net groundwater inputs are elevated, weathering and ion supply are higher, nutrient cycling is faster, and coniferous leaf litter is present in the catchments. These climate-related environmental conditions affect lake

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water chemistry, typically resulting in higher concentrations of major ions (DIC, alkalinity, conductivity), nutrients, and DOC in boreal forest lakes compared to arctic tundra lakes. Hence, climate-induced changes in lake catchment characteristics will affect the chemical and physical limnology of lakes, and therefore also influence the biota. Diatom-based calibration studies from Québec-Labrador (Fallu et al. 2002), northern Québec (Fallu and Pienitz 1999), the central Canadian Subarctic (Pienitz and Smol 1993; Rühland and Smol 2002), the Yukon Territory (Pienitz et al. 1995a), and Alaska (Gregory-Eaves et al. 1999) have consistently identified the above limnological parameters as important explanatory variables in describing diatom distributions across circumpolar treeline regions. Although the North American Subarctic is a geographically vast region, in many cases similar distribution trends have been recorded in these diatom-based calibration studies, which should eventually allow for the development of large or ‘supra-regional’ calibration sets (sensu Lotter et al. 1999). There are numerous indirect properties affecting lake ecosystems that are more challenging to model, but are important in determining diatom assemblage composition. For example, these factors include the length of the growing season (e.g., duration of seasonal ice cover on the lake), the amount of solar radiation, and the degree of lake water stratification. Changes in these physical properties affect lake water characteristics, aquatic habitat availability and, ultimately, the biota found in the lake. For example, in deeper, subarctic lakes from various circumpolar regions, an increase in the duration and strength of thermal stratification influences a suite of other limnological properties for diatom growth, such as the amount of photosyntheticallyavailable light, nutrients, and mixing regimes, to name a few. Some studies have suggested that diatom assemblages may be used to provide estimates of temperature. For example, summer surface water temperature was identified as an important environmental variable explaining the distribution of diatom assemblages preserved in the surface sediments of 59 lakes located between Whitehorse in the southern Yukon Territory and Tuktoyaktuk near the coast of the Arctic Ocean in the Northwest Territories (Pienitz et al. 1995a). A weighted-averaging temperature reconstruction model was developed, although the authors acknowledged that many indirect influences of temperature may affect diatom assemblage composition (e.g., thermal stratification, timing of ice break-up, lake depth, biological factors), and so these models must be used cautiously. Pienitz et al. (1995a) contended, however, that regardless of whether diatoms were affected directly or indirectly by summer surface water temperature, these models have the potential to provide evidence for relative paleotemperature changes. Surface sediment calibration studies have also been undertaken to develop diatombased models for the reconstruction of past hydrological conditions through climatically-induced changes in lake depth in Wood Buffalo National Park, Alberta and the Northwest Territories (Moser 1996). However, lake-level inferences using paleolimnological models are challenging and inference results must be treated cautiously. For example, in lakes that are not closed basins, the response to climatic change is complicated as changing lake-level signals may be dampened by changes in outflow (Moser et al. 2000). As well, maximum lake depth is restricted by a lake’s morphometry so that it is possible for an inference model to reconstruct a lake-level that

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is not physically possible (Moser et al. 2000). Nevertheless, these models show promise for tracking the overall trends and the directions of lake-level changes. In an attempt to model flooding frequency in the Mackenzie River Delta, Hay et al. (1997) examined a suite of floodplain lakes along a gradient of differing river influences including: (a) lakes that have a continuous connection with the Mackenzie River; (b) lakes that lose their connection during the summer but flood in the spring; and (c) lakes that flood only during extreme flood events. Winter methane measurements from these floodplain lakes were found to explain a statistically significant portion of the diatom species distributions among the lakes. The turbid waters of the Mackenzie River impede the growth of macrophytes in lakes having connections to the river resulting in reduced winter methane production, whereas macrophyte growth and winter methane production were greater in lakes with limited river connections. As the authors acknowledged, diatoms do not respond directly to changes in methane concentrations. However, models such as these may be used as indirect proxy for summer production levels that are associated with flooding frequency. Campeau et al. (1999) used diatom assemblages to develop a paleo-depth transfer function from 74 sites distributed from environments of different depths along the southeastern Beaufort Sea coast. The relationship between diatoms and the water depth gradient was believed to be a function of shoreface circulation. Shallower sites were dominated by epipelic and epipsammic diatom species, whereas deeper sites were dominated by planktonic taxa. This paleo-depth model was then applied to sediment cores as a quantitative estimate of past sea-level fluctuations during the late Holocene (Campeau et al. 2000; discussed below). New techniques are being assessed for refining diatom-based inference models. For example, Racca et al. (in press) re-analysed calibration data from Alaska and the northwestern Canadian Subarctic to explore the influence of removing certain taxa, and therefore only including species that appeared to have numerically useful relationships to modeled environmental variables. By restricting or “pruning” the diatom data in this way, they found that the predictive power of their models for DOC, water depth, and water temperature, as well as the predictive ranges of these models, were improved. They acknowledge that, by excluding diatom species, they may decrease the sensitivity of these models in some cases, as the distributions and abundances of these taxa may be strongly influenced by several variables that may have been important in past environments. Chrysophytes In subarctic North America, paleolimnological studies using scaled chrysophytes are uncommon and only a handful of investigations using chrysophycean cysts have been undertaken (Pienitz et al. 1992; Duff et al. 1995; Brown et al. 1997; Taylor 1997; Wilkinson et al. 2001). Nonetheless, there is considerable potential for future applications. Brown et al. (1997) concluded that the distribution of chrysophycean cysts across boreal ecotones in the Yukon and Northwest Territories was significantly related to lake water chloride concentrations, DIC, and temperature, and that the development of inference models provided a useful complement to other paleoecological indicators, such as diatoms. Taylor (1997) examined the distribution of chrysophycean cysts across

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the treeline in the Northwest Territories and in the boreal forest zone of Wood Buffalo National Park. The strongest inference models developed from the distribution of cysts were for lake water calcium concentrations and DOC. One of the continuing challenges in using chrysophycean cysts for paleolimnological reconstructions is the lack of autecological information on the cyst morphotypes, although advances in the field are being made (Duff et al. 1995; Wilkinson et al. 2001). Chironomids Fossil midge head capsules are being used increasingly in paleolimnological research (Walker 2001) and subarctic North America is no exception. This increased attention is often related to the statistically significant relationships found between the distributions of chironomid head capsule assemblages in lake surface sediments and summer surfacewater temperature, maximum lake depth, and salinity (e.g., Walker et al. 1991, 1997, 2003; Pienitz et al. 1992; Wilson et al. 1993; Walker and MacDonald 1995; Bennike et al., this volume), which are all important variables for paleoclimatic reconstructions. Across the treeline gradient in subarctic Canada, many chironomid taxa appear to have defined thermal ranges, particularly in more northern lakes (Walker et al. 1997, 2003; Bennike et al., this volume). The resulting transfer functions have provided quantitative tools for paleoclimatic reconstructions (e.g., Walker et al. 1991; Wilson et al. 1993; Levesque et al. 1997; Pellat et al. 2000). Lake depth has often been identified as a significant explanatory variable for chironomid assemblage distributions (Walker 2001; Walker et al. 2003). Walker and MacDonald (1995) and Walker et al. (2003) examined the spatial distribution of chironomids across treeline in the central Canadian Subarctic and in the Yukon and Northwest Territories, respectively, and both found strong relationships to lake depth, suggesting that past water-levels can be reconstructed from subarctic chironomid records. As past lake-level fluctuations may be linked to climatic change, these data may also provide important paleoenvironmental inferences. In addition to lake depth and surface water temperature, Walker et al. (2003) also found that chironomid distributions were strongly correlated to total Kjeldahl nitrogen (TKN) and pH, a trend not found in northern Labrador (Walker et al. 1991), likely because the gradient lengths for these variables in northwestern Canada were much longer. Similarly, the relationship between chironomid taxa distributions and lake water salinity (Walker et al. 1995) may provide important paleoclimatic information on past evaporationprecipitation regimes from cores collected from athalassic (inland saline) lakes (Pienitz et al. 1992, 2000). Chironomid indicators are now being used increasingly for a wide variety of applications in temperate regions. However, the potential of chironomid contributions to paleolimnological work in subarctic North America has not yet been fully exploited, and future developments are anticipated.

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Thecamoebians Thecamoebians (testate Protozoa) are commonly found in Sphagnum-dominated bogs and other wetland habitats (Booth 2002), as well as in other freshwater environments (Pienitz et al. 1995b; Dallimore et al. 2000). As moisture content is an important determinant of testate amoebae distributions (Ogden and Hedley 1980), these indicators have been used to infer moisture and surface wetness of substrates (Warner and Charman 1994; Burbridge and Schröder-Adams 1998; Charman and Hendon 2000; Booth 2001; Hendon et al. 2001). However, only a few studies using these paleoindicators have been completed for the North American Subarctic. For example, Dallimore et al. (2000) examined the distribution of thecamoebian species in 25 thermokarst lakes and their distributions in the Holocene sediments of two thermokarst lakes on Richards Island, Northwest Territories (ca. 69º10’- 69º43’N, 134º00’135º00’W). They concluded that climate appears to influence thecamoebian distributions and changes throughout the Holocene, but (as is true for all proxies) there are complex relationships between thecamoebian assemblages and environmental factors. Zooplankton Paleolimnological studies using Cladocera have been used to assist in reconstructions of climate, water-level changes, deforestation, acidification, and trophic status in more southern, temperate regions (e.g., Brodersen et al. 1998; Bos et al. 1999; Jeppesen et al. 2001; Korhola and Rautio 2001; Bredesen et al. 2002), as well as in some studies from the European Subarctic (e.g., Bennike et al., this volume), but have rarely been used in the North American Subarctic. Cladocera depend on algae and other particulate matter for their food, whereas Cladocera are preyed upon by fish and larger invertebrates (Korhola and Rautio 2001). As cladocerans occupy intermediate trophic positions in many food webs, they may provide important data on food web structure (e.g., Jeppesen et al. 2002; Sweetman and Finney 2003). For example, in Karluk Lake (57°25’N, 154°05’W), a sockeye salmon nursery lake on Kodiak Island, Alaska, salmon-derived nutrient-loading led to increased abundance of invertebrate predators (cyclopoid copepods) (Sweetman and Finney 2003; Sweetman et al., in press). As a consequence, Bosmina body size increased over time, presumably from invertebrate predation pressure. These morphological changes were used to track past salmon abundances and demonstrated some of the ecological insights concerning trophic interactions that may be gained from using cladocerans. Recently, there has been heightened interest in examining the effects of increased ultraviolet (UV) radiation on the ecology of shallow arctic and subarctic ponds (e.g., Hessen 1993; Zellmer 1996; Leavitt et al. 1997; Vinebrooke and Leavitt 1998; Rautio et al. 2001). For example, zooplankton populations from recently deglaciated lakes in Glacier Bay, Alaska (58º56’N, 136º15’W) varied considerably in species composition and abundance among lakes along a gradient of UV attenuation (Williamson et al. 2001). Based on the strong relationships between DOC and UV attenuation, as well as between DOC and zooplankton distributions, they concluded that UV radiation may be

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a more important factor than previously appreciated in determining the distribution and abundance of zooplankton in subarctic lake ecosystems. Ostracodes are also potentially sensitive indicators of changes in nutrients, salinity, temperature and habitat type (Holmes 2001), but few studies have used these indicators in subarctic lake sediments. Trace element content and stable isotope ratios of ostracode shells have been used to infer water temperature, water chemistry and productivity, and their shells have also been used for radiometric dating (Holmes 2001; Schwalb 2003). In addition, ostracodes have been used to reconstruct climatic change in Alaska based on their species assemblages (Forester et al. 1989), and on their trace element and stable isotopic geochemistry (Hu et al. 1998, 2001a,b). These studies demonstrated the potential of ostracode analyses which, to date, have not been fully developed in the North American Subarctic. Geochemical and isotopic analyses The use of geochemical and isotopic tracers in paleolimnological studies has significantly expanded since the 1990s (Edwards et al., this volume). Isotopic analyses of organic matter (į13C, į15N, į18O), carbonates (į13C, į18O), biogenic silica (į18O), and major and trace metal analyses of sediments and ostracodes have been conducted in subarctic Alaska and Labrador (Engstrom and Hansen 1985; Hu et al. 1993, 1998, 2001a,b; Finney et al. 2000, 2002; Anderson et al. 2001). Interpretations rely on universal geochemical principles, as process studies or surveys of these variables along subarctic environmental gradients are still lacking. Maps estimating broad patterns of į18O in precipitation are available (Bowen and Wilkinson 2002), and surveys of surface water į18O and įD are in progress (Coplen and Kendall 2000), but are not sufficiently developed for many parts of the Canadian Subarctic. Regional syntheses As with many other regions, the earliest studies on lake sediments from the North American Subarctic focused on fossil pollen (Gajewski and MacDonald, this volume). In Alaska, the Yukon and the Northwest Territories, pollen-based studies began in the 1950s (e.g., Livingstone 1955) and continued into the 1960s (e.g., Colinvaux 1964) and 1970s (e.g., Rampton 1971; Ritchie and Hare 1971; Ager 1975), to present times (Gajewski and MacDonald, this volume). In Labrador, palynological data are available from the 1940s (e.g., Wenner 1947). Although only a few studies were conducted in the 1950s and 1960s in northern Québec and Labrador (e.g., Grayson 1956), many new investigations were initiated in the 1970s (e.g., reviewed in Short and Nichols 1977; Richard 1979). Multiproxy approaches and advances in dating (e.g., accelerator mass spectrometry (AMS) radiocarbon techniques), aided by improvements in analytical techniques, have led to an increased understanding of the history of this region. Within the past decade, several major paleoenvironmental research programs had been initiated in this region, including the U.S. National Science Foundation funded Paleoclimates from Arctic Lakes and Estuaries (PALE), and Paleoenvironmental Arctic Sciences (PARCS), as well as the “Paleoecological Analysis of Circumpolar Treeline” (PACT)

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project, founded by the Natural Sciences and Engineering Research Council (NSERC) of Canada. As a result, studies using new and multiple indicators (e.g., diatoms, chrysophytes, isotopes, etc.) have become more common. Eastern Beringia For this review, we have adopted a broad definition of eastern Beringia that encompasses the region from east of the Bering Strait to the Mackenzie River. This area includes Alaska, the Yukon Territory, and the western part of the Northwest Territories (Figure 1). During periods of low sea-level in the Quaternary period, when the broad and shallow continental shelves of the Bering and Chukchi seas were exposed, eastern and western Beringia (the Chukotka Peninsula and the region of Siberia west of the Kolyma River) were united into a large, high latitude subcontinent (Hopkins 1967). This clearly facilitated the exchange of biota, including humans, between the old and new worlds (Hopkins et al. 1982; Dixon 2001). The limited ice extent during the last glacial period, as well as during earlier glaciations, distinguishes eastern Beringia from the Canadian subarctic regions farther east. Late Pleistocene In unglaciated regions of eastern Beringia, lakes have the potential to provide information on environmental conditions throughout the Last Glacial Maximum, and even earlier periods. However, long and continuous lacustrine sequences are uncommon in this region, as many extant lakes dried up completely during arid phases of the last glacial period (e.g., Abbott et al. 2000). Some lake records appear to span at least two glacial-interglacial cycles, or contain sediments dating to the last interglacial period, although accurate dating remains a challenge for periods older than the limits of the 14C technique (Colinvaux 1964; Eisner and Colinvaux 1990; Edwards and McDowell 1991; Berger and Anderson 1994). Environmental conditions for the previous interglacial (probably marine oxygen isotope stage 5e, occurring between 130,000 to 115,000 yr BP) inferred from these records are almost exclusively based on terrestrial pollen data, which indicated climatic conditions as warm or warmer than today. For example, lake sediment records north of the Brooks Range in Alaska suggest that Picea (spruce) treeline was substantially north of its Holocene position (Eisner and Colinvaux 1990). More studies are needed to elucidate the environmental conditions that predate the Last Glacial Maximum and that probably correspond to interstadial climates during marine oxygen isotope stage 3 (between ca. 60,000 to 25,000 yr BP). Records based mainly on pollen data for central and northern Alaska (Ager 1983; Ager and Brubaker 1985; Anderson 1985; Anderson et al. 1994) and the Yukon (Rampton 1971), as well as lake productivity indicators in Arolik Lake (59°28’N, 161°07’W) in southwestern Alaska (Kaufman et al. 2003), all suggest a climate intermediate between that of the Holocene and that of the full glacial period during marine oxygen isotope stage 3. Lake sediment records are available for the full and early deglacial periods in Alaska from the central and eastern regions (e.g., Ager and Brubaker 1985; Anderson et al. 1988; Lamb and Edwards 1988; Hu et al. 1993; Abbott et al. 2000; Bigelow and Edwards 2001), the northern and northwestern regions (e.g., Edwards et al. 1985;

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Anderson 1988; Anderson et al. 1994; Mann et al. 2002a), and the southern regions (e.g., Engstrom et al. 1990; Peteet and Mann 1994; Hu et al. 1995; Hansen and Engstrom 1996; Brubaker et al. 2001; Hu and Shemesh 2003), as well as the Yukon Territory (e.g., Cwynar 1982; Ritchie and Cwynar 1982; Ritchie et al. 1983) and the Northwest Territories (e.g., Ritchie 1984; MacDonald 1987; Szeicz and MacDonald 2001). Once again, most paleoenvironmental information originates from pollen data, which generally suggest a variety of tundra vegetation types across the region until about 12,000 to 14,000 yr BP. In some higher elevation regions, or at sites proximal to glaciers, tundra vegetation existed until later into the Pleistocene. Lake-level reconstructions for Birch (64º18’N, 146º38’W; Abbott et al. 2000), Jan (63º34’N, 143º54’W; Barber and Finney 2000) and Windmill (63º65’N, 148º13’W; Bigelow and Edwards 2001) lakes, in conjunction with the pollen data, indicated cool and arid conditions, with precipitation in central Alaska less than half of modern-day values (Barber and Finney 2000) for the period just prior to 12,000 yr BP. Many early studies relied on 14C dating of bulk sediment; however subsequent dating of terrestrial macrofossils using AMS techniques suggested that bulk dates were often too old by more than 1000 years (Abbott et al. 2000). As with other arctic regions, the lack of a spatially coherent network of well-dated sites currently hinders our ability to determine regional trends and to make comparisons to hemispheric and global records (Duval et al. 1999). Nonetheless, pollen, lake-level, geochemical, and isotopic records suggest that a complex set of abrupt climatic changes occurred during the deglacial period ca. 14,000 to 10,000 yr BP in Alaska (Figure 3; Abbott et al. 2000; Hu and Shemesh 2003; Kaufman et al. 2003). Lake-level and pollen data suggest significant climate amelioration, including increased precipitation beginning ca. 14,000 to 12,000 yr BP. Because of uncertainties in chronology owing to both differences in dating methods and, more recently, to the calibration of radiocarbon dates to calendar years (Stuiver et al. 1998), it is not clear if these changes represent regional climatic variability. Further, it is not well understood how these changes related to broader northern hemispheric climate shifts. Some records suggest that this first major deglacial climatic step is correlative to the Bølling Period (ca. 15,000 to 13,000 yr BP) (e.g., lakelevels from Lake of the Pleistocene, 68°36’N, 156°16’W; Mann et al. 2002a), while other records (e.g., biogenic silica į18O data from Grandfather Lake, 60°48’N, 158°31’W; Hu and Shemesh 2003) suggest a later initiation. Paleolimnological evidence of the Younger Dryas (YD) cold period (ca. 11,000 to 10,000 yr BP) in eastern Beringia has only recently been presented (e.g., Figure 3). For example, pollen and sedimentological records from southeastern Alaska (Engstrom et al. 1990), as well as pollen, sedimentological, isotopic and geochemical records from southwestern Alaska (Hu et al. 1995, 2002, 2003; Brubaker et al. 2001; Hu and Shemesh 2003; Kaufman et al. 2003) and from Kodiak Island (Peteet and Mann 1994; Hajdas et al. 1998) all suggest cooler and/or drier conditions that were consistent with the YD event. Pollen records from central Alaska (Bigelow and Edwards 2001) and lake-level records from northern Alaska (Mann et al. 2002a) also suggest drier conditions during the YD chronozone. The belated recognition of the existence of a YD interval in lake records from this region may be because the changes occurring in eastern Beringia during this period were subtle compared to the more dramatic changes before and after this event. In addition, other regional factors may have contributed to climate change during the late Pleistocene, which partially obscures or overprints the

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Figure 3. Paleolimnological record of the deglacial period for Alaska represented by time-series of organic į13C and C/N, organic matter (OM) % and biogenic silica (BSi) % for Arolik Lake, southwestern Alaska, compared to the į18O record from the Greenland ice core (GISP) and northern hemisphere solar insolation (from Kaufman et al. 2003).

YD signal in eastern Beringia. For example, sea-level rise during the deglacial period flooded the expansive continental shelves of the Bering and Chukchi seas, drastically changing the continentality of some sites in Alaska. There appear to have been minor changes in shelf area during the YD chronozone, but particularly large changes occurred immediately following the end of this period (Abbott et al. 2000; Mann et al. 2001). Moreover, southern coastal sites in Alaska may have registered stronger signals during the YD as they were less susceptible to the effects of sea-level rise due to the relatively

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narrow continental shelf of the Gulf of Alaska, and were more strongly influenced by changes in atmospheric circulation in the North Pacific during this period. Early to mid-Holocene climate-related paleolimnological records An increasing number of multiproxy lake records are available for understanding environmental change during the Holocene. The early to mid-Holocene period in eastern Beringia was significantly different than the late Holocene in that more dramatic environmental changes occurred. In the interior, pollen records document a complicated vegetation history leading to the development of the modern boreal forest. Likely environmental controls were almost certainly related to climatic changes, including temperature, precipitation, and seasonality, as well as ecological factors such as migration, fire, and vegetation and edaphic changes (e.g., Lynch et al. 2002). The Holocene records of many interior lakes indicate that water-levels were at their lowest during the earliest part of the period ca. 10,000 to 8,000 yr BP, and generally deepened during the mid-Holocene (Abbott et al. 2000; Barber and Finney 2000; Edwards et al. 2000b). Estimates of precipitation based on hydrologic modeling of these low lakelevels suggest that precipitation was about 25 to 40% less than modern values in the early Holocene (Barber and Finney 2000). Trace metal data from ostracodes preserved in Farewell Lake (62º55’N, 153º63’W; Hu et al. 1998), į18O records from aquatic organic matter in Meli Lake (68°41’N, 149°04’W; Anderson et al. 2001), diatom-based salinity reconstructions in Lake U60 (62°45’N, 136°38’W; Pienitz et al. 2000) (Figure 4), as well as lake-based records of sand influx reflecting aeolian activity in Wolverine Lake (67°06’N, 158°55’W; Mann et al. 2002b) all support lake-level inferences of a relatively dry early Holocene period. There is, however, some uncertainty in the temperature reconstructions for the early Holocene, but glacial, pollen, and macrofossil data suggested warmer conditions (Mann et al. 1998; Calkin et al. 2001). As well, ostracode trace metal ratios in Farewell Lake suggest temperatures warmer than present during the earliest Holocene (Hu et al. 1998). The Holocene thermal optimum occurred earlier in this region than in subarctic sites to the east, and the timing in eastern Beringia is generally consistent with forcing from changes in solar insolation, coupled with the deglaciation history (Ritchie et al. 1983). Most records are in agreement that climatic conditions generally became cooler and wetter through the mid-Holocene (e.g., Abbott et al. 2000; Pienitz et al. 2000; Anderson et al. 2001; Mann et al. 2002b), and that the severe early arid phase ended between 9000 and 7000 yr BP (Figure 4). Some records suggest a significant change to a cooler and wetter climate mode between about 5500 to 3000 yr BP (Abbott et al. 2000; Anderson et al. 2001). This onset of the Neoglacial is generally associated with the re-advancement of mountain glaciers in the Brooks and Alaska ranges, and in the mountains along the southern Alaska coast (Calkin et al. 2001). Given the present network of sites, it is unclear whether differences in the expression of the Neoglacial represent regional climate variability, dating uncertainties, differing sensitivities of environmental proxies, and/or changes in different components of the climate system. Except for hydrologic models of lake-level based precipitation, and diatom-based salinity reconstructions, quantitative studies are rare for this period.

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Figure 4. (A) Diatom-inferred salinity reconstruction for Lake U60 in the central Yukon (from Pienitz et al. 2000). Upper panel is on a log scale with a 3-point running average, and lower panel is the deviation from the mean log salinity. (B) Oxygen isotope records from Meli (lower) and Tangled Up (upper) lakes in the central Brooks Range, northern Alaska (from Anderson et al. 2001). The į18O on sediment cellulose from Meli Lake, which has a relatively long water residence time, is interpreted in terms of the evaporation-to-input ratio. The calcite į18O from Tangled Up Lake, which has a shorter water residence time, is interpreted as being controlled more strongly by temperature.

Early to mid-Holocene non-climatic paleolimnological records As previously mentioned, studies of paleolimnological changes that are not climatically induced are rare in subarctic North America. However, a recent paleolimnological study of the sediment records from a suite of 33 lakes from Glacier Bay, Alaska, tracked changes in chemical and biological trends during the lakes’ developments within the past 10,000 years (Engstrom et al. 2000). Comparisons of limnological trends inferred from a chronosequence of lakes, in combination with diatom-based reconstructions from 10 sediment cores (Engstrom et al. 2000), revealed that there was a directional change in all lakes towards more dilute and acidic water chemistry, coupled with the accumulation of dissolved organic carbon over time. Trends in nutrients, however, were more variable. The authors linked these results to a tight hydrological coupling between terrestrial and aquatic environments. Primary terrestrial succession (i.e., the rapid revegetation of the newly deglaciated landscape), together with associated changes in soil

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development and humus accumulation, have led to directional changes in the aquatic chemistry and have ultimately changed the biological components of these lakes. Late Holocene records and climate warming There has recently been an increasing emphasis on reconstructing environments at higher temporal resolution during the late Holocene so as to place recent observations into a longer term context. Instrumental climate records for parts of eastern Beringia extend back ca. 100 years, and suggest significant warming in recent decades (Chapman and Walsh 1993). In interior Alaska, for example, the trend in warming over the 20th century is greater than that for the Northern Hemisphere as a whole, and has had significant impacts on permafrost and terrestrial ecosystems (e.g., Barber et al. 2000). Over most of eastern Beringia, pollen data indicate that vegetation had reached its modern-day distribution by at least ca. 2000 yr BP, and pollen variability has generally been relatively low since this time. However, records of glacial advance (Calkin et al. 2001) and tree-ring analyses (Overpeck et al. 1997) suggest significant environmental and climatic changes during this period. Several recent studies have used lacustrine sediment records at relatively high resolution (decadal scale) to address climate change questions during the late Holocene. A ca. 2000-year record of temperature, reconstructed from į18O of ostracodes and abiotic carbonates from Farewell Lake, suggested significant changes in temperatures during the growing season (Hu et al. 2001b) (Figure 5). Warm periods were recorded between 0 to 300 AD, 850 to 1200 AD, and post-1800 cal. yr AD. Interestingly, data representing the 20th century do not indicate that this recent warm period is unusual relative to the preceding warm periods, although this may partially reflect the sampling resolution. Anderson et al. (2001) compared į18O records from two lakes from the central Brooks Range in northern Alaska (Tangled Up, 67°40’N, 149°04’W; Meli, 68º41’N, 149º04’W) with markedly different water residence times to estimate relative changes in temperature and moisture (Figure 4B). These records indicated generally cooler and wetter conditions over much of the last ca. 2000 years, relative to the present. The sand influx record from Wolverine Lake from the Kobuk sand dunes in northwestern Alaska suggests an overall trend of increasing moisture coincident with the Neoglacial (Mann et al. 2002b). Since 2600 cal. yr BP, two arid phases were indicated at 700 to 1300 AD and 1700 to 1900 cal. yr AD. Similarly, a diatom-based lake water salinity reconstruction for Lake U60 in the central Yukon Territory also suggested relatively arid and unstable conditions during the last ca. 2000 years, relative to the previous ca. 6000 years (Pienitz et al. 2000). A trend of increasing salinity occurred from ca. 0 to 1500 AD, after which salinity decreased to present-day values, which were much higher than those reconstructed for the mid-Holocene (Figure 4A; Pienitz et al. 2000). It is difficult to draw conclusions about broad patterns of past climates from the available lake records, as study sites are widely spaced and use different approaches. Analysis of 20th century climate data for the interior of eastern Beringia (Mock et al. 1998; Edwards et al. 2001) showed that two widespread summer weather patterns (warm/dry and cool/wet) dominated this region. However, the paleolimnological records suggested regional variability within eastern Beringia, and different synoptic

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patterns than those that have been observed historically (Edwards et al. 2001; Barber et al. 2004). The lake records that are currently available, as well as recent work on varved sediments from nearby British Columbia (Cockburn 2003), indicate the potential of this region for high-resolution climatic reconstructions. As a more widespread network of study sites develops, a better understanding of past climatic patterns should be possible. Diatoms are well known to be sensitive indicators of nutrient levels and lake trophic status. Past changes in nutrient levels may be linked to climatic changes (e.g., MacDonald et al. 1993), but may also be related to other sources in the Subarctic, including anthropogenic influences, or inputs from biological pathways, such as Pacific salmon migrations. For example, recent studies have attempted to reconstruct past Pacific salmon dynamics from sediment core analyses (e.g., Finney et al. 2000, 2002; Gregory-Eaves et al. 2003). As sockeye salmon utilize lake habitats (i.e., nursery lakes) for spawning and rearing, they are well suited for paleolimnological studies. Past salmon abundance can be reconstructed from į15N analysis, as sockeye salmon deliver significant quantities of marine-derived nutrients to lakes when they return to spawn and die. Salmon carcasses are enriched in 15N relative to other terrestrial sources, and thus į15N is a suitable proxy for past levels of salmon-derived nutrients (Finney et al. 2000). Similarly, diatom assemblages are also sensitive indicators of past changes in nutrient loading from sockeye salmon (Figure 6A; Finney et al. 2000; Gregory-Eaves et al. 2003). Reconstructions (e.g., Figure 6B) from Kodiak Island lakes (i.e., Karluk,

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57°25’N, 154°05’W; Akalura, 57°11’N, 154°12’W; Frazer, 57°15’N, 154°08’W) and Bristol Bay region lakes (Becharof, 57º50’N, 156º07’W; Ugashik, 57°41’N, 156°39’W; Tazimina, 59°59’N, 154°32’W) of Alaska provide new insight into salmon population dynamics, as they reveal variability over timescales much longer than those provided by historical records (ca. 100 years). The similar patterns in reconstructed sockeye salmon trends for Alaska provided strong evidence that large-scale climatic changes in Alaska have influenced past salmon populations (Finney et al. 2000, 2002). This variability occurred on a decadal timescale, similar to historical records, as well as over several centuries (Finney et al. 2002). The timing of inferred changes in salmon stocks generally coincided with changes in paleoclimatic records.

Figure 6 . (A) Paleolimnological evidence of sockeye salmon and lake ecosystem variation in Karluk Lake, Alaska, over the past ca. 300 years. The historical information for Karluk sockeye salmon includes commercial catch data beginning in 1882, and escapement counts from a weir on the Karluk River beginning in 1921. The paleolimnological data consist of į15N profiles (two replicated cores), concentrations of cladoceran Bosmina longirostris microfossils, relative abundances of dominant diatom taxa, and the relative abundance sum of all benthic diatoms. The data are divided into three different zones (from top to bottom): the fertilization period, the commercial fishing period and the pre-commercial fishing period. From Finney et al. (2000).

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Figure 6 ( continued). (B) Sedimentary į15N profiles and archeological phases from Kodiak Island, Alaska. Both Karluk and Akalura lakes are natural sockeye salmon nursery lakes, and the similar patterns for these lakes indicate similar histories of salmon abundance (higher į15N: more salmon). Frazer Lake, the non-salmon control lake, was barren until the 1950s. The importance of changes in sockeye salmon abundance to the indigenous people of Kodiak is reflected in archeological deposits and the delineation of cultural phases (which have been independently established). A shift towards greater abundance of fishing tools, greater housing densities and larger multi-room houses is evident during the Koniag phase. From Finney et al. (2002).

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Paleoecological data from freshwater sockeye salmon lakes further demonstrated the importance of marine-derived nutrients (derived from salmon carcasses) to freshwater ecosystems, which has implications for management and long-term sustainability because juvenile salmon are reared in nursery lakes before migrating to the ocean (Schmidt et al. 1998). In systems where salmon-derived nutrients (SDN) represent a significant proportion of annual nutrient loadings, there are strong positive correlations between adult salmon returns, lake nutrient levels, primary productivity, and zooplankton productivity (Finney et al. 2000; Gregory-Eaves et al. 2003; Sweetman and Finney 2003). Thus, the carrying capacity of nursery lakes for juvenile sockeye salmon is not constant, and there is potential for positive feedbacks between SDN and carrying capacity. However, there appeared to be natural gradients in the magnitude of SDN loading between lakes, and the productivity of some lakes was less dependent on SDN (Finney et al. 2000). Paleoproductivity in some sockeye salmon lakes was highly variable relative to nearby control lakes during the last ca. 2000 years, demonstrating the amplifying effects of climate change through biological nutrient transport vectors (Finney et al. 2000; Gregory-Eaves et al. 2003). Central Canadian Subarctic Early to mid-Holocene paleolimnological records Paleolimnological studies that extend back to pre-Holocene sediments are not available in the central Canadian Subarctic, as this region is believed to have been completely covered by the Laurentide Ice Sheet until ca. 8000 to 7000 yr BP (Fulton and Andrews 1987). Many terrestrial-based paleoecological studies (e.g., Nichols 1976; Kay 1979; Moser and MacDonald 1990; MacDonald and Gajewski 1992) have inferred a climate warmer than present for the period ca. 6000 to 5000 yr BP. Pollen, peat and paleopodzol evidence have indicated an increase in vegetation density and a northward displacement of treeline during the mid-Holocene, and a subsequent cooling and treeline retreat between ca. 3500 to 3000 yr BP. The few paleolimnological studies where complete postglacial sedimentary histories of lakes have been retrieved (e.g., Queen’s Lake, 64º07’N, 110º34’W; Toronto Lake, 63º43’N, 109º09’W; TK-20, 64º09’N, 107º49’W) have basal sediments that have been radiocarbon dated to between 8500 and 7500 yr BP. The close agreement in the timing of the early developmental stages of these lakes is likely due to their geographic proximity (all are within ca. 150 km of each other; Figure 1). As discussed below, diatom analyses (MacDonald et al. 1993; Pienitz et al. 1999; Rühland 2001) recorded substantial changes over the postglacial periods in these lakes, indicating that significant limnological shifts have occurred as a result of climate change. Diatom, pollen, isotopic, and geochemical analyses from Queen’s and Toronto lakes provided strong evidence for a rapid northward expansion of the forest-tundra ecozone between ca. 5000 and 4000 yr BP, when Picea mariana pollen, lake productivity indicators, and diatom-inferred lake water DOC increased sharply, and G18O and į13C analyses inferred a wetter climate, a more open hydrology, and a more productive environment (MacDonald et al. 1993; Wolfe et al. 1996; Pienitz et al. 1999; Gajewski and MacDonald, this volume). Application of bio-optical models (i.e., models that used diatom-inferred DOC concentrations to reconstruct past ultraviolet radiation

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penetration) found that, corresponding to an increase in diatom-inferred DOC in Queen’s Lake as coniferous trees moved northward into the catchment, there was a decrease by two orders of magnitude in biologically effective ultraviolet (UV) radiation exposure (Pienitz and Vincent 2000; Douglas et al., this volume). After ca. 3500 to 3000 yr BP, a cooler climate was marked by a rapid reversion back to tundra conditions, indicated by a decrease in Picea pollen, a return to diatom assemblages that existed prior to the treeline advance, and a decrease in diatom-inferred DOC. In contrast, changes in the abundances of pollen and diatoms of nearby Lake TK-20 were less pronounced during the mid-Holocene warm period, but were nevertheless approximately synchronous with changes reported in Queen’s and Toronto lakes (Rühland 2001; Huang et al. 2004). Pollen percentages suggest that the catchment of Lake TK-20 was likely occupied by Picea mariana and Alnus during regional treeline advance ca. 6300 yr BP, but not by Pinus (Huang et al. 2004). Coincident with the Picea pollen increases, a more diverse and complex diatom community had developed (ca. 6300 to 3300 yr BP) that now included planktonic taxa, likely triggered by a moister and warmer climate. Unlike nearby lakes, the TK-20 diatom profile did not record a rapid reversion back to pre-forest-tundra conditions during the onset of Neoglacial cooling (ca. 3300 to 3000 yr BP), but rather a different assemblage had developed that was unlike any of the previous communities recorded in the lake’s history. Non-climate-based early to mid-Holocene studies Campeau et al. (2000) applied a diatom-based inference model developed for reconstructions of water depth (Campeau et al. 1999) to five sediment cores retrieved from the Beaufort Sea coast and the Tuktoyaktuk Peninsula (69º54’N, 131º21’W) (Figure 1). They found that diatom assemblage changes were related to the transgression of the Beaufort Sea at specific periods throughout the late Holocene. Using this transfer function, they were able to determine when a lagoon-type environment transgressed over a thermokarst lake, and therefore provided a detailed sea-level history for the past ca. 2000 years. This quantitative approach represented a significant improvement over traditional sea-level reconstruction methods, which were previously limited to qualitative estimates of paleo-depth fluctuations. Late Holocene climate warming Recent syntheses of available instrumental and paleoclimatic proxy sources indicate that the average surface air temperature has increased by 0.3qC per decade in the 20th century in the North American Arctic and the boreal forest zone (Chapin et al. 2000; Keyser et al. 2000), representing the highest temperature increases recorded in the past few centuries (Overpeck et al. 1997; Mann et al. 1999; Briffa 2000; Keyser et al. 2000; Serreze et al. 2000). Increases in annual temperatures are expected to continue, perhaps by as much as 4qC in central subarctic Canada, over the next century (Schlesinger and Mitchell 1987; Foley et al. 1994; Houghton et al. 1996, 2001). High-resolution chronologies provided by varved lake sediment records (Anderson et al. 1996; Lamoureux and Bradley 1996; Gajewski et al. 1997; Hughen et al. 2000) and calibrated tree-ring studies (MacDonald et al. 1998; D’Arrigo et al. 1999; Briffa 2000)

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have provided evidence that the climate of Canada’s Arctic over the last few centuries has been both the coldest and the warmest over the last millennium (Bradley 2000). Although there are distinct spatial and temporal variations in the data (Bradley and Jones 1993; Barlow 2001; Jones and Briffa 2001; Ogilvie and Jónsson 2001), the colder period of this record is often referred to as the “Little Ice Age” (LIA) with the coldest conditions in North America of the last ca. 600 years occurring in the early 19th century (Bradley and Jones 1993; Jones and Briffa 2001; Ogilvie and Jónsson 2001). In subarctic North America, dendroclimatological and historical data provide evidence for colder conditions prevailing in the early 1700s and the early to mid-1800s, followed by a pronounced warming since ca. 1850 (Bradley and Jones 1993; MacDonald et al. 1994, 1998; Szeicz and MacDonald 1996; D’Arrigo et al. 1999; Briffa 2000). MacDonald et al. (1998) concluded from historical and dendrochronological records that the treeline region in central Canada (Kazan River region) is responsive to recent temperature variations associated with the end of the LIA. Their findings generally concur with other tree-ring studies in both eastern and western subarctic Canada (Lavoie and Payette 1994, 1996; MacDonald et al. 1994; Payette and Lavoie 1994), tracking an increase in the recruitment of trees at the northern treeline and large structural changes in forests after 1880 AD. However, unlike the rapid and pronounced estimates of vegetation shifts in the early Holocene (Moser and MacDonald 1990; MacDonald et al. 1993; MacDonald 1995), there was no evidence of a significant treeline movement during this recent period of warming. The resistance of trees to northward migration during this most recent warming period suggests that changes in the limits of spruce range and the increase in stand density may take centuries in the central Canadian Subarctic (MacDonald et al. 1998). As well, modeling of future subarctic vegetation responses to continued climatic warming showed that significantly warmer summer temperatures would result in the expansion of forest into treeless tundra regions, although there would be a lag between temperature increase and migration (Chapin et al. 2000; Kittel et al. 2000; Rupp et al. 2000, 2001). Much like the northward expansion of treeline in the mid-Holocene, future treeline migrations may affect subarctic lake characteristics in similar ways, including an increase in DOC and nutrients, a decrease in wind-induced turbulence, and ultimately changes in the aquatic biota. Paleolimnological studies focusing on recent (post-19th century) environmental changes in arctic regions have shown that this was, and is, a period of unprecedented ecological changes (e.g., Douglas et al. 1994). There is growing evidence from numerous arctic and subarctic lake studies that longer ice-free periods, as a result of 19th century warming, has noticeably affected lake water characteristics, habitat availability and, ultimately, aquatic biota (e.g., Smol 1988). However, in the central Canadian Subarctic, only a few paleolimnological studies with sufficient sampling resolution and an established chronology specific to this time period (i.e., 210Pb dating or varved sediments) have been undertaken (e.g., Stoermer et al. 1990; Rühland 2001; Moser et al. 2002; Rühland et al. 2003c). Changes in fossil diatoms from two tundra lakes (Slipper Lake at 64q35’N, 110q50’W and TK-20 at 64º09’N, 107º49’W) revealed substantial limnological shifts that occurred over the last ca. 150 years, following relatively stable assemblage composition during the last ca. 5000 years (Rühland 2001). Planktonic Cyclotella taxa (primarily C. stelligera and C. pseudostelligera) increased in the 19th century sediments of both lakes,

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concurrent with decreases in benthic Fragilaria species (F. pinnata, F. brevistriata, F. construens, F. constr uens var. venter) (Figure 7A). This recent (19th century) shift in diatom assemblages from primarily benthic to planktonic forms was unprecedented in the history of these two lakes (ca. 8000 to 5000 years), and approximately coincided with the mid-19th century warming period. To determine whether these recent environmental changes were occurring over a wide geographic gradient, Rühland et al. (2003c) examined differences in diatom assemblage composition between modern and pre-industrial sediments in a regional study of 50 lakes across the central Canadian treeline. Using this “top-bottom” approach, they found that the present-day aquatic habitats were markedly different from those during preindustrial times (pre-1850 AD). Similar to more detailed downcore analyses discussed previously, the most pronounced change in these 50 lakes was a shift to higher percentages of planktonic diatoms (mostly Cyclotella species) in the modern sediments, compared to the pre-industrial assemblages where small benthic Fragilaria taxa were more common (Figures 7B and 7C). In Wood Buffalo National Park to the south of Great Slave Lake, Moser et al. (2002) examined a ca. 200-year diatom record from subarctic Rainbow Lake A (RLA) (59q48’N, 112q10’W). As this lake contained varved sediments, it provided a rare opportunity for a high-resolution study of the timing of recent limnological changes. Based on changes in diatom assemblages, biogenic silica concentrations, and trends in diatom-inferred total phosphorus (TP), they concluded that RLA had become increasingly eutrophic beginning ca. 1830 AD, coincident with a period of warmer summer temperatures in the central Canadian Subarctic. Diatom assemblage composition underwent an abrupt change ca. 1830 to 1840 AD marked by increases in eutrophic taxa (e.g., Stephanodiscus parvus, S. mi nutulus, and Asterionella formosa ). Moser et al. (2002) explored several scenarios that could explain this 19th century eutrophication. However, human activities in the catchment, increased incidence of local fires, wetter conditions, atmospheric transport of pollutants, and nutrient addition from pollen grains could not satisfactorily explain the timing of nutrient enrichment. Rather, their results suggested that increased epilimnetic phosphorus concentrations were derived from internal loading of nutrients. Warmer summer temperatures starting ca. 1830 likely weakened meromictic stability and strengthened thermal stratification in RLA, resulting in enhanced internal cycling of phosphorus. Although most North American subarctic lakes (and those discussed thus far) are relatively small in size, the central Canadian Subarctic also contains some of the world’s largest and deepest lakes. For example, Great Slave Lake in the Canadian Northwest Territories is the sixth deepest lake in the world and tenth largest lake by volume (Herdendorf 1982). A diatom-based study of recent changes in the sediments of McLeod Bay (62º50’N, 109º46’W) reported that the most substantial diatom changes over the last ca. 5000 years occurred within the past ca. 200 years (Stoermer et al. 1990). Based on their previous paleolimnological work on the Laurentian Great Lakes (Schelske et al. 1983; Stoermer et al. 1985a,b, 1987; Wolin et al. 1988), Stoermer et al. (1990) concluded that the Great Slave Lake diatom changes were likely indicative of an increase in nutrient supply through atmospheric transport from remote sources. However, they acknowledged that McLeod Bay had remained ultra-oligotrophic and that other causative mechanisms may be working in concert with this hypothesis.

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Figure 7. (A) Diatom profile of the three most common and abundant diatom groups from Slipper Lake, Northwest Territories, in the central Canadian subarctic tundra. Dates at top of profile are 210 Pb (years AD) and dates on lower profile are in 14C calibrated age BC. The relationship between pre-industrial and modern relative abundances of (B) Cyclotella stelligera complex (C. stelligera, C. pseudostelligera ), and (C) small, benthic Fragilaria species (F. brevistriata , F. construens, F. construens var. venter, F. pinnata) from a 50-lake regional analysis in the central Canadian Subarctic. From Rühland et al. (2003c).

If we revisit the Stoermer et al. (1990) study today and compare it to more recent paleolimnological studies from the central Canadian Subarctic and to studies from other circumpolar regions, some parallel trends become evident. For example, similar to Slipper Lake, TK-20, and the 50-lake top-bottom study discussed earlier, the McLeod Bay diatom assemblages recorded a distinct increase in the relative abundances of planktonic diatoms (mostly Cyclotella stelligera ) with a concurrent decrease in small, benthic diatoms (mostly Fragilaria and Achnanthes species) over its recent history (Figure 8). The present-day ultra-oligotrophic nature of McLeod Bay, together with the shift to a more planktonic diatom assemblage, is consistent with the climatic warming scenario in terms of changes in the length of the ice-free period, the length of the growing season, enhanced thermal stratification patterns, and/or overall increased

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primary production (e.g., Smol 1988; Rühland et al. 2003c; Korhola and Weckström, this volume). Paleolimnological studies from the central Canadian Subarctic have documented considerable evidence for climate-related shifts, and a conspicuous lack of evidence for Holocene changes induced by non-climatic mechanisms, perhaps more so than in other parts of the North American Subarctic. Given the characteristics of this geographic region, such as negligible marine influences, a relatively uniform topography (no alpine zones) combined with the strong latitudinal vegetation/climatic gradient, it is not surprising that climate has played such a prominent role in influencing lake ecosystem changes in this region.

Figure 8. Relative abundances versus core depth of (A) benthic diatoms, (B) planktonic diatoms, (C) Cyclotella stelligera, (D) Cyclotella species, (E) benthic Fragilaria species. The date of ca. 1750 AD is based on uncertain age estimates and is therefore an approximation. From Stoermer et al. (1990).

Québec and Labrador Early to mid-Holocene paleolimnological records Similar to the central Canadian Subarctic, paleolimnological studies extending beyond the Holocene are not available for the eastern Canadian Subarctic, as the Laurentide Ice Sheet covered the region until ca. 8000 yr BP (Clark et al. 2000; Jansson 2003). Pollen evidence from northern Québec and Labrador indicates that Holocene warming was not as strongly expressed in these regions as it was in central and western Canada

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(Lavoie and Payette 1996). For example, in northwestern Québec, only moderate peaks in Picea pollen abundances were recorded ca. 3700 and 2900 cal. yr BP (Gajewski and Garalla 1992; Gajewski et al. 1996; Ponader et al. 2002), suggesting only modest changes in the position or stand density of treeline (Payette and Lavoie 1994). Paleolimnological studies similarly track only minor changes in limnic assemblages throughout the Holocene, with only subtle shifts recorded during the proposed Neoglacial period (between ca. 3000 and 1000 yr BP). The few sediment profiles that have recorded substantial Holocene changes are lakes that have undergone a transition from a marine to a freshwater phase as a result of postglacial isostatic adjustment. However, once the lacustrine phases were established in these lakes, biotic assemblages remained relatively unchanged until the present (e.g., Pienitz et al. 1991; SaulnierTalbot and Pienitz 2001). The first diatom-based paleolimnological study of subarctic Labrador was undertaken by Kingston (1984) who examined postglacial profiles from Leech Lake (53º10’N, 58º27’W) and an adjacent fen. He found that, unlike other parts of the North American Subarctic, the diatom assemblage composition over the last ca. 9000 years from both the lake and the adjacent fen had remained remarkably stable. Although the sampling resolution of this pioneering study was quite coarse (ca. every 15 cm), it did provide the first indication that parts of the eastern Arctic may have experienced relatively little climatic change over the Holocene. Changes in diatom and pollen assemblages from Square Lake (58º38’N, 63º36’W) in northern Labrador indicated that soil development began ca. 8000 to 7500 yr BP (Clark et al. 1989). Based on changes in diatom concentrations and assemblages, they found that the Holocene Thermal Maximum (HTM) occurred in this region ca. 6500 yr BP. A subsequent decrease in diatom concentration was associated with a cooling trend following the Hypsithermal. However, from ca. 6500 yr BP to the present, diatom assemblage composition remained relatively unchanged. Diatom-inferred changes in dissolved organic carbon (DOC) and loss-on-ignition (LOI) trends (Ponader et al. 2002), together with abundances of chironomid-inferred temperature (K. Swadling, pers. comm.), from a lake in northwestern Québec (Lake Karinbou, 57q44’N, 76q09’W), similarly recorded little change in limnological conditions over the last ca. 3000 years. Picea pollen suggested that spruce trees had reached their maximum local abundance at ca. 2900 cal. yr BP (Ponader et al. 2002). Although not pronounced, the most notable change in diatom assemblages over the last three millennia occurred ca. 1300 cal. yr BP, coinciding with a decrease in fire activity and/or increased periglacial activity in the catchment. Several paleolimnological studies of Lake Kachishayoot (55°20’N; 77°37’W), northwestern Québec have been undertaken to better understand the potential impact of future warming in the eastern Canadian Subarctic (Saulnier-Talbot and Pienitz 2001; Miousse et al. 2003; Saulnier-Talbot et al. 2003). By applying existing diatom-based inference models from the region (Fallu and Pienitz 1999; Fallu et al. 2002), SaulnierTalbot and Pienitz (2001) reconstructed lake water alkalinity, DOC, and water colour. Isolation of this lake from Hudson Bay occurred ca. 4600 yr BP (5400 cal. yr BP), at which time there was an abrupt shift from marine to freshwater diatom taxa. As well, diatom-inferred alkalinity was high, whereas DOC and colour inferences were low during this transitional phase. Immediately following this transition, there was a rapid decrease in inferred alkalinity during the initial lacustrine phase. Similar to other

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subarctic paleolimnological studies from eastern Canada (Pienitz et al. 1999), the initial establishment of spruce in the lake’s catchment ca. 2800 cal. yr BP (ca. 3700 yr BP) (Miousse et al. 2003) resulted in a gradual decrease in diatom-inferred alkalinity, concomitant with an increase in DOC and water colour. DOC and colour reached modern values in Lake Kachishayoot ca. 1000 cal. yr BP and remained relatively stable to the present. Chironomid-based temperature reconstructions showed no major changes during this lacustrine phase (K. Swadling, pers. comm.).

Figure 9. Changes in the catchment vegetation and limnological properties of Lake Kachishayoot (subarctic Québec) over the 5000 years subsequent to deglaciation. Concomitant with the gradual development of organic soils (a) and the invasion and establishment of spruce (Picea) (e) there was a rise in diatom-inferred DOC (b) and lake water colour (c), and a sharp decrease in underwater UV radiation (depth of paleo-UVR (h) and biological transparency weighted for UVphotoinhibition, T*PI (i), or for UV-photodamage of DNA, T*DNA (j). Reproduced from SaulnierTalbot et al. (2003).

The first paleo-optical study in northern Québec was also carried out on Lake Kachishayoot by Saulnier-Talbot et al. (2003). Their multi-proxy investigation revealed abrupt increases in diatom-inferred DOC concentrations and water colour that coincided with the retreat of postglacial marine waters and the arrival of spruce trees within the landscape and catchment of the study site (Figure 9). Their investigation also revealed large changes in the underwater irradiance environment over the course of the postglacial period, from extremely high UV exposure following the initial formation of

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the lake and its isolation from the sea, to an order-of-magnitude lower exposure associated with the development of spruce forests in the catchment. Furthermore, the use of additional macrofossil markers revealed that UVR penetration remained low even following forest retreat due to the development of alternate DOC sources in the catchment such as Sphagnum wetlands (Figure 9f). Similar to paleo-optical studies in the central Canadian Subarctic (Pienitz and Vincent 2000), these results suggest that the aquatic ecosystem had undergone DOC/UV fluctuations in the mid-Holocene and will likely do so again under future warming scenarios. Using multiple proxy indicators, Fallu (2003) examined the Holocene environmental history of two lakes in subarctic Québec (Lake K2, 58º44’N, 65º56’W and Lake Oksana, 54º49’N, 66º50’W). In Lake K2, changes in the relative abundances of diatoms, chironomids and pollen recorded only subtle changes throughout the last ca. 6700 cal. yr BP. Not surprisingly, diatom-based reconstructions of DOC, alkalinity and water colour inferred only minor changes throughout the sequence (Fallu et al., in press). Although muted, slightly increased abundances of Picea pollen suggested that conditions had become warm enough to allow for the growth of spruce (krummholz) in the valleys surrounding Lake K2. The installation of krummholz at this time is a likely explanation for the subtle decrease in reconstructed alkalinity and minor increases in reconstructed DOC and water colour between ca. 3000 and 2000 cal. yr BP. Chironomid-based reconstruction of summer surface water temperature also suggested that temperatures had remained relatively stable until ca. 1500 cal. yr BP, at which time conditions cooled to present-day values (Fallu et al., in press). Similar to Lake K2, a lack of a pronounced change during the Holocene was also clearly evident in Lake Oksana, where chironomid-inferred summer surface water temperatures and changes in diatom assemblages showed no notable shifts throughout the ca. 6300-year sediment sequence (Fallu 2003). Non-climate-based studies In addition to the climate-based paleolimnological studies reviewed thus far, several studies in the eastern Canadian Subarctic have examined the isolation of coastal lakes from the marine environment (Pienitz et al. 1991; Cameron et al. 1998; Saulnier-Talbot and Pienitz 2001; Miousse et al. 2003). Given the distinct differences between marine and freshwater environments, diatoms are powerful indicators of the retreat of postglacial seas. In northern Québec, shifts in diatom assemblages suggested that lake isolation commenced ca. 4850 yr BP in Lake Tasirlaq (58º14’N, 68º27’W) and ca. 4300 yr BP in Lake Hendry (58º07’N, 68º14’W) (Pienitz et al. 1991). Once the lacustrine phase had been established, diatom assemblages remained stable to the present. Similarly, diatom assemblage changes, together with other types of paleoindicators, in Lake Kachishayoot (55°20’N; 77°37’W) showed that the lake’s isolation phase occurred between ca. 4600 and 4000 yr BP (Saulnier-Talbot and Pienitz 2001; Miousse et al. 2003). Similar to Tasirlaq and Hendry lakes, little change was recorded in the diatom assemblages once the freshwater phase had been established. Diatom changes in the sediments of Isiurqutuuk Lake (60º49’N, 77º58’W) suggested that this lake became isolated from the sea at a much later date (ca. 2000 yr BP) (Cameron et al. 1998), likely due to its proximity to the coast.

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Late Holocene Climate models have revealed that, over the last few centuries, climate has cooled in eastern subarctic Canada (Chapman and Walsh 1993; Serreze et al. 2000). Based on tree-ring width chronologies in Labrador, D’Arrigo et al. (2003) inferred increasing warmth from the late 1800s through to the mid-20th century and, in agreement with observational data, cooler temperatures over the last several decades. Consistent with these instrumental and dendrochronological data, paleolimnological studies in eastern subarctic Canada have recorded no evidence of recent warming (Laing et al. 2002; Ponader et al. 2002; Paterson et al. 2003; Saulnier-Talbot et al. 2003; Fallu et al., in press). These studies from northern Québec and Labrador demonstrate that chironomidbased surface water temperature reconstructions from four lakes, together with diatom records from 16 lakes, did not show a major warming trend over the last few centuries (Pienitz et al. 2003). This is in direct contrast with paleolimnological studies that infer recent warming trends in western and central subarctic Canada (e.g., Rühland et al. 2003c), the Canadian Arctic Archipelago (e.g., Douglas et al. 1994; Overpeck et al. 1997; Wolfe and Smith, this volume), Alaska (e.g., Anderson et al. 2001; Hu et al. 2001b), as well as in other circumpolar regions mentioned in this volume. Collectively, the data from subarctic Québec and Labrador suggest that, unlike other subarctic regions, climate has remained relatively stable over the last few decades. Laing et al. (2002) examined recent changes in diatom assemblages from seven unnamed lakes in northeastern Québec (58q17’N, 65q40’W; 58q19’N, 65q38’W; 57q34’N, 65q31’W; 57q28’N, 65q21’W; 57q36’N, 65q24’W; 58q30’N, 65q45’W; 58q43’N, 65q57’W) to assess the impact of the world’s largest caribou herd in the Rivière George region on lake water quality. They found no evidence for marked changes in lake conditions as a result of the fluctuating herd populations. Mathieu (2003) examined changes in macrofossils and pollen from many of the same lakes and also found that assemblages had remained unchanged over the last ca. 200 years, suggesting not only a negligible impact by the caribou herd, but also no evidence for recent warming in this region. Paterson et al. (2003) reached similar conclusions from their multiproxy analyses of indicators from a PCB-contaminated lake (Lake Saglek 2, 58q23’N, 62º35’W) in northeastern Labrador. Diatom and chrysophyte assemblages over the last ca. 150 years revealed little change, despite the elevated concentrations of PCBs (Figure 10). Further, this lack of change is consistent with instrumental and paleolimnological studies from the eastern Subarctic that have shown minimal change in climate over recent decades. These findings further support the position that the pronounced changes observed in arctic lakes to the west of northern Québec and Labrador are not a function of longdistance transport of persistent pollutants, but that recent warming is the underlying mechanism. The general absence of Holocene climatic changes in northern Québec and Labrador is not yet fully understood, however several important observations have been made. For example, it is believed that northern Québec is strongly influenced by oceanographic factors due to its peninsular shape (Pienitz et al. 2003). Cold ocean currents from Hudson Strait and from Labrador encircle the peninsula and, together with cooling of the prevailing westerly winds by sea ice on Hudson Bay, have dampened the warming trend that has been recorded in paleolimnological studies from other parts of subarctic North America. Regional variations of this trend are subtle and

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likely due to the proximity of most study sites to the coast where cooling is particularly acute. The seasonal melt of sea ice on Hudson Bay follows a similar pattern in that the ice breaks up earlier on the western (Keewatin) side and the resulting floating ice is pushed eastward towards northern Québec on the eastern shore of Hudson Bay, strengthening the cooling effect in these regions.

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Late Holocene non-climate-based studies High-resolution diatom analyses of Isiurqutuuk Lake (60º49’N, 77º58’W), northern Québec, were used to examine the probable causes for the disappearance of fish populations from this lake through a massive fish kill reported by the local Inuit population in the early 1930s (Cameron et al. 1998). Their results showed that there was a substantial increase in lake acidity commencing ca. 1000 yr BP which continued to increase to the present pH of 4.5. They determined that the acidification of this lake was a natural process (i.e., not a result of human-induced acid deposition) related to the glacio-isostatic uplift of the region and subsequent leaching of the sulfide-rich graphitic shale present in the catchment. The recent accentuation of this acidification process was caused by the depletion of marine shells, deposited when the lake basin was influenced by the marine waters of the postglacial Tyrell Sea, and which had served as buffer. The decrease in the lake’s buffering capacity led to rapid acidification and the elimination of the fish population through the failure of fish eggs to survive low pH conditions. Challenges and future directions Accurate dating of sediments in northern regions remains a critical issue in assessing temporal patterns and gradients, and relationships to large-scale climatic controls (Wolfe et al., this volume). Dating of high latitude lakes is often problematic, and in some ways this challenge increases as an inverse function along the climate gradient, both spatially across the region and temporally within a core. Many lakes from the northern part of the Subarctic are similar to high arctic lakes in having slow sedimentation rates and low carbon and 210Pb fluxes (e.g., Laing et al. 2002). Southern subarctic lakes tend to be more productive and generally have faster sedimentation rates and higher carbon and 210Pb fluxes, and thus may potentially provide sediment profiles with higher temporal resolution (e.g., Naidu et al. 1999). However, the glacial and deglacial sediments of these lakes were generally deposited under arctic-like conditions, thereby increasing some other types of dating challenges for these periods (Wolfe et al., this volume). The advent of AMS radiocarbon techniques on specific carbon-bearing components provided a significant improvement over conventional dating, and should ultimately lead to refining our understanding of late Quaternary changes in this region (e.g., Abbott et al. 2000). Currently, the lack of accurate radiocarbon chronologies within the North American Subarctic limits precise inter- and intra-regional comparisons. Improvements could be achieved by revisiting key sites that were conventionally dated, increasing the density of sites, and by taking multiple and higher volume sediment cores to increase the likelihood of finding suitable macrofossils. The use of new datable materials with AMS radiocarbon techniques, such as chironomid head capsules, has also been proposed as an alternative when terrestrial macrofossils are absent (Fallu et al. 2004). Accurate dating of recent sediments for high-resolution studies in arctic regions is another substantial challenge. Varved lake sediments, although rare in this region, have not been well exploited. The limitations of 210Pb-dating for arctic lake sediments stem from the low rates of sediment accumulation coupled with the commonly low fluxes of these nuclides to lake sediments (Hermanson 1990). As a result, important paleoindicator changes over the last ca. 150 years are usually contained within the upper five

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to ten centimetres of the sediment core, or even less, resulting in a relatively coarse dating resolution. A finer sediment sampling resolution (e.g., 0.5 cm to 0.25 cm) may help increase the number of dated samples, but results in new challenges related to sufficient material to acquire reliable 210Pb counts. The surface sediments of most lake cores (including those from temperate regions) characteristically have a very high water content, as materials are not as consolidated as in lower levels. This often results in the dilution of the 210Pb signal, particularly when the sampling resolution is increased. A possible compromise to this challenge is to retrieve replicate cores and combine the top sediment interval from each core to obtain sufficient surface sediment material for dating. Although there are clearly difficulties in dating arctic sediments, even coarsely dated material may still provide important data on overall directions of environmental change. Considerable advances have recently been made in understanding variability in both limnological changes and the distributions of paleoindicators along environmental gradients in many parts of the North American Subarctic. Process-based and ecological studies are certainly still needed to improve our understanding of these lake systems and species distributions, but a basic framework is now available. With some of this knowledge now in hand, future sites for calibration studies can be selected, so as to increase the accuracy of inference models for key variables. Thus, a varied and powerful array of tools is now available for paleolimnologists to exploit, but this “toolkit” has not yet been well utilized in this region. Currently, spatial networks of key proxies (other than pollen), analysed and interpreted in a consistent manner, are sparse and limit robust evaluations of spatial and temporal gradients in past environments. In the future, study sites should be selected based more on criteria such as their limnological characteristics (and not simply on geographic location, which, thus far, has been a primary factor in lake choice) to increase the likelihood of attaining the most informative data. In addition, we have now sufficient information to strategically select sites with regard to localities sensitive to changes in climatic patterns. As discussed in this chapter, there appears to be significant regional variability in climate history across this region, which highlights the need to develop a spatial network of sites. This spatial perspective is necessary to determine past synoptic climate patterns, assess their relationship to current patterns, and determine controlling factors and teleconnections. Major questions to be more fully addressed by future paleolimnological studies include: (1) What are the magnitudes, patterns, and causes of past warm and/or cool episodes in this region?; (2) How did these episodes influence subarctic ecosystems and biogeochemical processes, and subsequently influence the global climate system?; (3) How does the instrumental-based record of climate for this region compare with the long-term perspective?; and (4) How might subarctic ecosystems respond to future climatic and environmental changes? The framework and methods are in place to address these important and complex questions for the North American Subarctic.

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Summary Results from the various limnological and paleolimnological studies from the North American Subarctic show a clear indication of the sensitivity of these lake ecosystems to environmental factors. Thus, paleolimnological techniques can be used to assess the natural variability of these lake systems, and their responses to past climatic and environmental changes. Independent indicators of past climate, as well as knowledge of factors such as fire history and watershed processes, are needed to draw robust conclusions, and several such examples are presented in this review. For example, in the central Canadian Subarctic, marked changes in lake plankton communities and productivity, and water quality variables such as DOC and nutrients, resulted from midHolocene climatic change to warmer and wetter conditions. Some of these limnological changes can be directly attributed to climatic shifts, but interactions of lake systems with their watersheds were also clearly important. However, other factors may influence lake ecosystems. For example, in southern coastal lakes in Alaska, diatom communities, zooplankton assemblages, lake productivity, and nutrient levels have changed dramatically (more than a factor of two) over decadal timescales due to changes in the number of returning sockeye salmon (e.g., Finney et al. 2000). Subarctic lakes are dynamic systems, and results such as these indicate that major disturbances/stresses are likely to occur in response to rapid, human-induced environmental changes in the future. Sufficient paleolimnological data for meaningful regional and global comparisons commence in the last deglacial period in eastern Beringia. Records for the deglacial to early Holocene period show complex influences from both hemispheric and regional processes. The major climatic steps during this period began with significant climatic amelioration ca. 12,000 yr BP, followed by a period of major fluctuations, including changes during the Younger Dryas chronozone, and ended in a dry and probably warm early Holocene. The relative magnitude and perhaps timing of these changes may vary within eastern Beringia, due to different regional expressions of large-scale factors, including solar insolation (Bartlein et al. 1992) and ocean/atmospheric circulation variability originating in the North Altantic, and regional factors related to sea-level rise and the response of atmospheric circulation to the retreating Cordilleran and Laurentide Ice sheets. Broader cross-regional comparisons for subarctic North America are possible from about 8000 yr BP for the central Canadian Subarctic, and about 7000 yr BP for the eastern Canadian Subarctic, due to the patterns of deglaciation in North America. The paleoenvironmental data for eastern Beringia suggested that the Holocene thermal optimum occurred between about 10,000 to 8000 yr BP, and evidence for aridity is strong for this period. In contrast, paleoenvironmental studies from the central Canadian Subarctic provide evidence that the mid-Holocene (ca. 6000 to 3000 yr BP) was the warmest part of the Holocene (warmer than today) and probably wetter than present-day conditions. The timing of climate amelioration in the central Canadian Subarctic does not coincide with eastern Beringia, where Holocene warming is attributed to the Milankovitch thermal maximum (Ritchie et al. 1983), or in eastern subarctic Canada, where evidence for a warm period is typically not found. Rather, it may be more accurate to say that in the eastern Subarctic, a slow and subtle cooling trend commenced ca. 1500 cal. yr BP.

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During the 20th century, temperature trends across this region were not uniform. Strong warming trends were observed for much of eastern Beringia and central subarctic Canada, whereas climate stability or even slight cooling were observed for eastern subarctic Canada. Given these contrasting patterns, it is perhaps not surprising that high-resolution records from these regions also suggest different trends over the late Holocene. Records from both eastern Beringia and central subarctic Canada suggest that recent conditions are warm relative to the past few hundred to thousand years, while records from eastern subarctic Canada suggest relatively little change. The North American Subarctic is a vast and diverse region. Although considerable progress in deciphering past environmental conditions has been achieved over the last 20 years, the list of pending questions still far outnumber any answers attained to date. Nonetheless, based on the foundation of the paleolimnological research completed to date, the stage is set to further advance our quantitative understanding of paleoenvironmental change in subarctic North America. Acknowledgements Research in our laboratories is supported primarily by the National Science Foundation (USA), the Natural Sciences and Engineering Research Council of Canada, and the Canadian Polar Continental Shelf Project. We thank our reviewers (Irene GregoryEaves and an anonymous reviewer), as well as W.M. Last and É. Saulnier-Talbot for improving the quality and clarity of the manuscript. References Abbott M.B., Finney B.P., Edwards M.E. and Kelts K.R. 2000. Lake-level reconstructions and paleohydrology of Birch Lake, central Alaska, based on seismic reflection profiles and core transects. Quat. Res. 53: 154-166. Ager T.A. 1975. Late Quaternary environmental history of the Tanana valley, Alaska. Institute of Polar Studies, Report no. 54, Ohio State University, Columbus, 117 pp. Ager T.A. 1983. Holocene vegetation history of Alaska. Late Quaternary environments of the United States. In: Wright Jr. H.E. (ed.), The Holocene, Vol. 2. University of Minnesota Press, Minneapolis, pp.128-140. Ager T.A. and Brubaker L.B. 1985. Quaternary palynology and vegetation history of Alaska. In: Bryant Jr. V.M. and Halloway R.G. (eds), Pollen Records of Late-Quaternary North American Sediments. American Association of Stratigraphic Palynologists, Dallas, pp. 353-384. Anderson L., Abbott M.B. and Finney B.P. 2001. Holocene paleoclimate from oxygen isotope ratios in lake sediments, central Brooks Range, Alaska. Quat. Res. 55: 313-321. Anderson N.J., Odgaard B.V., Segerström U. and Renberg I. 1996. Climate-lake interactions recorded in varved sediments from a Swedish boreal forest lake. Global Change Biology 2: 399-405. Anderson P.M. 1985. Late Quaternary vegetational change in the Kotzebue Sound area, northwestern Alaska. Quat. Res. 24: 307-321. Anderson P.M. 1988. Late Quaternary pollen records from the Kobuk and Noatak River drainages, northwestern Alaska. Quat. Res. 29: 263-276.

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Stuiver M., Reimer P.J. and Reimer R.W. 2000. Calib 4.3 (WWW program and documentation – http://www.calib.org.-). Sweetman J.N. and Finney B.P. 2003. Differential responses of zooplankton populations (Bosmina longirostris ) to fish predation and nutrient loading in an introduced and a natural sockeye salmon nursery lake on Kodiak Island, Alaska, USA. J. Paleolim. 30: 183-193. Sweetman J.N., Honnold S.G., Sagalkin N. and Finney B.P. Trophic interactions with zooplankton communities of Alaskan sockeye salmon nursery lakes: Implications for sockeye salmon production. Can. J. Fish. Aquat. Sci. (in press). Szeicz J.M. and MacDonald G.M. 1996. A 930-year ring-width chronology from moisturesensitive white spruce (Picea g lauca Moench) in northwestern Canada. The Holocene 6: 345-351. Szeicz J.M. and MacDonald G.M. 2001. Montane climate and vegetation dynamics in easternmost Beringia during the Late Quaternary. Quat. Sci. Rev. 20: 247-257. Taylor S.J. 1997. Taxonomy and ecological characterization of chrysophycean stomatocysts from lakes extending from the boreal forest to the arctic tundra, Canada. M.Sc. thesis, Queen’s University, Ontario, 181 pp. Timoney K.P., La Roi G.H. and Dale M.R.T. 1993. Subarctic forest-tundra vegetation gradients: the sigmoid wave hypothesis. J. Vegetation Sci. 4: 387-394. Vinebrooke R.D. and Leavitt P.R. 1998. Direct and interactive effects of allochthonous dissolved organic matter, inorganic nutrients, and ultraviolet radiation on an alpine littoral food web. Limnol. Oceanogr. 43: 1065-1081. Walker I.R. 2001. Midges: Chironomidae and related Diptera. In: Smol J.P., Birks H.J.B. and Last W.M. (eds), Tracking Environmental Change Using Lake Sediments, Volume 4: Zoological Indicators. Kluwer Academic Publishers, Dordrecht, The Netherlands, pp. 43-66. Walker I.R. and MacDonald G.M. 1995. Distributions of Chironomidae (Insecta: Diptera) and other freshwater midges with respect to treeline, Northwest Territories, Canada. Arct. Alp. Res. 27: 258-263. Walker I.R., Smol J.P., Engstrom D.R. and Birks H.J.B. 1991. An assessment of Chironomidae as quantitative indicators of past climatic change. Can. J. Fish. Aquat. Sci. 48: 975-987. Walker I.R., Wilson S.E. and Smol J.P. 1995. Chironomidae (Diptera): quantitative paleosalinity indicators for lakes of western Canada. Can. J. Fish. Aquat. Sci. 52: 950-960. Walker I.R., Levesque A.J., Cwynar L.C. and Lotter A.F. 1997. An expanded surface-water palaeotemperature inference model for use with fossil midges from eastern Canada. J. Paleolim. 18: 165-178. Walker I.R., Levesque A.J., Pienitz R. and Smol J.P. 2003. Freshwater midges of the Yukon and adjacent Northwest Territories: a new tool for reconstructing Beringian palaeoenvironments? J. N. Am. Benthol. Soc. 22: 323-337. Warner B.G. and Charman D.J. 1994. Holocene changes on a peatland in northwestern Ontario interpreted from testate amoebae (Protozoa) analysis. Boreas 23: 271-279. Wenner C.G. 1947. Pollen diagrams from Labrador. Geogr. Ann. 29: 137-373. Wilkinson A.N., Zeeb B.A. and Smol J.P. 2001. Atlas of Chrysophycean Cysts, Volume II. Kluwer Academic Publishers, Dordrecht, The Netherlands, 180 pp. Williams K.M., Short S.K., Andrews J.T., Jennings A.E., Mode W.N. and Syvitski P.M. 1995. The Eastern Canadian Arctic at ca. 6 ka BP: a time of transition. Géogr. phys. Quat. 49: 13-27. Williamson C.E., Olson O.G., Lott S.E., Walker N.D., Engstrom D.R. and Hargreaves B.R. 2001. Ultraviolet radiation and zooplankton community structure following deglaciation in Glacier Bay, Alaska. Ecology 82: 1748-1760. Wilson S.E., Walker I.R., Mott R.J. and Smol J.P. 1993. Climatic and limnological changes associated with the Younger Dryas in Atlantic Canada. Clim. Dynam. 8: 177-187. Wolin J.A., Stoermer E.F., Schelske C.L. and Conley D.J. 1988. Siliceous microfossil succession in recent Lake Huron sediments. Arch. Hydrobiologie 114: 175-198.

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Wolfe B.B., Edwards T.W.D., Aravena R. and MacDonald G.M. 1996. Rapid Holocene hydrologic change along boreal treeline revealed by į13C and į18O in organic lake sediments, Northwest Territories, Canada. J. Paleolim. 15: 171-181. Wright Jr. H.E., Kutzbach J.E., Webb III T., Ruddiman W.F., Street-Perrott F.A. and Bartlein P.J. (eds) 1993. Global Climates Since the Last Glacial Maximum. University of Minnesota Press, Minneapolis, 544 pp. Zellmer I.D. 1996. The impact of food quantity on UV-B tolerance and recovery from UV-B damage in Daphnia pulex. Hydrobiologia 319: 87-92.

11. HOLOCENE PALEOLIMNOLOGY OF GREENLAND AND THE NORTH ATLANTIC ISLANDS (NORTH OF 60ºN)

N. JOHN ANDERSON ([email protected]) DAVID B. RYVES ([email protected]) Department of Geography Loughborough University Loughborough, Leicestershire LE11 3TU United Kingdom MARIANNE GRAUERT ([email protected]) Department of Geography University of Copenhagen Øster Voldgade 10 DK-1350 Copenhagen K Denmark and SUZANNE McGOWAN ([email protected]) School of Geography The University of Nottingham University Park Nottingham, NG7 2RD United Kingdom

Key words: Diatoms, Pigments, Macrofossils, Chironomids, Isotopes, Greenland, Faeroe Islands, Svalbard, Iceland

Introduction The North Atlantic is becoming a major focus of paleoenvironmental research due to its central role in natural climatic variability (Bond et al. 1993, 1997; Viau et al. 2002). The land masses surrounding the North Atlantic are profoundly influenced by the ocean’s effect on climate. As a result, Greenland and the other islands in the area (north of 60ºN) are considered to be optimal sites to study the development of Holocene climate and the natural variability of phenomena such as the North Atlantic Oscillation (NAO). Greenland is home to the Greenland Ice Core Project (GRIP) and Greenland Ice Sheet Project Two (GISP2) ice core records, two of the northern hemisphere’s premier

319 R. Pienitz, M.S.V. Douglas and J.P. Smol (eds), 2004. Long-term Environmental Change in Arctic and Antarctic Lakes. Springer. Printed in the Netherlands.

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paleoclimate records (Greenland Ice Core Project Members 1993; Hammer et al. 1997; Dahl-Jensen et al. 1998) and it is of considerable interest to compare local and regional lake-based records with the ice core records which reflect hemispheric climate and high altitude climatic processes. However, as well as their role in addressing climate variability, sediment records from the northern North Atlantic islands are also used to answer questions regarding postglacial refugia and pathways of species immigration. The northern North Atlantic islands include Iceland, Svalbard, Bear Island (Bjørnøya), Jan Mayen and the Faeroes (Figure 1), although paleolimnological studies in the region are dominated by those from Greenland, which has been the subject of much research by Danish Quaternary scientists since the late 19th century. We include the Faeroes here, although technically subarctic, as a large part of the Faeroese landscape over 200-300 m above sea-level has been shaped by geomorphological processes typical of the Arctic, e.g., patterned ground and solifluction features (Humlum and Christiansen 1998b), and many of the lakes are at higher altitudes. If the absence of trees is accepted as a defining characteristic of the Arctic (Humlum and Christiansen 1998b), then this further qualifies the Faeroe Islands for inclusion, although the role of cultural impacts and deforestation on the present treeless state of both the Faeroe Islands and Iceland should not be overlooked (see below). The variety of landscapes and processes found on Greenland and the northern North Atlantic islands (hereafter “the North Atlantic islands”) in fact represents the full range of arctic environments. Despite its long history, paleolimnology on Greenland and the North Atlantic islands is not very well developed. Early lake-based paleoecological studies on Greenland, the Faeroes and Iceland from the 1950s were pollen-based and focused on vegetation history. There were no surface sediment calibration data sets available until recently, although diatom and chironomid data sets have now been constructed for western Greenland and Svalbard (Brodersen and Anderson 2002; Ryves et al. 2002; Jones and Birks 2004), and are in progress for the Faeroes (M. Grauert, unpublished) and Iceland (Langdon 2002; Karst-Riddoch 2003). The need for more regionally sensitive climate reconstructions, however, has generated considerable interest in lake sediment records (Battarbee 2000), especially those that are located along the western seaboard of the Atlantic, but also on Greenland, Iceland and other islands in the North Atlantic region. Lake sediment records have also been widely used to reconstruct the glacial history (especially the timing and sequence of deglaciation and event stratigraphic records) on Greenland, Svalbard and Iceland (e.g., Björck et al. 1994). Although of considerable climatic significance, especially regarding local Holocene landscape development, the ecological and limnological content of these papers is often minimal and they are not considered in this review. Similarly, only passing reference is made to studies concerned primarily with terrestrial vegetation history. This review focuses on Holocene lacustrine records of lake ontogeny, climate inferences and cultural disturbance.

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r No th

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Figure 1 . Location map of Greenland and northern North Atlantic islands. The Arctic Circle is shown as a dotted line. Main surface ocean currents: NAC - North Atlantic Current (warm); IC - Irminger Current (warm); EGC - East Greenland Current (cold); WGC - West Greenland Current (warm); LC - Labrador Current (cold). Paleolimnological records from Greenland mentioned in the text are numbered: 1 West Greenland area (includes lakes SS2: 67ºN, 50º58’W, Braya Sø: 67ºN, 51ºW, SFL4: 67º05’N, 50º17’W, SFL6: 67º05’N, 50º21’W, NAUJG1: 66º40’N, 51º58’W, Lake 31: 67º03’N, 50º27’W, Lille Saltsø: 66º59’N, 50º38’W, and Store Saltsø: 66º59’N, 50º36’W); 2 Godthåbsfjord; 3 Qipisarqo Lake; 4 Raffles Sø; 5 Basaltsø and Lake B1: 72º43’N, 22º28’W. The locations of the GRIP, GISP2 and Renland ice cores are also shown (Greenland). ZERO = Zackenberg Ecological Research Operations; P. Ch. Sund = Prins Christian Sund.

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Geographic and climatic setting The northern North Atlantic is an important area for the formation of dense, saline deep water (North Atlantic Deep Water; NADW), which has a major role in the thermohaline circulation. Formation of NADW is in turn the motor for poleward heat transfer and hence affects global climate (Cronin 1999; Ruddiman 2001). Ocean circulation in the North Atlantic north of 60ºN is dominated by five main currents (Figure 1). The warm north-east flowing North Atlantic Current continues into the Norwegian Sea and Arctic Ocean, and has a major branch, the Irminger Current, that flows south of Iceland and northward into the Denmark Strait between Iceland and East Greenland. The cold East Greenland Current flows southwards along the eastern seaboard of Greenland, and mixes with the Irminger Current to become the West Greenland Current. The West Greenland Current itself flows north-westwards into the Davis Strait, where it joins the southward flowing Labrador Current. Although the role of the North Atlantic Current in ameliorating the climate of north western Europe and Scandinavia is well-known, its role in global climate is now being re-assessed (Bond et al. 1997). Greenland covers 22 degrees of latitude (ca. 2700 km) (Figure 1) with a corresponding range in climate conditions, from subarctic in the maritime-influenced south to the extreme continental, high arctic conditions of Peary Land and Inglefield Land in the far north and north east. The Greenland ice sheet, which covers ca. 85% of the island and reaches over 3200 m in altitude, has a major influence on both local weather patterns as well as depression tracks in the North Atlantic. Much of Greenland north of ca. 67º N has continuous permafrost today and vegetation is mainly dwarf shrub tundra and fell-fields with sparse vegetation development at higher altitudes. Vegetation is very poorly developed above 80ºN. The climate of the south of the island is considerably milder and more maritime than either the east and west coasts further north: at Qaqortoq the mean annual temperature is above freezing (Table 1). Precipitation around Tasiilaq (Ammassalik) in the south east can reach 2500 mm yr-1 with large amounts of snow, falling to 750 mm yr-1 at Nuuk (64ºN), while at Kangerlussuaq (Søndre Strømfjord) it is only 150 mm yr-1, reflecting the rainshadow effect of the Sukkertoppen icecap on this area. The coastal zone of the western side of the island up to ca. 68ºN is generally more “maritime”, reflecting the influence of the West Greenland Current on sea-ice conditions, although winters are still quite severe (Table 1). The south western part (Disko to Kap Farvel; Figure 1) is characterized by dwarf shrub tundra (Salix, Betula nana, Ericaceae). Lying immediately south of the Arctic Circle, Iceland has a maritime climate that largely reflects the interplay between cold air masses from the north and warmer air from the south. This variability is clearly reflected in the NAO, so that in negative mode years the westerly depression track is pushed further south, with corresponding high pressure over Greenland. Depending on location on the island, mean summer temperatures range from 7-11ºC while mean winter temperatures range from 1-2ºC to 5-6ºC. Precipitation ranges from 800-1500 mm yr-1 in the lowland areas but as much as 4000 mm yr-1 over the glaciers. Jan Mayen, 600 km north of Iceland (ca. 71º05’N; 373 km2; Figure 1) is also a mountainous and actively volcanic island with a maximum altitude of almost 2300 m. Climate is maritime arctic, with frequent storms and persistent fog.

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The Faeroe Islands lie between 61º20’N and 62º24’N (ca. 6ºW; Figure 1) and have a very windy, temperate, maritime climate for localities close to sea-level (Table 1). Rainfall ranges from ca. 900 mm yr-1 on the westward islands to nearly 3000 mm yr-1 in the northeast. Seasonal temperature variations are relatively small due to their oceanic setting, greatly influenced by the North Atlantic Current. Lying some 350 km off the northern coast of Norway, the Svalbard archipelago (74º-82ºN; Figure 1) is indisputably arctic (mean annual air temperature at Longyearbyen is -6ºC; Table 1), although even here, the fading effects of the North Atlantic Current ameliorate the local climate compared to similar latitudes in arctic Canada. In terms of its landscape, extensive ice caps, glaciers and fell-field vegetation, Svalbard is similar to northern and eastern Greenland and the Canadian High Arctic, with the highest point over 1700 m. At the southern tip of the Svalbard archipelago, ca. 220 km south of Spitsbergen, Bear Island (Bjørnøya; ca. 74º30’N; Figure 1) has a similar climate, with altitudes in the southern part generally from 360 to 440 m, reaching a maximum of 563 m. Although the island is only about 20 km long and 15.5 km at its widest point, it contains an estimated 600 lakes. Table 1 . Major climatological parameters for selected locations on Greenland and the northern North Atlantic islands. See Figure 1 for locations. Source: Norwegian Meteorological Institute, Danish Meteorological Institute, Icelandic Meteorological Office. Data from 1961-1990 (a19 year record).

Location Greenland Ilulissat Kangerlussuaq Nuuk Qaqortoq Prins Christian Sund Tasiilaq Iceland Reykjavik Akureyri Svalbard Longyearbyen Faeroe Islands Tórshavn Vágar airport

Annual Temperature (ºC) Annual average Month precipitation (mm) Mean Max. Min. Warmest Coldest 266 149 752 858 2474 984

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-8.3 -10.7 -3.9 -2.9 -1.9 -5.2

10.3 16.3 9.9 11.1 9.8 10.4

-24.7 -26.7 -10.7 -9.2 -6.3 -12.3

799 490

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-8.3a

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-15.2

1284 1555

6.5 6.0

8.6 8.0

4.4 3.9

12.8 12.3

1.2 0.5

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Limnology of the North Atlantic islands Although a good understanding of contemporary lake functioning is vital to the interpretation of sediment records, the limnology of the North Atlantic islands is poorly understood. Despite early limnological surveys made by Hansen in the 1960s in Greenland (e.g., Hansen 1967), research derived from arctic North America formed the basis for understanding limnological functioning across the North Atlantic region. Assumptions about arctic limnology are now changing rapidly as more lakes are being investigated, highlighting the geographical and temporal variability of lake systems across a wide range of environments (e.g., on the Faeroes (Jeppesen et al. 2002) and Bear Island (Skjelkvåle et al. 2001)). Coordinated research efforts involving biological, physical and climatic interactions are being developed (e.g., the NORLAKE project within the Nordic Arctic Research Programme (NARP) 1999-2003: http://thule.oulu.fi/narp/index.htm) and interest in cultural impacts on arctic lakes, such as pollution studies, is growing (e.g., the Arctic Monitoring and Assessment Programme (AMAP); Lien et al. 1995; Skotvold et al. 1999). A number of arctic lakes have now been monitored continuously for a longer period, both within North America (e.g., Toolik Lake, Alaska) and in East Greenland at the Zackenberg Ecological Research Operation (ZERO; Figure 1) (Jeppesen et al. 1998; Christoffersen and Jeppesen 2000; Vadeboncoeur et al. 2003). The lakes of Greenland and the North Atlantic islands are incredibly diverse, ranging from dilute, species-depauperate, low dissolved organic carbon (DOC) and ultraoligotrophic systems (Ellis-Evans et al. 2001) to true athalassic subsaline, chemically concentrated lakes (with high DOC) in the arid inland areas of western Greenland, some of which are meromictic (Anderson et al. 2001). There are a huge number of lakes on the ice-free margins of Greenland, particularly on the west coast, which is perhaps an under-researched lake district when seen in global terms. As well as this geographic spread in lake typology, altitudinal gradients also enhance the “arctic” effect. There are also local and regional spatial gradients: maritime influence varies from sea-spray effects on major ion chemistry (Anderson et al. 2001) to the creation of coastal meromictic systems (associated with isostatic rebound and basin isolation) with strongly chemically stratified water columns (as elsewhere in the Arctic, e.g. Ouellet et al. 1989). Jónasson and colleagues provide an introduction to the limnology of Iceland (Jónasson et al. 1977). In general, however, limnology on Iceland has been dominated by studies of lakes Thingvallavatn (64º10’N, 21º07’W) and Mývatn (65º35’N, 17º00’W) (e.g., Jónasson and Lindegaard-Petersen 1975; Jónasson and Adalsteinsson 1979; Lindegaard and Jónasson 1979; Kairesalo et al. 1989). Lakes on the Faeroe Islands are typically wind-stressed because of the local climate and their exposed nature. The majority are isothermal during the summer with only short periods of limited thermal stratification. Lakes at higher altitudes regularly freeze during the winter. Few data are available on the ecology of Faeroese lakes (Christoffersen 2002). A recent study of five oligotrophic lakes (summarized in Jeppesen et al. 2002) has suggested that species richness at a variety of trophic levels (fish, zooplankton, phytoplankton) was lower than expected from their position at the climatic border between the Arctic and mainland Europe, perhaps due to the oceanic dispersal barrier. Species diversity was intermediate between Greenland and Scandinavia/UK, but comparable with Iceland, although several key structuring organisms (e.g., invertebrate predators and certain fish

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species) found elsewhere on other North Atlantic islands were absent (Jeppesen et al. 2002). Nonetheless, across the northern North Atlantic the generalities of arctic limnology still apply: seasonally near-continuous, but low intensity insolation during the summer with a short ice-free season of up to ca. 3 months of the year (some lakes only developing a littoral, ice-free moat). Shallow lakes (< 2 m) often freeze to the bottom in winter. Paleolimnological themes across the northern North Atlantic Despite the prevalence of lakes on Greenland, Iceland, the Faeroes and elsewhere (e.g., Bear Island), paleolimnological studies are relatively limited, certainly in comparison with arctic Canada. Even within the northern North Atlantic, by far the most work has been carried out on Greenland to date (though this is now starting to change). As a result, we have chosen to base our synthesis around a number of themes identified in the literature rather than by geographic area. These themes reflect the development of important research questions in Holocene paleolimnology across the Arctic. Long-term limnological development (lake ontogeny) Given the relatively low anthropogenic impact on the majority of lakes in western and northern Greenland and Svalbard, their Holocene records largely reflect long-term developmental processes (especially the influence of climate on catchment soil processes and hence surface water chemistry). With sparse vegetation and lack of soil development in many areas of northeastern Greenland and Svalbard, the overriding factor affecting lake and catchment flora and fauna is perhaps that of indirect climate forcing, together with immigration-dependent processes that affect the development of the lake biota. Most of these records are, however, primarily from low arctic locations. There are very few studies from high arctic regions on either Greenland or Svalbard. The most obvious indications of lake ontogeny are the rapid changes following deglaciation. In many lakes in central West Greenland, sediments deposited immediately following deglaciation are carbonate-rich (because of calcite deposition; Fredskild 1977, 1983a, 1992b; Willemse 2002) indicating high alkalinity water. The period after deglaciation is characterized by an abrupt increase in sediment organic content in many lakes (Figure 2). Timing of the increase is dependent on lake position relative to the ice sheet, ranging from ca. 10,000 cal. yr BP (NAUJG1; 90 km west of present ice sheet margin; Figure 1), ca. 8000 to 8500 cal. yr BP for lakes in Søndre Strømfjord (Lake SS2 and Braya Sø; 40 km west of the ice sheet; Figure 1) and ca. 7400 cal. yr BP (SFL4; 5 km west of the ice sheet; Figure 1). In eastern Greenland, only Lake B1 (Figure 1) shows a clear early postglacial rise in organic content (Figure 2). There is ample sedimentary fossil evidence suggesting elevated lake productivity immediately after deglaciation. During this period, sediments often contain abundant alkaliphilous Characeae remains and elevated relative abundances of Pediastrum spp. (Fredskild 1983a, 1992b). Analyses at Lake SS2 indicate that Oscillatoria spp. pigments

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Figure 2. Loss-on-ignition (LOI) and total organic content (TOC) as % of dry weight from sediment records in Greenland. Data for lakes SFL4 and NAUJG1 from Willemse and Törnqvist (1999), Raffles Sø from Cremer et al. (2001a), Basaltsø and Lake B1 from Wagner et al. (2000) and Lake SS2 and Braya Sø from K. Brodersen et al. (unpublished). Ages given as calibrated years BP (present = AD 1950). Data for SFL4 and NAUJG1 from NOAA/NGDC Paleoclimatology Program, Boulder, Colorado, USA (IGBP PAGES/World Data Center A-Paleoclimatology Data Contribution Series, West Greenland Lake Sediment %LOI Data, #1999-034). Raffles Sø, Basaltsø and Lake B1 data from the PANGAEA database (www.pangaea.de/data; Dr. B. Wagner, pers. comm.).

Calibrated years before AD 1950

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(S. McGowan, unpublished) and diatoms indicative of mesotrophic conditions (Achnanthes minutissima and Cyclotella radios a; D. Ryves, unpublished) were abundant in the period following deglaciation when the lake was precipitating calcite. Diatom assemblages in Lake SFL4 near Søndre Strømfjord (Willemse 2002) shifted from disturbance-tolerant and cosmopolitan Fragilaria brevistriata immediately after deglaciation, to alkaliphilous and mesohalobous forms (ca. 7500 cal. yr BP) together with Ach nanthes minu tissima, a cosmopolitan taxon also commonly associated with early stages of lake development in alkaline-rich, glacial landscapes in more southerly latitudes (e.g., Ryves et al. 1996). In many lakes, after the initial period of high productivity following deglaciation, alkalinity declines and oligotrophication occurs as ions (and nutrients) are lost from the lakes or trapped within the developing soils of the terrestrial system. Development of humic acids in soils as organic matter accumulates exacerbates the decrease in soil alkalinity. Declines in the relative abundance of Pediastrum spp. (Fredskild 1983a) and a change from Potamogeton to Isoëtes spp. (ca. 7000 14C yr BP) in lakes near Nuuk and Ilulissat (Fredskild 1977, 1983a) indicate lake oligotrophication. In Lake SS2, a switch to littoral, benthic diatoms after ca. 7600 cal. yr BP (D. Ryves, unpublished) indicates a decline in lake productivity while subjective interpretation of diatom analyses from eastern and western Greenland (Foged 1989) also suggest natural acidification. Similarly, in Lake SFL4 (Søndre Strømfjord region) after ca. 6500 cal. yr BP, the alkaliphilous diatom assemblage is replaced by a relatively diverse range of cosmopolitan epiphytic diatoms (notably Gomphonema spp.), suggesting oligotrophication (Willemse 2002). Lake development in the later Holocene varies spatially, reflecting differences in local geology, geomorphology, climate influences and lake-specific characteristics. Some lakes become less productive throughout the Holocene (e.g., Basaltsø and the adjacent Lake B1 in eastern Greenland; Figures 1 and 2; see also Wagner et al. 2000; Cremer et al. 2001a). Lakes in southern and western Greenland examined by Fredskild (1985) became increasingly oligotrophic throughout the Holocene, with the most intense oligotrophication taking place after the Hypsithermal (ca. 4000 yr BP). However, local geology influenced the severity of oligotrophication (Fredskild 1983a), which was more pronounced in lakes with barren rock catchments than in those on glacial tills. Several lakes in eastern Greenland and around the head of Søndre Strømfjord in western Greenland show slight sedimentary organic enrichment during the middle to late Holocene (e.g., lakes SS2, Braya Sø, Raffles Sø; Figures 1 and 2). Pollen and macrofossil analyses of a 5000-year record from Lake 31 (Figure 1; Eisner et al. 1995), a small lake close to the ice margin, indicate a diverse macrophyte flora that has oscillated between eutrophic and mesotrophic conditions (Chara, Tolypella, Myriophyllum and Potamogeton filiformis). Similar shifts in the abundance of Chara at nearby Lake SFL6 (Figure 1; Willemse 2000) may also be indicative of nutrient enrichment. However, the interpretation of the macrofossil data primarily in terms of trophic status contrasts with that based on diatoms from the same lakes, highlighting the importance of multiproxy approaches. Substrate effects, lake depth and major ion chemistry are all important for macrophytes (especially Chara), and therefore inferences about “nutrient enrichment” based on their occurrence should be made cautiously. In shallow lakes, the development of a rich macrophyte flora is likely driven by sediment infilling processes that lead to

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altered internal nutrient cycling and expansion of littoral areas. The high abundance of Characeae at their northern limit in shallow thermokarst lakes on Svartenhuk Halvø (71ºN, 53ºW; Bennike 2000) supports the notion that lake morphometry is an important determinant of macrophyte abundance. Biological changes may also reflect ionic concentration associated with variable effective precipitation, as much as nutrient enrichment. Climate influences also likely play a role at Lake SFL6 and Lake 31 as these are situated in the exceptionally arid area around Søndre Strømfjord. Although probably a feature of the unusual climate in this region, evaporative concentration of lake water may be an important control on ontogeny. The effect of evaporative concentration is most apparent in closed-basin sites, which become increasingly saline in the late Holocene, such as Braya Sø (Figure 2; McGowan et al. 2003), Store Saltsø and Lille Saltsø (Figure 1; Bennike 2000). Evaporation could also explain why several lakes near the head of the fjord have elevated levels of total phosphorus (TP; maximum 34 µg L-1) and total nitrogen (TN; maximum 1860 µg L-1; Brodersen and Anderson 2002) which might be classed as mesotrophic or eutrophic (e.g., Vollenweider and Kerekes 1982). Climate forcing and paleoproductivity Given the problems of deriving unambiguous climate signals from a single biological proxy in lake sediments (Anderson 2000; Battarbee 2000; Brodersen and Anderson 2000), attention has turned to other proxies in an effort to identify a suitable variable. Despite the complex nature of factors controlling organic production in lakes, and preservation in lake sediments, sedimentary organic content and biogenic silica have been used as indirect measures of past climate (Wagner et al. 2000; Kaplan et al. 2002). Sediment dry weight loss-on-ignition at 550ºC (%LOI; Dean 1974) or total organic content (%TOC) have been used as a proxy for climate-driven lake productivity in several studies from Greenland (Willemse and Törnqvist 1999; Wagner et al. 2000; Cremer et al. 2001b) (Figure 2). It is apparent, however, that such an interpretation cannot be generalized among all lakes, as %LOI or %TOC show distinct differences in both value and pattern even within the same region (e.g., among western Greenland or eastern Greenland lakes; Figure 2; cf. Bennike 2000). While part of the reason may be an artefact of sampling frequency (e.g., coarser sampling resolution in eastern Greenland lakes; Figure 2), differences in lake characteristics (e.g., volume, depth, mixing regime, chemistry) as well as location (e.g., proximity to ice sheet, geological setting, altitude) are also important. Sediment organic content (as %) varies in response to productivity, preservation changes (diagenesis), recycling and dilution by both biogenic and inorganic mineral fractions (which may be autochthonous or allochthonous). Biogenic silica (BSi; itself related to lake productivity) may dilute the %LOI signal in lakes such as NAUJG1 (West Greenland; Figure 1), to the extent that the residual-on-ignition (%ROI) provides the most convincing signal of lake productivity (and in this case may be linked to climate; Willemse and Törnqvist 1999). In larger, deeper lakes in southern and eastern Greenland, there is good agreement between %TOC (%LOI) and biogenic silica estimates, such as at Qipisarqo Lake (Kaplan et al. 2002; Figure 1) and Raffles Sø (Cremer et al. 2001b). Studies at Qipisarqo Lake (Kaplan et al. 2002) and Basaltsø (Wagner et al. 2000) concluded that

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BSi can be used as a measure of paleoclimatic variability, based on a relatively good agreement of the paleoproductivity proxies and independent paleoclimate estimates. However, in Basaltsø (Figure 1; Wagner et al. 2000), there is a much weaker correlation between %TOC and BSi, leading Wagner and colleagues to dismiss the early Holocene %TOC profile as reflecting enhanced decomposition and other diagenetic processes. Nevertheless, organic carbon preservation tends to increase with higher productivity and anoxic conditions associated with thermal stratification. Increased ice cover may also lead to anoxia and hence better preservation. At a nearby smaller lake (Lake B1), the lack of agreement between BSi and %TOC is argued to be due to the dissolution of diatoms in the early Holocene (ca. 8500 to 6500 cal. yr BP) based on a weak inverse relationship between the BSi and carbonate profiles (Wagner et al. 2000). Some common general trends in %LOI and %TOC curves are nonetheless apparent, and reflect both ontogenetic development (which may be non-synchronous among lakes) and climatic forcing (which may be both longitudinally and latitudinally asynchronous; Ritchie et al. 1983; MacDonald et al. 1993). In particular, neighbouring lakes SS2 and Braya Sø (West Greenland; Figure 1) show striking synchronology (within dating errors) in %LOI trends throughout their records (Figure 2). Both show elevated %LOI from 7000 to 8000 cal. yr BP, with a sudden drop followed by a longterm rise to a mid-Holocene peak at ca. 3600 cal. yr BP. Fluctuations continue until a second maximum at ca. 1000 cal. yr BP, after which values fall to the present day. Such agreement is all the more striking as these lakes have followed different developmental pathways. Lake SS2 has remained fresh throughout the Holocene, after a period of high alkalinity soon after deglaciation (K. Brodersen et al., unpublished), while the closed lake Braya Sø has been characterized by salt accumulation and meromixis since the early/mid-Holocene, after an initial freshwater phase (Figure 3; McGowan et al. 2003; see below). Some of these trends are seen in the other western Greenland lakes (Figure 2), with the period from 7000 to 8000 cal. yr BP registered as a trough in %LOI at NAUJG1 (presumably due to high biogenic silica concentrations). High %LOI values were recorded at Lake SFL4 (but low values at NAUJG1) during the period ca. 3000 to 4000 cal. yr BP. Correlations with eastern Greenland are less convincing, although there are elevated %TOC values at Raffles Sø from ca. 4000 to 5000 cal. yr BP (though there is speculation that this reflects enrichment by sea bird colonies; Cremer et al. 2001b; Wagner and Melles 2001), and again at ca. 800 cal. yr BP, with low values in more recent sediments. Such regional agreement between widely different lake types points to a common regional climatic causal mechanism driving changes in lake productivity. In southern Greenland, at Qipisarqo Lake, where %BSi and %LOI curves correlate well, %LOI values increase just prior to 9000 cal. yr BP, remaining between 10 to 20% before declining around 2000 cal. yr BP (Kaplan et al. 2002). While the Qipisarqo Lake %LOI curve broadly agrees with the %TOC curve at Raffles Sø (Figure 2), the respective BSi profiles do not. This implies that interpretations of such bulk sediment parameters purely (or even largely) in terms of regional paleoproductivity and paleoclimate need to be made with caution, even where there is internal consistency between “productivity indicators” within an individual sediment record (e.g., organic C and BSi). As Figure 2 suggests, care must also be taken in choosing sites for comparison of regional signals, as there is a risk that single sites are not representative of the regional trend. Ideally, records from several sites within a region should be

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examined (e.g., Anderson et al. 1999; Willemse 2002), as each lake will differ in its response, and the fidelity of its sedimentary archive. There is some evidence, for example, that Raffles Sø is an unusual lake and its paleolimnological record might be more complex than other lakes in the region (Cremer et al. 2001b; Wagner and Melles 2001). The problems of using %LOI and %TOC as regional indicators of climate-driven paleoproductivity are further demonstrated by Bennike’s study of four lakes in western Greenland (Bennike 2000). These profiles show little similarity even when agecorrected, perhaps a reflection of their different lake types. It is perhaps more revealing to compare organic and BSi (or diatom) accumulation rates, although values are sensitive to processes of sediment focussing, and Si dissolution (cf. Ryves et al. 2001). Eisner et al. (1995) show that long-term changes in %LOI in a sediment core from Lake 31 in western Greenland are better explained by dilution of the organic fraction by aeolian silts rather than changes in lake productivity (although there is debate as to the amount of aeolian material that enters lakes, especially over the shorter term; Willemse and Törnqvist 1999; Willemse 2002). However, this long-term effect is clearly revealed when %LOI is plotted as an accumulation rate (g cm-2 yr-1) rather than a dry weight proportion (Eisner et al. 1995). Geochemical and magnetic studies on sediment cores from two neighbouring oligotrophic lakes in northwestern Iceland, Thidriksvallavatn (65º41’N, 22º45’W) and Vatnsdalsvatn (65º36’N, 23º05’W), covering the last 500 and 1000 yr respectively, have found evidence for 80 to 100 yr erosion cycles for most of this period (Doner 2003), possibly linked to the NAO. If such cycles have had major impacts on lake productivity, sedimentary total carbon content (%TC, almost all of which is organic C) does not reflect them (mean 2-3% for both lakes). However, neither geochemical/magnetic parameters nor %TC respond strongly to the Little Ice Age (LIA), possibly due to the oceanic setting and human activity (sheep grazing) in the region. Sedimentary properties other than organic content might be better guides to climatic and environmental impact in arctic lakes, while local setting (e.g., maritime influence) may override the expression of otherwise strong regional forcing signals. Diatom assemblages and climate Use of diatom assemblages in the inference of climate often relies on current knowledge of the autecology of individual taxa. This can be problematic when Fragilaria spp. dominate the diatom assemblage, as is often the case in arctic environments, because these taxa have near global distribution over a wide range of habitats (Anderson 2000). In Kap Inglefield Sø in the far northwest of Greenland (78ºN; Figure 1; Blake et al. 1992), the diatom assemblage is extremely species-depauperate (with only 20 species in total), and the entire 8000-year sediment record is dominated by Fragilaria construens var. venter, together with F. pinnata which increases after 5500 14C yr BP. Short core studies on three shallow lakes in Svalbard (Jones and Birks 2004) show a dominance of Fragilaria spp. with no change over the last 700 years. The diatom record from Mývatn, Iceland (Einarsson 1982; Einarsson and Haflidason 1988), a large (37 km2), shallow (3-4 m water depth) lake formed some 2300 years ago, has been completely dominated by Fragilaria construens and F. pinnata following a brief period of co-dominance with

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Aulacoseira spp. immediately after the lake formed. It has been suggested that dominance of Fragilaria is the result of extended periods of ice cover and variable moating (Smol 1988; Douglas and Smol 1999). However, based on early Holocene changes in Staurosirella (= Fragilaria) pinnata and Staurosira (F.) construens at Basaltsø, eastern Greenland (Figure 1), Cremer et al. (2001a) concluded that Fragilaria are temperature-sensitive species. The near total loss of planktonic diatoms since 2000 cal. yr BP at Raffles Sø, eastern Greenland was interpreted as being the result of increased ice cover (Cremer et al. 2001b), and correlated to a Neoglacial cooling period beginning ca. 3000 cal. yr BP in southern Greenland (Kaplan et al. 2002). However, at Basaltsø, some 280 km to the north of Raffles Sø, a sustained cooling since ca. 5000 cal. yr BP (inferred from lithological, biogeochemical and palynological evidence; Wagner et al. 2000) was associated with a significant and sustained increase in the abundance of planktonic Cyclotella spp. (Cremer et al. 2001a). At two deep sites on Svalbard, substantial changes in the diatom assemblages since ca. AD 1200 were asynchronous between lakes and uncorrelated with records of atmospheric pollution (Jones and Birks 2004), indicating the problems of identifying regional patterns, when in-lake responses to climate are mediated through catchment processes. These studies highlight a need for further research both on the basic ecology of diatom taxa, and limnological responses to climate forcing in order to strengthen paleoclimatic interpretations. Records of changing effective precipitation Many of the papers reviewed above are concerned with inferences of past temperature changes. However, as large areas of the Arctic have low amounts of precipitation (< 300 mm yr-1) and are classified as polar deserts (Orvig 1970), precipitation change may be equally important. Reconstructions of precipitation are necessary to complement temperature records from the ice cores and other proxies. The effects of long-term evaporation on arctic lakes are not fully understood, although well-known subsaline lakes (0.5-3 g L-1 total dissolved solids), such as those in the Søndre Strømfjord area of western Greenland (Anderson et al. 2001), lie above the marine limit and have thus not formed as a result of trapping of sea water (e.g., Ouellet et al. 1989). The presence of fossil shorelines around many of the saline lakes today suggests that evaporation has lowered lake-levels (Anderson et al. 1999). A strong climatic (largely effective moisture) gradient exists today from the coast to the ice sheet in western Greenland, which has a major influence on lake water chemistry, and in particular conductivity (salinity; Anderson et al. 2001). The influence of evaporative processes on lake water chemistry is also supported by stable isotope analyses of modern water samples from western Greenland (Leng and Anderson 2003). Conductivity is the dominant control on the diatom flora found in these lakes (Ryves et al. 2002) but not on the chironomid fauna (Brodersen and Anderson 2002). Diatom transfer functions have been developed which can be used to reconstruct conductivity (Ryves et al. 2002) from sedimentary diatom assemblages, and may be interpreted in terms of lake ontogeny and/or climate change (McGowan et al. 2003).

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In Braya Sø near Søndre Strømfjord (Figure 3), conductivity changes throughout the Holocene indicate that the lake was fresh (ca. 200 µS cm-1) before 7000 cal. yr BP. Salt accumulation between ca. 6800 and 5500 cal. yr BP caused conductivity to increase to at least ca. 1800 µS cm-1, during a period of poor diatom preservation. A subsequent return to freshwater conditions is implied by a rapid fall in diatom-inferred conductivity until ca. 3500 cal. yr BP, after which time there was a long-term trend of increasing but variable conductivity with periods as high as 4500 µS cm-1 inferred. Conductivity probably responded to both short-term, within-lake processes connected with lake stratification (see below), as well as long-term shifts in climate (effective moisture) at Braya Sø. Sustained long-term evaporation also results in carbonate deposition in these saline lakes, after increasing ionic concentration reaches the bifurcation point where calcite precipitation occurs (Anderson et al. 2001). Stable isotope analyses (G18O) of carbonate-rich lake sediments also suggest that evaporative concentration was important in Braya Sø (Leng and Anderson 2003). Diatom analyses on nearby shallow (5.3 m deep) Lake SFL4 suggest an early Holocene increase in effective precipitation somewhat prior to the arid period recorded at Braya Sø (Willemse 2002). Alkaline and mesohaline conditions after ca. 7500 yr BP (Denticula k uetzingii var. r umrichae, Mastogl oia smit hii var. lac ustris, Achnanthes delicatula and Na vicula halophila) are followed by a decline in ionic concentrations from ca. 6500 cal. yr (Cymbella and Gomphonema spp.). Increased abundance of A. delicatula between 3800 and 2000 cal. yr BP indicates a substantial increase in the lake water ionic concentration, in agreement with elevated diatom-inferred conductivity values at Braya Sø (Figure 3) and neighbouring lake SS6 (Figure 1; McGowan et al. 2003), but contrasting with the concurrent decline in ionic concentration at Lake 31 (some 25 km from Braya Sø; Eisner et al. 1995). However, later increases in ionic concentration in the last 1000 years in Lake 31 are in good agreement with the Braya Sø record. Gastropod fossils from other saline lakes in the Søndre Strømfjord region (Store and Lille Saltsø; Figure 1) indicate that the lakes were saline at ca. 4000 cal. yr BP (Bennike 2000). Periods of low effective precipitation will also lead to increased aeolian transport of loess deposits from the ice sheet outwash plains, that may contribute to enrichment of lake water chemistry (see Eisner et al. 1995). Mineral magnetic analyses of lake sediments (Clarke 2002) and stratigraphic studies of peaty soils along the ice margin (Willemse 2000) suggest that aeolian loess deposition began ca. 5000 yr BP following deflation of the plains, further supporting the hypothesis of increased late Holocene aridity. Biological structure To date there have been remarkably few attempts to examine the effects of changing biological structure on the ontogeny of arctic lakes during the Holocene from multiproxy analyses of faunal and floral remains. In the areas included in this review, very few studies have tried to consider lake development in terms of autogenic processes (trophic interactions) and their interplay with climate-controlled catchment processes (such as soil-microbial processes).

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Given the difficulty of interpreting bulk sedimentary organic content simply in terms of climate-driven lake productivity and the inferential ambiguities inherent with any single biological proxy (Figure 2; see earlier text), examining the response of whole biological communities may be a better guide to lake development. There is little doubt that in more extreme environments such as northern Greenland and Svalbard, climate is probably the overriding factor controlling biological structure (and perhaps productivity) in lakes. At lower latitudes, however, biological interactions are probably an important factor in determining the signal observed in lake sediment records. Figure 4 summarizes the variation among diatom assemblages using detrended correspondence analysis (DCA) axis 1 scores for four Greenland lakes (explaining between 18 to 34% of diatom data), with widely differing diatom communities. There is surprising agreement among the DCA trajectories, however, implying some common trends in lake development in eastern and western Greenland, which is not so apparent from gross parameters such as sediment organic or BSi content (cf. Figure 2). The diatom assemblages of Lake SS2 and Braya Sø closely follow diatom-inferred water chemistry (Figure 3), supporting the interpretation of assemblage shifts in terms of dominant water chemistry, but the same general pattern of DCA scores is found at all four sites. In particular, diatom assemblage shifts from ca. 8000 to 5000 cal. yr BP (coincident with the early Holocene climatic optimum) suggest a common pattern of limnological development, although the timing and extent of this and subsequent developments vary between lakes. Although the long-term climate signal in arctic saline lakes (Figure 3) is simplified by the minimal effect of groundwater, in-lake processes also have to be considered, particularly in relation to the interplay between climate, thermal stratification and biological productivity. Multi-proxy studies can help separate the confounding effects of internal biological processes from climate change per se , as well as long-term development. Pigment studies using High Pressure Liquid Chromatography (HPLC) are now relatively common in paleolimnology (Leavitt and Hodgson 2001) but there have been few studies in Greenland or elsewhere in the North Atlantic islands. Pigment analyses of recent (past ca. 1100 yr BP) sediments from Braya Sø (Figure 5; S. McGowan, unpublished) demonstrate how complex in-lake processes may cause abrupt changes in photoautotrophic communities. The fluctuating presence of purple sulphur bacteria (indicated by okenone) suggests that there have been periods of intense stratification or meromixis in the lake over the last ca. 1100 years, associated with elevated concentrations of pigments from siliceous (diatoxanthin), cryptophyte (alloxanthin), green (chlorophyll-b, pheophytin-b) and green/blue-green (luteinzeaxanthin) algae (Figure 5). During meromictic periods, mean surface diatom-inferred lake water conductivity was significantly lower than in the intervening periods (mean 2679 µS cm-1 compared to 3441 µS cm-1; p < 0.002, t-test) and stratification was probably stabilized by the density separation between overlying fresher water and more saline bottom waters. In such a meromictic system, increasing aridity would lead to lower lake-levels and evaporative concentration of upper lake waters, both tending to promote the breakdown of stratification. This would both eliminate the anoxic habitat for purple sulphur bacteria (cf. lower okenone concentrations in Figure 5) and further increase surface conductivity as more saline water from the hypolimnion was mixed throughout the water column. The decline in water depth would cause increased light penetration to the sediment surface and allow the development of benthic blue-green

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mats (oscillaxanthin, myxoxanthophyll; Figure 5). The development and breakdown of meromixis across critical thresholds can help explain the high frequency fluctuations observed in the diatom-inferred surface conductivity at both Braya Sø (Figure 3) and SS6 (McGowan et al. 2003) over the later Holocene. Such in-lake processes can thus modify the direct effect of climate on lake functioning and biotic development. Primary production in many shallow and nutrient-poor lakes is dominated today by mosses, benthic diatoms and non-siliceous filamentous algae. In oligotrophic lakes in eastern Greenland, benthic production of algae comprises > 80% of total algal production (Vadeboncoeur et al. 2003). Although aquatic mosses are important in many shallow arctic lakes to both primary production and lake functioning, to date there have been few attempts to examine the changing biomass of mosses over time from the sediment record. However, moss layers were used to aid in the interpretation of lateglacial climate-environmental change in a small coastal lake situated in southern Greenland (Björck et al. 2002).

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Figure 4 . Diatom DCA axis 1 sample scores for four Greenland lake sediment records. Ages given as calibrated years BP (present = AD 1950). Unknown or unidentified taxa were deleted prior to recalculating percentages to aid comparison between data sets. Percentage data from Raffles Sø were square-root transformed as species data were dominated by only 6 taxa (Cremer et al. 2001b), and were untransformed in all other cases. All final data sets had axis 1 gradient lengths from 2.6 to 4.2 S.D. units (and thus unimodal models were appropriate; ter Braak 1995). DCA performed on percentage data within CANOCO 4 (ter Braak and Šmilauer 1998), with downweighting of rare species. Amount of variation in data sets accounted to by DCA axis 1 is given for each lake (significance estimated using a broken stick model; Jolliffe 1986). Data for Raffles Sø (Cremer et al. 2001b) and Basaltsø (Cremer et al. 2001a) courtesy of Dr. H. Cremer (pers. comm.) and the PANGAEA database (www.pangaea.de/data). Data for Braya Sø from McGowan et al. (2003) and Lake SS2 from D. Ryves (unpublished).

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Macrofossil studies on Svalbard and Iceland have tended to be focussed on terrestrial remains with associated inferences about catchment development and climatic controls. At Stevatnet on Bear Island, preliminary analyses of aquatic remains were inconclusive, in part due to the poor preservation of diatoms (Wohlfarth et al. 1995). Rundgren (1998) analysed plant macrofossils at two sites on the Skagi Peninsula, northern Iceland, while Birks (1991) provided a similar record for Skardtjørna (Spitsbergen, Svalbard) covering the last 8100 cal. yr BP. As these records are primarily of terrestrial vegetation succession they are not considered further here. One of the earliest studies of the faunal remains contained in Greenland lake sediments was that of Fredskild et al. (1975) but this primarily consisted of an analysis of bulk samples with no stratigraphic constraints. The results were considered in relation to the known distribution today. In Iceland, as a result of the history of limnological research on the productivity and ecology of the chironomid-dominated benthic fauna of Mývatn and Thingvallavatn, it is not surprising that most paleolimnological research has concentrated on the Holocene development of the benthic fauna (Einarsson 1982; Gardarsson et al. 1988; Caseldine et al. 2003). Research has also been carried out examining the fossil pigment stratigraphy of Lake Mývatn in relation to tephra deposition (Einarsson et al. 1993). Bennike’s macrofossil work in western Greenland has provided a regional history of the aquatic invertebrate fauna (e.g., Bennike 1995, 2000). The tadpole shrimp, Lepidurus arct icus, is a dominant component of the lake fauna in the initial phases following deglaciation in sites around Søndre Strømfjord (Lille Saltsø and Store Saltsø; Bennike 2000 and Lake SS2; K. Brodersen et al., unpublished), Svartenhuk Halvø (Bennike 2000), Godthåbsfjord (southwestern Greenland; Figure 1; Fredskild 1983b) and elsewhere in western Greenland (Bennike 1995). In eastern Greenland, L. arcticus is abundant in both the early and late Holocene (Bennike et al. 1999), while at Nuussuaq (Disko Bay; Figure 1) only in the late Holocene (Bennike 2000), possibly reflecting sediment infilling at this shallow site which would have increased the availability of habitat for this benthic detrital scavenger. Chironomids were not differentiated by Bennike but were early colonisers in a number of the lakes at the same time as the cladoceran Chydorus arcticus . Another general trend is the declining abundance of Alona spp. in the late Holocene (dates ranging from ca. 2000 to 4000 cal. yr BP) which, depending on the species present, may reflect declining productivity or habitat loss (reduced macrophyte abundance). At Mývatn (Iceland), remains of the filamentous alga Cladophora aegagr opila show a substantial increase after AD 1600, which Einarsson (1982) related to improved light conditions, perhaps due to reduced cyanobacterial blooms. Cladoceran remains were dominated by Chydorus sphaericus and a number of Alona spp. prior to the increase in Cladophora. There is, however, a clear increase in the abundance of Eurycercus lamellatus and the chironomid Psectrocladius bar bimanus, both of which are commonly associated with Cladophora today (Gardarsson et al. 1988). Daphnia l ongispina is only important in the period immediately following the formation of the lake (Einarsson and Haflidason 1988). A chironomid training set was developed for western Greenland (Brodersen and Anderson 2002) with the primary aim of creating a water temperature inference model. Forty-seven lakes were analysed and 24 taxonomic groups identified, of which the dominant species/forms found were Micropsectra spp. (33%) and Psectrocladius spp. (18%). Only 21 lakes had both adequate water temperature and chironomid data that

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could be used in a weighted-averaging (WA) calibration, but the model has reasonable error statistics compared to other published chironomid-temperature models. However, a major problem with this model, as with many others, is that of the covariance of water temperature and nutrient concentrations (especially nitrogen in western Greenland), that may compromise the application of the model to Holocene profiles for temperature reconstructions. Cold lakes are also nutrient poor lakes and there is no distinct cold fauna as a result. Nonetheless, chironomid assemblages do discriminate among major lake types, and so can be used to reconstruct broad changes in lake development over time (Brodersen and Anderson 2002). On Svalbard, chironomid assemblages from 23 lakes have been analysed (Brooks and Birks 2004). Although there was an insufficient gradient to develop a WA model, four faunal assemblages were identified, primarily influenced by pH, nutrient concentrations, water temperature and water depth, respectively. Short cores (covering the last ca. 700 years) from three lakes were also analysed and at two lakes there is evidence for synchronous changes, indicative perhaps of the end of the Little Ice Age (Birks et al. 2004). Cultural disturbances Long-term cultural catchment disturbances and their impacts on lakes are generally associated with northwestern Europe. However, areas of southern Greenland, the Faeroe Islands and Iceland have all experienced cultural disturbance to varying degrees (Hannon and Bradshaw 2000; Hannon et al. 2001). Dates of human colonisation and the start of agriculture are ca. AD 650 to 970 for the Faeroes, ca. AD 950 to 1000 for Iceland and AD 950 for Greenland. The Viking occupation of southwestern Greenland has become a focus of research activity (Buckland et al. 1996; Barlow et al. 1997) because of the possible role of climatic deterioration (during the Little Ice Age) in causing the abandonment of the settlements. Substantial increases in soil erosion following the start of agriculture in southern Greenland ca. AD 1000 have been reported (Sandgren and Fredskild 1991; Fredskild 1992a). After a settlement period of ca. 500 years, vegetation recovered with a subsequent reduction in rates of soil loss. The introduction of sheep in the 20th century, however, resulted in over-grazing and further major problems of soil erosion. The response of the lake biota to such catchment disturbance was however not addressed in these studies. It is estimated that some 30% of Iceland was covered by birch woodland prior to the settlement of the country around AD 900 (Loftsson 1993). Deforestation, following the start of agriculture, resulted in a substantial increase in the rate of soil erosion, which can be clearly identified in the sediment record (Bradshaw and Thompson 1985). The extent of forest coverage on the Faeroes is a matter of dispute (Hannon and Bradshaw 2000), but like Iceland, it is today a treeless landscape. At both locations, anthropogenic disturbance was superimposed on a finely-balanced system that was being increasingly stressed by climate (Humlum and Christiansen 1998a; Olafsdottir and Gudmundsson 2002). This was enough to result in major ecological changes in terrestrial vegetation and soil loss, with paleolimnological evidence from Faeroese lakes testifying to associated impacts on lake chemistry and biota (Hannon et al. 2001). Despite the widespread interest in cultural development and its environmental impacts, the effect on

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lakes in southern Greenland and Iceland, interpreted from a more overtly limnological perspective, is only just being examined. More recent anthropogenic disturbance is indicated by sediment records of atmospheric pollution (Hg, Pb, spheroidal carbonaceous particles (SCPs) and persistent organic pollutants) on Greenland and SCPs on Svalbard (Bindler et al. 2000, 2001; Birks et al. 2004). At most lakes there is unambiguous evidence for late 19th to mid-20th century pollution. For Greenland, the enrichment factors suggest that this substantial increase is part of the general atmospheric pollution in the Arctic. At Svalbard, the situation is slightly more complex because of distinct local pollution sources, such as coal mines (Birks et al. 2004). Birks et al. (2004) conclude that the Svalbard lakes, which show "considerable lake individualism", are insensitive indicators of long-range atmospheric pollution because of the natural variability of the sediment record and their remoteness. Muir and Rose (this volume) provide a more detailed synthesis of arctic atmospheric pollution records. Tephra layers and lake sediments The relationship of Iceland to the North Atlantic rift zone, and the associated volcanic activity in the area, provide a source for ash dispersal that can be deposited in lake and bog sediments throughout the North Atlantic region. Their main use has been for dating and the effect of these eruptions on the lakes themselves has not been fully addressed in the Faeroes or Iceland. Assuming a historically known age of the eruption event or through careful dating of the sediment enclosing the ash layer, these (sometimes visible) layers can be used for precise dating and chronostratigraphical correlation of records over very large areas of the North Atlantic. Examples of widely used marker horizons are the Saksunarvatn tephra (ca. 9000 14C yr BP) (Mangerud et al. 1986; Merkt et al. 1993; Birks et al. 1996), the Hekla-4 tephra (ca. 3800 to 4000 14C yr BP) (Persson 1968; Dugmore et al. 1995) and the Landnam tephra (AD 871 ± 2 cal. yr) (Thórarinsson 1944; Grönvold et al. 1995; Hannon et al. 1998; Wastegård et al. 2001). Icelandic volcanic glass shards, found as ice-rafted debris in marine sediments from the North Atlantic, have also provided convincing evidence of maritime circulation changes at subMilankovich scales during the late Quaternary (Bond et al. 1997). There is also scope for further paleolimnological studies on the impact of tephra deposition on lake ecology and development, especially on Iceland (see Einarsson et al. 1993) as has been explored elsewhere in Europe (e.g., Lotter and Birks 1993). As far as we are aware, no tephra has been reported from Holocene sediments in Greenland lakes, although it has been recorded in Greenland ice cores (Palais et al. 1992; Grönvold et al. 1995). Synthesis and areas for further research Paleolimnology on Greenland and Svalbard has progressed rapidly in the last 5 to 10 years, and similar work is now ongoing on the Faeroes. Holocene studies on Iceland are more limited as the focus has been on the late-glacial and event stratigraphy. The majority of work to date has been concerned with lake ontogeny, paleoclimate

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inferences or cultural disturbances. Most of the studies on full Holocene sequences have been carried out on Greenland lakes. Despite the general assertions that lake sediment records from Greenland contain clear climate signals (Wagner et al. 2000; Kaplan et al. 2002), the evidence from multiple proxies suggests that this is perhaps an oversimplification. Lakes differ in their sensitivity to environmental change (e.g., according to size, depth, setting, and chemistry), and thus different aspects of environmental change are better recorded and preserved in some lake types than others (Birks et al. 2004). Effects of immigration constraints on species assembly are difficult to disentangle unambiguously from climate effects. As a result, autogenic biological processes are super-imposed on those changes associated with both lake ontogeny (i.e., climate forcing mediated via catchment processes) and direct climate forcing. As the importance of both UV-radiation and DOC concentrations on lake biological structure becomes more apparent in mid- and high latitudes (Schindler et al. 1996b; Williamson et al. 1996; Pienitz and Vincent 2000), there is a clear need for research to assess their role in arctic systems. Many lakes on Greenland and Svalbard have long retention times (due to low precipitation and closed basins) with associated build-up of refractory organic compounds during the Holocene. There are clear problems with the bioproductivity-climate approach: notably the lack of regional synchronology in most %LOI and BSi records, for example, between Lake B1 and Basaltsø (Figure 2; Wagner et al. 2000). These differences indicate the sensitivity of lakes to their local setting, such as the influence of local vegetation, lake area, volume and catchment-lake area ratios. There are also substantial problems with using either BSi or %TOC (%LOI) as indirect measures of paleoproductivity and paleoclimate, most notably that of dilution effects associated with non-constant accumulation rates and/or inputs of minerogenic matter, preservation and diagenesis. Few, if any of the papers reviewed here, use accumulation rate data to overcome dilution effects or changes in sediment texture. Eisner et al. (1995), however, did correct %LOI for dilution by aeolian silt, and used pollen accumulation rates to infer changes in vegetation despite there being little compositional change. Thus, there is a clear need for more comparison of accumulation rates, although adequate chronological control may preclude this at many sites. Percent LOI or TOC curves should also be compared with other proxies of lake productivity (e.g., aquatic macrofossils, diatoms, or pigments) across several sites within a region and including several lake types if possible. Major perturbations such as the late-Glacial (and the 8.2 ka BP event) have left clear imprints in the paleorecord but the more subtle shifts in climate during much of the Holocene are more difficult to detect, as not all the biota show a similar response. There is, however, the added problem of spatial variability of climate forcing: for example, there appears to be no “8.2 K” event recorded in eastern Greenland sediments, nor is it found in the Renland ice core (71ºN, 27ºW; Figure 1). In Efstadalsvatn, northwestern Iceland (66º55’N, 21º40’W), there is little biostratigraphic (pollen or chironomid) evidence of disturbance at 8.2 ka BP, but there are changes in sedimentary characteristics (organic content and magnetic susceptibility; Caseldine et al. 2003). The early to mid-Holocene Hypsithermal and later Neoglacial cooling is also clearly spatially variable. Cooling is inferred from 5000 to 6500 yr BP in eastern Greenland (despite Raffles Sø not showing obvious signals until 1900 yr BP), yet started ca. 3000 yr BP in southern Greenland (Kaplan et al. 2002) and 5000 to 4000 yr BP in western Greenland. Based on their studies at Lake 31 (western Greenland), Eisner et al. (1995)

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infer a glacial still-stand or re-advance after 4600 cal. yr BP, while diatoms indicating freshwater conditions reappear at Braya Sø and Lake SS6 from 5000 to 4000 yr BP (McGowan et al. 2003; Figure 3). There are also differences between coastal and inland lakes, perhaps reflecting local maritime effects. The evidence suggests climate change is latitudinally asynchronous (Wagner et al. 2000) but is regionally coherent. Sites in the north warmed last during the Holocene climatic optimum and cooled first subsequently, perhaps related to changes in oceanic circulation and variable radiation inputs. The initial assumption of many early paleolimnological studies of arctic lakes was that there should be a direct link between the lake sediment record and climate, demanding comparison with proxies from the Greenland ice sheet records. While this may be true at high latitudes (e.g., northern Greenland) and in the earliest stages of deglaciation, multi-proxy records show that this assumption does not hold true in more complex, less extreme systems. For example, based on analyses of six cores from both the high and low Arctic in western Greenland, Fredskild (1983a) concluded that direct temperature effects were only important for lake development immediately following lake formation. Within 1000 to 1500 years, the influence of catchment soil and vegetation processes on lakes became dominant and they developed along an oligotrophication pathway, although expressed individually among lakes. Generally, lake response is truly multivariate and biological interactions and withinlake processes become important (Anderson 2000; Birks et al. 2004). Climate will often have a major indirect role, especially on shorter timescales (Schindler et al. 1996a; Myneni et al. 1997; Quayle et al. 2002), and perhaps set the parameters within which lake trajectories develop, but effects are mediated through catchment-lake processes. The scale and timing of responses differ amongst biological and physical proxies, and between lakes with different geological, topographical and hydrological settings. Contemporary limnological monitoring within the North Atlantic islands (e.g., at Zackenberg) and ongoing coordinated physical, ecological and cultural studies in the region (e.g., the NARP/NORLAKE program) will provide a better insight into shortterm, individualistic environment/lake interactions. Increasing our understanding of biological and physical responses to external and internal processes will improve our interpretation of their signals preserved in the sediment records of these arctic systems, and their long-term reconstruction over the Holocene. It is perhaps more interesting to focus on among-lake differences rather than attempting to force the signals to agree both among lakes themselves and with the ice core records. As such, lakes give a better indication of regional climate variability. There is a clear need to synthesize records among lakes in a more objective manner but we are at the moment extrapolating from a small number of sites. Our understanding will be improved as more studies on Greenland and the other islands become available. It is clear from the ambiguities we have identified that there is a need for more high resolution, well-dated studies. Given the slow accumulation rates and the lack of macrofossils especially in the early Holocene at many arctic sites, good chronological control will, of course, remain a challenge (Wolfe et al., this volume). We should, however, begin by considering the differences among the different records rather trying to identify the Greenland ice core climate signal in every lake sediment record in the region.

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Summary This chapter reviews the available literature on paleolimnological research on Greenland, Iceland, the Faeroe Islands and Svalbard. Emphasis is placed on paleolimnological work and terrestrial paleoecological investigations are ignored. A brief review of relevant limnological research is given. Because of the limited number of studies, work to date is synthesized under four main areas: lake development, climate forcing and paleoproductivity, changes in biological structure and cultural disturbances. Most studies to date have been undertaken on Greenland but work is progressing rapidly elsewhere, e.g., on Iceland, Svalbard and the Faeroes. There are very few training sets for quantitative reconstruction available for any biota as yet. Many paleoclimate interpretations are perhaps oversimplistic in that they ignore a limnological perspective. Possible areas for future development are outlined. Acknowledgements We are grateful to Drs B. Wagner and H. Cremer for facilitating access to their published and unpublished data, and Emily Bradshaw for comments on an earlier version. We thank Nico Willemse for a thorough review. References Anderson N.J. 2000. Diatoms, temperature and climatic change. Eur. J. Phycol. 35: 307-314. Anderson N.J., Bennike O., Christoffersen K., Jeppesen E., Markager S., Miller G. and Renberg I. 1999. Limnological and palaeolimnological studies of lakes in south-western Greenland. Geology of Greenland Survey Bulletin 183: 68-74. Anderson N.J., Harriman R., Ryves D.B. and Patrick S.T. 2001. Dominant factors controlling variability in the ionic composition of West Greenland Lakes. Arct. Ant. Alp. Res. 33: 418-425. Barlow L.K., Sadler J.P., Ogilvie A.E.J., Buckland P.C., Amorosi T., Ingimundarson J.H., Skidmore P., Dugmore A.J. and McGovern T.H. 1997. Interdisciplinary investigations of the end of the Norse western settlement in Greenland. The Holocene 7: 489-499. Battarbee R.W. 2000. Palaeolimnological approaches to climate change, with special regard to the biological record. Quat. Sci. Rev. 19: 107-124. Bennike O. 1995. Palaeoecology of two lake basins from Disko, West Greenland. J. Quat. Sci. 10: 149-155. Bennike O. 2000. Palaeoecological studies of Holocene lake sediments from west Greenland. Palaeogeogr. Palaeoclim. Palaeoecol. 155: 285-304. Bennike O., Björck S., Bocher J., Hansen L., Heinemeier J. and Wohlfarth B. 1999. Early Holocene plant and animal remains from north-east Greenland. J. Biogeog. 26: 667-677. Bindler R., Renberg I., Anderson N.J., Appleby P.G. and Rose N.L. 2000. Mercury accumulation rates and spatial trends in lake sediments from West Greenland: a coast to ice margin transect. Environ. Sci. Technol. 35: 1736-1741. Bindler R., Renberg I., Anderson N.J., Appleby P.G., Emteryd O. and Boyle J. 2001. Pb isotope ratios of lake sediments in West Greenland: inferences on pollution sources. Atmos. Environ. 35: 4675-4685.

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MacDonald G.M., Edwards T.W.D., Moser K.A., Pienitz R. and Smol J.P. 1993. Rapid response of treeline vegetation and lakes to past climate warming. Nature 361: 243-246. Mangerud J., Furnes F. and Jóhansen J. 1986. A 9000 year old ash bed on the Faroe Islands. Quat. Res. 26: 262-265. McGowan S., Ryves D.B. and Anderson N.J. 2003. Holocene records of effective precipitation in West Greenland. The Holocene 13: 239-249. Merkt J., Müller H., Knabe W., Müller P. and Wieser T. 1993. The early Holocene Saksunarvatn tephra found in lake sediments in NW Germany. Boreas 22: 93-100. Myneni R.B., Keeling C.D., Tucker C.J., Asrar G. and Nemani R.R. 1997. Increased plant growth in the northern high latitudes from 1981 to 1991. Nature 386: 698-702. Olafsdottir R. and Gudmundsson H.J. 2002. Holocene land degradation and climatic change in northeastern Iceland. The Holocene 12: 159-167. Orvig S. 1970. Climates of the polar regions. World Survey of Climatology. Landsberg H.E. Amsterdam, Elsevier Publishing Company 14: 355. Ouellet M., Dickman M., Bisson M. and Pagé P. 1989. Physico-chemical characteristics and origin of hypersaline meromictic Lake Garrow in the Canadian High Arctic. Hydrobiologia 172: 215-234. Palais J.M., Germani M.S. and Zielinski G.A. 1992. Interhemispheric transport of volcanic ash from a 1259 AD volcanic-eruption to the Greenland and Antarctic ice sheets. Geo. Res. Lett. 19: 801-804. Persson C. 1968. Försök till tefrokronologisk datering i fyra färöiska myrar. Geologiska Föreningens i Stockholm Förhandlingar. 90: 241-266. Pienitz R. and Vincent W. 2000. Effect of climate change relative to ozone depletion on UV exposure in subarctic lakes. Nature 404: 484-487. Quayle W.C., Peck L.S., Peat H., Ellis-Evans J.C. and Harrigan P.R. 2002. Extreme responses to climate change in Antarctic lakes. Science 295: 645. Ritchie J.C., Cwynar L.C. and Spear R.W. 1983. Evidence from northwest Canada for an early Holocene Milankovitch thermal maximum. Nature 305: 126-128. Ruddiman W.F. 2001. Earth's Climate: Past and Future. W.H. Freeman, New York, 465 pp. Rundgren M. 1998. Early-Holocene vegetation of northern Iceland: pollen and plant macrofossil evidence from the Skagi peninsula. The Holocene 8: 553-564. Ryves D.B., Jones V.J., Guilizzoni P., Lami A., Marchetto A., Battarbee R.W., Bettinetti R. and Devoy E.C. 1996. Late Pleistocene and Holocene environmental changes at Lake Albano and Lake Nemi (central Italy) as indicated by algal remains. Mem. Ist. Ital. Idrobiol. 55: 119-148. Ryves D.B., Juggins S., Fritz S.C. and Battarbee R.W. 2001. Experimental diatom dissolution and the quantification of microfossil preservation in sediments. Palaeogeogr. Palaeoclim. Palaeoecol. 172: 99-113. Ryves D.B., McGowan S. and Anderson N.J. 2002. Development and evaluation of a diatomconductivity model from lakes in West Greenland. Freshwat. Biol. 47: 995-1014. Sandgren P. and Fredskild B. 1991. Magnetic measurements recording Late Holocene maninduced erosion in S. Greenland. Boreas 20: 315-331. Schindler D.W., Bayley S.E., Parker B.R., Beaty K.G., Cruikshank D.R., Fee E.J., Schindler E.U. and Stainton M.P. 1996a. The effects of climatic warming on the properties of boreal lakes and streams at the Experimental Lakes Area, northwestern Ontario. Limnol. Oceanogr. 41: 1004-1017. Schindler D.W., Curtis P.J., Parker B.R. and Stainton M.P. 1996b. Consequences of climate warming and lake acidification for UV-B penetration in North American boreal lakes. Nature 379: 705-708. Skjelkvåle B.L., Henriksen A., Jónsson G.S., Mannio J., Wilander A., Jensen J.P., Fjeld E. and Lien L. 2001. Chemistry of lakes in the Nordic region - Denmark, Finland with Åland, Iceland, Norway with Svalbard and Bear Island, and Sweden. Acid Rain Research Report 52/2001. Nordic Council of Ministers. SNO 4391-2001.

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Skotvold T., Wartena E.M.M., Christensen G.C., Fjeld E. and Schlabach M. 1999. Organochlorine contaminants in biota and sediment from lakes on Bear Island. Akvaplan-niva report 510.1443.1. Smol J.P. 1988. Paleoclimate proxy data from freshwater arctic diatoms. Verh. Internat. Verein. Limnol. 23: 837-844. ter Braak C.J.F. 1995. Ordination. In: Jongman R.H.G., ter Braak C.J.F. and van Tongeren O.F.R. (eds), Data Analysis in Community and Landscape Ecology. Cambridge University Press, pp. 91-173. ter Braak C.J.F. and Šmilauer P. 1998. CANOCO Reference Manual and User's Guide to Canoco for Windows. Software for Canonical Community Ordination (version 4). Microcomputer Power, Ithaca, NY, USA, 352 pp. Thórarinsson S. 1944. Tefrokronologiska studier på Island. Geogr. Ann. XXVI: 1-217. Vadeboncoeur Y., Jeppesen E., Van der Zanden M.J., Schierup H.H., Christoffersen K. and Lodge D.M. 2003. From Greenland to green lakes: cultural eutrophication and the loss of benthic pathways in lakes. Limnol. Oceanogr. 48: 1408-1418. Viau A.E., Gajewski K., Fines P., Atkinson D.E. and Sawada M.C. 2002. Widespread evidence of 1500 yr climate variability in North America during the past 14 000 yr. Geology 30: 455-458. Vollenweider R.A. and Kerekes J. 1982. Eutrophication of waters. Monitoring, assessment and control. Organization for Economic Co-Operation and Development (OECD). Wagner B. and Melles M. 2001. A Holocene seabird record from Raffles Sø sediments, East Greenland, in response to climatic and oceanic changes. Boreas 30: 228-239. Wagner B., Melles M., Hahne J., Niessen F. and Hubberten H.W. 2000. Holocene climate history of Geographical Society Ø, East Greenland - evidence from lake sediments. Palaeogeogr. Palaeoclim. Palaeoecol. 160: 45-68. Wastegård S., Björck S., Grauert M. and Hannon G.E. 2001. The Mjáuvøtn tephra and other Holocene tephra horizons from the Faroe Islands: a link between the Icelandic source region, the Nordic Seas, and the European continent. The Holocene 11: 101-109. Willemse N.W. 2000. Arctic natural archives. Nederlandse Geografische Studies 272: 1-178. Willemse N.W. 2002. Holocene sedimentation history of the shallow Kangerlussuaq lakes, West Greenland. Medd. Grønland, Geosci. 41: 1-48. Willemse N.W. and Törnqvist T.E. 1999. Holocene century-sale temperature variability from West Greenland lake records. Geology 27: 580-584. Williamson C.E., Stemberger R.S., Morris D.P., Frost T.M. and Paulsen S.G. 1996. Ultraviolet radiation in North American lakes: attenuation estimates from DOC measurements and implications for plankton communities. Limnol. Oceanogr. 41: 1024-1034. Wohlfarth B., Lemdahl G., Olsson S., Persson T., Snowball I., Ising J. and Jones V. 1995. Early Holocene environment on Bjørnøya (Svalbard) inferred from multidisciplinary lake sediment studies. Pol. Res. 14: 253-275. Wright S.W., Jeffrey S.W., Mantoura R.F.C., Llewellyn C.A., Bjornland T., Repeta D. and Welschmeyer N.A. 1991. Improved HPLC method for analysis of chlorophylls and carotenoids from marine phytoplankton. Mar. Ecol. Prog. Ser. 77: 183-196.

12. PALEOLIMNOLOGICAL RESEARCH FROM NORTHERN RUSSIAN EURASIA

GLEN M. MACDONALD ([email protected]) Departments of Geography and Organismic Biology, Ecology and Evolution University of California Los Angeles, California 90095-1524, USA THOMAS W.D. EDWARDS ([email protected]) Department of Earth Sciences University of Waterloo Waterloo, Ontario N2L 3G1, Canada BRUCE GERVAIS ([email protected]) Department of Geography California State University Sacramento, California 95819-6003, USA TAMSIN E. LAING ([email protected]) Environmental Sciences Group Royal Military College of Canada PO Box 17000 Stn. Forces Kingston, Ontario K7K 7B4, Canada MICHAEL F.J. PISARIC ([email protected]) Department of Geography and Environmental Studies Carleton University Ottawa, Ontario K1S 5B6, Canada DAVID F. PORINCHU ([email protected]) Department of Geography Ohio State University Columbus, Ohio 43210-1002, USA

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G.M. MACDONALD ET AL. JEFFREY A. SNYDER ([email protected]) Department of Geology Bowling Green State University Bowling Green, Ohio 43403, USA NADIA SOLOVIEVA ([email protected]) Environmental Change Research Centre (ECRC) University College London London, WC1H 0AP United Kingdom PAVEL TARASOV ([email protected]) Department of Geography Moscow State University Moscow, 119899 Russia and BRENT B. WOLFE ([email protected]) Department of Geography and Environmental Studies Wilfrid Laurier University Waterloo, Ontario N2L 3C5, Canada

Key words : Paleolimnology, Chironomids, Diatoms, Isotopes, Palynology, Climate change, Quaternary, Arctic, Eurasia, Russia

Introduction As access to northern Russia has become easier, and as more powerful analytical tools and interpretative approaches have been developed for arctic paleolimnology (e.g., MacDonald et al. 2000a), the pace and geographic extent of paleolimnological research in the far north of Russian Eurasia has increased markedly. Recent work by joint Russian-foreign teams has included analysis of biological proxies such as diatoms and chironomids to reconstruct past limnological and climatological conditions (e.g., Kienel 1999; Laing et al. 1999a; Snyder et al. 1997, 2000; Laing and Smol 2000; Solovieva 2000; Porinchu and Cwynar 2002; Solovieva and Jones 2002). Additional new studies have also examined sedimentological and geochemical changes in lake sediments to reconstruct sea-level changes, limnological and hydrological conditions and climate (e.g., Snyder et al. 1997; Wolfe et al. 2000; Jones et al. 2004). Other recent work has consisted of the collation and interpretation of existing data from previously published

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and unpublished Russian studies (e.g., Davydova and Servant-Vildary 1996; Harrison et al. 1996; Tarasov and Harrison 1998). In this chapter we present some examples of research that have led to the creation of new biological and stable isotope records from lakes in far northern Russian Europe and Siberia. We also present the results of synthetic analyses of existing Russian data. The examples are largely from the Kola Peninsula in the west and along the Lower Lena River in the east. These examples are admittedly biased towards studies in which the data are available in English language journals and reflect also the contributions and interests of this chapter’s authors. Despite this bias, the examples provide a reasonable picture of the analytical approaches and geographic scope (Figure 1) of recent paleolimnological studies in the far north of Russian Eurasia.

Figure 1. Location of study sites described in this chapter. Location of modern treeline, maximum Holocene extension and 9000 yr BP shoreline (after MacDonald et al. 2000c).

Recent analyses using biological evidence The Kola Peninsula, Russia The Kola Peninsula (Figure 1) is an excellent region for paleolimnological studies as it contains numerous small lakes of glacial origin and displays steep climatic and vegetation gradients. The presence of the Barents Sea, relatively warm in winter and cool in summer, produces a general moderating impact on climate, but increases the climatic differences between the coast and the interior. The mean January and July temperatures for the Barents coastline and the central interior of the Kola are -9ºC and -13ºC, and 9ºC and 11ºC, respectively (Arctic Atlas 1985). The mean annual total precipitation ranges from 500 to 700 mm. The ecotone between Scots pine (Pinus sylvestris) dominated northern coniferous forest and birch forest-tundra runs across the peninsula roughly parallel to the northern coast. A tundra zone of variable width lies between the northern limits of the birch forest-tundra zone and the coastline.

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Most earlier paleoecological studies in the Kola Peninsula used pollen analysis to reconstruct major changes in regional vegetation and climate dynamics during the postglacial period (e.g., Aario 1943; Sorsa 1965; Lebedeva and Pervuninskaya 1973; Lebedeva et al. 1987; Yelina et al. 1995; Kremenetski and Patyk-Kara 1997). Both pollen and diatom analyses were used by Lebedeva et al. (1989), Kagan et al. (1992) and Davydova and Servant-Vildary (1996) to reconstruct the Holocene environmental history of the Kola Peninsula. However, the analytical resolution in the above studies is low, most records are undated, and important paleolimnological indices, such as the diatom and lake ecosystem responses to Holocene climate changes, were not investigated. The general pattern of climate change suggested by these previous studies includes relatively warm conditions during the early Holocene and cooling from ca. 6000 to 4000 yr BP (all ages are reported as 14C years Before Present with present taken to be AD 1950). However, during the Holocene, vegetation and climate changes in the Kola Peninsula and Fennoscandian region may have been diachronous. For instance, there is some evidence of a considerable time difference between spruce and pine expansion in the east and west of Fennoscandia (Kremenetski et al. 1999; Seppä and Hammarlund 2000). A steep north-south humidity gradient may also have existed in the midHolocene (Eronen et al. 1999). Modern surface sample studies Solovieva (2000) analysed surface sediments from 25 Kola Peninsula lakes located along the vegetation gradient between 67º30’N and 69º14’N and 28º40’E and 36º40’E. The sediments were analysed for pollen and diatoms in order to generate pollen-climate and diatom-pH inference models. In another study, a similar set of samples was taken from 31 lakes along a larger transect extending from the northern to southern coast of the peninsula to study the geochemistry, pollen, stomate, and chironomid content of lake sediments (Blom et al. 1999; Gervais and MacDonald 2001). A diatom-pH model based on weighted-averaging regression was generated using 24 of 25 lakes sampled by Solovieva (2000). Although the predictive statistics of the resulting KOLA WA model are comparable to those from other studies (RMSEP is 0.37), the correlation coefficient r2 is somewhat lower (0.52) than in the other diatom-pH models. However, the diatom optima generated by the KOLA model are generally in agreement with the optima derived by other models and the Kola inference model can be used to reconstruct pH for a relatively narrow range between 6.0 and 7.0 (Solovieva 2000). The detailed analysis of pollen and stomata from Kola Peninsula surface sediments using principal components analysis (PCA) showed that the lakes located in the tundra, forest-tundra transition and forest zones could generally be differentiated on the basis of differences in their pollen assemblages (Gervais and MacDonald 2001) (Figure 2). In addition, there was a significant northward decline in the ratio of arboreal to nonarboreal pollen in the lake sediments. Finally, it was found that stomates of Pinus were consistently present in lakes within the modern distribution of pine trees, but absent from lakes that lie north of the pine treeline (Gervais and MacDonald 2001).

Figure 2. The pollen and stomate content of modern lake sediments from a transect of sites across the Kola Peninsula (after Gervais and MacDonald 2001).

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The modern sediment studies outlined above show the potential of Kola lakes to provide sensitive records of past limnological conditions and terrestrial vegetation. These surface sediment studies provide useful comparative data for the paleolimnological investigations that are described below. The paleolimnological record from coastal lakes on north-central Kola Peninsula Abundant coastal lakes, emergent above sea-level due to glacio-isostatic rebound, occur on the Murman coast of the northern Kola Peninsula. The study site discussed here represents such a setting situated on the central Murman coast near Yarnyshnoe Bay (69q04’N, 36q04’E). The lakes in this region are surrounded by tundra vegetation with scattered stands of Betula pubesce ns present. The modern Pinus treeline occurs about 50 km to the south. The region’s proximity to the treeline ecotone and apparent sensitivity of the local climate to North Atlantic circulation make the site particularly well suited for paleoclimatic investigations. Compared to similar settings on the northern Norwegian coastline (e.g., Donner et al. 1977; Vorren et al. 1988; Corner and Haugane 1993), the sediment records from lakes of the northern Kola Peninsula have been less extensively studied. Recently, a series of emerged lakes on the northwestern Kola Peninsula near the Norwegian border were studied to decipher the sea-level history of that area (Corner et al. 1999). Also, Grönlund and Kauppila (2002) used a core record from Soldatskoe Lake on northeastern Kola Peninsula near Drozdovka Bay to relate the diatom-inferred nutrient status of the lake to the occupation of nearby Neolithic dwellings. Snyder et al. (1997, 2000) sampled lakes along an elevational transect from modern sea-level to the limit of postglacial marine submergence (locally 65 m a.s.l.). Coring was conducted from a lake ice platform using a percussion-driven piston coring system (Nesje 1992). Cores recovered from each lake contained marine sediment overlain by between 1 and 3 m of lacustrine sediment. Basal radiocarbon ages indicate that the region was deglaciated and marine sedimentation in the basins commenced by 11,000 yr BP. The transition from marine to lacustrine sediment is identified in the cores by analysis of sediment physical properties and diatom assemblages. The records from lakes at 54, 41 and 7 m a.s.l. indicate a relatively abrupt transition from marine to lacustrine conditions. A lake located at 22 m a.s.l. exhibits a more gradual transition from marine to lacustrine conditions, indicating an extended period of meromixis in the lake. The lakes at 54 and 41 m a.s.l. bracket the elevation of the prominent late-glacial shoreline on the Murman coast (locally 48 m a.s.l.). Radiocarbon ages from these lakes provide the first direct dating of this shoreline at ca. 10,500 to 10,300 yr BP, confirming its correlation to the Younger Dryas Stade (between ca. 11,000 to 10,000 yr BP). The longest of the lacustrine core records (Lake Yarnyshnoe-3 at 54 m a.s.l) has been analysed for diatoms, pollen, stomates, stable isotopes, and chironomids. The diatom record will be the focus here (Figure 3). Abrupt transitions in the diatom flora occur at ca. 10,500 yr BP at the marine-lacustrine transition and between ca. 9500 and 9000 yr BP coincident with the transition from organic-poor to organic-rich sediment in the core. Above this interval, alkaliphilous Fragilaria spp. decline in favour of oligotrophic species such as Achnanthes minutissima and Tabellaria flo cculosa. In addition to the general trend toward more oligotrophic conditions, a more pronounced change occurs

Figure 3. Diatom stratigraphy from Lake Yarnyshnoe-3 on the northern coastline of the Kola Peninsula (after Snyder et al. 2000).

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between ca. 5000 and 4000 yr BP. During this interval Fragilaria pseudoconstruens and F. pinnata decline to near zero, while acidophilous taxa such as Tabellaria flocculosa, Cymbella grac ilis, C. gaeumannii , and Frustulia rh omboides all increase. A similar transition also has been observed in the diatom records from Soldatskoe Lake (Grönlund and Kauppila 2002) and from Chuna Lake on the western Kola Peninsula (Solovieva 2000; Solovieva and Jones 2002). The changes in the diatom assemblage approximately coincide with the decline in pine inferred from pollen, stomate, and macrofossil records from the core and a nearby site (MacDonald et al. 2000b; Snyder et al. 2000; Gervais et al. 2002). The diatoms may indicate a decrease in nutrient delivery to the lake associated with inferred cooler and moister climate and associated development of peat in the lake's catchment (Snyder et al. 2000; Wolfe et al. 2003). The paleolimnological record from lakes on the central Kola Peninsula Recently, the Holocene treeline dynamics in the central Kola, including the Khibiny Mountains, have been studied by Kremenetski et al. (1999), MacDonald et al. (2000b) and Gervais et al. (2002) using pollen records from peats and lakes, and pine macrofossils recovered from lakes. Poteryanny Zub Lake (68º48’N, 35º19’E) is a key site located in the birch forest-tundra zone approximately 25 km north of the modern pine treeline. Pollen, stomate and wood macrofossil records from the lake provide an environmental history spanning from ca. 10,000 yr BP to the present (MacDonald et al. 2000b; Gervais et al. 2002). The lake provides a well dated sequence of climate and vegetation change that includes the initial establishment of herb and small shrub tundra following deglaciation at around 10,000 yr BP, the establishment of birch tundra or birch forest-tundra between 9500 and 9000 yr BP, the development of pine forest or forest-tundra between 8000 and 3500 yr BP, and the establishment of the modern birch forest-tundra following 3500 yr BP. The changes in vegetation can likely be attributed to orbital-induced warming following the close of the last glaciation which resulted in warmer summers than today and a northward extension of treeline until around 3500 yr BP (MacDonald et al. 2000b,c; Gervais et al. 2002). Increasing amounts of pollen and spores from taxa such as Ericaceae, Alnus and Sphagnum, and the fact that a number of the pine macrofossils from the site are stumps recovered below the water line from submerged peat banks that are eroding, suggest an increase in moisture from the mid-Holocene to the present. More explicit paleolimnological evidence comes from a sediment core from Chuna Lake (67°57’N, 32°29’E) (Solovieva 2000; Solovieva and Jones 2002). The lake is currently located in an area of alpine tundra above the regional treeline. The core was studied for pollen, diatoms and sediment geochemistry in order to infer postglacial environmental changes and to investigate associated responses of the lake ecosystem (Solovieva 2000; Solovieva and Jones 2002). The main features of the Chuna pollen stratigraphy (Figure 4) include a decline from maximum Betula abundance between ca. 8000 and 7000 yr BP, a high abundance of Pinus between ca. 7000 and 4200 yr BP, increase in Picea from ca. 3800 yr BP, and a rise in Sphagnum during the last 2000 years. These changes approximately correlate with the record from Poteryanny Zub Lake (68°49’N, 35°19’E) and other regional pollen records. A Betula forest was likely present in the vicinity of Chuna Lake during the early to mid-Holocene as Betula pollen

Figure 4. Pollen, loss-on-ignition (LOI), geochemistry and diatom stratigraphy from Chuna Lake on the central Kola Peninsula (after Solovieva 2000; Solovieva and Jones 2002).

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accumulation rates were above modern values between 7000 and 5000 yr BP. The absence of pine macrofossils from the Chuna sediments suggests that the Betula foresttundra was not replaced by the Pinus forest in the catchment of the lake even at the height of Holocene warming. The Chuna Lake data imply that the late Holocene climate cooling may have been evident slightly earlier at this site than at more westerly sites in Fennoscandia or other areas of the Kola Peninsula. The limnological changes evident in the diatom record from the Chuna Lake core (Figure 4) may be attributed to the major patterns of Holocene climate changes evident in the Kola Peninsula pollen, stomate and macrofossil records (Solovieva 2000; Solovieva and Jones 2002). The period between 8000 and 7000 yr BP reflects a recently deglaciated alpine lake dominated by the pioneer diatom taxa Fragilaria and Stauroforma. Bare, highly minerogenic soils and undeveloped vegetation were characteristic for the Chuna catchment causing relatively high catchment erosion, which might have affected water transparency and therefore diatom composition. Between 7000 and 4500 yr BP, benthic Fragilaria taxa (not shown) were replaced by a succession of planktonic and mero-planktonic Aulacoseira taxa. The occurrence of planktonic and mero-planktonic diatom taxa and increased diversity in Chuna Lake correlates with the development of a Betula forest. The development of Betula in the lake catchment probably contributed to decreased erosion, whereas leaching (reflected by Ca/K ratio) and sediment organic content increased. Abundant remains of bryophytes, which grow on the bottom of the lake at present, appeared in the Chuna sediment between 7000 and 4700 yr BP. All of the above paleolimnological indices point to a more productive lake ecosystem, likely generated by warmer climate, at that time. After ca. 4300 yr BP, planktonic diatoms and Stauroforma disappeared, and from 3800 yr BP the lake became dominated by benthic Brachysira, Frustulia and Pinnularia taxa (Solovieva and Jones 2002). During the last 3800 years, diatom-based proxies show no clear response of the lake to the vegetational and erosional changes suggested by the pollen record. The most plausible explanation for this may be the combination of different climate-related and ecological factors affecting the terrestrial and aquatic environments. Progressive natural acidification may have favoured the more acidophilous Brachysira and Pi nnularia taxa, whilst the retreat of forest-tundra led to a decrease in organic supply to the lake as tundra soils are generally poorer in organics. The development of Sphagnum mosses around the lake may be one of the reasons why Brachysira, Pinnularia and Frustulia taxa out-competed Stauroforma in the second half of the Holocene, as Sphagnum is the microhabitat preferred by Brachysira and Frustulia taxa in Chuna Lake at present. Lower Lena River Delta The fact that portions of northeastern Siberia were not glaciated during the Last Glacial Maximum (LGM), coupled with the presence of lakes formed by tectonic activity and other non-glacial processes, makes the area of the lower Lena River of particular interest for paleolimnological studies. The climatically controlled ecotone between closed boreal forest and tundra is situated in the region, providing the potential for securing records that are sensitive to past changes in climate and enabling assessment of

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the impact of such changes on terrestrial and aquatic environments. The climate of the Lena River region is characteristic of a cold continental regime. Mean January temperatures range from -35°C to -40°C whereas July temperatures range from 4°C to 12°C. Mean annual precipitation is generally less than 300 mm (Arctic Atlas 1985). There has been little earlier paleolimnological work along the lower Lena River. Some lake cores were obtained to the east of the Lena River near Tiksi Bay to date the basal organics in an attempt to determine a minimum age for deglaciation (Grosswald et al. 1992). These samples dated to the Holocene. Two high-latitude diatom records from peat cores also exist from northeastern Siberia: one from the Labaz Lake area in the southeastern part of the Taymyr Peninsula (Kienel et al. 1999), and the other from the northern boreal forest of the Lena River region, approximately 330 km south of the Lena River Delta (Rühland et al. 2000). Diatom shifts in these records mostly appear to reflect moisture availability and hydrological change associated with peatland ontogeny and climatic change. For both records, diatom assemblages indicative of higher waterlevels were linked to warmer and wetter climatic conditions in the early Holocene. Modern surface sample studies Surface sediment samples were obtained from a transect of 31 small lakes, extending across ca. 250 km from the tundra zone of the Lena River Delta in the north to the closed coniferous forest zone of the lower Lena River in the south. The sediments were analysed for fossil pollen, stomates, diatoms and chironomids. The pollen assemblages (Pisaric et al. 2001a) from the tundra lakes in the calibration set were dominated by shrubs, herbs and graminoids (Betula, Alnus, Cyperaceae and Poaceae). Forest-tundra lakes were also dominated by Betula and Alnus, but contained lower percentages of Artemisia. Although dominated by Betula and Alnus pollen, lakes from the forested region in the south contained small, but consistent percentages of Larix pollen. A PCA of the pollen data indicates that forest and tundra sites can be clearly distinguished from one another based on their modern pollen spectra, but foresttundra sites cannot. Larix stomates were abundant in almost all samples from regions where the trees currently grow. However, some Larix stomates were found in tundra lakes, likely due to the redeposition of mid-Holocene aged material from eroding peat banks. The diatom assemblages from modern lake sediments were analysed as part of a series of such studies that included sites across treeline near the Pechora River in northern Russia, the western Taymyr Peninsula in western Siberia and the Lena River region (Duff et al. 1998; Laing et al. 1999b; Laing and Smol 2000). Diatoms were recovered from 21 lakes distributed along a transect from the tundra to coniferous forest zones along the Lena River (Figure 5). The modern assemblages included a total of 292 species. It was found that small benthic species, particularly Fragilaria spp., such as Fragilaria pinnat a, were more common in tundra lakes. As the mean depth of the sampled lakes was 2.2 m, there was generally a higher abundance of benthic species in the assemblages. Some planktonic taxa, such as members of the genus Cyclotella, were more common at the deeper, forested sites. Comparison of the Lena assemblages with others from across treeline in northern Russian Eurasia indicates that the Lena lakes are typified by higher numbers of Achnanthes minutissima. The diatom assemblages from

Figure 5. Diatoms from modern lake surface samples from a transect of lakes across the treeline zone near the Lena River (after Laing and Smol 2000).

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the forest, forest-tundra and tundra zones displayed clear differences, and canonical correspondence analysis (CCA) of the Lena assemblages indicated that depth-adjusted lake temperature, along with particulate organic matter, dissolved inorganic carbon, Fe2+, SO42- and Cl-1 were correlated with the differences in the diatom assemblages (Laing and Smol 2000). Analysis of the chironomid remains from the surface sediments using redundancy analysis (RDA) (Figure 6) revealed that the chironomids were responding to a climatevegetation gradient, resulting in two distinct groupings of taxa. Taxa such as Paracladius, Stictochironomus, Abiskomyia, and Parakiefferiella nigra comprise the first group. These taxa were associated with shallow, oligotrophic tundra lakes. The second group consists of taxa such as Microtendipes, Glyptotendipes, Psectrocladius and Zalutschia za lutschicola. These taxa were more commonly found in deeper,

Figure 6. An RDA ordination of chironomid taxa and environmental variables from a transect of lake surface sediment samples taken across treeline near the lower Lena River. The ordination diagram is based on redundancy analysis (RDA) of a 30 lake data set. Environmental variables that exert significant influence on chironomid distribution are illustrated with arrows. Qualitative variables describing vegetation classes that were included passively in the ordination analysis, are positioned based on their centroid value and are represented by open squares. Abbreviations for chironomid taxa: AB - Abiskomyia; CH - Chironomus; CL - Cladopelma; CA - Corynocera ambigua; CI - Corynocera oliveri ; CT - Corynoneura/Thienemanniella; CO - Cricotopus/ Orthocladius; DI - Dicrotendipes; GL - Glyptotendipes; HE - Heterotrissocladius; HO Hydrobaenus/Oliveridia; LP - Limnophyes/Paralimnophyes; MP - Micropsectra; MI Microtendipes; PA - Paracladius; PB - Parakiefferiella cf. bathophila; PN - Parakiefferiella nigra; PI - Pentaneurini (Other); P1 - Psectrocladius grp. Allopsectrocladius/Mesopsectrocladius; P2 - Psectrocladius grp. Monopsectrocladius; P3 - Psectrocladius grp. Psectrocladius; PR Procladius; SE - Sergentia; SZ - Stempellinella/Zavrelia; ST - Stictochironomus; TA - Tanytarsina; ZA - Zalutschia; ZZ - Zalutschia zalutschicola.

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productive forest lakes. A distinct change in the distribution of two Corynocera species, C. ambigua and C. oliveri, corresponds to northern treeline, with C. ambigua present only in warm forested sites south of treeline and C. oliveri essentially restricted to the colder tundra lakes. C. oliveri has been associated predominantly with cold, northern lakes and ponds and cold, oligotrophic, high-elevation lakes in the Sierra Nevada of California (Oliver and Roussel 1983; Olander et al. 1999; Porinchu and Cwynar 2000; Porinchu et al. 2002). Walker and Mathewes (1989) note that coastal, upper subalpine lakes and ponds of British Columbia (Canada) are dominated by head capsules resembling those attributed in Wiederholm (1983) to C. oliveri. The paleolimnological record from Dolgoe Ozero Paleoenvironments have been reconstructed from pollen, stomate, chironomid, and diatom assemblages preserved in a long sediment core from a lake informally named Dolgoe Ozero (71°52’N, 127°04’E), near the delta of the Lena River (Figure 1). Isotope analysis has also been conducted on the core and will be discussed later. Dolgoe Ozero is a relatively small (surface area = 84 ha), shallow (maximum depth = 4 m), circumneutral (pH = 7.4) and dilute (conductivity = 25 µS/cm) lake (Laing et al. 1999). No evidence of glacial deposits or processes was found within the catchment, suggesting that the region was not glaciated during the Last Glacial Maximum or perhaps at any time in the Pleistocene. Dolgoe Ozero is located within the continuous permafrost zone (Davydova and Rakovskaya 1990). The lake is presently surrounded by herb and graminoid tundra with some shrub birch (Betula nana) and extremely sparse larch (Larix dahurica) forest-tundra in the vicinity, with continuous forest consisting of larch and spruce found about 100 km to the south. The pollen and stomate stratigraphy from the Dolgoe Ozero core provides a detailed record of climate and vegetation change spanning the period of at least 12,000 yr BP to the present (Pisaric et al. 2001b) and is useful for the interpretation of past limnological conditions. The record suggests that the late-glacial vegetation was dominated by a sparse cover of small shrubs, herbs and graminoids (Salix, Artemisia, Dryas, Poaceae, Cyperaceae). By about 12,000 yr BP, climate had begun to warm and the earlier vegetation was replaced by shrub birch tundra. The Younger Dryas Stade is clearly marked by a sharp decline in birch and a reversion to the earlier small shrub, herb and graminoid vegetation between ca. 11,000 and 10,000 yr BP (Pisaric et al. 2001b). Climate warmed again following the Younger Dryas and shrub birch tundra was the dominant vegetation until ca. 8500 yr BP when the pollen and stomate record from the lake and a wood macrofossil record from the vicinity (MacDonald et al. 2000c) indicate the establishment of larch and spruce forest. At ca. 3500 yr BP, birch, small shrub, herb and graminoid tundra typical of modern conditions replaced the coniferous forest vegetation around the site. Zones for the diatom stratigraphy were identified by stratigraphically constrained cluster analysis (Laing et al. 1999a). Past changes in lake water alkalinity were reconstructed with weighted-averaging techniques using a diatom-alkalinity model developed from the Northwest Territories, Canada (Rühland 1996). The lack of close analogs for some sections of the core suggests that the absolute values for reconstructed alkalinity should be regarded with caution.

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Early diatom assemblages (Figure 7) prior to 8500 yr BP were dominated by small benthic Fragilaria taxa, which are commonly found in tundra lakes (e.g., Laing and Smol 2000), suggesting cool climatic conditions at this time. Diatom-inferred alkalinity was generally high during this period, most likely reflecting mineral weathering (e.g., carbonates) from poorly developed soils in the catchment. The establishment of a larch forest in the catchment basin between ca. 8500 and 7000 yr BP (Pisaric et al. 2001b) corresponded with an initial period of fluctuating limnological conditions, as evidenced by a series of shifts in dominant diatom taxa. The shift in catchment vegetation corresponded with a decline in diatom-inferred alkalinity, most likely reflecting soil development and increased organic inputs from the catchment. Stable diatom assemblages dominated by Achnanthes species and consistently low diatom-inferred alkalinity prevailed between 7000 and 3500 yr BP, reflecting the influence of organic runoff associated with a fully developed forest. Concurrent with the return to cooler conditions and a shift back to shrub tundra after 3500 yr BP, diatom communities reverted back to a Fragilaria-dominated assemblage and diatom-inferred alkalinity increased. Recent conditions are associated with declining alkalinity and a minor change in diatom assemblages, most likely reflecting an influx of humic substances from the remnants of forest soils and peats in the catchment basin. Overall, the diatom record for Dolgoe Ozero indicates close coupling between terrestrial and aquatic environments throughout the Holocene, most likely as a result of changes in organic runoff associated with vegetation change and soil development in the catchment basin. The period of declining diatom-inferred alkalinity at the onset of the forested period suggests that soil development and vegetation establishment at Dolgoe Ozero occurred over approximately 1400 years, indicating a gradual transition to a fully forested catchment. A diatom record from Lama Lake (Kienel 1999), located in the forest-tundra zone of the western Taymyr Peninsula, also indicates that significant changes in the diatom assemblages occurred concurrent with a mid-Holocene shift to a forested catchment (Hahne and Melles 1997). The forested period coincides with an increase in planktonic diatom taxa, most likely reflecting increases in surface water temperature corresponding with early to mid-Holocene climatic warming (Kienel 1999). The difference in limnological responses between the two records can be explained by the much larger size of Lama Lake, rendering the lake less sensitive than Dolgoe Ozero to shifts in the chemical composition of catchment runoff as a result of environmental change. Chironomid analysis of the Dolgoe Ozero core supports the pollen, stomate and diatom evidence indicating that significant climate fluctuations occurred during the last 12,000 yr BP (Figure 8). Cooling during the Younger Dryas is evidenced by decreases in temperate chironomid taxa such as Microtendipes and Corynocera ambi gua, and a concurrent increase in cold water taxa such as Hydrobaenus/Oliveridia and Parakiefferiella nigra. The early Holocene, between 10,000 and 6400 yr BP, was marked by rapid climate amelioration as indicated by an increase in temperate taxa such as Microtendipes and C. am bigua and the disappearance of essentially all cold water chironomids. However, a dramatic increase in the abundance of cold water types at approximately 8000 yr BP suggests that changes in the paleoenvironment, possibly related to decreasing summer insolation and a heightened response by lake ecosystems, may have occurred even earlier. The chironomid record suggests that climate

Figure 7. Diatom stratrigraphy from Dolgoe Ozero (after Laing et al. 1999).

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Figure 8 . Chironomid percentage diagram from Dolgoe Ozero. Influx of cold water taxa and temperate taxa in head capsules/cm2/yr also indicated. Psectrocladius is split into three groups; Psectrocladius subgenera Allopsectrocladius/Mesopsectrocladius, Psectrocladius subgenus Monopsectrocladius and Psectrocladius subgenus Psectrocladius , which are identified in the diagram as Psectrocladius subg. Allo./Meso., Psectrocladius subgenus Mono and Psectrocladius subgenus Psectro, respectively. Values for subtribe Tanytarsina do not include Micropsectra, Corynocera ambigua or C. oliveri.

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amelioration was terminated by 6400 yr BP when the first stage of a two-stage cooling began. The first stage is marked by an increase in Paracladius and Hydrobaenus/Oliveridia, with increases in cold water chironomid taxa, such as Abiskomyia and Microspectra, at approximately 3500 yr BP marking the initiation of the second stage. The modern chironomid assemblage was established at approximately 1500 yr BP with the disappearance of Paracladius from the record. The stomate (Pisaric et al. 2001b) and the macrofossil evidence (MacDonald et al. 2000c) suggest forest developed in the area between 8000 and 3500 yr BP. Inferences made from the chironomid record suggest that climate was significantly warmer than present and, possibly, sufficiently warm to support trees beginning roughly at 10,000 yr BP. It is not known if the later development of boreal forest was retarded due to differences between lakes and terrestrial vegetation in terms of response to early Holocene climate changes or a lag in vegetation response due to non-climatic factors. A slow rate of soil development (see diatom evidence above) has been invoked to explain the apparent lag in vegetation response to late glacial warming in Britain (Pennington 1986). Stable isotope studies of lake sediments from across northern Russian Eurasia An active area of research in physical and chemical approaches to paleolimnology has been the application of stable isotope analysis to examine past changes in arctic lake productivity and hydrologic balance (e.g., MacDonald et al. 1993; Wolfe et al. 1996, 1999, 2000; Hammarlund et al. 1997; see Edwards et al., this volume). Analysis of bulk organic carbon and nitrogen content, stable isotope composition, and cellulose oxygen isotope composition from several lakes along the northern boreal treeline in Russia, have provided the basis for local and regional reconstruction of Holocene paleohydrology and moisture conditions, as well as changes in watershed carbon and nitrogen cycles (Wolfe et al. 1999, 2000, 2003). Research has focused on three transects in northern Russia including the Kola Peninsula, the lower Yenisey River, and the lower Lena River. The biological records of some of the sites (Poteryanny Zub Ozero and Dolgoe Ozero) have already been discussed in this chapter. Organic matter from the lakes has provided an alternative substrate to the limited distribution of lacustrine biogenic and inorganic carbonates, the more traditional archives of isotope-derived hydrological information in temperate and tropical environments. Stratigraphic oxygen isotope analysis of fine-grained cellulose has been central to these studies, serving as a proxy for the oxygen isotope history of lake water and a record of paleohydrological change.

G18O analysis Cellulose-inferred lake water G18O (G18Olw) profiles have been developed from lake sediment cores using aliquots of the fine (< 500 µm) size-fraction of acid-washed bulk sediment that underwent solvent extraction, bleaching and alkaline hydrolysis in order to concentrate cellulose and remove other organic constituents, plus hydroxylamine leaching to remove iron and manganese oxyhydroxides where necessary (Wolfe et al.

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2001). Stable isotope values are expressed in G-notation as deviations in per mil (‰) from international standards. For oxygen, G18O values are expressed relative to the Vienna Standard Mean Ocean Water (VSMOW) standard, such that G18O = ((Rsample/RVSMOW)-1)1000, where R is the 18O/16O ratio (see Edwards et al., this voume). Reconstructed Holocene cellulose-inferred lake water G18O (G18Olw) histories, generated using a fixed cellulose-water fractionation factor (Wolfe et al. 2001), have been derived for Lake Yarnyshnoe-3 (69°04’N, 36°04’E) and Poteryanny Zub Lake (68°49’N, 35°19’E) on the Kola Peninsula, Derevanoi Lake (69°14’N, 86°34’E) in the lower Yenisey River region, and Dolgoe Lake (71°52’N, 127°04’E) on the Lena River Delta (Figure 1). A late Holocene G18Olw record spanning the past 4000 14C yr BP has also been generated for Middendorf Lake (70°22’N, 87°33’E) in the lower Yenisey River region (Wolfe et al. 1999, 2000). The G18Olw records are considered to be mainly a function of: (1) changing source water (i.e., precipitation) 18O composition, which is largely controlled by temperature-dependent fractionation during condensation of atmospheric vapour, seasonal distribution of precipitation, and various aspects of air mass history; and (2) hydrological effects, which are frequently dominated by evaporation resulting in lake water 18O-enrichment. To deconvolute these two major paleo-hydrometeorological parameters that control the G18Olw records, coarsely resolved profiles of precipitation G18O (G18Op) have been estimated for each of the three regions (Figure 9). For the Kola Peninsula, G18Op has been estimated at 500-year intervals by fitting a curve through the most negative G18O data points from the more variable and 18 O-depleted Lake Yarnyshnoe-3 record, which are interpreted to represent intervals of minimal evaporative G18O-enrichment (Wolfe et al. 2003). For the Yenisey and Lena study regions, supplemental isotopic information from modern precipitation and surface water, frozen peat porewater from peatlands near the study lakes, other paleoclimatic data, and the G18Olw profiles, have been used to constrain the G18Op reconstructions, also at 500-year intervals (Wolfe et al. 2000). Notably, shifting G18Olw-G18Op offsets provide a measure of changing lake water evaporative 18O-enrichment, which is interpreted to be a primary function of changing local water balance in response to shifting moisture regimes. As described in more detail below, early Holocene inferred changes in moisture conditions from these records closely correspond with boreal forest expansion (represented by macrofossil (MacDonald et al. 2000c) histogram plots in Figure 9) along the northern Eurasian coast. Kola Peninsula On the Kola Peninsula, increasing G18Olw-G18Op offset at Lake Yarnyshnoe-3 and Poteryanny Zub Lake from 9500 to between 8000 and 7500 yr BP is interpreted to reflect increasing evaporative 18O-enrichment in response to drier conditions that rapidly developed in the earliest Holocene. Conversely, reduced G18Olw-G18Op offset at Lake Yarnyshnoe-3 and Poteryanny Zub Lake at about 7000 yr BP is considered to signal a change to reduced evaporative 18O-enrichment and is correlated with the main interval of pine expansion on the Kola Peninsula (MacDonald et al. 2000b; Gervais et al. 2002). The change to wetter conditions may have been associated with reduced winter sea ice cover, an enhanced winter/summer precipitation ratio and increased input of 18O-depleted snowmelt on the basis of a similar decline in the reconstructed G18Op

Figure 9. West-east transect along northern Russia of Holocene lake sediment cellulose-inferred lake water G18O (raw data and three-point running mean plotted), reconstructed precipitation G18O (estimated at 500-yr intervals), and bulk organic G13C (raw data and three-point running mean plotted) records versus 14C yr BP. Sites include Poteryanny Zub Lake and Lake Yarnyshnoe-3 on the Kola Peninsula, Derevanoi Lake in the lower Yenisey River region, and Dolgoe Lake on the Lena River Delta (see Wolfe et al. 1999, 2000, 2003). Also shown are tree macrofossil records for each region, representing the number of samples found beyond conifer treeline (MacDonald et al. 2000c).

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profile (Wolfe et al. 2003). After 6000 yr BP, the G18Olw-G18Op offset increases once again at Lake Yarnyshnoe-3, suggesting that drier conditions similar to the early Holocene developed near the coast of the Kola Peninsula, while the corresponding record at the more inland Poteryanny Zub Lake indicates that relatively stable hydrological conditions characterized this location since the mid-Holocene. A trend towards reduced G18Olw-G18Op offset at Lake Yarnyshnoe-3 after 5000 yr BP suggests a gradual shift to wetter conditions. Lower Yenisey River region In the lower Yenisey River region, the early Holocene G18Olw-G18Op offset record at Derevanoi Lake attains a minimum similar to the Kola sites at about 7000 yr BP, reflecting reduced evaporative 18O-enrichment and moister conditions. This corresponds with the arrival of Picea in the catchment (Clayden et al. 1997), an important boreal tree genus that is favoured by warm and moist conditions. From 7000 to 1500 yr BP, increasing G18Olw-G18Op offset is probably related to evaporation becoming a more important component of the lake water balance of Derevanoi Lake and its large watershed in response to increasing aridity (Wolfe et al. 2000). At about 1000 yr BP, a shift to a lower G18Olw-G18Op offset indicates a return to reduced evaporative 18Oenrichment and a wetter climate. Lower Lena River region At Dolgoe Lake on the Lena River Delta, high G18Olw-G18Op offset from before 9000 about 6000 yr BP suggests dry conditions prevailed near the lower Lena River during the early Holocene, in contrast to the equivalent record from Derevanoi Lake. Reduced G18Olw-G18Op after 7000 yr BP suggests reduced evaporative 18O-enrichment. Wetter conditions developing after about 7000 yr BP are also suggested by a decline in a diatom-inferred alkalinity record from Dolgoe Lake (Laing et al. 1999a) and the local presence of Picea (Pisaric et al. 2001b). After 4000 yr BP, close correspondence between the Dolgoe Lake G18Olw record and the late Holocene G18Olw record from Middendorf Lake in the lower Yenisey River region suggests coherent submillennialscale changes in moisture conditions in these two sectors of northern Russia (Wolfe et al. 2000).

G13C analysis Additional analysis of the bulk organic carbon stable isotope composition (G13Corg) has also been used as an indicator of changes in carbon cycling in northern Russian Eurasian lakes (Figure 9). Although a number of factors can control the carbon isotope composition of lacustrine organic matter (e.g., Håkansson 1985; Meyers and LallierVergès 1999), centennial- to millennial-scale changes in Holocene carbon cycling at lakes near northern treeline appear to be mainly driven by changes in watershed vegetation, as well as delivery of dissolved inorganic carbon from soil decomposition and hydrology (Wolfe et al. 1996, 1999, 2003). Similar conclusions have been drawn

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from other mid- and high-latitude lake records, whose catchments experienced significant changes in vegetation and soil maturation following deglaciation (Hammarlund 1993; Hammarlund et al. 1997). Similar trends are evident in the G13Corg records from these lakes during the early to mid-Holocene, overall characterized by a decrease in G13Corg values in the lower strata followed by an increase in G13Corg values after 6000 yr BP. The decrease in G13Corg values is probably a reflection of increased influx of 13C-depleted dissolved CO2 from soil decomposition associated with rapid hydrological flushing, given that this shift is generally coupled with reduced G18OlwG18Op offset and boreal forest expansion. Subsequent 13C-enrichment post-6000 yr BP likely reflects generally drier conditions and declining soil decomposition in relation to boreal treeline retreat. At Derevanoi Lake, a brief interval of 13C-enrichment occurs between 8000 and 7000 yr BP, which may reflect the initial effects of higher lake productivity associated with warming, moister conditions, and elevated nutrient influx to the lake. Analysis of bulk organic nitrogen stable isotope composition has also been shown to vary in association with changes in hydrology and catchment terrestrial vegetation (Wolfe et al. 1999, 2003), indicating that integration of cellulose G18O, bulk organic G13C, and bulk organic G15N analyses on lake sediment cores offers useful insights into linkages between past hydrological processes and nutrient cycling in boreal treeline watersheds. Summary of isotope paleolimnology The lake sediment isotope-inferred paleohydrological records lay the framework for the potential development of a coherent regional reconstruction of early to mid-Holocene moisture histories in northern Russia, which appears to have been driven mainly by varying oceanicity (Wolfe et al. 2000, 2003). According to Koç et al. (1993), optimal Holocene conditions developed at 7000 yr BP as a result of the dominance of warm Atlantic waters in the central and eastern portions of the Greenland, Iceland, and Norwegian Seas. Warm sea surface temperatures for the southeastern Barents Sea between about 7500 and 4500 yr BP have also recently been inferred from cyst assemblages in marine sediment cores (Voronina et al. 2001), with the earliest part of this interval marked by maximum influx of North Atlantic water into the northern margin of the Barents Sea (Duplessy et al. 2001; Lubinski et al. 2001). These changes in North Atlantic circulation are thought to have led to increased cyclonic activity and the eastward flow of warm and moist air (Koç et al. 1993; Kerwin et al. 1999; Lubinski et al. 1999; MacDonald et al. 2000c). Enhanced supply of moisture-laden air masses to the Kola Peninsula at 7000 yr BP and at least as far as the lower Yenisey River region is suggested by reduced lake water evaporative 18O-enrichment at Poteryanny Zub Lake, Lake Yarnyshnoe-3, and Derevanoi Lake. Further to the east on the Lena River Delta, large areas of continental shelf exposed by glacial sea-level draw-down during the early Holocene may have suppressed maritime climatic influence in what are now coastal areas, resulting in increased evaporative 18O-enrichment at Dolgoe Lake. Subsequent reduced North Atlantic influence in the Nordic seas after 6000 yr BP contributed to drier (and cooler) conditions on the Kola Peninsula and eastward along the northern

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Siberian coast, while a cool, maritime climate was established on the Lena River Delta mainly as a result of rising sea-level. Regional lake status data bases and lake-level records from northern Russian Eurasia Compilations of data on late Quaternary changes in lake-levels analysed at a regional to continental scale have been used to reconstruct changes in the hydrological budget and in the atmospheric circulation patterns for a number of regions (e.g., Street-Perrott and Harrison 1985; Magny 1992; Harrison and Digerfeldt 1993). In the case of arctic Russia, available geological, geomorphological and biostratigraphic records from a number of lake basins have been similarly compiled and used to reconstruct late Quaternary lake-level changes. The European Lake Status Data Base (ELSDB: Yu and Harrison 1995) contains records from 118 sites and the Former Soviet Union and Mongolian Lake Status Data Base (FSULSDB: Tarasov et al. 1994, 1996) contains records from 103 sites, including some from the arctic region. For each site a relative water depth or lake-level (“lake status”) has been qualitatively estimated and coded from existing physical stratigraphic and biostratigraphic data. The shallowest phase in the lake history is attributed to status category 1 with each deeper phase successively placed into categories 2, 3, 4, etc. until the maximum reconstructed depth has been reached. The number of status categories may vary from site to site, depending on the quality of the record and sensitivity of the lake basin to the lake-level fluctuations. The chronology of the records is based on non-calibrated radiocarbon dates from the lake sediments, or on pollen-based correlation with radiocarbon-dated sites within the region, or on the basis of correlation with a regional pollen-based chronostratigraphy. Since the data bases were constructed to document climate changes, the authors tried to avoid those records or parts of records where fluctuations in water depth or lake-level were influenced by non-climatic factors (e.g., human and tectonic activity, natural or artificial damming and isostatic effects) or by indirect reasons, such as sea-level rise or glacial fluctuations. The lake status records from northern Eurasia have been used to reconstruct and map temporal and spatial variations in the late Quaternary moisture conditions and associated changes in the atmospheric circulation (e.g., Harrison et al. 1996; Tarasov and Harrison 1998; Yu 1998). The Russian arctic region (north of 65qN) is represented in the FSULSDB (Tarasov et al. 1996) by only six basins which are lakes today. Among these, three sites are from the Kola Peninsula and the three others come from Karelia, the eastern European Plain and western Siberia (Figure 10). Lake-level reconstructions suggest that all studied lakes changed their status during the Holocene. However, each record is different, making it difficult at present to reconstruct and determine mechanisms for general regional changes in past hydrology. Lake-levels from the Kola Peninsula (Kovdor, Kanent’yavr and Pasmlambina) appear to have been low around 9500 yr BP and high from 8200 to 7200 yr BP. In Lake Paanayarvi (Karelia), the deepest phase occurred between 10,000 and 8000 yr BP, and rather shallow conditions existed between 6500 and 7000 yr BP and after 2500 yr BP. Lake Mitrofanovskoe, situated further east close to the Ural Range, fluctuated around intermediate levels during the last 6500 years. The record

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from Lake Jamylimujaganto in western Siberia is rather short, suggesting the lake was shallow at 3000 yr BP, becoming deeper toward the present. It is worth noting that, in addition to lake-to-lake differences in the FSULSDB data set, the lake status reconstructions based upon existing stratigraphic data do not always agree with the paleohydrological inferences provided by the recent pollen, diatom, chironomid and isotope studies outlined in previous sections of this chapter. For example, evidence from the multi-proxy analysis of the records from Poteryanny Zub Lake and Lake Yarnyshnoe-3 on the Kola Peninsula suggest generally dry conditions from ca. 9000 to 7000 yr BP and increased effective moisture following 7000 yr BP. However, the FSULSDB lake status reconstructions from existing data suggest that moister conditions may have been initiated as early as 8200 yr BP. The site by site differences in the lake status changes reconstructed for the Russian Arctic (Figure 10) and the divergence between some of the lake status reconstructions and the paleolimnological inferences mentioned above may reflect a number of factors. First, there may be reasons that objective measurements of the FSULSDB lakes are in error (e.g., poor dating control, loss of information due to low time resolution of the record and coarse sampling of the lake sediment, and limited number of analytical methods used to reconstruct water depth or lake-level changes). Second, there may be weaknesses and errors in the somewhat subjective status interpretation methodology. For example, the diatom and lithology records are often used as a rough measure of relative water depth or lake-level, although those records, as most of the other biostratigraphic evidence, are often susceptible to alternative interpretations (Harrison and Digerfeldt 1993; Duck et al. 1998). Thus, the diatom plankton/non-plankton ratio could reflect lake-level changes, but lakeshore morphometry, ice cover, and changing water chemistry may also influence this ratio (Battarbee et al. 1998). Sand accumulation may reflect lake-level lowering, but can also be explained by water or wind erosion from the surrounding slopes. A combination of different types of sedimentary evidence increases the possibility of yielding reliable reconstructions, however it does not guarantee a correct interpretation. The risk of an incorrect interpretation is higher when the reconstruction is based on the records from a single core (Digerfeldt 1998). Moreover, there are many local non-climatic factors that may affect lake-level and sedimentary records (Dearing and Foster 1986). These changes may potentially be misidentified as consequences of moisture and temperature change (Duck et al. 1998). Comparison of the FSULSDB data with lake status records from the ELSDB (Yu and Harrison 1995; Figure 11) may be helpful in generating a more robust record of former lake status in northern Eurasia. The ELSDB data from the west European sector of the Arctic (north of 65qN) is also scarce, but the dating control and the time resolution of the records is much better than in the FSULSDB data set (Figure 11). However, changes in the relative water-level in eight lakes from Norway, Sweden and Finland are also very different from each other, with the exception of the two small neighboring lakes from Finland (Isohattu (68°38’N, 23°36’E) and Jierstivaara (68°40’N, 23°44’E)). Visual comparison of the west European and Russian data sets indicate that the Endletvatn (Norway (69°44’N, 16°05’E)) and Jierstivaara (Finland) records show patterns of change similar to that of Paanayarvi in Karelia, while the Skinkevatna (74°29’N, 18°55’E), Pølsa (74°30’N, 18°56’E) and Ellasjöen (74°23’N, 19°03’E) records from northern Norway are more similar to three FSULSDB records from the Kola Peninsula. Attempts to correlate reconstructed changes in Russia or to explain them in terms of

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Figure 10. Lake status based on the Former Soviet Union and Mongolian Lake Status Data Base (FSULSDB) analysis of existing Russian data (after Tarasov et al. 1996).

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Figure 11. Lake status based on the European Lake Status Data Base (ELSDB) analysis of existing western European data (Yu and Harrison 1995).

regional climate changes, as it has been done in the European and Asian mid-latitudes (Tarasov and Harrison 1998; Yu 1998), would be speculative without additional sites. The addition of new sites, such as those described in the preceding sections of this chapter, may greatly enhance the reliability and usefulness of the FSULSDB data set. Summary The pace, methodological sophistication and geographic scope of paleolimnological research by Russian and joint Russian-international scientific teams have increased in recent years and have produced important new records of environmental change across northern Eurasia. From the subset of studies presented in this chapter, it is clear that the paleolimnological approach of coupling intensive surface sampling and the

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development of environmental calibration sets with the procurement of long core records provides an optimal approach for tracking the environmental and climatic history of the region. The biologically-based studies discussed here demonstrate that, in general, the climate across northern Eurasia was relatively warm during the period of at least 9000 yr BP to about 3000 yr BP. The impact of these warm conditions is reflected in the limnological indices and in evidence of terrestrial vegetation that is found in lake sediments (macrofossils, stomates and pollen). The warming that occurred during the early and mid-Holocene was likely generated by high summer insolation caused by the earth’s orbital geometry, coupled with the disappearance of the Fennoscandian Ice Sheet and the concurrent warming of the North Atlantic (MacDonald et al. 2000c). The cooling in the late Holocene was likely a direct result of decreasing summer insolation due to the continued changes in orbital geometry. The stable isotope studies discussed in this chapter have provided evidence of moisture balance that has been difficult to derive from the biological indicators. The stable isotope results suggest that both large scale changes in circulation-precipitation patterns caused by changing insolation, glacial ice cover and sea surface temperatures, and more regional to local changes in moisture balance, caused by eustatic and isostatic sea-level changes, have led to regional differences in the history of moisture balance in the northern Eurasian Arctic. Evidence of such differences is also apparent from the analysis of the Russian lakelevel data sets. The stable isotope records suggest that the arctic areas of northern Russia and northwestern Siberia were moister in the early to middle Holocene, while the Lena River region was drier at that time. Much remains to be done in terms of applying paleolimnological studies to far northern Eurasia. In the spatial domain, large tracts remain unsampled or analysed. This is particularly true for far eastern Siberia and the arctic islands. In the temporal domain, it is crucial to start developing truly high resolution records that might address events or changes in environmental variability that occur over decades or centuries. Significant problems still remain in meeting these goals. The logistics of working in areas such as eastern Siberia are in some ways becoming more difficult as prices escalate and the human population in these regions declines. In non-glaciated areas there are often many thermokarst lakes but, due to the dynamic nature of the periglacial landscape, it is often difficult to find lakes that have sediment records dating back more than a few thousand years. Different groups working in the region do not always exchange data and ideas, or share logistics as often as might be hoped. However, the need to understand the dynamics of arctic climate and environmental change are pressing enough to behove us to overcome the difficulties outlined above. Acknowledgements We thank Les Cwynar, Marianne Douglas, Reinhard Pienitz and John Smol for a number of constructive comments and suggestions on an earlier draft of this chapter. The research reported here was supported by a number of agencies including the Russian Foundation for Fundamental Research, the Royal Society of Canada, the Natural Sciences and Engineering Research Council of Canada, NERC of the United Kingdom and the National Science Foundation of the United States. This is a PACT and a PARCS contribution.

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Rühland K. 1996. Assessing the Use of Diatom Assemblages as Paleoenvironmental Proxies in the Slave and Bear Geological Provinces, N.W.T., Canada. M.Sc. thesis, Queen’s University, Kingston, Ontario, Canada. 143 pp. Rühland K., Smol J.P., Jasinski J.P.P. and Warner B.G. 2000. Response of diatoms and other siliceous indicators to the developmental history of a peatland in the Tiksi forest, Siberia, Russia. Arct. Ant. Alp. Res. 32: 167-178. Seppä H. and Hammarlund D. 2000. Pollen-stratigraphical evidence of Holocene hydrological change in northern Fennoscandia supported by independent isotopic data. J. Paleolim. 24: 69-79. Snyder J.A., Forman S.L., Mode W.N. and Tarasov G.A. 1997. Postglacial relative sea-level history: sediment and diatom records of emerged coastal lakes, north-central Kola Peninsula, Russia. Boreas 26: 329-346. Snyder J.A., MacDonald G.M., Forman S.L., Tarasov G.A. and Mode W.N. 2000. Postglacial climate and vegetation history, north-central Kola Peninsula, Russia: pollen and diatom records from Lake Yarnyshnoe-3. Boreas 29: 261-271. Solovieva N.A. 2000. A Palaeoecological Study of Holocene Environmental Change in a Small Upland Lake from the Kola Peninsula, Russia. Ph.D. thesis, University College London. London, UK, 408 pp. Solovieva N.A. and Jones V.J. 2002. A multiproxy record of Holocene environmental changes in the central Kola Peninsula, northwest Russia. J. Quat. Sci. 17: 303-318. Sorsa P. 1965. Pollenanalytische Untersuchungen zur spätquartären Vegetations- und Klimaentwicklung im östlichen Nordfinnland. Ann. Bot. Fenn. 2: 301-413. Street-Perrott F.A. and Harrison S.P. 1985. Lake levels and climate reconstruction. In: Hecht A.D. (ed.), Paleoclimate Analysis and Modeling. John Wiley, New York, pp. 291-340. Tarasov P.E. and Harrison S.P. 1998. Lake status records from the former Soviet Union and Mongolia: a continental-scale synthesis. Paläoklimaforschung/Palaeoclimate Research 25: 115-130. Tarasov P.E., Harrison S.P., Saarse L., Pushenko M.Y., Andreev A.A., Aleshinskaya Z.V., Davydova N., Dorofeyuk N.I., Efremov Y.V., Khomutova V.I., Sevastyanov D.V., Tamosaitis J., Uspenskaya O.N., Yakushko O.F. and Tarasova I.V. 1994. Lake status records from the former Soviet Union and Mongolia: data base documentation. NOAA Paleoclimatology Publications Series Report No. 2. Boulder, USA, 274 pp. Tarasov P.E., Pushenko M.Y., Harrison S.P., Saarse L., Andreev A.A., Aleshinskaya Z.V., Davydova N., Dorofeyuk N., Efremov Y.V., Elina G.A., Elovicheva Y.K., Filimonova L.V., Gunova V.S., Khomutova V.I., Kvavadze E.V., Neustrueva I.Y., Pisareva V.V., Sevastyanov D.V., Shelekhova T.S., Subetto D.A., Uspenskaya O.N. and Zernitskaya V.P. 1996. Lake Status Records from the Former Soviet Union and Mongolia: Documentation of the Second Version of the data base. NOAA Paleoclimatology Publications Series Report No. 5. Boulder, USA, 224 pp. Voronina E., Polyak L., de Vernal A. and Peyron O. 2001. Holocene variations of sea-surface conditions in the southeastern Barents Sea reconstructed from dinoflagellate cyst assemblages. J. Quat. Sci. 16: 717-726. Vorren T.O., Vorren K.-D., Gulliksen T. and Løvlie R. 1988. The last deglaciation (20,000 to 11,000 B.P.) on Andøya, northern Norway. Boreas 17: 41-77. Walker I.R. and Mathewes R.W. 1989. Chironomidae (Diptera) remains in surficial lake sediments from the Canadian Cordillera: analysis of the fauna across an altitudinal gradient. J. Paleolim. 2: 61-80. Wiederholm T. (ed.) 1983. Chironomidae of the Holarctic region. Keys and diagnoses. Part I Larvae. Ent. Scand. Suppl. 19., 457 pp. Wolfe B.B., Edwards T.W.D., Aravena R. and MacDonald G.M. 1996. Rapid Holocene hydrologic change along boreal treeline revealed by į13C and į18O in organic lake sediments, Northwest Territories, Canada. J. Paleolim. 15: 171-181.

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13. PALEOLIMNOLOGICAL STUDIES IN ARCTIC FENNOSCANDIA AND THE KOLA PENINSULA (RUSSIA)

ATTE KORHOLA ([email protected]) JAN WECKSTRÖM ([email protected]) Environmental Change Research Unit (ECRU) Department of Biological and Environmental Sciences University of Helsinki P.O. Box 65 (Viikinkaari 1) FIN-00014 Helsinki Finland

Key words: Paleolimnology, Lake ontogeny, Natural acidification, Lake-level fluctuations, Thermal development, Aquatic succession, DOC, Arctic warming, Fennoscandia, Kola Peninsula

Introduction The undulating landscape of the Fennoscandian Arctic, including northwestern Russia, is littered with lakes and ponds that are fed by precipitation. In the Kola Peninsula alone there are over 100,000 lakes larger than 1 ha, while the highest lake density in northern Fennoscandia can be found in northeastern Finnish Lapland, with more than 800 lakes per 100 km2 (Raatikainen and Kuusisto 1990) (Figure 1). The majority of these arctic freshwaters are small (area usually from 0.05 to 12.0 km2), shallow (Zmax usually < 10 m), oligotrophic (total phosphorus (TP) < 10 Pg L-1), extremely dilute (total dissolved solid (TDS) concentrations usually 15-30 mg L-1), and clear (platinium color units < 30 mg L-1) (Korhola et al. 2002a). There are only two large water bodies in the area: Lake Inari (1116 km2) in northeastern Finnish Lapland and Lake Imandra (881 km2) on the Kola Peninsula. Due to their small size, remote location and resultant logistical constraints, the arctic and subarctic lakes in the Eurasian Arctic have been far less intensively studied for scientific purposes than their more southerly counterparts. For example, the first description of a complete annual cycle of physical and chemical variables in an arctic lake in Finnish Lapland appeared only recently (Sorvari et al. 2000). The arctic lakes of the region are, however, in many respects interesting and unique. Due to limited nutrient availability, low light levels, low temperatures, ice presence and persistence, and a short growing season, the lakes are generally characterized by low productivity, low species diversity, and relatively simple food webs. Much of the primary production occurs on the lake bottom, whilst the production maximum of plankton commonly occurs during the autumn overturn period (Sorvari et al. 2000). Attached diatoms in the rocky littoral zones and on lake bottom mosses are important primary producers throughout the open

381 R. Pienitz, M.S.V. Douglas and J.P. Smol (eds), 2004. Long-term Environmental Change in Arctic and Antarctic Lakes. Springer. Printed in the Netherlands.

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Oo

A rc t ic

c ir

cle 6

Study area 6.5ooN

1:20000 Figure 1. Aerial photograph from northeastern Finnish Lapland (black dot in the index map) with 1960 lakes per 100 km2 (Raatikainen and Kuusisto 1990). Most of the lakes are of glacial origin. The index map shows the location of the area discussed in this paper (Copyright Maanmittauslaitos 2003).

water period. Phytoplankton is typically dominated by chrysophytes and diatoms, with dinoflagellates and cryptophytes of lesser importance. The lakes are at least partly frozen for 7 to 9 months per year, for which reason they commonly suffer from hypolimnetic oxygen depletion in late winter (winter anoxia). Lake mixing patterns are largely dependent upon the morphometry of the lake basins, including volume and depth, solar radiation regime, and wind patterns. Two types of thermal structures are typical for arctic lakes in Fennoscandia and the Kola Peninsula: dimictic lakes, in which the water circulates freely all the way to the bottom twice a year in spring and fall with a stagnation period in between, and monomictic or isothermal lakes (usually < 10 m deep)

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with no or only weak stratification during summer (Figure 2). Stratification regimes can determine the thermal, physical and chemical characteristics of the lake and its overall productivity (Forsström et al. 2002).

A.

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Ice

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July

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J F M A M J J A S O N D Figure 2. Schematic diagram of two types of thermal structures typical for the arctic lakes in the study region. In an isothermal lake (A) no or only weak stratification occurs during summer, while in a dimictic lake (B) water circulates freely to the bottom twice a year in spring and fall with a short stratification period in between.

During the past few years, considerable development has taken place in limnological studies of arctic and subarctic waters of northern Fennoscandia and the Kola Peninsula. A broad range of environmental, water chemistry, lake physical, and lake biological data have been collected and analysed from these waters in association with international and national research programs (e.g., Kortelainen 1993; Henriksen et al. 1997a,b; Moiseenko et al. 1997; Weckström et al. 1997a; Blom et al. 1998, 2000; The MOLAR water chemistry group 1999; Mannio et al. 2000; Rautio et al. 2000; Solovieva 2000; Sorvari et al. 2000; Rautio 2001; Weckström and Korhola 2001; Korhola et al.

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2002a). Collectively, these studies have contributed to an improved understanding of limnological and hydrobiological features of these fascinating aquatic environments. However, little data exist concerning the longer term limnological development, ecosystem dynamics, species successions, and biotic interactions of arctic and subarctic lakes in northern Fennoscandia and the Kola Peninsula. The purpose of the present review is to compile and summarize such paleolimnological data that exist from the region dealing with the postglacial limnological development of freshwaters in this area. We focus on paleolimnology and lake environments and discuss the general regional environmental history (as interpreted in many cases from the lake sediments) only to the extent that is needed in order to understand the past changes in lake development. For the present purpose we define “Arctic” pragmatically as every region inside the latitude at which the sun does not set in the summer – that is the Arctic Circle. However, some of our examples come from slightly more southerly sites. Most of the published paleolimnological studies in the area deal with long-term chemical trends (e.g., Korhola et al. 1999; Korsman 1999; Seppä and Weckström 1999; Grönlund and Kauppila 2002; Solovieva and Jones 2002; Weckström et al. 2003), while some studies focused on lake-level variations (Hyvärinen and Alhonen 1994; Eronen et al. 1999; Virkanen 2000; Korhola and Rautio 2001), hydrological characteristics (e.g., Hammarlund et al. 2002), or recent changes in biological associations in response to some external factors (Sorvari and Korhola 1998; Korhola et al. 2002b; Sorvari et al. 2002). During the past few years, particular efforts have been made to quantitatively infer some key limnological variables (e.g., lake temperature, pH, dissolved organic carbon) from biological remains in sediments, using transfer functions based on modern reference data sets (training sets) of biological and environmental data. The key Holocene sites where such quantitative data are available are shown in Figure 3. The questions dealing with the natural species succession and long-term change in aquatic communities have only recently been tackled (e.g., Weckström et al. 2001; Solovieva and Jones 2002). Such baseline data about past ecosystem composition and function would be of great value in understanding the factors that control the modern limnology, biodiversity, and fragility of these systems, and might also help to predict the future trajectories of lake development. Waters in the far north of Europe are increasingly threatened by multiple environmental pressures, including atmospheric pollution, climate warming, and increased levels of harmful ultraviolet (UV) radiation (Rautio and Korhola 2002). To predict the likely future direction of change, it is essential to assess variance and rate of change of natural communities and to separate that natural component from any human-induced driver of ecosystem change. Origin and sedimentological characteristics of lakes The whole of Fennoscandia, as well as the Kola Peninsula, were ice-covered during the last glacial epoch, for which reason most lake basins in the region are of glacial origin. These include lakes that originated in the carvings caused by the retreating ice as well as lakes formed due to damming by moraines and various other landforms generated by the moving ice. Ice-scour lakes are the principal types found in arctic Fennoscandia. Moraine lakes, formed where terminal or recessional moraines have blocked the stream that replaces the retreating glacier, are particularly common in northern Norway.

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Hundreds of small kettle lakes are found in Finnish and Swedish Lapland formed in the depressions left by melting blocks of ice in glacial drift or in the outwash plains. Numerous water bodies have originated also via cryogenic processes (e.g., thermokarst and thaw lakes in the Russian Arctic), by deflation processes, by ice damming, or through emergence from the sea as a result of glacio-isostatic rebound. There are numerous cirque-lakes (tarns) on the highest summits of the Norwegian and Swedish mountains. Old deflation surfaces can be very extensive in the Fennoscandian Arctic and their deepest points are nowadays mostly under water as a result of the rise in the groundwater table some 4000 to 5000 years ago. Many rivers in this region meander extensively and, as a consequence, meanders may be cut off from the main stream of the river creating oxbow lakes. The sediments deposited in these lakes of different origin offer typically continuous records to the present day; most of the paleolimnological studies consequently concentrate on the late Quaternary and the Holocene.

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Sediment accumulation rates Sediment accumulation rates in lakes vary according to the input of both organic and inorganic material from the watershed and air, and biogenic material from the lake itself. Lakes all over the European Arctic are characterized by low sediment accumulation rates typical of northern oligotrophic lakes with rocky catchments. Calculated long-term average sedimentation rates vary from ca. 0.01 cm yr-1 in arctic tundra sites to 0.05 cm yr-1 in more southerly treeline lakes (e.g., Hyvärinen and Alhonen 1994; Korhola et al. 1999; Barnekow 2000; Snyder et al. 2000; Virkanen 2000; Rosén et al. 2001; Shemesh et al. 2001; Bigler et al. 2002; Sorvari et al. 2002). Whereas the total postglacial sediment thickness in the high-altitude tundra lakes seldom exceeds one meter, it is not exceptional to find a sediment sequence up to 3 m in a Fennoscandian treeline lake (Eronen et al. 1999; Seppä and Weckström 1999; Rosén et al. 2001). Applying the sub-sampling of 1 to 2 cm (or 5 to 20 cm in most studies prior to 1990), the typical resolution of Holocene records from homogenous non-laminated sediment sequences is thus approximately 50 to 100 years. The last 100 years is usually recorded in the top 2 to 4 cm of the core, requiring a particularly high sampling resolution if recent changes are to be studied. Recognizing this, Sorvari et al. (2002) subsampled five cores from Finnish Lapland at fine intervals of 2 to 5 mm, equivalent to a temporal resolution of about 3 to 10 years. This allowed them to explore recent changes in biotic assemblages in these arctic lakes. Fairly constant sediment accumulation rates during the Holocene are typical for lakes in the region (Barnekow et al. 1998; Barnekow 1999; Seppä and Weckström 1999; Rosén et al. 2001; Shemesh et al. 2001; Seppä et al. 2002; Bigler et al. 2003). However, according to Eronen et al. (1999), in some closed-basin lakes with considerable fluctuations in their water-level (see below), sediment deposition was generally slow until the mid-Holocene, but increased around 5000 cal. yr BP (calibrated radiocarbon years before present where present is AD 1950). Not all records from northern Fennoscandia mimic this pattern, and an opposite trend has been found for some lakes (Barnekow 2000; Virkanen 2000). Because estimation of the Holocene sediment accumulation rates is largely based on radiocarbon dating, it should be assessed with some caution. Often, no terrestrial macrofossils are found in the sediments, for which reason less reliable bulk sediment samples or aquatic moss remains have been used as dating material. Furthermore, in some cases sedimentation rate estimation is based on fairly thick (15 cm) bulk sediment samples in the frame of ‘conventional’ radiocarbon dating (Hyvärinen and Alhonen 1994; Eronen et al. 1999). As shown by Barnekow et al. (1998), bulk sediment dates are more prone to errors and may give considerably older ages than dates based on carefully selected terrestrial macrofossils from the same sediment layers. Until now, few lakes with a continuous Holocene record of annually laminated sediments have been found in the European Arctic despite many attempts. The sediment is laminated on an annual basis in the deepest part in Lake Vuolep Njakajaure, a relatively large lake located in the Abisko valley in northern Sweden (68q20’N, 18q45’E) at an altitude of 409 m a.s.l. (Barnekow et al. 1998). Furthermore, there are several varved lake sediment sequences quite far north in Sweden at a latitude of about 64qN (Snowball et al. 2002). Recently, Swedish scientists have recovered a lake sequence from the Abisko area that contains continuous laminations (I. Renberg, pers.

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comm.). However, it is not yet resolved whether these laminations are varves, i.e., represent annual layers. Variations in the sediment organic content The organic matter content of the sediment can be estimated by measuring weight loss in sediment core sub-samples after burning at selected temperatures (loss-on-ignition (LOI); Heiri et al. 2001). However, the percentage weight lost on ignition gives only a crude measure of the organic content of the sediment; the results are accurate to 1-2% for organic matter in sediments with over 10% organic matter (Håkanson and Jansson 1983). According to Bengtsson and Enell (1986), organic carbon accounts for about 40% of the weight lost on ignition. In the sediment of a typical Fennoscandian tundra lake with a barren catchment, LOI values usually vary between 10-20% during the Holocene (Rosén et al. 2001; Seppä et al. 2002), implying highly minerogenic soils in the catchments and low primary production in the lakes. However, in some shallow treeline lakes in southern Lapland, sediment organic matter content up to 70-80% has been reported (Korhola et al. 1999; Virkanen 2000). Although sediment organic matter contents this high are commonly related to unique local environmental conditions, as discussed below, high amounts of organic matter in the sediments of arctic and subarctic lakes are not exceptional, as cold waters also typically have lower decomposition rates of organic matter. Results of some published LOI measurements covering the entire postglacial period are assembled in Figure 4. There are marked differences in the shapes of the LOI profiles, reflecting the dominant aspects of the local environment. For example, the unusually high LOI values during the mid-Holocene in Lake Kuttanen (69q25’N, 22q53’E), a groundwater-controlled kettle-hole lake in Finnish Lapland, are likely due to a marked lowering in the water-level that resulted in the spread of telmatic vegetation across the lake basin (Virkanen 2000; Korhola and Rautio 2001). The sudden drop in LOI from 15% to 5% in the sediment of Toskaljavri (69q12’N, 21q28’E) around 5000 cal. yr BP is associated with an exceptionally strong snow avalanche or slush flow from the surrounding steep slopes (Seppä et al. 2002). This distinct sediment layer, with high minerogenic content, is further recognized by an increase in median mineral grain size and distinctly anomalous pollen and spore assemblages (Seppä et al. 2002). Certain common features can nevertheless be discovered from the LOI profiles. The oldest sediments reflecting recently deglaciated lake environments are regularly low in organic matter (< 5%). They often contain glacial clays and silty material, presumably derived from inflow streams. As a consequence, the lowermost sediment is characterized by the higher amplitude and frequency of LOI fluctuations than the more recently deposited sediment. In sites located at lower latitudes where vegetation cover was rapidly developing (e.g., Sjuodjijaure in Figure 4C), a sharp increase in LOI values takes place around 9000 to 8000 cal. yr BP after which the values stay relatively constant. In some profiles a slight decrease can be noted during the latter half of the Holocene, probably as a result of both increased erosion rates during the past few millennia (Snowball et al. 1999) and late Holocene retreat of forest-tundra that led to a decrease in organic supply to the lakes. However, in some sites an increasing trend in

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Calibrated years BP (x10 ) Figure 4. Holocene loss-on-ignition (LOI) profiles of selected northern Fennoscandian lakes. (A) Jeknajaure, (B) Niak, (C) Sjuodjijaure, (D) Seukokjaure, (E) Toskaljavri, (F) Tsuolbmajavri, (G) Dalmutlad’do, (H) Kuttanen. Lakes (A-D) and (E-H) are arranged according to decreasing altitude from top to bottom. Plots (A-D) are after Rósen et al. (2002), (H) is taken from Seppä et al. (2002), (F) from Seppä and Weckström (1999), and (H) from Virkanen (2000). (G) is unpublished.

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LOI values can be seen during the last 4000 to 3000 years that may be related to mire development in their catchments (e.g., Tsuolbmajavri; Seppä and Weckström 1999). In lakes located on erosion-sensitive soils (e.g., Kuttanen) or surrounded by steep mountains (e.g., Dalmutlad’do, 69q18’N, 20q72’E), a gradual increase in LOI and corresponding decline in minerogenic matter until ca. 6000 cal. yr BP can be observed, after which the values seem to stabilize. This trend can presumably be attributed to decreasing erosion rates due to gradually declining precipitation and spring snowmelt in Fennoscandia from the early to mid-Holocene. Many recent studies (e.g., Shemesh et al. 2001; Hammarlund et al. 2002; Snowball et al. in press) suggest a transition from relatively moist, oceanic conditions in the early Holocene to a much drier climate after ca. 6000 cal. yr BP. The increasing continentality finally led to very dry conditions between 6000 and 4000 cal. yr BP during which many lakes experienced a marked drop in their water-levels and when pine (Pinus sylvestris L.) expanded up to 300 m higher altitudes than today (Barnekow 2000). In contrast to the trends described above, the LOI values in some high-altitude lakes located on organically-poor, mineral-rich soils of the arctic tundra remain more or less constant (e.g., Toskaljavri) or show irregular distribution (e.g., Jeknajaure; 67q13’N, 17q48’E) throughout the Holocene. These results suggest that the broadscale variations in the sediment organic matter content are predominantly linked to soil and vegetation development and should not be interpreted, without further evidence, as indication of lake productivity. The likely allogenic control of organic versus minerogenic matter in these systems is further confirmed by geochemical analyses that show a close relationship between the LOI variations and the elemental geochemistry indicative of catchment erosion, weathering, and leaching (Kauppila and Salonen 1997; Virkanen 2000; Solovieva and Jones 2002). In general, the records show there is a tendency towards lower LOI values with increasing altitude (Figure 4). A statistically significant inverse relationship between LOI and altitude (r = -0.58, p < 0.01) was notable in the surface sediment (top 1 cm) data set collected from 53 small lakes spanning the Fennoscandian treeline (Korhola et al. 2002a). A particularly pronounced decrease in LOI occurred when moving from forested catchments to treeless tundra, with the median LOI value dropping from 61 to 30%, respectively. The overall LOI values in this contemporary data set ranged between 17 and 88% with the median of 48.3%, these values being surprisingly high. However, due to incomplete mineralization in surficial sediments, organic matter content can be much higher than in the lower sediment sections. In summary, the interpretation of the gathered LOI data is quite complex because the values are affected by both inputs of inorganic and organic matter, the preservation of which can be incomplete (Battarbee et al. 2002). Owing to the simultaneous operation of many environmental agents, such as climate, catchment development, groundwater fluctuations and various short-term disturbances, it is difficult to form a coherent and systematic picture of the changes in the organic flux that the lakes have undergone during their ontogeny. It is evident, however, that the driving force behind the general long-term patterns in sediment organic matter deposition has been climate, especially changes in humidity, and resulting variability in erosion rates (Nesje et al. 2000; Virkanen 2000). To what degree the LOI fluctuations are indicative of lake productivity is difficult to assess because of the lack of detailed studies on sediment organic geochemistry. However, studies on mountain lakes in general demonstrate that the fine-

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scale variability in the sediment organic content may indeed be attributed to within-lake processes and primary productivity (Battarbee et al. 2002). The C:N ratio in the top sediment sequence of Lake Saanajärvi (69q05’N, 20q52’E), northwestern Finnish Lapland, points to a prevalent autochthonous origin of the organic matter since the values are just above or very near to 10 (Korhola et al. 2002b). Holocene changes in physical attributes Changes in water-level Interpreting past lake-level fluctuations is predominantly based on plant and animal micro- and macrofossil analyses, although variations in sediment accumulation rates and geochemical records have also been used in these interpretations. The studies are mostly based on single cores taken from the middle of the small and shallow closed-basin lakes that are considered sensitive to regional changes in the hydrological balance, and the results are interpreted in a qualitative and descriptive manner. However, Barnekow (2000) used sedimentary features and macrofossil records from multiple cores taken along a transect over a small water body in northern Sweden in an attempt to quantify water-level fluctuations in accordance with the approach used by Digerfeldt (1988) in southern Sweden. Furthermore, Korhola and Rautio (2001) recently applied a cladoceran-based transfer function to quantitatively predict past lake-level changes from subfossil cladoceran assemblages in a groundwater-fed lake in Finnish Lapland. Meanwhile, Moser et al. (2000) have developed a transfer function for predicting lakelevel changes from diatom assemblages in Finnish Lapland, yet that model has thus far not been used to reconstruct past water-levels in the region. Hyvärinen and Alhonen (1994) used diatom and cladoceran assemblages to infer changes in water-levels in two closed basins in northwestern Finnish Lapland. The changing planktonic–littoral (P/L) proportion of the cladoceran and diatom remains in the sediment cores was used as an indicator of the areal changes between the shallowwater and open-water zones. The authors suggested, based on the low occurrence of planktonic cladocerans and diatoms between about 8000 and 4000 yr BP, that waterlevels were considerably lower during the entire early to mid-Holocene than they are today. The plankton component then rose at approximately 4000 yr BP, suggesting a late Holocene rise in water-levels, apparently in response to increased humidity. Further confirmation to their interpretation was provided by sediment accumulation rates that were noted to be low during the early Holocene but which increased during the late Holocene. However, the determination of accumulation rates was based on a low number of conventional bulk sediment radiocarbon dates, not converted to calendar years, which render the interpretation more difficult. Eronen et al. (1999) reported Holocene sediment accumulation rates from five closedbasin lakes in Finnish Lapland. After deglaciation of the lake basins, the rate of sediment deposition was initially slow, but accelerated after 5000-4000 14C yr BP. The authors postulated that the early Holocene slow deposition rates were due to shallow water-levels, whereas the increased rates later in the Holocene probably resulted from “intensified surface runoff, which carried more allochthonous humus and plant detritus into the basins”. Further support for the late Holocene water-level rise came from

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subfossil pine megafossils that are preserved in large quantities in small lakes in Lapland. As demonstrated by Eronen et al. (1999), the position of the subfossil trunks and stumps indicated that many of these trees had been growing on dry land along the former lake shores, which are today inundated. Also in this study, sediment accumulation rates were estimated by conventional radiocarbon dating of fairly few and thick (10 to 15 cm) sediment slices, which should be taken into account when interpreting the results. Virkanen (2000) used sediment physical properties and geochemical tracers to investigate water-level fluctuations in the closed-basin Lake Kuttanen in northwestern Finnish Lapland. Minimal mineral matter accumulation, low elemental accumulation rates and a shift in the redox state towards more reducing conditions all suggested a marked lowering in the local water table between ca. 6000 and 5000 cal. yr BP. The elemental accumulation rates again increased since ca. 5000 cal. yr BP, suggesting higher water-levels and dilution by Fe supply from the catchment soils. Similarly, using a 9000-year varved lake sediment sequence in northern Sweden (Lake Sarsjön), Snowball et al. (1999) found that the mid-Holocene was characterized by sediments with a low concentration of detrital minerals and a high organic carbon content, whereas the late Holocene was marked by more frequent and intense periods of detrital sedimentation. The concentration of detrital minerogenic material in the sediment was shown to be controlled to a large extent by the intensity of the spring snow melt; thus the sediment structure in this case seemed to closely track variations in winter precipitation. To assess the magnitude of lake-level fluctuations more systematically, Barnekow (2000) collected eight sediment cores along a transect at 5 to 10 m intervals from the small groundwater-fed Lake Badsjön (68q20’N, 18q45’E), northern Sweden. Differences in the sediment stratigraphy along the transect were interpreted in terms of lake-level changes (Figure 5A). The sedimentation limit is seen in cores I-IV in the form of coarser material overlain by fine detritus gyttja. As radiocarbon dates were available only from the central master core, the chronology was transferred to the more nearshore cores on the basis of variations in the macrofossil record and changes in the concentration of magnetic minerals (Figure 5B). The extremely low sediment accumulation rate around 5000-6000 cal. yr BP in core I was interpreted as a result of a low lake-level. By following the position of the sedimentation limit in various cores, in addition to the information derived from macrofossil counts, mineral magnetic properties and sediment accumulation rates, Barnekow (2000) was able to estimate that the lake-level lowering in this particular lake was in the order of 1 to 1.5 m. A combination of high summer temperatures and reduction in winter precipitation during the local pine maximum was considered the major drivers for these changes. Using the species changes in the subfossil cladoceran assemblages in conjunction with the transfer function developed from the surface-sediment calibration data set (Korhola 1999; Korhola et al. 2000a), Korhola and Rautio (2001) were able to provide quantitative estimates of past water-level fluctuations for Lake Kuttanen, previously studied by Virkanen (2000). These inferences showed that the initially ca. 8 m deep lake experienced a marked water-level decrease of 5 to 6 m during the mid-Holocene, with the lowest lake-level phase between 6000 and 4000 cal. yr BP, followed by a gradual water-level increase of 2 to 3 m during the latter half of the Holocene. The inferred values of ca. 8 m for the lowermost sample and that of 5.8 m for the uppermost sample

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Figure 5. (A) The stratigraphy of Lake Badsjön recorded in a transect from south to north. Note that depth is recorded from the water-level. 1 - Fine gyttja; 2 - Coarse detritus gyttja; 3 - Drift gyttja; 4 - Calcareous gyttja/lake marl; 5 - Silty calcareous gyttja (transitional zone); 6 - Clay; 7 - Peat, and (B) stratigraphy and (A-C) SIRM (Saturated Isothermal Remanent Magnetisation) curves (m A m2kg-1) from Lake Badsjön. The ages are transferred from core point V to the core points I and IV. Correlations between the core points are based on changes in the macrofossil records; m1-m3 and a-g corresponds to peaks in the SIRM curves. Sedimentation rate is expressed in mm/yr. (modified from Barnekow 2000).

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correspond to the initial (sediment thickness = 1.5 m) and present measured depth (6 m) of the lake, giving credibility to the model. A similar pattern in lake-level fluctuations was detected when the cladoceran - lake depth inference model was applied to sediment core data from the original Hyvärinen and Alhonen (1994) study; a distinct lowering in water-levels of some 3 to 5 m was noted in all cases between 6000 and 4000 cal. yr BP (unpublished data). Collectively, these inferences suggest that water-levels in many closed-basin lakes in Finnish Lapland were high during the early Holocene, but declined up to 4-6 m during the mid-Holocene dry period around 6000-4000 cal. yr BP. Seppälä (1971) had previously estimated, on the basis of deflation hollows now filled with water, that in some places the water-level has risen as much as 5 m during the past 5000 to 4000 years, which matches our results. During the culmination of the Holocene dryness, many shallower water bodies of the region decreased greatly in size or may have dried up entirely. The overall pattern of the Holocene lake-level variations generally follows the regional changes in climate humidity as reconstructed by means of other sedimentary proxy indicators, such as pollen (Seppä and Birks 2001, 2002; Bigler et al. 2002) and oxygen- and hydrogen-isotopic compositions (Snowball et al. 1999; Shemesh et al. 2001; Hammarlund et al. 2002). For example, the recently published pollen-based quantitative reconstruction of Holocene precipitation variability (Seppä and Birks 2002) suggests that annual precipitation was high (ca. 600-800 mm) in northwestern Finnish Lapland during the earliest Holocene, decreased markedly between ca. 6000 and 4000 cal. yr BP (ca. 300-500 mm), and then gradually increased towards the present, largely governed by North Atlantic oceanic and atmospheric circulation patterns (Yu and Harrison 1995). Due to a high zonal index more oceanic conditions prevailed in northern Fennoscandia during the early Holocene, which meant more rain in particular during winter. Vassiljev (1998) has shown that winter precipitation has greater influence on lake-levels than variation in summer precipitation. Although much information on Holocene humidity changes and associated waterlevel variations for northern Fennoscandia now exists, the assessment of their impacts on aquatic ecosystems still remains unfinished. One can assume, however, that these marked fluctuations in water tables will certainly have affected many aspects of freshwater ecosystems, including water renewal rates, thermal characteristics, habitat availability, oxygen regimes, concentrations of most chemical constituents, and compositions of aquatic assemblages. Even small changes in precipitation and temperature can cause large changes in water chemistry, particularly in lakes where groundwater is the predominant source of chemical inputs. On the other hand, it should be stressed that water-level changes only seem to be the norm for small closed-basin lakes in the region, whereas there is no indication of any substantial water-level fluctuations in the tectonic basins located on less permeable soils. Changes in thermal features During the past few years, numerous temperature transfer functions have been developed from modern ecological data that allow inferences of summer temperatures from remains of aquatic organisms preserved in sediments of northern Fennoscandian lakes (for a complete list of the existing data sets, see Snowball et al. in press).

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However, the existing transfer functions are mostly based on air temperatures and hence do not allow direct hindcasting of lake thermal features. In Figure 6 we show preliminary results of a water temperature reconstruction for Lake Tsuolbmajavri based on fossil diatoms. Modern calibration is based on surface-sediment diatom assemblages and the corresponding measurements of summer epilimnion (surface mixed waters) temperatures from 29 lakes in northwestern Finnish Lapland. Water temperatures were recorded at two hour intervals by automatic mini-thermistors installed in the lakes; the July mean epilimnion temperature for the year 2000 was used in the calibration. Because of the very close positive correlation between summer air and water temperature in Finnish Lapland (see Korhola et al. 2001) it was not surprising that the obtained water temperature curve closely tracks the air temperature reconstruction of the same lake (see Korhola et al. 2000b). Lake Tsuolbmajavri (68q41’N, 22q05’E) is a shallow, well-mixed water body with interpolated July mean surface-water temperature of ca. 14qC. Reconstructed July mean water temperature has varied within the range of 2qC during the lake’s history (Figure 6). Diatom-inferred summer lake temperature was high during the mid-Holocene ‘thermal maximum’ (ca. 8000-5500 cal. yr BP), when summer air temperatures in the North Atlantic region were likely to have been 1 to 2qC higher than at present owing to orbital changes (Barnekow 1999; Johnsen et al. 2001; Korhola et al. 2002c; Snowball et al. in press). This warm interval was followed by a decline in lake temperature to a local minimum between ca. 4000 and 3000 cal. yr BP. Lake temperature rose again from ca. 3000 cal. yr BP and peaked around 1000 cal. yr. BP, roughly corresponding to the period of the Medieval Climate Anomaly (MCA). There is a distinct drop in inferred lake temperatures again between ca. 300 and 600 cal. yr BP associated with the cooling event of the Little Ice Age (LIA). Diatoms suggest again an increase of ca. 1.5qC in the water temperature since the termination of the LIA, which corresponds to the estimation of the post-LIA air temperature increase recently presented by Overpeck et al. (1997) for the Arctic in general. The changes in both air and water temperature, as described above, must have influenced the thermal and stratification regimes, duration of ice cover and oxygen regimes of lakes during postglacial times, thereby significantly affecting their ecosystem dynamics. For example, from the established empirical relationship between temperature change and duration of ice-free season, where a 1qC increase corresponds to a 7 to 9 day increase in open water duration (e.g., Hobbie et al. 1999; Magnuson et al. 2000), we can infer, on average, a two to three week increase in open water duration in Tsuolbmajavri during the Holocene Hypsithermal. Such a decrease in ice cover duration would in turn have reduced the winter anoxia common at these latitudes. The generally warmer climate conditions may have led to the establishment of a stronger thermal gradient in the water column of this presently monomictic (isothermal) lake. An opposite development has presumably taken place during cool periods such as the LIA, with more ice, shorter open water season, more effective water column mixing, and more frequent late-winter anoxia. To which degree can we extrapolate the data from Lake Tsuolbmajavri to other arctic freshwaters in the region? It is known that the magnitude and duration of the physical changes are largely dependent on basin characteristics such as depth, surface area, and effective fetch. At a detailed level every lake is thus unique. Yet the same processes

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occur in all freshwaters and are controlled by the same principles; much of the understanding can therefore be extrapolated from general limnological knowledge. Accordingly, our assumption of more stable summer stratification and reduced winter anoxia during the Holocene thermal optimum on one hand, and the potential destratification during the LIA on the other hand, are likely applicable to many shallow systems in the region. To some extent, this is confirmed by Snowball et al. (1999) who interpreted from the sediment mineral magnetic data that anoxic conditions in some lakes in northern Sweden were more common during cool climate phases of the Holocene. In addition, changes in ice cover can radically influence the underwater UV field and biological UV exposure in arctic waters (Belzile et al. 2001; Rautio and Korhola 2002). Later we demonstrate how even quite moderate changes in temperature along with changes in wind conditions may significantly affect the onset, stability, and break-up of thermal stratification. Changes in optical environments Several factors can affect the underwater light climate of arctic lakes, most importantly the chromophoric dissolved organic matter (CDOM), i.e., the humic and fulvic materials derived from terrestrial soils and vegetation or from microbial autochthonous production. CDOM plays a major role in biological ultraviolet (UV) radiation exposure in Arctic lakes because of its strong UV-absorbing properties and effects on transparency (Gibson et al. 2000; Rautio and Korhola 2002). Another important factor affecting penetration of UV radiation is water turbidity. Marked changes have occurred

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in both CDOM and water turbidity during the ontogeny of arctic and subarctic lakes in the region that are discussed in more detail below. Dissolved organic carbon (DOC) comprises the most important form of CDOM. An increasing volume of paleolimnological research on past DOC concentrations in arctic treeline regions of Europe has been accumulating. Diatoms have been used to reconstruct past variations in DOC (e.g., Bigler et al. 2002; Solovieva and Jones 2002). The transfer functions were developed from modern diatom-DOC relationships, as there are marked south–north gradients in current DOC concentrations in waters of arctic Fennoscandia (Korhola et al. 2002a). The northward reduction of terrestrial vegetation with latitude is accompanied by decreases in DOC that diatom distributions seem to follow. In some cases total organic carbon (TOC) has been reconstructed. However, in oligotrophic systems of arctic regions, TOC concentrations generally reflect DOC concentrations (Davis et al. 1985). Seppä and Weckström (1999) and Korsman (1999) have also reconstructed past lake water colour values. In general, there is a high positive correlation between total organic carbon concentrations and colour in waters of the Fennoscandian Arctic (r = 0.74, p < 0.01; n = 53; Korhola et al. 2002a), suggesting that lake colour is to a large extent determined by the supply of allochthonous dissolved organic matter. It is well known that shifts in vegetation and hydrology caused by climate warming or cooling affect the quantity of DOC that is exported from catchments to their receiving waters, in turn affecting underwater UV (Schindler et al. 1996; Pienitz and Vincent 2000). Thus, it is not surprising that the predicted long-term trends in DOC/colour in northern Fennoscandian lakes are closely linked to Holocene vegetation and soil development. The inferred organic carbon and colour values are usually low in the earliest Holocene reflecting the organically-poor mineral soils that were present in the lake catchments (Solovieva and Jones 2002; Figure 7). The rapid climate warming and the subsequent Holocene thermal maximum between ca. 8000 and 5000 cal. yr BP were accompanied by a northward movement of the treeline up to 200 km, and a corresponding increase in organic supply to the lakes reflected by the slight increase in inferred lake water DOC concentrations (Seppä and Weckström 1999; Solovieva and Jones 2002) (Figure 7). However, at high altitude sites with no or only sparse occurrence of forests during the warm mid-Holocene, reconstructed organic carbon concentrations remained unchanged (e.g., Toskaljavri in Figure 7). The late Holocene pattern in organic carbon and colour development is less systematic depending on lake altitude, soil characteristics, surrounding vegetation cover, acidification status and, in particular, development and occurrence of peatlands in the lake catchments. For example, in Korsman’s (1999) study covering the last ca. 4000 years, lake water colour increased in one lake, remained constant in three lakes, and fluctuated inversely with the pH in another lake. In high elevation areas where mountain birch/pine forests were replaced by alpine tundra some 5000 to 6000 cal. yr BP, there was a general tendency towards declining DOC/TOC concentrations up until the present (e.g., Toskaljavri and Lake Chuna in Figure 7). However, where extensive paludification occurred, rising DOC/TOC and colour values are noted during the latter half of the Holocene (e.g., Lake Tsuolbmajavri in Figure 7). The close coupling between TOC and the occurrence of peatlands was expected on the basis of modern limnological surveys, where a high positive correlation between TOC and catchment peatland

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Figure 7. Holocene diatom-inferred organic carbon (TOC and DOC) and Pinus sylvestris pollen influx values for two lakes in northwestern Finland and one on the Kola Peninsula, Russia. Toskaljavri from Seppä et al. (2002); Tsuolbmajavri from Seppä and Weckström (1999); Lake Chuna from Solovieva and Jones (2002). For Tsuolbmajavri, the diatom-inferred colour values are also shown.

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percentage is observed (Korhola et al. 2002a). In Fennoscandia, peatlands are considered to be the main source of humus to the lakes (Kortelainen 1993). Turbidity of the open water zone, especially in eutrophic waters, is often caused by phytoplankton, whereas particulates from catchment erosion (e.g., clays and silts), resuspended bottom sediments, and organic detritus from stream discharges also affect the water turbidity. An increased amount of particulates can smother benthic habitats resulting in a reduction of benthic organisms. Limited light penetration can, in turn, reduce the rate of photosynthesis resulting in a lower oxygen release into the water. Thus, turbidity is an important factor in controlling the structure of aquatic ecosystems. Geochemical data, sediment mineral magnetic data and diatom floristic compositions all suggest that turbid conditions prevailed in arctic lakes of our study region during the early Holocene (e.g., Snowball and Sandgren 1996; Korhola et al. 2000b; Virkanen 2000; Solovieva and Jones 2002). After regional deglaciation, high allochthonous input of clastic material from the bare catchments, together with stronger wind, maintained turbid conditions. High catchment erosion is suggested by high concentration of erosion indicators, such as Ca, K and Ti in the sediment geochemistry (Virkanen 2000), which may have affected water transparency and therefore underwater light and UV penetration. Erosion rates often remained high during the entire early Holocene until ca. 6000 cal. yr BP (Virkanen 2000) reflecting high winter precipitation and heavy spring snowmelt. A striking feature of lakes in the far north of Europe is that initial diatom communities are usually heavily dominated by pioneering benthic Fragilaria species, sometimes constituting more than 80% of the diatom flora (Seppä and Weckström 1999; Bigler et al. 2002; Solovieva and Jones 2002; Rosén et al. 2003). Fragilaria spp. are commonly associated with high environmental instability and are known to tolerate broad environmental gradients (from brackish to freshwater) and poor light conditions (Haworth 1976; Smol 1988; Denys 1990; Korhola 1995). The ability of small Fragilaria species to quickly reproduce and to tolerate short-term environmental fluctuations makes them very competitive, which is of vital benefit especially in unstable limnological conditions that prevailed during the early Holocene. On the basis of the Fragilaria dominance, turbid conditions during the early Holocene prevailed perhaps for some hundreds of years. However, Bigler et al. (2002) have shown that Fragilaria populations can maintain their dominance for quite a long time where late snow-beds have remained. In contrast, shorter term fluctuations in turbidity seem to be insufficient to cause any re-establishment of Fragilaria. For example, no changes in the diatom communities were found in Lake Toskaljavri during the period of a sudden increase in mineral erosion around 5000 cal. yr BP (Seppä et al. 2002). High levels of turbidity for a short period of time may therefore not be as crucial for aquatic organisms as lower levels that persist longer. Northern Fennoscandia still lacks the studies in which the past DOC or turbidity levels are connected to present-day bio-optical models in order to quantitatively infer past underwater light climates (paleo-optics), as has been done in North America. For example, Pienitz and Vincent (2000) estimated using their paleo-optical model that the diatom-inferred increases in DOC concentrations during a warming period some 6000 to 3500 cal. yr BP in Queen’s Lake in the Canadian Subarctic were equivalent to about a 100-fold decrease in biologically effective UV exposure. In contrast, decreasing DOC concentrations during the late Holocene cooling, from about 3500 years ago to the

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present, resulted in increased biological exposure that was two orders-of-magnitude greater than that associated with moderate (30%) ozone depletion. In all, the paleo-optical records show large regional differences in the magnitude and direction of change in underwater light conditions. These paleolimnological inferences suggest that, near the Fennoscandian treeline, past radiation environments of lakes have changed quite radically by climatic change and the associated changes in vegetation cover and soils. The records of DOC and turbidity are helpful in this regard, linking changes in UV exposure to past changes in climate, which may serve as a reference to predicted changes. Holocene trends in chemical attributes The overall chemical trend The current concept that lakes become more productive as they age has recently been challenged by a growing body of paleolimnological evidence. According to these data, many lakes have actually followed the opposite trend and became more dilute and unproductive over time (e.g., Digerfeldt 1971; Ford 1990; Korhola and Tikkanen 1991; Renberg et al. 1993; Saulnier-Talbot et al. 2003). Recently, Engstrom et al. (2000) addressed more specifically this question by studying a suite of recently formed lakes in Glacier Bay, Alaska. Their results showed that the lakes had become more dilute, more acid and more coloured with dissolved organic matter over time (and, by inference, less rather than more productive), with the terrestrial surroundings including the groundwater seepage exerting a major control over lake development. Apart from information about long-term development in lake water pH and dissolved organic carbon, there exist little data on long-term patterns in other chemical attributes in northern Fennoscandia. Grönlund and Kauppila (2002) recently applied the diatomtotal phosphorus (TP) transfer function developed by Kauppila et al. (2002) for southern Finland to reconstruct the Holocene TP history of Lake Soldatskoje, a small tundra lake located in the northern coastal area of the Kola Peninsula and surrounded by numerous archaeological dwelling sites. They discovered a progressively declining trend in TP concentrations towards the present. However, slightly more nutrient-rich conditions were inferred between ca. 5000 and 4000 BP (non-calibrated dates), presumably due to combination of local climate, watershed characteristics, and the influence of Stone Age humans (Grönlund and Kauppila 2002). Because the environment represented by the training set differed markedly from their study area on the Kola Peninsula, the results were used only qualitatively. In order to explore the direction and rate of change in overall chemical limnology, we applied the Engstrom et al. (2000) approach to our study lakes (unpublished data). Our previous investigations on surface sediment diatoms from 64 lakes throughout Lapland suggested that changes in water chemistry over space significantly controlled distributional changes in algal communities, although physical factors (altitude, temperature, depth) also were important forcing factors (Weckström and Korhola 2001). After elimination of the physical variables from the set of the environmental parameters, canonical correspondence analysis (CCA) of the diatom assemblages and associated water chemistry variables identified five significant water quality parameters, with the

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+4.0

major axis of variation in the diatom data being closely aligned with pH, Ca, and Mg, while the second axis was most strongly correlated with total nitrogen (TN) and DOC (Figure 8). CCA was then used to assess the effects of environmental factors on Holocene diatom succession in two lakes in northernmost Finland (Tsuolbmajavri, Toskaljavri) and one in Norway (Dalmutlad’do) (Figure 3) by plotting fossil diatom samples as supplementary samples in the CCA ordination bi-plot defined by modern diatom assemblages. A negative lake trajectory along the primary axis (defined by pH and base cations) suggested that the dominant limnological trend in all cases is a decline of pH, alkalinity and base cations. In Toskaljavri some dilution occurred with respect to DOC and TN. However, the overall results show that the lakes have experienced only minor changes with respect to the measured water chemistry variables during their history, as indicated by dense clustering of the core samples in the CCA space.

pH

Db Tsb

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Figure 8. Canonical correspondence analysis (CCA) ordination bi-plot defined by modern diatom assemblages and the five most significant water quality parameters from surface sediments of 64 lakes in northwestern Finnish Lapland. The sample scores of downcore diatom assemblages from three Holocene sites are shown; the core samples were treated as supplementary samples in the CCA. The shaded ellipses highlight the direction and magnitude of the trajectories along the temporal gradient. Tob = bottom sample of Toskaljavri, Tsb = bottom sample of Tsuolbmajavri, Db = bottom sample of Dalmutlad’do.

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Age (cal. BP)

Bigler et al. (2002) also used diatom assemblages and numerical analyses to assess limnological and environmental trends in five subarctic study sites in northern Sweden. In this case, chemical and physical explanatory variables were not separated, which to some extent complicates the interpretation because of the high inter-correlations between these two sets of variables. Although the diatom assemblages showed similar directional succession during the initial phases of the lake development, the timing and scale of development differed substantially between lakes. The authors concluded that site-specific features, such as catchment vegetation, hydrological conditions and in-lake processes, appear to control lake development in northernmost Sweden. They further postulated that the influence of long-term natural acidification on diatom assemblages progressively declined during the Holocene, whereas the influence of climatic factors increased. The comparison of limnological trends indicates that arctic lakes in northern Fennoscandia may indeed have developed along a common trajectory: the lakes have become more acidic and dilute with time, driven largely by broad-scale changes in the environment, such as climate variability, forest succession, and soil development. However, the rate and magnitude of these changes differ considerably between lakes depending on local geological and hydrological conditions. In Figure 9, we present a summary of the main Holocene environmental changes in aquatic and terrestrial ecosystems in the Fennoscandian arctic region. Below we discuss the changes in water quality and biology of the lakes in more detail.

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Figure 9. Chart summarizing major environmental changes and lake conditions since deglaciation in Fennoscandia and the Kola Peninsula. Note that the Shannon diversity (H´) classification is based on a limited number of diatom floristic studies referred to in this chapter.

Long-term changes in lake acidity Natural long-term acidification is a common feature of lakes situated in cold environments with thin soils and acid bedrock in the catchment. Excluding northern Norway and the southern part of the Kola Peninsula, where alkaline rocks predominate,

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the bedrock of arctic Fennoscandia and the Kola Peninsula is dominated by slightly acid rocks (i.e., granite and gneiss). The lakes in the region are characterized by low buffering capacity and are principially sensitive to acidification. Several studies concerning the Holocene acidification histories of lakes in arctic Fennoscandia and on the Kola Peninsula have been completed (Figure 10). Most of the pH reconstructions are based on changes in the diatom flora. Seppä and Weckström (1999) studied the acidification history of Lake Tsuolbmajavri, located slightly above the current pine treeline in northwestern Finnish Lapland. The current pH of the lake is 7.4. Although pH was noted to decrease almost gradually throughout the Holocene, the overall change was very small with the reconstructed pH lying within the error estimate of the diatom model (= 0.36 pH units). However, two periods with a slightly more evident drop in pH could be distinguished (Figure 10). The first decline in pH occurred between 8500 and 6000 cal. yr BP, which the authors related to the immigration of pine into the area, whereas the second acidification period about 3000 cal. yr PB was proposed to be due to the initiation of paludification. In both cases, changes in the catchment area likely induced increased leaching of humic and fulvic acids to the lake (e.g., Engstrom 1987). Organic acids derived from peatlands are known to strongly contribute to the acidity of lakes in Finland (Kortelainen and Mannio 1990; Kortelainen 1993). Several paleolimnological studies from southern Finland have demonstrated a clear relationship between the decline in lake water pH and the spread of peat in the catchment area (e.g., Tolonen et al. 1986; Korhola 1992). However, the influence of peatlands on the acidification status of lakes is probably less clear in northern Fennoscandia than in more southerly boreal environments owing to the less acidic nature of the mire water in the northern fens (Korhola et al. 2002a; Sjörs and Gunnarsson 2002). Since vegetation in the treeline area reflects primarily the prevailing climate, Seppä and Weckström (1999) concluded that Holocene climate dynamics have ultimately controlled the limnological succession of Lake Tsuolbmajavri. The first period of slight acidification was driven by warmer and drier climate favoring pine to spread, whereas the second episode was induced by cooler and moister climate promoting paludification. Thus, climate change may affect water quality via catchment processes, no matter the direction of change. These findings have some relevance also with regard to the expected future climate warming. It can be postulated that the predicted increase in air temperature in this area by some 4-5ºC during the 21st century (Boer et al. 1990) will have significant impacts on the water chemistry of arctic and subarctic lakes. A corresponding pH trend was also found in the nearby Lake Toskaljavri, located in the barren tundra, about 60 km north of Lake Tsuolbmajavri and some 100 m above the local pine treeline (Seppä et al. 2002). Although the bedrock of the catchment area partly consists of readily weathered dolomitic and calcium- and magnesium-rich limestones increasing the buffer capacity of the lake, the reconstructed pH has still decreased by ca. 0.3 to 0.4 pH units during the Holocene (Figure 10). The processes controlling the gradual decrease and small-scale fluctuations in pH are most probably the same as in the case of Lake Tsuolbmajavri, namely climatic variation and associated changes in soil-forming processes and catchment vegetation patterns (Figure 9). Korsman (1999) reconstructed the late Holocene pH history from five currently acidic lakes in northern Sweden (lakes N-R in Figure 13) in order to assess the exact timing of acidification. Excluding one lake, which was noted to have been acidic throughout the

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ca. 4000-year period covered by the sediment core, all remaining lakes had experienced slight natural acidification beginning thousands of years ago, as a result of soil-forming processes and natural changes in vegetation. Soil development would have resulted in an increased accumulation of organic acids, while the base cation utilization during vegetation growth would have resulted in a lowering of soil base saturation. Bigler et al. (2002) studied the Holocene acidification history of the small subarctic lake Vuoskkujávri near Abisko in northern Sweden. The lake, located at 348 m a.s.l., acidified according to the diatom-based pH reconstruction by about 0.4 pH units during the first half of the Holocene, after which the pH remained relatively constant (Figure 10). The early Holocene slow natural acidification trend seemed to correlate with the arrival of trees in the catchment, changes in erosion, and sediment organic content. No evidence was found that changes in land-use or reindeer herding would have affected the pH history during the last 1000 years.

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Figure 10. Diatom-inferred Holocene pH histories of lakes in arctic Fennoscandia and the Kola Peninsula. Toskaljavri is redrawn from Seppä et al. (2002), Tsuolbmajavri from Seppä and Weckström (1999), Vuoskkujávri from Bigler et al. (2002), and Lake Chuna from Solovieva and Jones (2002). Note the different scales for pH.

Solovieva and Jones (2002) studied the Holocene environmental history of an upland lake (Lake Chuna) on the Kola Peninsula, located above the present treeline at 475 m a.s.l., using pollen, diatom and geochemical records. The water of Chuna Lake is currently dilute, clear and slightly acidic with a pH around 6.4. Between ca. 9000 and 4200 cal. yr BP, the lake experienced a slow natural acidification resulting from humus accumulation and vegetation development. From 3800 cal. yr BP, the acid-base balance in the lake was apparently achieved and in the late Holocene no indication of further acidification was found. In all, inferred pH showed little change during the past 9000 years, with the overall pH fluctuation being around 0.4 pH units (Figure 10). By applying the diatom models developed by Weckström et al. (1997b) and Solovieva

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(2000), Grönlund and Kauppila (2002) observed a similar trend of progressively declining pH towards the present from Lake Soldatskoje, a small tundra site located in the northern coastal area of the Kola Peninsula. In summary, the spatial evaluation of changes in lake water acidity in arctic Fennoscandia and the Kola Peninsula shows slow change or no change at all during the Holocene. Excluding the initial transient alkaline period following deglaciation evident in some sites, we can estimate that the long-term natural rate of pH decline in the arctic lakes of the region is ca. 0.005 to 0.01 pH units in 100 yr. This is generally less than in more southerly sites in boreal and temperate Fennoscandia where rates between 0.01 to 0.03 pH units per 100 yr were observed (e.g., Tolonen et al. 1986; Renberg 1990; Korhola and Tikkanen 1991). The slow rate of pH change demonstrates the remarkable chemical stability of the arctic freshwater systems in the region. Holocene trends in aquatic communities Although aquatic organisms have been widely used to reconstruct past lake environments, surprisingly little data exist on aquatic succession itself. In particular, the mechanisms driving floral and faunal developmental patterns in remote arctic and subarctic lakes are poorly understood. However, critical examination would involve independent indicators for assessing cause and effect relationships, and to avoid circular reasoning. Table 1. The length of gradient of the first detrended correspondence analysis (DCA) axis in 12 study lakes from northern Fennoscandia and the Kola Peninsula expressed in standard deviation (SD) units of species turnover (S = Sweden, F = Finland, N = Norway, R = Russia).

Site Jeknajaure (S) Niak (S) Sjuodijaure (S) Seukojaure (S) Lake Njulla (S) Vuoskkujavri (S) Vuolep Njakajaure (S) Lake 850 (S) Tsuolbmajavri (F) Toskaljavri (F) Dalmutlad’do (N) Chuna Lake (R)

Diatoms 3.30 2.79 3.39 2.62 3.97 3.79 2.81 2.75 2.86 1.33 3.41 2.40

Chironomids

Cladocera

1.71 2.66 1.98 3.52 2.15 2.26 1.85 1.00 2.50

1.20

In Table 1, gradient lengths are presented for biological data sets from some Fennoscandian arctic lakes covering the entire Holocene. In each case, lengths of gradient were estimated by detrended correspondence analysis (DCA) with detrendingby-segments and downweighting of rare taxa. Basal samples were deleted if they

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contained only a few taxa that dominated the assemblages. Although we are aware of problems that are associated with the interpretation of such data (e.g., different sampling resolution and count sums), the results nevertheless yield some approximation of the overall changes that occurred in biotic communities over time. The average gradient length is about 2.6 standard deviation (SD) units of species turnover suggesting a modest amount of unimodality in the species response. Hence, the assemblages have gone through distinct changes over time, although no complete turnover of species assemblages has taken place at any site. Nevertheless, biological gradients are longer than would have been expected on the basis of rather stable chemical environments that seem to have been characteristic of the sites (see above), suggesting that other factors in addition to chemical variability have been responsible for shaping the aquatic communities. Tsuolbmajavri

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Figure 11. Temporal variations in diatom diversity (Shannon-Weaver diversity index H´) in three study sites. LOESS smoother is drawn with a span of 0.45.

There appears to have been a slight overall increase in algal diversity during the Holocene in the northern Fennoscandian lakes (Weckström et al. 2001; Figure 11). In their survey of modern algal communities across a latitudinal/altitudinal transect in northwestern Finnish Lapland, Weckström and Korhola (2001) found diatom-algal species richness to be highest in the mountain birch woodland zone (transition zone between spruce-pine-birch forest and treeless tundra), from where the diversity declined towards the cold, arctic tundra lakes. Major biotic boundaries with more distinct

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changes in species richness occurred at water temperatures around 13ºC and at around 9.5ºC. Thus, the species diversity of diatoms seems to be regulated to some extent by temperature and other climate-related factors. Sorvari et al. (2002) found diatom diversity to be negatively correlated with altitude/temperature in their study of the recent development of some arctic lakes in Finnish Lapland. However, these patterns have yet to be examined more carefully using more robust estimations of diversity.

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Figure 12. Zone limits of biotic stratigraphies of nine subarctic lakes. The zones are identified in each case by optimal sum of squares partitioning (Birks and Gordon 1985), and the broken-stick model (Bennett 1996) was used to test if the proposed zonations were statistically significant. Where six or more of the biostratigraphic boundaries overlap, a gray ‘summary’ zone is drawn. TSU = Tsuolbmajavri, TOS = Toskaljavri, SJU = Sjuodjijaure, JEK = Jeknajaure, NIA = Niak, SEU = Seukokjaure, VUO = Vuoskkujávri, NJU = Lake Njulla, CHU = Lake Chuna.

In an attempt to identify the time periods with the most significant shifts in the fossil species assemblages, we present the zone limits of biotic stratigraphies in nine study sites in Figure 12. The zones were identified in each case by optimal sum of squares partitioning (Birks and Gordon 1985); the broken-stick model was used to test if the proposed zonations were statistically significant (Bennett 1996). Our hypothesis was that any isochronism in the position of zone boundaries between the sites would point to

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a regional (external) cause, while stochastic structure would refer to local factors as being responsible for significant changes in lake biology. Dates where six or more of the biostratigraphic boundaries coincided in the nine lakes were found at around 1000, 3500-4000, 5000-6000, 7000-7700, and 9200-9800 cal. yr BP. The more or less synchronous change in aquatic assemblages possibly points to an external cause which, in the absence of human influence, most probably is regional climate. Indeed, the observed periods of change correspond with the time periods of significant cooling or warming in the temperature reconstruction by Korhola et al. (2000b) from Lake Tsuolbmajavri. The recent study by Seppä et al. (2002) also concludes that the ultimate determinant behind the changes in the Holocene chironomid succession in Lake Toskaljavri was changing climate, as assessed by canonical ordination analyses. Recent limnological changes Industrial acidification Several recent studies have applied diatom transfer function techniques to lakes of northern Fennoscandia and the Kola Peninsula with the aim of providing temporal perspectives on possible recent surface water acidification. The available pH reconstructions dealing with the recent acidification history are summarized in Figure 13. Collectively, these studies indicate that no substantial changes in acidification status of the lakes have taken place; the majority of the lakes that are presently acidic have long been acidic due to natural processes. This is not surprising, as most of the arctic Fennoscandian lakes receive very low levels of atmospheric deposition of environmental contaminants (e.g., sulphur and heavy metals) compared with most other regions in Europe (Hettelingh et al. 1991, 1992; Lükewille 1994), and have generally negligible cultural activity in their catchment areas. Some of these studies are reviewed in more detail below. Korsman (1999) applied the so-called ‘top–bottom’ approach, in which one sediment sample was taken from the top of the core representing present assemblages, and another sample deeper down in the core representing pre-industrial assemblages, to infer changes in lake acidity in 118 northern Swedish lakes. According to his results, pH had decreased significantly in eight lakes, five lakes had a significant increase, while the majority of the lakes had remained unchanged. Results from his study did not support the hypothesis of large-scale modern acidification in northern Sweden. Similarly, Sorvari et al. (2002) did not find any indications of pH changes in five arctic tundra lakes that they investigated using high-resolution sampling for the last ca. 200 years in northwestern Finnish Lapland (Figure 13H-L). In comparison to little polluted northwestern Fennoscandia, the Kola Peninsula is exceptional in the arctic pollution context due to its huge sulphur emissions that originate from the large non-ferrous smelters in the area. The area is characterized by strong environmental contrasts; large areas are almost in pristine condition whereas human impacts have created ‘technogenic wastelands’ around the smelting and mining industries, with some of the largest heavy metal and sulphur emissions in the world. After the start of the smelting activities (1930s), the amount of accumulated sulphur emissions are estimated to have exceeded 20,000,000 t (Kashulina and Reimann 2002).

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Contrasting views have been presented about the effects of the Kola smelter pollution on the water quality in the surrounding arctic lakes. For example, Henriksen et al. (1997a,b) report that the critical loads of acid deposition have been exceeded in 50 to 70% of the lakes on the Kola Peninsula and in the Norwegian-Russian border areas, while Moiseenko et al. (1995) report that extensive acidification has taken place in the mountain lakes of the Peninsula. In contrast, Reimann et al. (2000) recently argued that the emissions do not seem to present a major environmental threat on a regional scale, that the major effects of the emissions occur in the near vicinity of the smelters, and that “unexpected/natural processes (e.g., sea-salt deposition) may be more dominant than more obvious (pollution-related) processes”.

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Figure 13. Diatom-inferred recent pH histories of 16 lakes and two ponds in northern Fennoscandia. Note the different scales for age, depth and pH. Figures (A-C) are from Korhola et al. (1999), figures (D-E) from Erola (1999), figures (F-G) from Bigler and Hall (2003), figures (H-L) from Sorvari et al. (2002), figure (M) from J. Weckström (unpublished), and figures (N-R) from Korsman (1999).

Several paleolimnological studies concerning the impacts of Kola smelter pollution have been completed recently. Korhola et al. (1999) examined the pH history of the last ca. 150 years of small acid-sensitive lakes in eastern Lapland, close to the Kola

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emission sources. Stable pH reconstructions suggested no substantial changes in the acidification status of the lakes despite the high acid deposition. They concluded that the watershed and in-lake alkalinity-generating processes are still effectively countering the acidification. Furthermore, the alkaline nature of the lakes appears to have suppressed the environmental availability of the toxic metals, which are found in increasing quantities in the surficial sediments of some Lapland lakes (Dauvalter 1998). Weckström et al. (2003) applied the ’top–bottom’ approach to 32 lakes distributed around the smelter industries on the Kola Peninsula. Using a quantitative diatom-based transfer function, they were able to reconstruct the pre-industrial lake water pH and to compare it with the modern measured values. They discovered significant changes in the acidity status in only seven lakes of which five were located close to the metallurgical complexes. They concluded that most of the studied lakes have undergone long-term natural acidification, while human-induced acidification was restricted to the immediate vicinity of the smelters. Thus, their study contradicted the idea of widespread acidification of arctic lakes due to sulphur pollution from the Kola smelting and mining industry. However, this situation may not persit for long if emissions continue at present high levels. Recent changes in bio-assemblages and their possible causes Despite the small changes in lake acidity, noticable changes in floral and faunal communities have occurred in lakes of northern Fennoscandia within the recent past. Sorvari and Korhola (1998) studied the recent (ca. 150 years) environmental history of Lake Saanajärvi, located in the barren tundra at 679 m a.s.l. in the northwestern part of Finnish Lapland. They found distinctive changes in the diatom community composition with increasing occurrences of small planktonic diatoms such as Cyclotella come nsis and C. glomerata starting about 100 years ago. Since no changes in lake water pH could be inferred (Figure 13H), and because both airborne pollution and catchment disturbances are known to be almost non-existent in the region, they postulated that recent warming in the Arctic has been the main reason behind the observed ecological change. Meanwhile, Erola (1999) studied two ponds located at 520 and 930 m a.s.l. in the same region, but found only a slight change in diatom assemblages in just one of these ponds. The weaker response was probably related to the shallow nature (1.5 m and 3.3 m, respectively) of these ponds compared to Lake Saanajärvi and the resulting dominance of the diatom flora by periphytic communities. As shown by Moser et al. (2000), the abundance of planktonic diatoms never exceeded 20% in arctic Fennoscandian lakes with < 5 m depth, highlighting the sparse availability of favorable habitats in shallow lakes for the growth of planktonic algae. To test the climate warming hypothesis further, Korhola et al. (2002b) analysed additional sedimentary proxy indicators, such as cladocerans, chrysophycean stomatocysts, plant pigments and chironomids, in addition to mineral magnetics and various sediment quality indices, from Lake Saanajärvi. The biological and sedimentological records were contrasted with a 200-year long climate record specifically reconstructed for the region using a compilation of measured meteorological data and various proxy sources. They found synchronous changes in

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lake biota and sedimentological parameters that seemed to occur in parallel with the increasing mean annual and summer temperatures starting around the 1850s. It was hypothezised that the rising temperature had increased the metalimnion steepness and thermal stability in the lake, which in turn supported increasing productivity by creating more suitable conditions for the growth of plankton. In order to test the spatial representativeness of these ecological changes, Sorvari et al. (2002) analysed recent diatom communities in four other remote and unpolluted lakes in northwestern Finnish Lapland at high resolution. They were able to trace a distinct change in algal floral assemblages that paralleled the post-19th century arctic warming as indicated by examination of long-term instrumental series, historical records of ice cover and tree-ring measurements. Despite the different physico-chemical conditions between the lakes, the change was synchronous in all lakes and affected the overall diatom diversity. In most cases, the change was from benthic to planktonic assemblages, characterized by the increase of small centric Cyclotella diatoms. When the change was summarized as a principal components analysis (PCA), all study sites were found to have statistically significant correlations between spring (March to May) mean air temperature and PCA axis 1 sample scores (Figure 14). As there is no direct human activity near the lakes, and as no changes were found in the diatom-inferred lake pH (Figure 13H-L), they concluded that the change in aquatic communities was caused predominantly by the 19th century warming of the Arctic. The longer duration of the icefree season and increased water column stability, both heavily affected by spring temperatures in Lapland, were assumed to be the most probable causative mechanisms behind the change. It was postulated that during the Little Ice Age many of the lakes in the region were only weakly stratified thermally, while steeper stratification developed in the lakes as a result of the subsequent climate warming (Sorvari et al. 2002). That thermal stability in the lakes of the region can be affected by even minor changes in climate is well documented by our contemporary monitoring. For example, in 2001, average summer air and water temperatures in northwestern Finnish Lapland were 1 to 2qC lower than usual, resulting in a development of an unusually weak thermocline in Lake Saanajärvi (Zmax = 24 m), a typically well-stratified and dimictic subarctic lake (unpublished data). The situation eventually led to a complete break-up of thermal stratification in late July in association with a temporary drop in temperature and high winds. In contrast, an exceptionally sharp metalimnion and stable stratification occurred in the summer of 2002 when summer air and water temperatures were about 3qC higher than on average (unpublished data). Since lakes and their catchments behave individually in many respects, and even closely located lakes can respond differently to environmental changes, further work on high latitude lakes is needed to get a more comprehensive picture of past ecological changes and processes affecting them. Hence, far-reaching general conclusions of past limnological conditions should not be made without a reasonably large data set. Nevertheless, the results from arctic Fennoscandia are consistent with records of shifts in algal communities over the past 150 years observed elsewhere in the Arctic (e.g., Douglas et al. 1994; Gajewski et al. 1997; Wolfe and Perren 2001; Rühland et al. 2003) that suggest a regional rather than a local cause. If the rate of warming accelerates in the future, as has been predicted by climate models, we should expect biological and ecological changes in arctic and subarctic lakes to continue.

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Figure 14. Comparison of the diatom assemblage changes with the regional temperature anomalies. (A) Principal component analysis (PCA) primary axes scores derived from the correlation matrices of diatom percentage counts of five subarctic lakes and (B) Spring temperature anomalies (ºC) for northwestern Finnish Lapland, smoothed using a 10-year low-pass filter (modified from Sorvari et al. 2002).

Summary The arctic region of northern Fennoscandia and the Kola Peninsula is littered with small lakes and ponds that are predominantly shallow, clear, dilute and oligotrophic. Despite growing interest concerning their water quality and ecological status, the environmental histories of these arctic and subarctic water bodies are still relatively poorly known. Most of the previous paleo-studies have focused on the climatic and environmental history of the region in general, not on the water bodies themselves. This review chapter

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aimed to compile and summarize the data that are available on the formation, ontogeny, and limnological trajectories of these interesting and unique ecosystems. Some unpublished material was also presented. Most of the lakes in the region are of glacial origin, although numerous water bodies have also originated via deflation and cryogenic processes, from former river meanders, or through emergence from the sea as a result of glacio-isostatic rebound. The sediment accumulation rates are very slow, varying from 0.01 cm yr-1 in the arctic tundra sites to 0.05 cm yr-1 in more southerly treeline lakes. The sediments are generally poor in organics, with most of the organic matter being of allochthonous origin from outside the water body. Many closed-basin lakes located on permeable soils have experienced marked fluctuations in their water-levels, which are reflected in the sediment as changes in sedimentation rates, variations in organic matter flux, and compositional changes in subfossil fauna and flora. The lake-levels were high during the early Holocene, faced marked reduction up to 4-6 m in the mid-Holocene (ca. 6000 to 4000 cal. yr BP), and rose again during the latter part of the Holocene. Distinct changes have also taken place in the thermal features of the lakes. During the ‘Holocene thermal maximum’ (ca. 8000 to 5000 cal. yr BP), epilimnion temperatures were about 1 to 2ºC higher than today, resulting in increased thermocline depth and length of the ice-free season. It is assumed that some of the presently monomictic (isothermal) water bodies may have become stratified in summer during the Holocene warm interval. A decrease in ice cover duration would in turn have reduced the winter anoxia common at these latitudes. An opposite development has presumably taken place during the cool periods such as the Little Ice Age, with more ice, shorter open water season, more effective water column mixing, and more frequent late-winter anoxia. In contrast to often quite distinct changes in physical limnology, changes in chemical limnological conditions have been relatively moderate and smooth during lake development. As a result of changing climate and successional changes in surrounding vegetation and soils, lakes close to the present treeline are typically characterized by a progressive decline in pH, alkalinity and base cations, and a corresponding increase in dissolved organic carbon (DOC). In contrast, lakes in the barren arctic tundra at higher altitudes manifest remarkable chemical stability throughout the Holocene. No evidence of widespread recent acidification can be found on the basis of quite extensive paleolimnological assessments in arctic Europe. However, fine-resolution studies from a number of remote lakes in the region demonstrate that aquatic bio-assemblages have gone through distinct changes that parallel the post-19th century of the Arctic warming. A growing body of evidence suggests that the driving force behind the shifts in physical, chemical and biological conditions during lake histories has been the changing climate, apparently in conjunction with complex interactions of local and regional factors, such as erosion intensity, weathering rates, soil and vegetation succession, and human activity. This review points out the need for further paleolimnological studies of the freshwaters in the area in order to obtain a more complete regional picture of lake development, natural variability and the role of humans in shaping these systems.

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Rosén P., Segerström U., Eriksson L., Renberg I. and Birks H.J.B. 2001. Holocene climatic change reconstructed from diatoms, chironomids, pollen and near-infrared spectroscopy at an alpine lake (Sjuodjijaure) in northern Sweden. The Holocene 11: 551-562. Rosén P., Segerström U., Eriksson L. and Renberg I. 2003. Do diatom, chironomid, and pollen records consistently infer Holocene July air temperature? A comparison using sediment cores from four alpine lakes in northern Sweden. Arct. Ant. Alp. Res. 35: 279-290. Rühland K., Priesnitz A. and Smol J.P. 2003. Paleolimnological evidence from diatoms for recent environmental changes in 50 lakes across Canadian arctic treeline. Arct. Ant. Alp. Res. 35: 110-123. Saulnier-Talbot É., Pienitz R. and Vincent W.F. 2003. Holocene lake succession and palaeo-optics of a Subarctic lake, northern Québec, Canada. The Holocene 13: 517-526. Schindler D.W., Bayley S.E., Parker B.R., Beaty K.G., Cruikshank D.R., Fee E.J., Schindler E.U. and Stainton M.P. 1996. The effects of climatic warming on the properties of boreal lakes and streams at the Experimental Lake Area, northwestern Ontario. Limnol. Oceanogr. 41: 1004-1017. Seppä H. and Weckström J. 1999. Holocene vegetational and limnological changes in the Fennoscandian tree-line area as documented by pollen and diatom records from Lake Tsuolbmajavri, Finland. Écoscience 6: 621-635. Seppä H. and Birks H.J.B. 2001. July mean temperature and annual precipitation trends during the Holocene in the Fennoscandian tree-line area: pollen-based climate reconstructions. The Holocene 11: 527-539. Seppä H. and Birks H.J.B. 2002. Holocene climate reconstruction from the Fennoscandian treeline area based on pollen data from Toskaljavri. Quat. Res. 57: 191-199. Seppä H., Nyman M., Korhola A. and Weckström J. 2002. Changes of tree-lines and alpine vegetation in relation to post-glacial climate dynamics in northern Fennoscandia based on pollen and chironomid records. J. Quat. Sci. 17: 287-301. Seppälä M. 1971. Evolution of eolian relief on the Kaamasjoki-Kiellajoki river basin in Finnish Lapland. Fennia 104: 1-88. Shemesh A., Rosqvist G., Rietti-Shati M., Rubensdotter R., Bigler C., Yam R. and Karlén W. 2001. Holocene climatic change in Swedish Lapland inferred from an oxygen-isotope record of lacustrine biogenic silica. The Holocene 11: 447-454. Sjörs H. and Gunnarsson U. 2002. Calcium and pH in north and central Swedish mire waters. J. Ecol. 90: 650-657. Smol J.P. 1988. Paleoclimate proxy data from freshwater arctic diatoms. Verh. Int. Ver. Limnol. 23: 837-844. Snowball I.F. and Sandgren P. 1996. Lake sediment studies of Holocene glacial activity in the Kårsa valley, northern Sweden: contrasting opinions. The Holocene 6: 367-372. Snowball I.F., Sandgren P. and Petterson G. 1999. The mineral magnetic properties of an annual laminated Holocene lake sediment sequence in northern Sweden. The Holocene 9: 353-362. Snowball I.F., Zillén L. and Gaillard M.-J. 2002. Rapid early-Holocene environmental changes in northern Sweden based on studies on two varved lake-sediment sequences. The Holocene 12: 7-16. Snowball I.F., Korhola A., Briffa K. and Koç N. Holocene climate dynamics in high latitude Europe and the North Atlantic. In: Battarbee R.W., Gasse F. and Stickley C.E. (eds), Past Climate Variability Through Europe and Africa. Kluwer Academic Publishers (in press). Snyder J.A., MacDonald G.M., Forman S.L., Tarasov G.A. and Mode W.N. 2000. Postglacial climate and vegetation history, north-central Kola Peninsula, Russia: pollen and diatom records from Lake Yarnyshnoe-3. Boreas 29: 261-271. Solovieva N.A. 2000. A palaeolimnological study of Holocene environmental change in a small upland lake from the Kola Peninsula, Russia. Ph.D. Thesis, University College London, 413 pp.

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Solovieva N.A. and Jones V.J. 2002. A multiproxy record of Holocene environmental changes in the central Kola Peninsula, north-west Russia. J. Quat. Sci. 17: 303-318. Sorvari S. and Korhola A. 1998. Recent diatom assemblage changes in subarctic Lake Saanajärvi, NW Finnish Lapland, and their palaeoenvironmental implications. J. Paleolim. 20: 205-215. Sorvari S., Rautio M. and Korhola A. 2000. Seasonal dynamics of subarctic Lake Saanajärvi in Finnish Lapland. Verh. Int. Ver. Limnol. 27: 507-512. Sorvari S., Korhola A. and Thompson R. 2002. Lake diatom response to recent Arctic warming in Finnish Lapland. Global Change Biol. 8: 171-181. The MOLAR Water Chemistry Group 1999. The MOLAR project: atmospheric deposition and lake water chemistry. J. Limnol. 58: 88-106. Tolonen K., Liukkonen M., Harjula R. and Pätilä A. 1986. Acidification of small lakes in Finland documented by sedimentary diatom and chrysophycean remains. In: Smol J.P., Battarbee R.W., Davis R.B. and Meriläinen J. (eds), Diatoms and Lake Acidity. Jonk Publishers, Dordrecht, pp. 169-199. Vassiljev J. 1998. The simulated response of lakes to changes in annual and seasonal precipitation: implication for Holocene lake-level changes in northern Europe. Clim. Dyn. 14: 791-801. Virkanen J. 2000. The effects of natural environmental changes on sedimentation in Lake Kuttanen, a small closed lake in Finnish Lapland. The Holocene 10: 377-386. Weckström J. and Korhola A. 2001. Patterns in the distribution, composition and diversity of diatom assemblages in relation to ecoclimatic factors in Arctic Lapland. J. Biogeogr. 28: 31-45. Weckström J., Korhola A. and Blom T. 1997a. The relationship between diatoms and water temperature in thirty subarctic Fennoscandian lakes. Arct. Alp. Res. 29: 75-92. Weckström J., Korhola A. and Blom T. 1997b. Diatoms as quantitative indicators of pH and water temperature in subarctic Fennoscandian lakes. Hydrobiologia 347: 171-184. Weckström J., Korhola A. and Seppä H. 2001. Climate-driven aquatic ecosystem changes in Finnish Lapland during the Holocene as documented by diatoms. PAGES-PEPIII conference: Past climate variability through Europe and Africa, 27-31 August, Aix-en-Provence, France. Abstract book, p. 159. Weckström J., Korhola A., Snyder J.A., Laing T.E. and MacDonald G.M. 2003. Diatom inferred acidity history of 32 lakes on the Kola Peninsula, Russia. Wat. Air Soil Pollut. 149: 339-361. Wolfe A.P. and Perren B.B. 2001. Chrysophyte microfossils record marked responses to recent environmental changes in high- and mid-Arctic lakes. Can. J. Bot. 79: 747-752. Yu G. and Harrison S.P. 1995. Holocene changes in atmospheric circulation patterns as shown by lake status changes in northern Europe. Boreas 24: 260-268.

14. PALEOLIMNOLOGICAL STUDIES FROM THE ANTARCTIC AND SUBANTARCTIC ISLANDS

DOMINIC A. HODGSON ([email protected]) British Antarctic Survey Natural Environment Research Council High Cross, Madingley Road Cambridge, CB3 0ET United Kingdom PETER T. DORAN ([email protected]) University of Illinois at Chicago Department of Earth and Environmental Sciences 845 W. Taylor St. Chicago, Illinois 60607-7059, USA and DONNA ROBERTS ([email protected]) ANDREW MCMINN ([email protected]) Institute of Antarctic and Southern Ocean Studies University of Tasmania Private Bag 77 Hobart, Tasmania 7001 Australia

Key words: Antarctica, Paleolimnology, Climate change, Sea-level, Glaciology, Ice cores, Marine cores, Geomorphology, Earth system science

Introduction The Antarctic is a remarkable continent – remote, hostile and uninhabited. Yet, it is of key importance in Earth system science. With an area of 13.2 million km2 it is about half the size of North America and 1.3 times the size of Europe. Almost the entire land surface of Antarctica is covered by a vast ice cap with an area more than six times larger than that in Greenland. In places, this ice cap is 4 km thick. It contains over 70% of the world’s freshwater and over 90% of the world’s ice. The surrounding ocean freezes in winter to cover an area 1.5 times the size of the continent. Antarctica is also distinguished by being

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the highest, windiest, coldest and driest of the continents. Being the main storehouse of the world’s ice, some of it close to the melting point, environmental processes taking place in the Antarctic affect the world’s climate and its oceans, linking the continent inextricably to what is experienced thousands of kilometres away. In helping us to understand global change, the Antarctic has a crucial role to play. The role of the Antarctic in Earth system science The Milankovich cycle of 100,000 years orbital eccentricity has been widely accepted as the ‘pacemaker’ of environmental change, controlling the amount of solar radiation reaching the Earth (Hays et al. 1976; Berger 1992). The other astronomic cycles of tilt and precession, in theory, should redistribute the radiation between seasons in the two hemispheres with glaciations occurring alternately in the northern and southern hemispheres marked by a warming or cooling of around 2°C. Evidence in ice cores from Antarctica (Vostok, Dome C, Figure 1) and Greenland (named GISP 2 and GRIP), however, has shown that this is not the case and that the polar glaciations are, in fact, approximately ‘in phase’ and involve temperature changes far greater than 2°C (Roberts 1998). Thus, astronomical cycles are not able to adequately explain the timing of glacial cycles and climate change, and other factors are clearly acting to amplify and, in some cases, over-ride orbital forcing. Of these factors the internal feedback effects between ice, ocean and climate in the polar regions play a key role. These often involve changes in the physical properties of water, specifically freezing point, density and state and ultimately result in an exchange of heat and mass between the ice sheets, oceans, land, atmosphere and space. An example of these feedback loops can be seen in one theory of Antarctic deglaciation at the beginning of the Holocene. This theoretical sequence started when a minor warming caused an increase in the melting of the ice resulting in a minor change in sea-level. This in turn destabilized the floating ice shelves, which broke up and no longer protected the grounded ice sheet margin from wave attack and disintegration. The result was that the ‘cork was out of the bottle’ and the glaciers behind the ice sheet were able to flow into the sea further raising sea-level until other feedbacks were able to slow down the system (Roberts 1998). It can be seen, from this example, how a small external forcing can precipitate a run-away series of events that result in a series of changes of much greater magnitude. Other dynamic factors which can cause feedback effects in the Antarctic include changing volumes of cold, fresh meltwater and the production of dense saline brines (Antarctic Bottom Water) during sea-ice formation which can influence global ocean circulation. Evaporation rates from the surface, and changing surface albedo effects associated with sea ice extent are also important (Roberts 1998). However, it is the concentration of greenhouse gases, in particular atmospheric CO2, that offers the most convincing explanation for the north-south synchronisation of glaciations (Alley 2000). Physical factors can also cause an imperfect relationship of cryosphere-driven climate change with orbital eccentricity. One of these is the observation that the ice sheets waste faster than they grow (Hays et al. 1976). This constrains the timing of the glaciations and the behaviour of the cryosphere. Humans are also playing an important role with greenhouse gases (from the burning of fossil fuels) and pollutants creating changes in the Earth’s atmospheric composition that may already be disrupting the climate through additional forcing processes.

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In understanding the scale and scope of natural climate change, its deviation from the theoretically predictable orbital forcing and the influence of human activities, the polar regions, particularly the Antarctic, are playing a pivotal role. This has been made possible by its remarkable stratigraphic archives of global change. The Antarctic as a recorder of environmental change Instrumental data from the Antarctic only cover the last four decades. However, archived in the 4 km thick ice sheet is a 420,000 year record of past climate (Petit et al. 1999) with new drilling at Dome C likely to extend it towards 900,000 years (Wolff 2002). Here water molecules, particles and gas bubbles trapped in the ice archive changes in atmospheric composition. Analysis of ratios of hydrogen and oxygen isotopes in water molecules, and concentrations of methane, carbon dioxide and other gases in trapped bubbles has enabled scientists to infer past temperatures and other environmental parameters (Lorius et al. 1990). Salt concentrations give information on sea ice production and extent (Wolff et al. 2002). The ice cores also contain evidence of global pollution by industry, agriculture and atomic bombs. Sediments accumulating in the seas around the Antarctic also provide an archive of oceanic and climate changes typically spanning the late Quaternary (Grobe and Mackensen 1992), but occasionally dating as far as the Cretaceous-Cenozoic period (Barrett 2001). These have enabled reconstructions of the extent of ice cover during the most recent glacial cycles, the growth and decay of the region’s ice shelves, changing oceanic temperatures, and the production of the Antarctic Bottom Water which drives global thermohaline circulation (Anderson 1999). Stratigraphic records have also accumulated in lakes occupying terrestrial depressions around the coast of Antarctica. With a few exceptions, these are typically of Holocene age and post-date the last deglaciation. With their water balancing on the physical interface between solid and liquid states, these lakes register terrestrial environmental changes and often amplify the environmental signals in the ice and marine cores and can thus provide detailed records of climate change and its influence on the biosphere, hydrosphere and cryosphere. Their living communities and sedimentology rapidly respond to local and/or regional changes in temperature, precipitation, evaporation, light, ice cover, deglaciation and ice shelf decay. They also record rates of isostasy and relative sea-level change. It is this latter attribute that makes them (along with raised beach deposits) the critical link between the records from the ice cores and marine cores in that they can show, unequivocally, how changes in the mass balance of the ice sheets have influenced sea-level. In combination the above-mentioned stratigraphic archives, together with glacial and geomorphological records, have provided much of the scientific data behind our understanding of global environmental and climate change. They have permitted scientists to begin evaluating current changes in the context of long-term natural variability and, ultimately, will enable better predictions of future trajectories. In particular, these stratigraphic archives are helping to identify the ‘pendulum’ of (orbitally forced) natural climate change and determine whether human activities are buffering the climate system or affecting the balance in a way that could precipitate a feedback event resulting in a rapid global climate change. Ice and marine cores have already shown that, even under natural

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conditions, the Earth has experienced rapid climate changes known as Dansgaard-Oeschger (Dansgaard et al. 1993) and Heinrich events (Heinrich 1988), on a scale that could devastate modern civilisations. Indeed, stratigraphic studies show that, although the Holocene period has been relatively stable, rapid climate changes are a common feature of Earth’s history with natural jumps between fundamentally different modes of operation occurring in as little as a few years (Alley 2000). In the following sections we describe the range of lake types found in the Antarctic. We then outline how the accumulated sediments have been used to address some important global questions and how Antarctic paleolimnology is playing an integral role enhancing and developing the knowledge gained from ice core, marine core and geomorphological records.

Figure 1 . Map of Antarctica showing the locations of paleolimnological studies and other sites mentioned in the text. Dark shaded areas indicate the Antarctic ice shelves.

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Antarctic limnology Only 0.35% of the Antarctic continent is ice-free (Fox and Cooper 1994). This mainly consists of mountain peaks (nunataks) protruding from the polar ice cap. However, 1-2% of this includes coastal oases that have been exposed, both through postglacial retreat of the ice cap and isostatic rebound of the Earth’s crust following the most recent deglaciation. These oases contain the majority of Antarctica’s lakes. The largest oasis is situated near the coast of the Ross Sea and is known as the McMurdo Dry Valleys. Other large oases are found in eastern Antarctica including the Vestfold Hills, Larsemann Hills, Bunger Hills, Schirmacher Oasis and Syowa Oasis. There are also large ice-free areas along the Antarctic Peninsula, for example, the east coast of Alexander Island (Figure 1). Numerous smaller oases, archipelagos and isolated islands make up the remainder of the ice-free area. Even though the Antarctic is the repository of more than 70% of the world’s freshwater, less of it is available in liquid form than on any other continent. Average precipitation (as snow) amounts to only 100-150 mm of water equivalent a year (Weller et al. 1987), making the area technically a desert. Despite this, water does accumulate in areas where solar radiation and advected heat promote melting. This process is self-perpetuating since, as bare ground is exposed, the albedo of the region drops and more radiation is absorbed. As a result, the ice-free oases fringing the continent contain some of the most unusual and interesting lake systems on Earth. Many of the lower altitude lakes have been formed as a direct result of isostatic rebound after postglacial retreat of ice (Zwartz et al. 1998a; Verleyen et al. 2003a). In most cases these become ‘closed’ (endorheic) lakes with no surface outflow, or ‘open’, connected to others by inflow and outflow streams. Both types of lake are common in Antarctic oases such as the Larsemann Hills, Vestfold Hills, Schirmacher Oasis, Bunger Hills, and on the Antarctic Peninsula and maritime islands. In general, the sea water trapped during isostatic uplift has been flushed out of the open lakes or diluted with fresh meltwater (Figure 2a). However, many of the closed lakes have subsequently experienced an excess of evaporation over precipitation and have gradually become saline (Figure 2b). Examples of these are found in the Rauer Islands (Hodgson et al. 2001b), Vestfold Hills (Roberts and McMinn 1996) and McMurdo Dry Valleys (Doran et al. 1994). Some of these saline lakes become stratified for part of the year (monomictic) and others are permanently stratified (meromictic), retaining a trapped layer of salt water below their fresh surface waters. Other lakes, located above the maximum limit of Holocene relative sea-level, occupy rock hollows formed by glacier and ice-cap retreat (Figure 2c). These are common in the maritime Antarctic islands such as Signy Island, King George Island and Livingston Island. They can also be open or closed and range from freshwater to saline. Lakes have also formed in marine embayments or fjords, dammed by advancing glaciers and ice shelves, and isolated from the sea. These are called epishelf lakes (Figure 2d). In these systems the marine water has been replaced over a period of time by meltwater from glaciers and snow. In some cases the ice shelf forms an incomplete seal and a hydrological connection to the sea persists under the ice shelf or glacier, sometimes more than 100 m below sea-level. When this happens the epishelf lake becomes stratified with freshwater overlying a layer of saline water. The seawater flows under the ice shelf or glacier, maintaining tidal regime, which can result in tide cracks appearing in the ice around the edges of the lakes. Two examples are found on the coast of Alexander Island (Moutonnée

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Figure 2. Photographs showing selected Antarctic lake types, (a) an isostatically uplifted ‘isolation basin’ lake on Pourquoi-Pas Island (in the foreground), (b) an open basin freshwater lake on Horseshoe Island, (c) a hypersaline closed basin lake in the Rauer Islands, (d) Moutonnée Lake – an epishelf lake on Alexander Island, formerly a marine basin, now a stratified lake impounded by George VI ice shelf (in the foreground), (e) Lake Hoare in the McMurdo Dry Valleys, and (f) ephemeral supra-glacial meltpools on the Sörsdal Glacier in the Vestfold Hills, each of these pools is approx. 250 m wide.

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and Ablation Lakes). Other epishelf lakes are currently isolated from the sea, formed during earlier periods of relative sea-level rise, lifting grounded ice masses and allowing sea water to penetrate into the lower part of the water column, or connection to marine waters during periods of reduced freshwater input (Gibson and Andersen 2002). Epishelf lakes have been described from the Bunger Hills (Doran et al. 2000; Gibson and Andersen 2002; Verkulich et al. 2002), Amery Oasis (Adamson et al. 1997; Laybourn-Parry 2001) and the Schirmacher Oasis (Korotkevich 1960) and Alexander Island (Heywood 1977). Lakes and pools are also found against glacier fronts (Figure 2e) or beside ice shelves and in depressions on glaciers, ice sheets and ice shelves (Figure 2f). These are called epiglacial and supra-glacial lakes, respectively. Supraglacial lakes are often ephemeral, forming during the summer melt. Nevertheless, they have been proposed as one of the main refugia for the eukaryotic biota in the controversial Snowball Earth hypothesis when the Earth underwent extreme freeze drying events between 750 million and 580 million years ago (Vincent et al. 2000). Epiglacial lakes may persist for several years often with frequent changes in water-level and morphology on account of glacial movements and changing meltwater inputs. Some, such as Lake Untersee, are older than 500 years (Wand and Perlt 1999). Kettle ponds can also form in depressions in ice-cored moraines. Examples of these can be found on Bratina Island (Fernandez-Valiente et al. 2001). Recent studies have also revealed an entire world of subglacial lakes in Antarctica (Siegert 2000). These lakes, which lie between 2036 to 4200 m under the eastern Antarctic ice sheet and range in size from 1 km up to 241 km long, were first detected by airborne radio echo sounding in the 1970’s. The largest and deepest of these is Lake Vostok. This lake covers an area of 14,000 km2 (about the same size as Lake Ontario, Canada), has a mean depth of 150 m and a water residence time estimated to be around 50,000 years. Others are found in the Dome C and Ridge B regions of eastern Antarctica. The paleolimnology of these lakes is considered by Doran et al. (this volume). Physical, chemical, and biological features of the lakes Physically, the lakes experience low temperatures (< 10°C) and have varying degrees of ice cover usually between 1.5 and 6 m thick during winter. In some lakes this ice either melts completely or peripherally, forming moats in summer. The ice cover forms a thermal barrier that prevents lakes from freezing to the bottom. A minority of shallow lakes are not sufficiently protected from the wind and can freeze to the bottom such as Lake Ferris in the Larsemann Hills, or freeze to within a few centimetres of the bottom where there is a thin layer of hypersaline brine such as Lake Vida in the McMurdo Dry Valleys (Doran et al. 2003). These are known as dry-based lakes. Other lakes are ice-free all year round on account of hypersalinity, examples being Deep Lake in the Vestfold Hills, Don Juan Pond in the McMurdo Dry Valleys (Takamatsu et al. 1998) and some lakes in the Rauer Islands, 30 km south of the Vestfold Hills (Hodgson et al. 2001b). Ice and snow cover results in low levels of annual photosynthetically available radiation in the lakes. However, when the ice melts in summer, the high transparency of the water column can transmit so much light that it has an inhibiting effect on photosynthesis (Goldman et al. 1963). However, this does not necessarily apply to the benthic algae and mosses, which can form thick felts on the bottom of the lakes.

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Chemically, the lakes range from freshwater to hypersaline and often exhibit different degrees of temporal or permanent stratification (Roberts and McMinn 1996). They are typically oligotrophic, with eutrophy usually restricted to those which are directly influenced by visiting marine mammals, birds or people (Butler 1999). The lakes are often classified as ‘coastal maritime’ or ‘coastal continental’ depending on the degree of chemical influence from the sea, versus interior climate regimes. Epishelf lakes such as Kakapon Bay in the Bunger Hills can exhibit full marine conditions together with marine flora and fauna including seal populations (Melles et al. 1994a). Others, like Moutonnée Lake on Alexander Island, have marine water restricted to the lower parts of their water column (Heywood 1977). Biologically most lakes, not in connection with the sea, contain truncated food chains characterized by the microbial loop (Roberts et al. 2000b). Benthic cyanobacteria and diatoms dominate the biomass, and aquatic mosses are amongst the highest forms of plant life. There are often few zooplankters and no fish. Small phytoplankton and bacterioplankton populations are present with large seasonal variations in their populations, often related to the light climate. Despite this, of all Antarctic ‘terrestrial’ habitats, the lakes contain the highest biomass and species diversity. A recent study of the species composition of benthic microbial mats from selected lakes in the Larsemann Hills, Vestfold Hills and McMurdo Dry Valleys found 1500 strains of bacteria (in a subset of 9 lakes), 60 strains of cyanobacteria (24 lakes), 230 strains of fungi (17 lakes), 91 strains of algae (3 lakes) and 50 protozoans (6 lakes) (A. Wilmotte, pers. comm.). These were identified using combinations of morphology, phenotypic, chemotaxonomic and molecular taxonomy. For bacteria, 320 clones were obtained from a single mat sample! Technological advances have enabled scientists to study species compositions ranging from aerobes to anaerobes, from primary producers to degraders and brought a new level of detail to the understanding of Antarctic food-web ecology. With such a diversity of physical, chemical and biological attributes, Antarctic lakes represent important ecosystems for the study of a wide range of science questions. Many of them accumulate sediment deposits that incorporate the products of glacial erosion together with biological and chemical fossils. These can document both the lake’s development and changes in the surrounding environment. Antarctic paleolimnology Antarctic paleolimnology has seen an increase in activity over the last two decades and a number of the national Antarctic operators have supported multi-year paleolimnology programmes. Early studies focused on lakes in the more accessible maritime Antarctic Islands and the main Antarctic Oases with the aim of building a history of deglaciation and environmental change (Appendix 1). Whilst many of these studies were exploratory surveys, together they have become valuable in piecing together environmental and climate changes through the Holocene. With the regional histories documented, the questions asked by paleolimnologists have become more sophisticated. For example, emergent lake basins have been used to construct a history of marine transgressions and hence the relationship between relative sea-level change, ice thickness, and isostatic rebound (Zwartz et al. 1998a; Verleyen et al. 2003a), and closed lake basins have been used to reconstruct precipitation

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and evaporation regimes (Roberts and McMinn 1999a), lake ontogeny and environmental conditions (Spaulding et al. 1997). Older lakes have been used to help interpret the information found in ice cores (Doran et al. 2002a) and newly formed lakes to answer biological questions concerning colonization and succession (Swadling et al. 2001). In geological terms, the lakes are transient features of the landscape formed by a variety of geomorphological and geological processes. The last glaciation peaked between 25,000 and 18,000 (calibrated) years ago (yr BP) with ice occupying many areas of coastal Antarctica until much later. This ice scoured many older lake basins removing the sedimentary record of earlier glacial cycles. This phenomenon is widespread in the Antarctic Peninsula region and limits the time window available to paleolimnologists. However, some lakes, such as those in the Larsemann Hills, do contain older records, dating from the last interglacial (Hodgson et al. 2001a). Others, including lakes in the McMurdo Dry Valleys, provide intermittent glimpses into past interglacials up to 300,000 U/Th yr BP, with ‘missing’ sediments removed by aeolian activity during dry periods (Hendy 2000). Further back, ancient lake sediments deposited during the Jurassic period, and now exposed in the Transantarctic Mountains, contain a fascinating record of a time when fish and conchostracans once inhabited Antarctic lakes (Tasch 1970). In the next section we will outline some of the methodological challenges met by paleolimnologists in the Antarctic and some of the imaginative ways in which standard paleolimnological methods have been developed and applied to a range of science questions. Synthesis of methodological aspects Logistics The Antarctic presents a number of unique methodological challenges, not least of which is access. For scientists the usual approach is through one of the national programmes each of which has bureaucratic mechanisms in place for the entry of their own and foreign scientists, usually on a peer-reviewed basis. Most of these programmes will have the logistic capability to access sites within a particular sector of the Antarctic. For paleolimnologists, those with helicopter support offer the greatest flexibility, while those with fixed-wing aircraft have a greater operational range and payload often (essential for deep coring equipment). Without air support the geographical range of scientists is limited to those coastlines accessible to ships and amenable to overland travel by vehicles from the main Antarctic bases (Figure 3). Health and safety are of paramount importance in Antarctic fieldwork. Most national programmes require visitors to undergo extensive training in remote area survival, safe ice travel and first aid. For limnologists, field operations are usually carried out in the spring and early summer when lake ice offers a safe platform from which to work. Indeed, once assessed, the ice is often thick enough to support landings by helicopters and fixed-wing aircraft (Figure 3). Inflatable boats and rafts have been used successfully in the Subantarctic (Rosqvist et al. 1999) and on saline lakes where the ice cover is less reliable. Lake ice can be drilled using a motorized ice drill (e.g., Jiffy Drill) usually with an auger blade greater than 20 cm in diameter (Figure 4a). Ice thickness may be in excess of 4 m so drill extension rods are often required. Ice hardness also varies considerably and the ice can contain gravel

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or windblown rock particles. For these reasons we recommend taking several spare auger blades. Some approaches use a vertical heating element or “hot finger” to widen a drilled hole to accommodate larger equipment. This is not encouraged for sediment sampling since sediment in the ice is released into the water column during melting.

Figure 3. Logistics and access, (a) British Antarctic Survey Twin Otter landing on Citadel Bastion Lake, Alexander Island carrying coring equipment and personnel, (b) a field party being deployed on Pourquoi-Pas Island by Lynx helicopters from HMS Endurance, (c) lightweight coring equipment being transported on a quad motorcycle from Watts Hut in the Vestfold Hills.

Sediment core collection Sediment coring equipment has ranged from highly portable systems that can be carried by hand or on man-hauled sledges, to heavy systems requiring air or motorized over-snow transport (Figure 3). Lightweight systems typically consist of gravity corers such as the Glew corer (Glew 1991) and piston corers such as the square rod operated Livingstone

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corer (Livingstone 1955) operated from lake ice or rafts (Figure 4). These have been used to collect reliable sequences from lakes at up to 34 m water depth (Hodgson et al. 2001a). Another frequently used portable system is the Russian chamber corer type with steel rods (Björck et al. 1993, 1996), often used from an inflatable rubber boat or ice which, in combination with a percussion head, can penetrate highly minerogenic sediments without disturbance. In deeper lakes (> 30 m) surface sediment corers that deploy automatically on retrieval have been used, for example the UWITEC short (1.2 m) gravity corer. Single-drive gravity corers have also been used in the Bunger Hills to collect cores of up to 1.5 m length (Verkulich et al. 2002) and can be operated in very deep water. Most of these systems require ‘sediment catching’ devices such as steel lamellae, rubber diaphragms or other stoppers to prevent sediment loss during recovery, especially when used in unconsolidated sediments or deep water. Surprisingly, there has not been much use of freeze coring on account of the difficulties of transporting liquid nitrogen or dry ice into the field. For deep sediments, several teams have used 2 m and 3 m UWITEC cable-operated percussion-type piston corers with considerable success (Melles et al. 1997; Hodgson et al. 2003) and others have used customized designs based on similar principles (Bird et al. 1991). These systems have proved capable of collecting long sediment cores up to 13 m long (Kulbe et al. 2001) at depths of up to 100 m in both freshwater and marine environments, and theoretically could be operated at much greater depths (Figure 4c). Their limit of penetration is determined by total sediment depth and friction on the wall of the core tube. Where multiple cores are taken, two adjacent holes are drilled through the ice (ca. 1 m apart) and successively deeper cores taken from alternate holes with an overlap of 1020 cm. For very consolidated sediments a geological-type drilling rig would be required such as that employed offshore at Cape Roberts (Barrett 2001). We recommend subsampling or freezing cores in the field to avoid further disruption by vehicle, ship or aircraft vibration during the often long transit of cores back to the laboratory. Some lake sediments are no longer under water but exposed in dry lake beds, such as lake terraces, shorelines and deltas formed when the water level was higher than today (Pickard et al. 1986b). Here the sampling equipment consists of hand-held excavators (spades!), or a Russian chamber corer (Björck et al. 1996). Extraordinary examples of past shorelines have been found in the McMurdo Dry Valleys, up to 480 m above present-day sea-level, revealing that vast lakes, covering up to 250 km2, inundated the valleys during the Last Glacial Maximum (LGM; Hall et al. 2001). Many Antarctic studies have benefited from employing a combination of physical mineralogical and inorganic geochemical methods, together with terrestrial, algal and siliceous fossils and organic geochemistry, to reduce the errors inherent in the interpretation of single markers. The following sections present examples of how some of these stratigraphic markers have been used in Antarctic paleolimnological studies. Foremost amongst these are the various methods used to establish a chronology.

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Figure 4. Sediment coring in the Antarctic, (a) UWITEC coring rig at Moutonnée Lake, Alexander Island used to collect a 5.5 m core at 51 m water depth. This core is being used to reconstruct the history of George VI Ice Shelf (visible in the background), which impounds this epishelf lake. Coring are Mike Bentley, Elie Verleyen and DH, (b) scientists James Smith and Elie Verleyen drilling through 4 m thick ice cover on Citadel Bastion Lake, Alexander Island, (c) DH and James Smith using a Livingstone corer on Horseshoe Island.

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Chronostratigraphic methods Dating Antarctic sediments is best achieved using a combination of dating technologies, allowing for cross calibration. In some cases dates can also be checked against independent physical stratigraphic markers such as volcanic tephra layers (Björck et al. 1991c; Hodgson et al. 1998a). Ages of younger Antarctic sediment samples have been determined using radioisotopes, including 137Cs (post-1950 AD) and 210Pb (post 1850 AD). Despite low concentrations of 210 Pb in atmospheric fallout around continental Antarctica, it has been used successfully to date sediments in the sub-Antarctic on Signy Island (Appleby et al. 1995), the maritime Antarctic (Ye and Cuihua 1997), and the east Antarctic (Bird et al. 1991; McMinn et al. 1994). Some of the strongest signals have been detected where there is a focussing effect from large catchment areas, for example White Smoke Lake, Bunger Hills (Doran et al. 2000). Several studies have also recorded well-defined 137Cs peaks that presumably record the 1964/65 fallout maximum from the atmospheric testing of atomic weapons (Appleby et al. 1995; Hodgson et al. 2001a), and a 137Cs and 238Pu peak in surface sediments from Livingston Island was believed to originate from the re-entry burn up of the SNAP-9A satellite generator in 1964 (Björck et al. 1993). These radiometric techniques are therefore viable dating tools in paleolimnological and geomorphological studies, although in some regions, such as the McMurdo Dry Valleys, the radioisotopes in lakes remain at levels below detection (Doran et al. 1999). Radiocarbon dating has been the principal method used to date Antarctic lake sediments. However, there are a number of problems associated with dating lacustrine deposits in the Antarctic that can result in inaccurate chronologies (Björck et al. 1991c; Melles et al. 1994b; Gore, 1997; Ingólfsson et al. 1998; Doran et al. 1999, 2001; Hodgson et al. 2001a). These include 14C-depleted carbon from soils or weathered carbon-bearing rocks, and reduced CO2 gas exchange between the lake water and the atmosphere under ice cover. This latter effect, sometimes referred to as the residence time effect, can be enhanced in lakes where there is long-term water column stratification. Where marine sources contaminate freshwater sediments (marine transgressions, visits by marine mammals or birds), dates can be influenced by a carbon reservoir effect, which yields older radiocarbon dates. This is due to upwelling of old bottom water resulting in 14C-depletion in Antarctic Ocean water masses (Omoto 1983). CO2 formerly trapped as bubbles in the ice sheet may also result in 14Cdepleted water entering lakes during times of glacial melting. In some lakes no carbon reservoir effect is found, for example where summer melting of lake ice has enabled organisms to fix dissolved inorganic carbon (DIC), which is well-mixed with atmospheric CO2 (Hodgson et al. 2001a). Where a carbon reservoir effect is present it can often be determined simply by dating surface sediments (Melles et al. 1994b; Roberts and McMinn 1999a) or comparing dates against 137Cs and 210Pb (Doran et al. 2000) and Uranium/Thorium chronologies (Hall and Henderson 2001). It is also possible to overcome the carbon reservoir effect by using high precision Accelerator Mass Spectrometry (AMS) dating of discrete organic remains, for example aquatic moss or cyanobacterial mats (Jones et al. 2000) and even freshwater crustacean eggs (Björck et al. 1996). Doran et al. (1999) dated eight microbial mat samples in a 37 cm core from Lake Hoare in this way, and showed that the surface age of the core was ca. 2600 14C yr BP, and the remaining downcore ages were correlated with sediment depth (r2 = 0.97). Therefore, using the surface sediment age to correct for downcore age

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seems reasonably valid in some systems. However, the carbon reservoir effect is not always constant. It can vary from lake to lake in the same region and be depth-stratified. For example, shallow lacustrine microbial mats in the Dry Valleys have been dated as modern, but deeper parts of the lake the sediments have been contaminated by very old lake water DIC (> 10,000 14C yr). The carbon reservoir effect can also change over time. For example in White Smoke Lake, Bunger Hills, Doran et al. (2000) used 137Cs and 210Pb chronology in conjunction with AMS 14C dates to establish that the modern surface correction for sediments is actually less at present (ca. 725 14C yr) than in the past (ca. 1950 14C yr) due to excess bomb-derived 14C entering the lake. Therefore, simply correcting to the sediment surface does carry some potential pitfalls. Zale (1994) has shown how 14C ages of lake sediments, heavily “polluted” by marine-influenced particulate carbon from penguins, may be corrected by geochemistry. The alkali soluble humic acid and alkali-insoluble humin fractions may also be isolated and separately dated in order to isolate sediment components derived from organic material from those containing old carbon such as graphite and coal (Björck and Wohlfarth 2001). The uranium series (U/Th) has been used to date older sediments (> 25,000 yr). To evaluate this method, a number of U/Th dates of paleolake deposits in the McMurdo Dry Valleys have been compared with 14C dates of the same carbonate sample (Hendy et al. 1979; Hendy 2000; Higgins et al. 2000). Results showed that the methods matched to within 3000 (Hendy et al. 1979) and 8600 yr (Hendy 2000). In other cases, mismatches of 5950 yr and ca. 18,000 yr have been found between 14C dates and U/Th ages. These have been attributed to a carbon reservoir resulting from lack of aeration due to perennial ice cover and/or strong density stratification, and the input of old CO2 from glacial meltwater, respectively (Hall and Henderson 2001). The U/Th technique, therefore, has the potential to determine not only the reservoir effect in recent lakes, but to give some degree of chronological control to samples which are beyond the limits of the radiocarbon dating method. Luminescence dating methods have also been attempted where quartz or feldspar grains are present. Thermoluminescence (TL) dating and optically-stimulated-luminescence (OSL) are both dosimetry dating techniques that differ in the way the accumulated dose in detrital minerals is measured. Tests of TL on stream, ice, and lake bottom sediments have shown that it is a potentially reliable technique in this environment (Doran et al. 1999). However, despite a relict signal on the bottom of Lake Hoare of between 1000-2000 yr BP, depth-age trends in samples from short box cores were not evident. Berger and Doran (2001) refined and expanded the TL work by employing OSL. In this case, OSL did reveal depth-age trends in short surface sediment cores, with relict signals as low as 500 yr BP. OSL is inherently more sensitive to the light exposure history of the sediment, as it lacks a lightinsensitive signal. Unlike 14C, this signal is unaffected by the age of the lake water, and can be used on time scales within the range of both 14C and U/Th dating, making comparisons possible. Paleomagnetism has been attempted in Lake Hoare sediments, and although there was a strong magnetic remanence, the grains were too coarse for torquing by the magnetic field. Finer-grained sediments may be datable with paleomagnetism (Doran et al. 1999), particularly in lakes with high (> 30 cm/ka) rates of deposition (King and Peck 2001).

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Physical, mineralogical and inorganic geochemistry methods Physical lithostratigraphy. Most studies have employed some form of descriptive lithology to aid in core correlation, comparison between records, to check for slumps and slides, and perhaps most importantly to evaluate the sedimentary environment. Photography and imagery of the core, quantitative analyses of dry mass, water content, sediment density, organic content, carbonates and magnetic susceptibility are used routinely in many studies. Most sedimentary structures can be seen by eye but, occasionally, X-ray analyses and chemistry can reveal additional features such as fine-scale tephra horizons (Björck et al. 1993). Measurements of particle size and abundance can be related to glacial and glaciofluvial activity within catchments (Noon et al. 2001). In general, the quantity, sorting and grain size of the sediment depends on the proximity of terrigenous sediment sources (such as glaciers or snowfields) to the lakes, and on the energy conditions during transport and sedimentation (Verkulich et al. 2002). Stable isotopes. In sediments, the study of stable isotopes of oxygen and carbon can provide useful stratigraphic data that is often not accompanied by any macroscopic or sedimentological change (Ito 2001). Analysis of oxygen isotopes in lacustrine carbonates (į18Ocarb) provides information on rates of meltwater input and evaporation and is often most informative in closed lake systems. Oxygen isotopes in biogenic silica (į18OSi) reflect climate change through changes in water temperature and/or hydrology (Rosqvist et al. 1999). Carbon isotopes of organic matter (į13Corg) can be a sensitive indicator of biogenic primary production, carbonate precipitation, respiration of organic matter and CO2 exchange between the water and the atmosphere (Menking et al. 1997). In the Vestfold Hills, į13Corg has been used as an indicator of CO2 fixation and high levels of microbial activity in the lakes (Bird et al. 1991). Here the major factor controlling į13Corg values of Antarctic lake cyanobacterial mat communities was CO2 availability, which is limited by the year round ice cover in some lakes. In an epishelf lake in the Bunger Hills, į13Corg was employed to estimate the export of production from the surface waters to the anoxic bottom water and interpreted as a signal of increased production (Melles et al. 1997; Kulbe et al. 2001). In South Georgia, į13Corg and į18OSi isotopes have been used as a paleothermometer. Results showed no evidence of a large thermal decline during the Younger Dryas and permitted speculation that small oceanic islands may be buffered from this type of climate shift, despite following broader scale climate changes indicated in ice cores (Rosqvist et al. 1999). In general, the factors controlling isotope fractionation in modern Antarctic lake catchments and in biological communities, particularly the effect of secondary factors such as light availability, temperature and salinity, are poorly understood and calibration work is needed on a site by site basis to best utilise isotopic data. For example, Doran et al. (1998) show that the extent of carbon isotope fractionation in microbial mats is strongly controlled by depth in Lake Hoare. In the shallow regions of the lake, a diffusion limitation is set up by the high productivity in the moats, and CO2 undersaturation beneath the ice cover. At greater depth in the lake, productivity is reduced and the water column is CO2-rich providing maximum carbon isotope fractionation via the carbon fixation pathway that uses carboxylating enzymes (RUBISCO type I or II, beta-carboxylases). In a number of McMurdo Dry Valley lakes, the water columns are so stable that their vertical geochemical and isotopic properties can be used to infer lake history over several

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millennia (Lyons et al. 1998). In these lakes the isotopic and chemical structure of their present-day water columns are the legacy of past environmental changes. For example, using a Cl- diffusion model, it is possible to calculate that Lake Vanda evaporated to near dryness ca.1200 yr ago, based on the size of a diffusion cell in its salinity profile (Wilson 1964). A similar salt diffusion model was used to examine steps in the salinity profile (paleoepilimnia) of meromictic Ace Lake in the Vestfold Hills and revealed that these structures are the legacy of periods of lower water-levels (Gibson and Burton 1996). The survival of this ‘evaporitic’ signal is largely preserved by the long period of ice cover that preserves the meromictic status of the lakes and allows for the diffusional rather than advective mixing (stimulated by wind currents). Total lake ages can also be calculated from į18O, įD (hydrogen), and chlorinity profiles in the water column. For example, Hendy et al. (1977) proposed that Lake Bonney had been flooded by sea water between 100,000 to 300,000 U/Th yr BP before drying up. Helium isotopes have provided further details on the history of Lake Bonney, suggesting that the west lobe may contain a record of at least 106 yr. (Poreda et al. in press). In a recent study of Lake Vanda and Don Juan Pond, ratios of 87 Sr/86Sr isotopes have been used to explain the development of their unusual chemistries and differentiate between sources of salts (precipitation, sea water and rock weathering) in the lakes (Friedmann et al. 1995). In general, mineralogy and geochemistry have remained under-utilized. However, lake bottom sediments from the McMurdo Dry Valleys continue to be studied to determine the influence of water chemistry on mineralogy and geochemistry, and to evaluate methods for remote spectral identification of biomarker minerals for future applications on Mars (Bishop et al. 2001). Mineralogical analyses have included observations of sediment colour; the Fe(II)/Fe(III) ratio; the presence of pyrite; the abundance of Fe, S, and some trace elements; C, N, and S isotope fractionation patterns and reflectance spectroscopy (0.3-25 µm); Mössbauer spectroscopy (77 and 4 K) and X-ray diffraction. Mineralogy and inorganic chemical studies to detect source areas or determine the petrology of ice-covered rocks in catchments have not yet been conducted systematically on lake sediments, although these studies have been successful in the marine environment (Ehrmann et al. 1992; Diekmann et al. 2000). Terrestrial, algal, and siliceous indicators and organic geochemistry Micropaleolimnology. The discovery of intact, finely-laminated ‘stromatolytic’ cyanobacterial mats in many Antarctic freshwater lakes has encouraged researchers to study biological processes in the top few centimetres of ‘sediment’. Here, micro-electrodes (e.g., oxygen) and pigment analyses have been used to measure the physical and chemical structure and photosynthetic characteristics of the microbial mat surface. Results from a series of ponds near McMurdo Sound showed that the mats have a carotenoid-rich surface layer that overlies a deep chlorophyll maximum reflecting strong self-shading within the mat systems and physiological dominance by the motile populations of trichomes. These are able to migrate to the surface within 2 hours in response to an increase in surface irradiance (Vincent et al. 1993). Other studies have cultured artificial mats in benthic gradient chambers using Antarctic innocula and then manipulated the temperature, oxygen and sulphide gradients in order to study species responses (Pringault et al. 2001). For example, microbial mat samples cultured from Lake Fryxell have revealed some important chemical

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processes affected by temperature changes, including the migration of microorganisms, locally resulting in higher biomass, and an increase in oxygen requirements for the oxidation of reduced compounds such as H2S. Results show that the metabolic rates of the different functional groups in these artificial microbial mats are fine-tuned to the prevailing temperature. This resulted in a close coupling of oxygen and sulphur cycles as has been previously shown for mesophilic hypersaline mats (Pringault et al. 2001). Laboratory cultivation of the same innocula from Lake Fryxell has also shown how the microstructure of the mats is formed revealing a vertical structure comprising an upper green layer dominated by cyanobacteria and an underlying pink layer including purple non-sulphur phototrophic bacteria. Pigment content indicated that both main groups of cyanobacteria, Oscillatoriaceae and Nostoc sp., adopted different vertical distributions (Buffan-Dubau et al. 2001). Pollen, charcoal and plant macrofossils. No pollen-producing plants live in Antarctica, and only a few grow on the islands of the maritime Antarctic. On Signy Island, where there are two flowering plants, Deschampsia antarctica (Gramineae) and Colobanthus quitensis (Caryophyllaceae), pollen profiles are dominated by Deschampsia (Jones et al. 2000). However, in these and other records there are long-distance components, which can be used to detect the differing dominance of main air masses originating from temperate latitudes in South America, Australia and New Zealand (Björck et al. 1991a, 1996). For example, researchers have found Nothofagus fusca, Compositae, Chenopodiaceae and Podocarpus pollen derived from South America and Mimosaceae and Eucalyptus (Myrtaceae) from Australia. Typically, pollen is found in extremely low concentrations, e.g., 130 to 200 grains per cm3 (Jones et al. 2000). Peaks in concentrations of the chlorophyte Pediastrum, preserved in pollen slide preparations, have also been used to indicate good light conditions (less suspended material in the water), good nutrient supply, ice-free conditions and relatively mild water temperatures during the growing season (Jones et al. 2000). The presence of aquatic mosses and moss spores can indicate similarly favourable conditions including substrate stability, low influxes of terrigenous material and little suspended material in the water column (Verkulich et al. 2002). Plant macrofossils have been found in a number of studies (e.g., Akiyama et al. 1990; Jones et al. 2000). Charcoal has rarely been used, but an increase in micro-charcoal in the top of two cores from Signy Island could be the result of long range airborne transport of charcoal derived from human activities in the 20th century, with a component possibly derived from the historic offshore whaling stations and the scientific research bases (Jones et al. 2000). Diatoms. Diatoms have been one of the key indicators in Antarctic paleolimnology on account of their sensitivity to changes in water chemistry and the physical environment such as salinity, nutrients and depth (Spaulding and McKnight 1999). Baseline data on the ecology and distribution of diatoms has been collected from the subantarctic islands (Van de Vijver and Beyens 1999, 2002), the maritime Antarctic (Jones and Juggins 1995), the Antarctic Peninsula (Wasell and Håkansson 1992), and on the continent in the McMurdo Dry Valleys (Spaulding et al. 1997), Windmill Islands (Roberts et al. 2001a), Vestfold Hills (Roberts and McMinn 1999c), Rauer Islands (Hodgson et al. 2001b), Larsemann Hills

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(Sabbe et al. 2003) and the Schirmacher Oasis (Pankow et al. 1987). The most up-to-date list of Antarctic diatoms is given in Kellog and Kellog (2002). However, the diatom floras of many Antarctic regions are still poorly known and biogeographical and ecological information is limited. In two recent reviews on species diversity and biogeography of Antarctic and subantarctic diatoms, it was stated that cosmopolitanism is the norm (Jones 1996; Van de Vijver and Beyens 1999). Despite this, recent detailed morphological comparisons of diatoms from the Larsemann Hills, against type collections, has found that 42% of diatom taxa are morphological endemics to the region, hinting at a much greater degree of endemism than previously thought (Sabbe et al. 2003). This latter study emphasises the need for continuing morphological and molecular studies of Antarctic diatom taxonomy. Despite the paucity of ecological data, Antarctic diatoms have been used very successfully to reconstruct past environmental changes (Spaulding and McKnight 1999). This is particularly the case where there are regional datasets that statistically relate modern species assemblages to environmental variables and applications of these relationships to interpret fossil data. This technique involves the development of a transfer function (Battarbee et al. 2001). In the Vestfold Hills and Larsemann Hills, Roberts and McMinn (1996) and Verleyen et al. (2003b) have examined relationships between diatoms and salinity and developed transfer functions that have been applied to stratigraphic reconstructions. Other researchers have examined diatoms of surface sediments in lakes and fjords to construct transfer functions for depth (Whitehead and McMinn 1997). For maritime Antarctic lakes, Jones and Juggins (1995) have constructed a transfer function to reconstruct chlorophyll-a concentrations, an indirect measure of lake production. These types of transfer functions are likely to remain specific to the particular regions for which they were created, but may become more widely applicable as Antarctic taxonomy improves. For example, a combined dataset comprising data from Roberts et al. (1996, 2001a) and Verleyen et al. (2003b) is now being applied over a large sector of eastern Antarctica from the Bolingen Islands to the Windmill Islands (Figure 1). An alternative approach that has been employed in the Antarctic Peninsula region is to use multivariate statistics to analyse diatoms in combination with other environmental proxies, for a more holistic view of Holocene climate history (Björck et al. 1993, 1996). Sedimentary pigments. Sedimentary pigments are beginning to be used more widely in Antarctic paleolimnology to provide information on a whole range of biota that do not leave reliable morphological fossils. Long after the morphological remains of algae and bacteria are lost due to various degradation processes, sedimentary carotenoids (carotenes and xanthophylls), chlorophylls, their derivatives, and other lipid-soluble pigments can be used to track past populations (Hodgson et al. 1997; Leavitt et al. 2003). This is particularly the case in the Antarctic where cold temperatures and limited grazing activity favour their preservation (Figure 5). With the advent of new technological developments such as High Performance Liquid Chromatography coupled with mass spectrometry (HPLC-MS; Airs et al. 2001; Squier et al. 2002; Walker et al. 2002), considerable progress has been made, with fossil pigments being used in applications where the abundance, production, and composition of past phototrophic communities are important response variables (Leavitt and Hodgson 2001). Some pigments, such as beta-carotene, are found in all algae and some

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Figure 5 . A sediment core from the Larsemann Hills showing intact layers of pigmented fossil cyanobacterial mats that preserve a remarkable record of the growing conditions at the time they were formed.

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phototrophic bacteria, whilst other pigments are much more specific (for example, alloxanthin is specific to the Cryptophyta, and peridinin is specific to the Dinophyta). A detailed study of sedimentary pigments carried out on sediments from Pup Lagoon in the Larsemann Hills charts the colonisation by cyanobacteria, diatoms and other protists following the lagoon’s isostatic isolation from the sea (Verleyen et al. 2003a). In this paper the authors provide a listing of the commonly recovered pigments from the lake sediments and their taxonomic affinities. In another study of Kirisjes Pond in the Larsemann Hills, advanced HPLC/MS methods were used to track chlorophyll-a and b-derived components which are indicative of eukaryotic primary producers, bacteriochlorophyll-c and d-derived components indicative of prokaryotes, steryl chlorin esters which provide evidence of grazing, and shifts in the esterifying sterol composition which record changes in the primary producer community (Squier et al. 2002; Walker et al. 2002). Research has also been carried out measuring pigment compositional changes with depth in Lake Hoare (Hawes and Schwarz 1999). Pigments can also be applied to study natural variations in ultraviolet radiation (UVR) flux in polar regions through the Holocene (Leavitt et al. 2003). Surveys of shallow lakes and ponds in eastern Antarctica have revealed that modern sedimentary environments preserve fossil pigments and that the abundance of photo-protective pigments increases as a function of algal exposure to UVR (Leavitt et al. 2003). Mycosporine-like amino acids (MAA’s) and phycobiliproteins also perform light screening functions but are water soluble and therefore poorly preserved in sediments and thus cannot be used as paleo-optical tracers. To reconstruct past UVR environments, scytonemin and its derivatives are summed and expressed as a ratio with algal carotenoids, an index that has been shown to be linearly related to the depth of UVR penetration with similar compounds (Leavitt et al. 1997, 1999, 2002). Other forms of organic geochemistry. Fossil pigments are just one group of organic biochemical fossils found in Antarctic lake sediments. Other geochemicals do persist, despite diagenetic processes, and can yield fascinating insights into the organisms that lived in the lakes and their catchments (Matsumoto et al. 1982, 1988; Matsumoto 1989; Akiyama et al. 1990). For example, sediment cores collected from a lake on King George Island have been found to contain a range of organic geochemicals that originate from penguin droppings (Sun et al. 2000). Relative abundances of these geochemicals and 13C isotopes were used as a proxy for past penguin population sizes and extended the record from a few decades (census data) to 3000 years. Similarly, in a study of Lake Boeckella, Hope Bay, Zale (1994) showed how concentrations of P and Cu could provide an estimate of the size of the local Adélie penguin rookery through the last 5000 years. The study of geochemicals in sediments continues to keep pace with enabling technologies, a recent example being the application of FT Raman spectroscopy to the analysis of organic and inorganic signatures in Antarctic lake sediments (Edwards et al. 2003). DNA. Fossil DNA is also being used to trace past populations of protists and cyanobacteria in lakes where physical and chemical stresses play an important role in structuring the communities (Hawes and Schwarz 1999). The potential of microbial mats as quantitative

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indicators of environmental stresses has been under exploited because of methodological problems in characterizing their biodiversity. However, paleolimnologists are now employing sterile molecular techniques to assess the biodiversity and community structure of microbial mats using molecular clone libraries of cyanobacteria and protists (A. Wilmotte, A. Taton and K. Sabbe, pers. comm.). Direct gradient gel electrophoresis (DGGE) and temperature gradient gel electrophoresis (TGGE) have been successfully applied to obtain a genetic fingerprint of prokaryotic and eukaryotic organisms in microbial mats. Previous experience in sulfide-rich lake sediments (Coolen and Overmann 1998) has shown that ancient DNA can be preserved for a long time in sedimentary structures (10,000 yr), but that the length of the molecules decreases with time due to various chemical degradations. A preservation limit for fossil DNA of 10,000 to 100,000 years has been suggested (Austin et al. 1997). Using material originating from the Larsemann Hills, the Vestfold Hills and the Dry Valleys, more than 60 strains of cyanobacteria and 200 strains of autotrophic protists have been isolated and characterized, and techniques to assess the in situ biodiversity of these organisms using microscopical and molecular methods, employing rDNA, have been developed (K. Sabbe and A. Wilmotte, pers. comm.). When applied to sediment cores, this method should permit identification of the most important primary producers in lake systems and the tracking of species succession and climateinduced environmental changes. In effect, with the latest technology, these scientists are poised to watch species composition and succession in action. Zoological indicators . There have been relatively few studies of zoological remains in Antarctic lake sediments. Jones et al. (2000) identified egg cases of the Anostracan Branchinecta gaini (fairy shrimp), carapaces of the free-living Acari (mites) Alaskozetes antarcticus and Halozetes belgicae, and egg cases of the copepod Boeckella poppei in two lakes on Signy Island. Both Jones et al. (2000) and Björck et al. (1996) have found the highest concentrations of Anostracan eggs during a mid-Holocene hypsithermal (MHH) period indicating that well-developed productive benthic cyanobacterial mats were present at that time. Björck et al.’s (1996) record of Anostracan eggs on James Ross Island is interesting since this freshwater shrimp is not found there today. B. gaini feeds on epiphytes present in these mats, as well as the mats themselves. Jones et al. (2000) also found a marked increase in the number of Acari during the same period. Autecological data for these mites suggested greater development of crustose lichens, or organic debris in the lake catchment as a consequence of the warmer and wetter conditions. The colonisation and succession of rotifers, copepods (identified by body parts and eggs), foraminifera, and tintinnids have also been studied in early Holocene lake sediments in the Vestfold Hills (Swadling et al. 2001). Other zoological indicators have included seal hairs which have been used as historical indicators of seal populations at Signy Island (Hodgson and Johnston 1997; Hodgson et al. 1998b). Here, seals congregate in the catchments of coastal lakes and deposit their hairs during the annual moult. Anthropogenic pollutants. To date the record of anthropogenic pollutants in Antarctic lake sediments remains largely unstudied. This is surprising, as lake sediments can provide highresolution reconstructions of long distance pollutant imports over long periods of time and,

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being in the Antarctic (away from sources of pollution), enable the global signal to be determined. The most studied of the anthropogenic pollutants is 137Cs derived from the atmospheric testing of atomic weapons (Appleby et al. 1995), which is used as a chronological marker. Similarly, anthropogenic lead (derived from fossil fuel combustion) has been found in the Antarctic. For example, in a study of a 3000 year lake sediment on King George Island affected by penguin droppings, Sun and Xie (2001) showed that lead concentrations had significantly increased during the last 200 years and, especially in the last 50 years, compared to low and stable lead levels prior to the industrial revolution. This clearly indicates that global environmental pollution has influenced the Antarctic ecosystem and documents how one pollutant may find its way into the food web, bio-accumulate, and be passed along the food chain to marine birds and mammals. Having reviewed how a range of paleolimnological methods has been applied in the Antarctic, we now present four case studies to show how Antarctic paleolimnology is being employed to address some fundamental questions in Earth system science. Antarctic paleolimnology and Earth system science – case studies In the introduction we described how, of the continents, the Antarctic arguably plays one of the greatest roles in regulating the major forcings and feedbacks in the Earth’s system. In this section we aim to show how Antarctic paleolimnological studies have become integral in addressing some global science questions through four case studies. Studying Earth system science in the Antarctic context has the potential to unravel some of the key environmental challenges facing humans in the coming decades. Case study 1: The Antarctic contribution to global sea-level Global sea-level change is primarily controlled by the cyclic growth and decay of the world’s ice sheets. As the repository of more than 70% of the world’s freshwater, Antarctic ice is a major dynamic factor influencing global sea-level. Global sea-level observations suggest that total land-based ice volume was at its maximum 22,000 to 19,000 years ago (Yokoyama et al. 2000). However, at present, estimates of the contribution of the Antarctic ice to the global sea-level rise of 120 ± 20 m since the Last Glacial Maximum vary between 0.5 to 37 m (Bentley 1999; Ingólfsson and Hjort 1999). This uncertainty accounts for some of the large errors associated with reconstructing global sea-levels (Rohling et al. 1998). Research has therefore focused on determining the extent and thickness of the Antarctic ice sheets during the last glacial. One reason for this is to determine how much the ice sheets could influence future sea-levels if global temperatures increase by the 1.4°C to 5.8°C, between now and 2100, predicted by the Intergovernmental Panel on Climate Change (IPCC) (Houghton 2001). One promising method for constraining ice-sheet history over particular regions, which can be applied on the exposed coastlines of Antarctica, is the study of glacio-hydroisostasy, in which precise sea-level records are used to infer regional ice-sheet history (Lambeck and Chappell 2001). During glacial cycles there are both sea-level and isostatic changes, the former resulting from changes in the mass balance of the global ice and the

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latter from the mass of the ice sheets on the continents themselves. As global (eustatic) sealevel changes are reasonably well understood from far field sea-level observations (Peltier 2002), the ice mass on the Antarctic continent can be estimated using geological evidence of changes in the relative height above sea-level of the Antarctic coastline. These data can be derived from the sediments of low altitude coastal lakes, which can record the exact timing of the emergence of the lake basins from the sea. By dating the transition from marine to lacustrine conditions in sediment cores from lakes of different present-day altitudes, it is possible to calculate the relative sea-level history at that point (Bentley 1999). If this is corrected for eustatic sea-level variations then the crustal rebound through time can be quantified, providing a measure of the weight of the former ice load. Studies of emergent lake basins around the coast of Antarctica are helping to build a regional picture. Indicators of marine-lacustrine transitions in lake sediments and in marginal terraces have included diatoms (Pickard et al. 1986a; Wasell and Håkansson 1992; Roberts and McMinn 1999a), fossil molluscs, polychaetes, į18Ocarb (Bird et al. 1991; Zwartz et al. 1998a), pigments (Squier et al. 2002; Verleyen et al. 2003a) and descriptive sedimentology (Zwartz et al. 1998b). In the Vestfold Hills, studies of former sea-levels have been determined using 14C-dated organic material in lacustrine-marine and marine-lacustrine transitions in lake sediments (Pickard et al. 1986b; Bird et al. 1991; Bronge 1992; Fulford-Smith and Sikes 1996; Roberts and McMinn 1998, 1999a) and in marine terraces at different altitudes (Adamson and Pickard 1983; Zhang and Peterson 1984; Pickard 1985). Collectively, these have revealed that eustatic sea-level rise was more rapid than Holocene isostatic rebound and resulted in marine incursions, for example in Watts Lake between 8000 to 4700 14C yr BP (Pickard et al. 1986b). The first study to specifically relate these changes in relative sealevel to ice sheet thickness was carried out by Zwartz et al. (1998a). To achieve this, they developed numerical models of the ice and its hydrostatic load. The models did not account for the influence of ice volume change elsewhere in the Antarctic or the influence of changes in the northern hemisphere ice sheets, and thus could only be used to estimate regional changes in the thickness of the East Antarctic Ice Sheet. Similar datasets are now being generated for the Larsemann Hills (Figure 6) where areas of the Broknes Peninsula have remained ice-free for at least the last 45,000 years and possibly back to the last interglacial (Hodgson et al. 2001a). These have the potential to extend the relative sea-level reconstructions for this region further into the past. In the Bunger Hills, data from lake and nearshore marine sediments have shown that parts of the oasis were unglaciated prior to the Late Weichselian and were then inundated by grounded ice masses during the glaciation (Melles et al. 1997). Radiocarbon dates indicate that these ice masses retreated and lacustrine sedimentation commenced at the Pleistocene/Holocene boundary. It has been proposed that this deglaciation was triggered by a rapid eustatic sea-level rise of less than 11 m, the altitude of Figurnoye Lake (Verkulich et al. 2002). OSL dating of glaciofluvial and glacial lake shoreline sediments does not agree with the radiocarbon-dated expansion of the ice sheet at the Last Glacial Maximum and suggests that the deglaciation of the Bunger Hills may have commenced as early as 30,000 yr BP (Gore et al. 2001). Early deglaciation dates have been proposed for other oases in the eastern Antarctic and there is zoological, geomorphological and lake sediment evidence suggesting some regions remained ice-free during the last glacial (Hiller et al. 1988; Adamson et al. 1997; Gore et al. 2001; Hodgson et al. 2001a). This debate is unresolved,

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Figure 6. Relative sea-level curve (dotted line) derived from data from the Vestfold Hills (Zwartz et al. 1998a) and Larsemann Hills (Verleyen et al. 2003a). The sill height is the height above sea-level of the lake sill or outflow.

but the outcome of it is extremely important in guiding how scientists interpret the glacial history of eastern Antarctica, how thick the ice sheet was and how the ice masses there have contributed to global sea-level change. On the Antarctic Peninsula, reconstructions of relative sea-level are less well developed. In the Marguerite Bay area, the diatom stratigraphy of a coastal lake at 3.5 m altitude on Horseshoe Island indicates that isolation took place some time after 1860 corr.14C yr BP, resulting from isostatic uplift following retreat of the ‘Marguerite Bay’ glacier from ca. 10,000 yr BP (Wasell and Håkansson 1992). Radiocarbon-dated penguin remains in archaeological excavations of nests are also providing minimum ages when nesting sites were isostatically raised sufficiently above the storm surge limit to permit the establishment of penguin colonies (M. Bentley, pers. comm.). Ages of raised beach deposits from adjacent areas will give maximum ages for sea-levels in the region. Further north, on King George Island (South Shetland Islands), there is a series of raised beaches and more than 50 lakes. Lakes above 47 m, such as Jurasee, have remained fresh throughout their history and provide an upper limit for relative sea-level. An isolation date has been determined for Kiteschsee with sea-level falling below its 16 m a.s.l. sill by 6200 corr.14C yr BP (Mäusbacher et al. 1989). Long Lake, a former marine inlet on the Fildes Peninsula, with a hydrological connection to the sea down to 12-10 m a.s.l. was isolated at 2460 14C yr BP (Martinez-Macchiavello et al. 1996). Other lakes, occupying terrestrial depressions below 16 m a.s.l. record a different story. Here, the sequence is from fresh to brackish to marine to

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brackish to freshwaters (Shen et al. 1998). This marine transgression was not well dated in Yanou Lake (Yang and Harwood 1997) but is believed to correspond with a climate optimum (Mäusbacher et al. 1989). Raised beach evidence includes an undated beach at 45 m and a marine sediment (deposited in Long Lake) dated at 9670 14C yr BP (MartinezMacchiavello et al. 1996). Lower down, at 15.3-16.7 m a.s.l., penguin bones in raised beaches have been dated at 4540-4450 corr.14C yr BP (del Valle et al. 2002). Although the lakes and raised beaches studied on King George Island include a range of elevations and distances from the present ice margin, to date no detailed modelling has been carried out to construct a sea-level curve for the maritime Antarctic region. In addition, the chronologies are limited by insufficient 14C dates. However, studies completed to date (Mäusbacher et al. 1989; Martinez-Macchiavello et al. 1996; Yang and Harwood 1997) demonstrate that it is possible. In general, deglaciation of the Antarctic Peninsula following the Last Glacial Maximum continued over a longer time than in the northern hemisphere. Initially the marine (wet) based glaciers retreated at the same time as their northern hemisphere counterparts, forced by eustatic rises in global sea-level. However, retreat of the dry-based glaciers is out of phase between the northern and southern hemispheres, a feature that has been attributed to increased precipitation counterbalancing ablation, and to delayed warming. This interpretation has been based on evidence from the onset of sedimentation in lakes on the islands in the northern Peninsula region (Hjort et al. 1998), which became deglaciated between 9000 to 5000 yr BP. Minimum dates for the deglaciation range between 8680 14C yr BP for the Hope Bay area and 3900 14C yr BP for Hidden Lake on James Ross Island (Zale and Karlen 1989). The above paleolimnological records of deglaciation, isostatic uplift and isolation, together with raised beach evidence, ages and altitude of penguin bones and guano deposits, exposure ages of aeolian and fluvial deposits etc. are helping to piece together the regional history of relative sea-level change through the Holocene. With more data, geophysical modellers will be able to use accurately dated relative sea-level reconstructions to calculate the mass balance of ice sheets based on specific regional data. This can be supported by data in marine sediment cores from the continental shelf, which contain evidence of the extent of local ice sheets (e.g., Windmill Islands (Kirkup et al. 2002; Cremer et al. 2003b; Hodgson et al. 2003), Lallemand Fjord (Domack and McClennen 1996), and Prydz Bay (Domack et al. 1998)), grounding lines, and data on paleo-glaciers and ice streams derived from swath bathymetry of the ocean floor. Such studies will not only reveal the spatial variability in crustal subsidence and rebound rates, but will provide a more accurate determination of Antarctica’s contribution to past global sea-level change. Coupled with studies of relative sea-level in other regions on Earth, these data will determine each region’s vulnerability to the future sea-level rises predicted by the IPCC (Houghton 2001). Case Study 2: The risk of rapid sea-level change In the last section we discussed how the Antarctic contribution to global sea-level is being determined over the last glacial cycle. In recent years the Antarctic Peninsula has been one of the fastest warming regions on Earth (Vaughan et al. 2003) and scientists have witnessed the decay of some of the peninsula ice shelves (Vaughan and Doake 1996). Break-up of the

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ice shelves has fuelled concerns that their absence could trigger changes in the dynamics of the ice streams of the West Antarctic Ice Sheet (WAIS) and increase their rate of discharge into the sea (De Angelis and Skvarca 2003). The WAIS contains 3.8 million km3 of ice that, theoretically, could cause a sea-level rise of up to 4-6 m continuing beyond the 21st century (Oppenheimer 1998). Although the risk of the entire WAIS collapsing has recently been calculated at about 5% probability (Vaughan and Spouge 2002), even changes of 0.5 m (the approximate volume contained in individual WAIS catchments) could have significant economic consequences for coastal communities. Glaciologists are attempting to determine if collapse of the WAIS could result from readjustments continuing since the last glacial maximum, from internal flow instabilities, or not occur at all in the present interglacial. However, the major effort has been to determine how sensitive the ice sheet is to warm periods and if, and by how much, a regional warming trend could influence the WAIS and global sea-levels. To answer this question, one approach is to look at analogous Antarctic temperature optima within the Holocene and to determine what caused them, what effect they had around the Antarctic, and if they were accompanied by any changes in global sea-level. This would provide useful data for understanding forcing mechanisms and predicting the response of the ice sheet to abrupt and non-linear climate change (Goodwin et al. 1999). Paleolimnologists have therefore been examining Holocene sediment records for evidence of past hypsithermal events. By 1998 one research consensus was that the Antarctic experienced a climate warmer than today approximately 4700 to 2000 14C yr BP (Ingólfsson et al. 1998). Currently, the best-dated records put it at between 4000 to 2700 14C yr BP and 3300 to 1200 14C yr BP in the Antarctic Peninsula region (Björck et al. 1993; Jones et al. 2000), ca. 4000 to 3000 corr.14C yr BP in coastal Victoria Land (Baroni and Orombelli 1994) and 3500 to 2500 cal.14C yr BP in the Bunger Hills, eastern Antarctica (Kulbe et al. 2001) (Figure 7). This mid-Holocene hypsithermal (MHH) appears to have been most marked in the Antarctic Peninsula and maritime Antarctic regions. It is detected as a period of rapid sedimentation, high organic productivity and increased species diversity in lake sediments in the South Shetland Islands (Schmidt et al. 1990; Björck et al. 1996) and the South Orkney Islands (Jones et al. 2000). In continental Antarctica, it has been detected as a period of increased biogenic production in lakes, and in the Bunger Hills as a period of increased salinity in (Jaw Lake) between ca. 4000 to ca. 2000 14C yr BP (Roberts et al. 2000a). In the McMurdo Dry Valleys wetter conditions resulted in high water-levels at Lake Vanda (Lyons et al. 1997). These limnological records of the MHH are also matched in other proxies such as peat bank records (Mäusbacher et al. 1989) and marine sediments (Domack et al. 1991). Inland, modelled temperatures from ice cores show that a noticeably warmer climate episode occurred around 4000 yr BP followed by cooling between 2000 to 1000 yr BP (Ciais et al. 1994). At Law Dome, higher than mean Holocene ice accumulation rates have been attributed to a warmer climate between 4000 to 2500 yr BP (Goodwin 1998) and at Dome C a warm phase was recorded 4000 to 3000 yr BP (Lorius et al. 1979). Discrepancies between the exact timing of these events in ice core records and 14C-dated marine and freshwater sediments, may be due to poor understanding of the Antarctic carbon reservoir effect (and inconsistencies in its application) and insufficient numbers of dates, rather than representing real time lags (Kulbe et al. 2001). Although the current period of Antarctic Peninsula warming (King 1994) is occurring under an anthropogenically modified atmosphere (increasing concentrations of CO2) for

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Figure 7. Multi-archive compilation of mid-Holocene hypsithermal (MHH) records in selected ice, lake, marine and glacial records. References: 1Ciais et al. (1994), 2Lorius et al. (1990), 3Goodwin (1993), 4Ingólfsson et al. (1998), 5Jones et al. (2000), 6Björck et al. (1993), 7Björck et al. (1996), 8 Roberts et al. (2000a), 9Smith and Friedmann (1993), 10Melles et al. (1997), 11Domack and McClennen (1996), 12Kulbe et al. (2001), 13Hodgson et al. (2003), Cremer et al. (2003), 14Rathburn (1997), 15Baroni and Orombelli (1994), 16Ikehara et al.(1997), 17Pudsey and Evans (2001), 18Higgins et al. (2000). Note: 14C dates are as cited in source as raw (14C), corrected for the carbon reservoir effect (corr.) or calibrated (cal.). The ice core data, presented in calendar years BP cannot be directly compared with the 14C dates.

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which there is no recent paleoclimatological precedent, the MHH is our most recent ‘natural’ analog with which to test a number of theories regarding the wider impacts of Antarctic Peninsula warming events. For example, if and how an Antarctic warming event interacts with continental and world weather and climate, and if there is any evidence for sympathetic events in the marine environment, such as changes in ocean circulation, sea-ice extent and global sea-level. Looking first at the marine environment, cores from Lallemand Fjord (western Antarctic Peninsula) indicate reduced sea ice coverage, and greater primary productivity, before a glacial advance prior to approximately 2000 yr BP (Domack and McClennen 1996). Peaks in penguin populations indicate periods of open water in the Ross Sea between 4000 and 3000 corr.14C yr BP (Baroni and Orombelli 1994). Nearshore marine cores from the Windmill Islands show peaks in diatom concentration during the mid- to late Holocene between 3990 and ca. 2000 cal.14C yr BP (Cremer et al. 2003a; Hodgson et al. 2003) and core reconstructions from Prydz Bay indicate a period of increased production between ca. 3400 and 2700 14C yr BP (Rathburn et al. 1997). This was associated with longer periods of open water detected by Labracherie et al. (1989) which, in turn, brought warmer and more humid air masses to the continent. The consequent increase in coastal precipitation caused local expansion of ice sheets, particularly in zones where it was over-steepened. In the Vestfold Hills this event, known as the Chelnok Glaciation, dated between ca. 3000 and 1500 cal.14C yr BP (Adamson and Pickard 1986). It resulted in the temporary isolation of Abel Bay between ca. 3200 and 1750 14C yr BP (McMinn 2000). Evidence from eastern Antarctic marine sediment cores also indicates that an expansion of the ice sheet occurred there during the MHH (Domack et al. 1991). Further offshore, the influence of the MHH may have extended into the Southern Ocean, where alkenone and į18O analysis of foraminifera indicate a maximum Holocene ocean warming at approximately 4000 to 3000 corr. 14C yr BP (Labracherie et al. 1989; Ikehara et al. 1997). The fate of the ice shelves during the MHH is also being unravelled. There is evidence that the northern Antarctic Peninsula ice shelves broke up during the MHH (Pudsey and Evans 2001). However, whether this resulted in a destabilization of the WAIS and changes in sea-level is still being researched. The geomorphological evidence for sea-level change at this time (in the form of raised beaches and isolation basins) is complicated by changes in relative sea-level associated with isostatic rebound and more sites need to be studied. For example, a marine transgression in a low-lying coastal lake precisely at the time of the MHH would give circumstantial evidence that it did have an effect on relative sea-level in Antarctica (tentative examples of this can be found in the work of Yang and Harwood (1997) on King George Island and Verleyen et al. (2003a) in the Larsemann Hills). Similarly, shorelines could be surveyed for indications of higher beach energies at the time of the MHH as there is some evidence in the Taylor Dome ice core of increased storminess with atmospheric circulation being vigorous enough to increase the transport of sea salt (Clconcentrations in the ice) to Antarctica (Mayewski et al. 1996; Lyons et al. 1997). Comparison of these marine transgressions with detailed reconstructions of global eustatic sea-level (e.g., Bard et al. 1996) should constrain the influence of the Antarctic MHH on global sea-levels. Evidence of ice extent in other areas of Antarctica (as deduced from lake cores) may also contribute to the debate about whether small increases in temperature result in a positive mass balance in Antarctica, due to increased snowfall (as predicted by many GCMs), or a negative mass balance due to basal melting of the ice shelves and grounded

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portions of the ice sheet and glaciers (Gornitz 1995). With the MHH and its effect established, it is possible to speculate on the causes. Whilst the trigger has not yet been determined, it is likely that one or more forcings or feedback loops were in operation at the time. Hjort et al. (1998) have proposed that it approximately coincides with the peaking of the Milankovitch summer insolation in the south which, especially in the Antarctic Peninsula region, was out-of-phase with insolation in the northern hemisphere (Budd and Smith 1987; Budd and Rayner 1990). In eastern Antarctica, ice core evidence from the interior shows a late Holocene climate optimum between 6000 and 3000 yr BP. This earlier signal for a warm period cannot be explained by lowfrequency changes in insolation or Holocene greenhouse gases, and may result from internal oscillations of the climate system possibly amplifying short-term forcings such as solar activity or volcanic activity (Masson et al. 2000). Thus, an example of a possible feedback loop causing a warming period to become established at this time would be that a warm perturbation (possibly in the interior) leads to less sea ice, an increased local hydrological cycle, less stratification of the ocean, and more heat advection, which enhances the initial warming. Following this logic, Björck et al. (1993, 1996) attribute the MHH in the Peninsula to the presence of warmer and more humid air masses resulting from a decrease in sea ice around Antarctica. Conversely, if there is a weak influence from these westerly storm tracks and an increased influence from the dry and cold high pressure cell over the ice sheet in the south, it causes a cool period, and starves the glaciers of moisture and hence triggers a deglaciation. Models have shown that a reduction of sea ice around Antarctica results in a decreased intensity of the westerly wind tracks (Heusser 1989). Alternative explanations involve oscillations in the strength of the high pressure atmospheric cell over the Antarctic ice sheet or time-dependent variations in the upper-air long wave (Rossby wave) pattern around the Antarctic which may cause circulation changes (Björck et al. 1996). The historic position of the climatic Polar Front also deserves attention. Evidence from the Last Glacial Maximum indicates that the front was 5-7° latitude further north than today and was paralleled by a 3-5° northward shift of the Antarctic Convergence (oceanic Polar Front) in the South Atlantic, Indian and Southern Oceans during the LGM (Heusser 1989). Disposition of the westerly wind belt is also central to modeling atmospheric circulation in the southern high latitudes. A shift of the wind belt to the south may involve an increase in moisture from a subtropical source and may explain the synchronous warm events experienced in South America ca. 3330 to 2230 14 C yr BP (Clapperton and Sugden 1988) and in Australia between ca. 4000 to 2000 14C yr BP (Harle et al. 1999). There are also many records that do not show the MHH, or have warm periods indicated at different times. For example, in the Palmer Deep marine core, Domack et al. (2001) infer an early Holocene climate optimum 9070 to 3360 cal.14C yr BP and, in Lallemand Fjord, Taylor et al. (2001) report evidence for a late Holocene (< 3850 yr BP) neoglacial cooling. Similarly, it is absent from some of the continental interior ice cores (Dome C). However, the marked shift to colder climate conditions in both the Vostok and Komsomolskaya ice cores, observed after 2500 yr BP, has been suggested as marking the end of the MHH (Kulbe et al. 2001). Despite the apparent synchrony of the MHH in the Antarctic, its link to global climate remains uncertain and, in fact, is out-of-phase with the higher global temperatures within the Holocene from 6500 to 4500 yr BP (Folland et al. 1990). This has led researchers to

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highlight our currently poor understanding of the link between Antarctic and global climates and its effect on ice decay and sea-level rise (Kulbe et al. 2001). In summary, further paleoclimate records from the Antarctic Peninsula, where the greatest increases in temperature have been recorded in recent years (King 1994; Doran et al. 2002b), and better efforts to link chronologies of climate change from ice, marine and lake sediment sequences, will help refine regional paleoclimate reconstructions. If this is coupled with an increase in resolution of our records and the rigour with which they are interpreted, then the MHH (its timing, impacts and rates of change) presents both an analog and an opportunity for scientists to constrain modelled scenarios of the impacts of contemporary Antarctic Peninsula warming. It may then be possible to begin to understand how Antarctic Peninsula warming events can relate to both continental and global climate systems (and the relative timing of climate changes between the polar hemispheres), if these can influence ice shelf and glacier retreat and, in turn, sea-level. Case Study 3: Greenhouse gases and southern hemisphere climate The possibility that human actions, resulting from burning fossil fuels and the release of radiatively selective ‘greenhouse’ gases from agriculture and industry, could be a factor in causing rapid changes in global climate was noted at the end of the last century (Arrhenius 1896). It has been taken seriously by scientists since about 1960, when measurements of increasing CO2 content in the global atmosphere led to calculations of the net effect on the planetary surface heat balance (Plass 1956; Revelle and Suess 1957). It is only since about 1977, however, that advances in climate modeling have confirmed that measured changes in atmospheric chemistry were showing significant worldwide trends. This provided convincing indications that rapid changes in global climate in the near future appeared very likely, and that human activities, if not a contributing cause, were at least having an effect that reinforced the tendency toward change (Roots 1989). These concerns led to the World Meteorological Organisation releasing a statement in 1986 saying that “As a result of the increasing concentrations of greenhouse gases, it is now believed that in the first half of the next century, a rise in global mean temperature could occur which is greater than any in human history.” The possibility of an irreversible anthropogenically induced climate change has focused attention towards the Antarctic as it is one of the most climatically sensitive ecosystems on Earth, and has stimulated scientists to examine if there is evidence of a recent climate change. While orbital forcing is now well established as the pacemaker behind the major glacial/interglacial cycles, higher resolution climate forcing mechanisms (of the order of decades) remain to be identified. Current predictive models suggest that increases in global anthropogenic greenhouse gases would be manifested as an increase in the influence of precipitation over evaporation in coastal Antarctic regions (Kattenberg et al. 1995). Warmer temperatures will lead to a more vigorous hydrological cycle, which also suggests a possibility for more extreme rainfall events (Houghton et al. 1996). At present, knowledge is limited and it is difficult to determine whether current changes in the precipitation/evaporation (p/e) patterns in eastern Antarctic areas are natural or influenced by anthropogenic processes. Scientists in eastern Antarctica are therefore examining lake sediment cores to reconstruct the coastal p/e

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balance at high resolution for the last ca. 1000 years and to compare this with long-term patterns revealed by lower resolution studies spanning the Holocene. Extreme cold and aridity characterize the Antarctic environment and profoundly influence the limnological properties of lakes throughout the south polar region (Vincent 1987). Most continental aquatic life is restricted to the coastal fringes where temperatures are above freezing for only a few days or weeks of each year. Total annual precipitation, falling as snow, is between 100 and 150 mm of water equivalent (Weller et al. 1987). Consequently, low temperatures and precipitation have a significant influence on coastal Antarctic lake ecosystems. Such lakes, particularly arid closed systems, respond rapidly to climate-driven hydrological change. Fluctuations in the balance between precipitation and evaporation result both in changes in lake-level and the concentration of dissolved salts. The Vestfold Hills extend over an area of 410 km2 and have a large number of lakes ranging from brackish to hypersaline (Roberts and McMinn 1996). The sediments of these lakes provide a record of the changing climate through the preservation of climate-sensitive diatom assemblages (Roberts and McMinn 1999a). Paleosalinity has been determined from the microalgal fossil assemblages in sediment cores using a lake diatom-salinity transfer function (Roberts and McMinn 1998) and precipitation and evaporation patterns have been determined. In a study of Ace Lake, a customized paleohydrological model linked the diatom-derived paleosalinity of the lake to the paleohydrology of the lake basin (Roberts and McMinn 1999b). Application of this model revealed a detailed history of changes in the p/e balance during the late Holocene (Figure 8) that has been linked to periods of open marine water and periods of higher coastal precipitation (Roberts et al. 2001b). Conversely, a lower evaporation period between 150 and 200 corr. 14C yr BP has been found to correspond with the ‘Little Ice Age’ signal detected in oxygen isotopes of the Law Dome ice core (Figure 1) despite a large geographic separation of 35° longitude (Delmotte et al. 1999). The strong of agreement between the Law Dome ice core and the Vestfold Hills lake sediment core permits speculation that the ‘Little Ice Age’ was a widespread event. In fact, the coolest period in the Law Dome ice core record also corresponds with a cool anomaly in dendrochronological data (Cook et al. 1992; Bradley and Jones 1993) from Tasmania, Argentina and New Zealand. The latitudinal band from 30° to 40°S has been identified by Petit et al. (1991) as the dominant moisture source for central and eastern Antarctica, providing an understanding of how this period of cooling influenced the Antarctic environment at that time. Inferences can now be made on the causes of shifts in the p/e balance. These involve changes in the strength of the circumpolar trough to the north of this region influencing cyclonic activity storm tracks and changes in circumpolar circulation. Moisture transport is limited by saturation vapour pressure of the cold air and, particularly in coastal Antarctic zones, the frequency of ‘precipitation’ events depends on the intensity of cyclonic activity (Bromwich 1990). Sea ice cover can also change as a consequence of temperature fluctuations and this influences evaporation and moisture transport. Just as clear annual layers of undisturbed snow make ice cores particularly useful for high resolution studies of temperature change in the Antarctic, the special nature of many Antarctic lakes (meromixis, ice cover, etc.) limits bioturbative destruction of sediment deposition making them particularly suitable for recording fine-scale changes in the p/e balance of the catchment (Roberts and McMinn 1999a). These lacustrine records are enabling scientists to trace changes in the p/e balance of Antarctic lakes on timescales

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Figure 8. Reconstructions of net paleoevaporation calculated from the precipitation/evaporation balance at Ace Lake (upper curve), and summer į18O ‘temperatures’ at the Dome Summit South site, Law Dome ice core (lower curve). Paleoevaporation data are derived from diatoms sampled at ca. 25 year intervals in an Ace Lake sediment core. The curve is constructed using a Gaussian smoothing function (s = 7 years) to give comparable resolution to the į18O data (smoothed with s = 13 years). The chronology for the sediment core is derived from radiocarbon dating (with a radiocarbon reservoir age of 1300 years determined from a core surface date). The į18O chronology is determined absolutely from layer counting.

ranging from millennia to decades. On shorter timescales, microalgal communities are responding rapidly to current changes in the p/e balance enabling scientists to directly measure the links with meteorological data. With sufficient data of this type, general circulation modellers will have a more accurate proxy for past p/e trends that can be built

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into larger scale reconstructions of past atmospheric conditions over the Antarctic continent. These trends have now been directly linked with changes in the oxygen isotope data from ice cores. Therefore, integrating lake-based climate data with ice core and historical data may permit a calibration of the ice cores in terms of p/e, and extend the length of the record. Together these historical reconstructions and data on modern weather systems will enable environmental scientists to better predict the types of environmental changes that might result from shifting weather patterns and oceanographic conditions in the future. With increasing research effort and confidence in Antarctic limnological paleoclimate records, scientists can begin to answer the fundamental questions related to the climate change debate, such as “What are the approximate extremes of change in temperature and precipitation that might be expected in Antarctica?”, “When and how rapidly might such change occur?”, “How will Earth’s environment be affected and how will ecosystems respond?”. Antarctic lakes therefore offer one key to understanding future climate regimes on Earth. Case Study 4: Lake sediments and the interpretation of 1500 yr periodicities in ice cores The Holocene has received renewed focus due to the recognition that substantial and possibly global climate changes have occurred during this period (Steig 1999). Holocene variation occurred under conditions similar to today, and therefore is relevant to understanding contemporary climate change, and constraining predictions into the future. However, many of the smaller scale changes in ice cores, such as Holocene warm events, are difficult to interpret. In some cases these same changes are found at greater amplitude at coastal locations suggesting positive feedbacks due to sea ice and/or ice shelves (Masson et al. 2000). Recent comparisons of mid-latitude southern hemispheric climate proxies and the Taylor Dome ice core record indicate periodicities of ca. 1500 yr, but with a phase shift after the mid-Holocene (Lamy et al. 2001). These ca. 1500 yr periodicities in the Taylor Dome record are poorly understood and so scientists have been examining the McMurdo Dry Valley lake sediment record for an explanation of this paleoclimatic phenomenon. The McMurdo Dry Valleys is the largest ice-free region in Antarctica covering an area of 4800 km2. Although the valleys lie in one of the most extreme deserts on Earth, they contain about 20 closed-basin, perennially ice-covered lakes. Similar to lakes in the Vestfold Hills, these have responded rapidly to changes in the moisture balance of the region by exhibiting changes in their water-levels and salinity. Here paleolimnologists have instrumental records of lake-levels dating back to the Scott expedition of 1903 (Scott 1905), with which to calibrate their longer term reconstructions. The history of the McMurdo Dry Valley lakes is long and complex. Although the deglacial history of the region is now well known, century to millennial-scale climate records for the Holocene are lacking, and existing records are often contradictory. For instance, the Taylor Dome isotopic record indicates warming until ca. 6000 yr BP and a 2°C cooling since that time (Steig et al. 1998). The warming until ca. 6000 yr BP corresponds to the retreat of the West Antarctic Ice Sheet (WAIS) past Ross Island at ca. 7400 corr. 14C yr BP (Licht et al. 1996). However, penguin remains suggest warming temperatures, with more open water in the Ross Sea from 4000 to 3000 corr. 14C yr BP (Baroni and Orombelli 1994), and geomorphological evidence suggests a maximum advance of the Taylor Glacier

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between 3500 and 2500 14C yr BP (Higgins et al. 2000), at the same time as circumantarctic evidence for a MHH and an increase in precipitation. Evidence from both penguin and marine diatom records indicates that colder conditions ensued after ca. 3000 corr. 14C yr BP (Baroni and Orombelli 1994; Cunningham et al. 1999). However, ice core bore hole measurements indicate an earlier 1.2°C cold dry period from 4000 to 1000 yr BP (Clow 1999). The occurrence of this cold and dry period ending ca. 1000 to 1200 yr BP in the Dry Valleys has been supported by a variety of lake studies (Wilson 1964; Lyons et al. 1998), but evidence is lacking for lake draw-downs earlier in the Holocene where one might anticipate them. There is evidence that at ca. 1000 yr BP Taylor Dome and the Ross Sea warmed (Leventer et al. 1993; Masson et al. 2000). Bore hole logging at Taylor Dome supports a 2.1°C increase since 1000 yr BP (Clow 1999), which is a time of alpine glacier retreat in the Terra Nova Bay region as well (Baroni and Orombelli 1994). There is little evidence of the Taylor Dome ca. 1500 yr periodicities in the lake record and, conversely, no Little Ice Age (LIA) event is seen in the Taylor Dome record, but it has been observed in the Ross Sea (Leventer et al. 1993; Berkman 1994). The main periods of agreement in these records are therefore for a MHH, followed by late Holocene cooling with only some records having evidence of the Medieval Warm Period and a Little Ice Age (Lyons et al. 1997). These inconsistencies in the different paleoenvironmental records from the Ross Sea region (including the Taylor Dome ice core) suggest that paleolimnological research in the region has yet to achieve a satisfactory paleoclimate reconstruction for the Holocene. One problem with comparing all these proxies is that some record summer temperatures (lakelevels) and others annual temperatures (ice core records). This is being addressed by accurate and precise dating, and a direct linkage of sediment composition to processes occurring in the modern lake systems. Recent research has focused on a number of closed-basin, perennially ice-covered lakes in Taylor Valley (Figure 9). Lake-levels and lake ice thickness are constantly fluctuating in response to changing climate, primarily summer temperature (Fountain et al. 1999; Doran et al. 2002a). Based on 14C dating of organic matter preserved in fossil deltas and ancient strand-lines, the current lakes may be remnants of a much larger glacial lake ‘Lake Washburn’ that existed during the Last Glacial Maximum and occupied most of Taylor Valley (Stuiver et al. 1981; Hall and Denton 2000). Sediments from this lake, stranded on the sides of the valley, provide useful information on paleoclimate and the history of the WAIS. However this sediment record is sporadic and ends at ca. 8700 14C yr BP, the age of the youngest paleolake deposit found. This suggests that lake-levels have been mostly near or below present levels since that time, and that any information about lake history during the Holocene must be retrieved from the lake water or sediments. Scientists are now studying the influence of climate on the Dry Valley lakes, specifically identifying proxies for lake-level and lake ice thickness in Taylor Valley sediment cores, since both these indices are controlled by climate and impact sediment accumulation in the lake environment. There has been speculation that snow cover on ablation zones will reduce melt (Hall and Denton 1996; Fountain et al. 1999) but there is little empirical evidence to support this as a significant long-term control on local hydrology. Wharton et al. (1993) show that a simple “summer degree-days above freezing” model explained lake-level fluctuations in Lake Hoare exceptionally well. Likewise, Doran et al. (2002b) found that degree-days above freezing were significantly (P < 0.003) correlated to total stream flow in

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the Lake Fryxell basin between 1990 and 2000. Doran et al. (2002b) have also shown that lake levels have decreased between 1986 and 2000, during a time when summer temperatures decreased by 1.2qC per decade and solar insolation has increased by 8.1 W m-2 per decade. Thus, while snow cover on glaciers may play a modifying role, summer temperatures appear to be the main driver of local melt and lake-level rise in the Dry Valleys. Doran et al. (2002b) further show that decreasing summer temperature leads to reduced primary productivity in the lakes due to the thickening ice cover over time. Lake ice thickness is also influenced by air temperature, but if sublimation rates get exceedingly low (e.g., if snow cover remains on the ice for extended periods), ice thickness can increase substantially (McKay et al. 1985).

Figure 9. Maps of (a) Antarctica, (b) the McMurdo Dry Valleys region, and (c) the Taylor Valley. A photograph of Lake Hoare is shown in Figure 1.

Meteorological network data show that current summer temperatures are strongly controlled by elevation and distance from the coast (Doran et al. 2002b). If summer temperature is normalized to sea-level using the dry adiabatic lapse rate, temperature increases with distance from the coast at a rate of 0.09qC km-1 (r2 > 0.99, P < 0.0001). This is because onshore winds dominate in the summer, and winds warm as they progress inland. The nature of the mechanism has not yet been completely defined, but it is clear that the

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relationship holds true even for Wright and Victoria Valleys where large glaciers occupy the area between the coast and the valleys. During the LGM, when the Ross Ice Shelf was blocking the eastern end of the Dry Valleys (Stuiver et al. 1981; Denton and Marchant 2000), the effective coastline was further seaward and, thus, Doran et al. (2002b) propose that temperature increased in the Dry Valleys as distance to the coast increased. This is in agreement with the Taylor Dome record (located even further inland than the Dry Valleys) that shows increased aridity during the LGM (Steig et al. 2000), which would imply more solar insolation warming the valleys. While Hall and Denton (1996) may be correct in speculating that less snow on the glaciers results in more melt during the LGM, Doran et al. (2002b) believe this to be a side effect to the increased temperature with greater distance from the coast. They also predict that, during the high stages of glacial Lake Washburn (the high-stand glacial lake occupying the Taylor Valley), ice thickness was relatively thin and possibly seasonal. The lake ice was at the same elevation as the glacier ice being melted to form the large glacial lake. Efforts to link meteorological data from the Taylor Dome site (which is 150 km from Taylor Valley) with the Dry Valley meteorological data and the lake record over the same period will allow scientists to make better use of both data sets. Linking the Taylor Dome ice core to the Dry Valley lakes in this way may permit the determination of a statistical relationship with which to generate a time-constrained record of atmospheric conditions in the region extending beyond 200,000 yr BP (Steig et al. 1998, 2000). Recent research now demonstrates that the ca. 1500 yr periodicities in mid-latitude southern hemispheric proxies and the Taylor Dome record experienced a phase shift after the mid-Holocene possibly, coinciding with the onset of the modern state of the El Niño-Southern Oscillation system (Lamy et al. 2001). With a calibrated climate model between the McMurdo Dry Valleys and Taylor Dome, it should be possible to utilise the ice record to refine lake chronologies and examine these climate periodicities in greater detail. Discussion The four case studies, presented above, demonstrate how Antarctic paleolimnology is providing answers to some of the major questions in Earth system science. In some cases, studies of lake sediments are providing direct geological evidence (Case study 1) of sealevel changes that can be incorporated into glaciological models, and in others it provides an historical climate analogue (Case study 2) that can be used to constrain responses in the Earth’s system to known parameters of change, such as a shift in temperature recorded in ice cores. In the remaining examples (Case studies 3 and 4), paleolimnology is providing one part of the evidence contributing to an understanding of the main forcings and feedbacks in the climate system. The case studies also highlight the need for continued efforts to improve the rigour with which paleolimnology is applied in the Antarctic. First, there is a need for improved transfer functions to quantitatively interpret the lake sediment archive, for example transfer functions for temperature, light, nutrients etc., using a greater range of biological and chemical indicators. Second, much finer resolution is required in establishing chronologies. This means more dates and the application of multiple dating technologies to address chronological problems. This is essential if scientists are to be able to determine the precise

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order of events and identify the timing of triggers, forcings and feedbacks that have changed the Earth system during glacial and interglacial cycles and are influencing climate changes in the present interglacial. It cannot be stressed enough that successful theories in Earth system science depend ultimately on good chronologies. Finally, all the stratigraphic disciplines need to continue to build close collaborations with paleoclimate and ice sheet modellers if scientists are to have the best possible opportunity to identify the critical thresholds and switches that may result in abrupt climate change. Outlook Antarctic paleolimnology is at a turning point. More than a decade of regional studies has helped define the environmental history of the Antarctic and brought the discipline to a point from which it can address more specific questions. Until recently, some of the climate change evidence from the coastal oases of Antarctica has been overlooked in climate model (GCM) reconstructions. However, these studies have highlighted that many of the smaller scale changes in ice cores, such as Holocene warm events, are characterized by greater amplitude at coastal locations, suggesting positive feedbacks due to sea ice and/or ice shelves (Masson et al. 2000). This direct link has again emphasized the importance of coastal lakes in linking different types of Antarctic paleoclimate records in order to understand forcing and feedback mechanisms. Similarly, Antarctic lake sediments usually offer higher resolution records than Antarctic marine sediments and can archive the ‘terrestrial’ responses to changes in sea ice extent, atmospheric and oceanic circulation. Antarctic paleolimnology is set to develop into new areas. There are a number of Antarctic lake types that have received little attention but can answer important questions. For example, epishelf lakes provide researchers with the opportunity to study the history of the ice shelves that currently isolate them from the sea. These ice shelves are especially vulnerable to global or regional warming because they are exposed to thinning, both from above and below, and to catastrophic break-up. Whilst melting ice shelves themselves make little or no contribution to sea-level change, they may help restrain the Antarctic ice streams that flow into them, on account of their being pinned to submerged rises in the ocean floor. One theory proposed by Hoffman et al. (cited in Gornitz 1995) is that, as the ice shelves disappear from basal melting and increased ice discharge, the ice streams behind them will flow into the ocean at a faster rate (e.g., De Angelis and Skvarca 2003) resulting in a rapid rise in sea-level. As stated earlier, this theory is not universally accepted and, in some models, such as those of the IPCC (Warrick et al. 1990), it is estimated that Antarctica will contribute negatively (0 to -0.6 mm/yr) to sea-level rise in the next 100 years under ‘global warming’ due to an increased mass balance (precipitation) on the continent. For example, on the Antarctic Peninsula there is evidence for a 20% increase in the annual precipitation since 1950 with an increased frequency of meso-scale cyclones from the west (Peel 1992), though ablation may offset this. However, to date, evidence ‘on the ground’ for rapid disintegration of the ice shelves (Vaughan and Doake 1996) is more substantial than that for increased mass balance. Indeed, many studies show a retreat and thinning of the coastal glaciers of the Antarctic Peninsula (e.g., Smith et al. 1998) and of the glaciers of the WAIS (Shepherd et al. 2001) in recent decades. Results of GCM model predictions also remain controversial (see examples from the Arctic in Schneider 1992), and ice streams in the Ross

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Sea sector of the WAIS continue to exhibit a neutral to positive mass balance (Joughin and Tulaczyk 2002). Continued loss of the ice shelves will modify the exchange of energy between the ocean, atmosphere and ice sheet. In particular, an increase in the area available for brine exclusion during sea ice formation could turn these newly exposed continental shelf areas into important sources of Antarctic Bottom Water. In a study of the Ross Ice Shelf, Joughin and Tulaczyk (2002) speculate that an area of 400,000 km2 of new shallow sea surface area, resulting from the disappearance of the Ross Ice shelf, could create an Antarctic Bottom Water source that could out-compete the North Atlantic source of bottom water and switch the global ocean into a new mode of thermohaline circulation. Sediments from epishelf lakes should therefore permit an understanding of long-term variability in ice shelf collapse, and allow of testing and constraining of glaciological models which are needed to predict the response of Antarctic ice sheets and ice shelves to regional or ‘global warming’ in the future. To date there has also been a paucity of paleolimnological research on epi-glacial and supraglacial lakes, though there are some exciting biological possibilities. For example, supraglacial lakes can be used to test the biological refugia theory of the Snowball Earth Hypothesis (Vincent et al. 2000). Work on the colonisation of lakes in the Vestfold Hills indicates that the flora may have occupied supraglacial or other local refugia during glaciations, from which they have been able to rapidly colonise newly available habitats (Swadling et al. 2001). The subglacial lakes in Antarctica have also remained unexplored except using remote technology (airborne radio echo sounding). Of these, the largest and deepest, Lake Vostok, has received most attention as ice coring has drilled to within 500 m of the lake surface, making the exploration of this lake and its sediments a real possibility. Exploring these systems will almost certainly require the resources of several national programmes and some considerable technological effort to ensure sampling procedures that avoid contamination. The intriguing question is what these sub-glacial lakes are likely to contain biologically. The evidence suggests that scientists will discover communities of micro-organisms, as the ice core itself has revealed a diversity of yeasts, Actinomycetes, algae and spore-forming bacteria (Karl et al. 1999; Priscu et al. 1999). Some of these organisms have remained viable in the ice for 36,000 years. Since the ice above Lake Vostok is over 500,000 years old, there is speculation that it will contain microbial genomes which have been isolated from the rest of the biological world for that period. The prospect of analysing these subglacial communities, their genetics and physiology, offers considerable excitement to polar biologists and opens up a wide variety of questions concerning Antarctic biogeography and microbial evolution. However, better definition of the molecular taxonomy of the Antarctic microbiota is required before it will be possible to assess the significance of biological entities in Lake Vostok. Whilst scientists are still constrained by logistics and funding, the groundwork to some of the biological questions might be developed by studying recently exposed subglacial lakes which have been appearing from under retreating ice masses along the ice margins of many Antarctic oases. Some of these have escaped scouring and have maintained an intact sedimentary record but, to date, have received remarkably little attention.

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Conclusions Despite the Milankovitch orbital pacemaker of climate change, the Earth’s system is nonlinear in its response and delivers its own scheme of climate changes governed by the secondary properties and internal feedback cycles of the geosphere, cryosphere, hydrosphere and atmosphere. Each of these contains elements with physical properties, that respond at different rates to external forcing. Of these elements, water is perhaps the most influential. At a first glance, the Earth appears to be disregarding orbital forcing with irregular and unpredictable anarchy but, slowly, the historical records, together with modern data, are providing knowledge of the physical principles governing this fascinating subject. Of the stratigraphic records, those from the Antarctic have been amongst the most pivotal, each type (lake, ice, marine) adding to our understanding of climate change. In this chapter we hope to have left the impression that Antarctic paleolimnology plays an integral role in piecing together the history of environmental change in the Antarctic and, in turn, the key role that the Antarctic plays in Earth system science. Paleolimnology occupies a distinctive niche in providing an understanding of terrestrial responses to environmental change and of changes at the interfaces between ice and land, and land and sea. Often, the paleolimnological record appears to be hampered by high frequency variations, which are often not convincingly related to the major climate forcing events detected in marine and ice cores. In effect, the detail can be too variable and the resolution too high for these types of comparisons to take place. This is best illustrated by examples from the northern hemisphere where the recent discovery by glaciologists of Dansgaard-Oeschger events, and by oceanographers of Heinrich events, has led to a revelation that these events have, for a long time, been available from pollen records (Behre and Lade 1986; Guiot 1997). So can paleolimnology lead debates in Earth system science? The answer is that the stratigraphic records must coexist; we cannot understand the paleolimnological data without determining the forcing mechanisms from marine and ice cores, and conversely the impact of mechanisms found in marine and ice cores cannot be translated into terrestrial, sea-level, ecological and environmental responses without recourse to paleolimnological and terrestrial records. In effect, each stratigraphic deposit provides a new opportunity to crosscheck data on climatic reconstruction and new challenges to match their geochronologies. The coexistence of ice core, marine core, atmospheric, glaciological, geomorphological and environmental response data (lake and terrestrial records) is therefore critical if we are to understand past climate changes and their impacts, and to have the best possible chance of making accurate predictions into the future. Summary Lake sediments are valuable stratigraphic archives of environmental change in the Antarctic region. Together with ice cores, marine cores, geomorphology and glaciology, they are providing answers to some fundamental questions in Earth system science. In this chapter we review the different lake types found in the Antarctic and the range of questions that can be addressed by examining their sediments. We then provide practical advice on logistics, sediment coring and chronostratigraphic methods. This is accompanied by a synthesis of physical, mineralogical and inorganic chemistry methods, together with terrestrial, algal and

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siliceous markers and organic geochemistry. Four case studies are presented to show how Antarctic paleolimnological studies are integral to addressing global science questions. We conclude with a discussion on the future potential of Antarctic paleolimnology. Acknowledgements This text has benefited from many interesting discussions with Mike Bentley, Elie Verleyen and James Smith during a field season in 2001/2. Thanks also to Wim Vyverman, Peter Leavitt, John Gibson and Svante Björck for their valuable contributions. Krystyna Saunders and Melissa Barrett helped with the bibliography and Johanna Laybourn Parry kindly loaned a draft review of Antarctic limnology, which has been extensively referred to. We acknowledge the financial support of the British Antarctic Survey (BAS) Signals in Antarctica of Past Global Changes Program and the Belgian Federal Office for Scientific, Technical and Cultural Affairs Project, Late Quaternary Climate History of Antarctic Coastal Environments, a multi-proxy approach (LAQUAN). Finally, we thank Martin Melles and the DPER Series’ editors for highly constructive reviews. References Adamson D.A. and Pickard J. 1983. Late Quaternary ice movement across the Vestfold Hills, East Antarctica. In: Oliver R.L., James P.R. and Jago J.B. (eds), Antarctic Earth Science. Aust. Acad. Sci., Canberra ACT, pp. 465-469. Adamson D.A. and Pickard J. 1986. Physiography and geomorphology of the Vestfold Hills. In: Pickard J. (ed.), Antarctic Oasis: terrestrial environments and history of the Vestfold Hills. Academic Press, Sydney, pp. 99-139. Adamson D.A., Mabin M.C.G. and Luly J.G. 1997. Holocene isostasy and late Cenozoic development of landforms including Beaver and Radok Lake basins in the Amery Oasis, Prince Charles Mountains, Antarctica. Antarct. Sci. 9: 299-306. Airs R.L., Atkinson J.E. and Keely B.J. 2001. Development and application of a high resolution liquid chromtographic method for the analysis of complex pigment distributions. J. Chromatog. A 917: 167-177. Akiyama M., Hayashi M., Matsumoto G.I. and Miura H. 1990. Plant remains and related substances in the past lacustrine sediments of the Mt. Riiser-Larsen area, Enderby Land, east Antarctica. Proc. NIPR Symp. Polar Biol. 3: 207-217. Alley R.B. 2000. The two mile time machine. Ice cores, abrupt climate change, and our future. Princeton University Press, Princeton and Oxford, 229 pp. Anderson J.B. 1999. Antarctic marine geology. Cambridge University Press, Cambridge, UK, 496 pp. Appleby P.G., Jones V.J. and Ellis-Evans J.C. 1995. Radiometric dating of lake-sediments from Signy Island (Maritime Antarctic) - evidence of recent climatic-change. J. Paleolim. 13: 179-191. Arrhenius S. 1896. On the influence of carbonic acid in the air upon the temperature of the ground. Phil. Mag. 41: 237-250. Austin J.J., Smith A.B. and Thomas R.H. 1997. Palaeontology in a molecular world: the search for authentic ancient DNA. Tree 12: 303-306. Bard E., Hamelin B., Arnold M., Montaggioni L., Cabioch G., Faure G. and Rougerie F. 1996. Deglacial sea-level record from Tahiti corals and the timing of global meltwater discharge. Nature 382: 241-244.

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Appendix 1. Stratigraphic paleolimnological studies of Antarctic freshwater lakes organized by region. Notes: In applications, all studies include chronological controls and considerations of lake development and history. Sample locations are indicated in Figure 1. Lake(s)

Indicators Continental Antarctic Vestfold Hills Isostasy Diatoms, zoological indicators Isostasy, relative sea-level Stable isotopes (į13C, į18O), geochemistry Holocene climate change Diatoms, physical lithostratigraphy Salinity, hydrology Water column structure Isostasy, relative sea-level Diatoms Salinity, hydrology Diatoms Salinity, isostasy, Diatoms deglaciation Relative sea-level Physical lithostratigraphy, survey

Reference

Roberts and McMinn 1999a

Ace Ace

Salinity, isostasy, deglaciation Hydrology Biogeography

Kirisjes Pond, Pup Lagoon Nella 11 lakes Kirisjes Pond Kirisjes Pond Nella

Salinity, isostasy, deglaciation Glacial history Glacial history Environmental change Organic geochemistry Ecosystem function

Watts Highway, Organic Nicholson Ace, Organic Ace Deep Anderson Organic, Watts, Highway, Druzhby, Anderson, Ace Ace

Pup Lagoon Reid 22 lakes Trowbridge Hoare Hoare Hoare, Fryxell 30 lakes

Application

Diatoms Diatoms Zoological indicators Larsemann Hills Diatoms

Geochronology, erosion Geochronology Sedimentary pigments Sedimentary pigments Sedimentary pigments, organic geochemistry Isostasy, relative sea-level, Diatoms, sedimentary MHH, colonisation pigments Past UV radiation Sedimentary pigments McMurdo Dry Valleys Review regional Review multiple methods paleolimnology Glacial history, hydrology, Stable isotopes (į18O) relative sea-level Diatom taphonomy, Diatoms sedimentology Antarctic lakes as Mars Stable isotopes (į13C) analogues Precipitation/evaporation Stable isotopes Late Quaternary lake Geomorphology, hydrology, glaciology geochronology, physical lithostratigraphy

Pickard et al. 1986b Bird et al. 1991 Bronge 1992 Gibson and Burton 1996 Fulford-Smith and Sikes 1996 Whitehead and McMinn 1997 Roberts and McMinn 1998 Zwartz et al. 1998a

Roberts and McMinn 1999b Swadling et al. 2001 Gillieson 1991 Burgess et al. 1997 Hodgson et al. 2001a Squier et al. 2002 Walker et al. 2002 Edwards et al. 2003 Verleyen et al. 2003a Leavitt et al. 2003 Doran et al. 1994 Clayton-Greene et al. 1988 Spaulding et al. 1997 Doran et al. 1998 Lyons et al. 1998 Hendy 2000

472

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Appendix 1. Stratigraphic paleolimnological studies of Antarctic freshwater lakes organized by region. Notes: In applications, all studies include chronological controls and considerations of lake development and history. Sample locations are indicated in Figure 1. (continued) Lake(s)

Application

Figurnoye

Paleoenvironment and paleoclimate Radiocarbon dating, reservoir calibration

Figurnoye, Burevestnik, Rybiy Khvost Bay Figurnoye, Kakapon Bay, Rybiy Khvost Bay, Izvilistaya Inlet, Polyanskogo Jaw White Smoke Rybiy Khvost Bay Figurnoye

Glubokoye, Zub, Untersee Richardson Southwest basin, Lakes X and Y L. Richardson

‘Jurassic’ lakes

Skua

Åsa, Midge, Chester Cone Midge Åsa, Midge, Chester Cone

Regional environmental change, deglaciation, relative sea-level

Indicators Bunger Hills Plant macrofossils, physical lithostratigraphy Geochronology

Reference Verkulich and Melles 1992 Melles et al. 1994b

Physical lithostratigraphy, Melles et al. 1997 geochronology, geochemistry

Paleosalinity Diatoms Lake processes, site history Physical lithostratigraphy, geochemistry, stable isotopes Holocene climate history, Lithology, diatoms, MHH inorganic geochemistry, isotopes Climate / environmental Physical lithostratigraphy, change, relative sea-level geochemistry, diatoms Schirmacher / Untersee Oases Late Quaternary climate Physical lithostratigraphy, and environment geochemistry, stable isotopes Mt. Riiser-Larsen Relative sea-level Sediment description

Roberts et al. 2000a Doran et al. 2000

Environmental change

Akiyama et al. 1990

Plant macrofossils, organic geochemistry Transantarctic Mountains Species assemblages in Zoological indicators Jurassic lakes

Kulbe et al. 2001 Verkulich et al. 2002

Schwab 1998

Zwartz et al. 1998b

Tasch 1970

Antarctic Peninsula and subantarctic islands Horseshoe Island Isostasy Diatoms, chemistry, Wasell and Håkansson 1992 lithology Livingston Island Tephrochronology Tephra Björck et al. 1991c Deglaciation, environmental change, MHH Deglaciation, environmental change, MHH

Physical lithostratigraphy, Björck et al. 1991a diatoms, Pollen and spores Physical lithostratigraphy, Björck et al. 1993 diatoms, Pollen and spores

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Appendix 1. Stratigraphic paleolimnological studies of Antarctic freshwater lakes organized by region. Notes: In applications, all studies include chronological controls and considerations of lake development and history. Sample locations are indicated in Figure 1. (continued) Lake(s)

Application

Indicators King George Island 6 lakes Landscape and evolution Physical lithostratigraphy of biota 210 Not named Pb measurement Geochronology (210Pb) Tiefersee, Kiteschsee, Deglaciation, sea-level Physical lithostratigraphy, Jurasee change, isostatic history geochemistry, diatoms Mondsee, Tiefersee Deglaciation, Physical lithostratigraphy, environmental change diatoms Physical lithostratigraphy, Long Isostatic uplift, marine / diatoms estuarine / lacustrine development Not specified Climate changes cf. Paleoclimate model Southern Chile Yanou Holocene environmental Diatoms change, MHH Lake at 62°12.45’S, Evaluation of 210Pb dating Geochronology (210Pb) 58°56.20’W Yan’ou Isostasy Diatoms, zoological indicators Hotel Holocene environmental Diatoms, lithostratigraphy change 10 lakes Tephrochronology Tephra Y2 Past penguin populations Organic geochemistry Xihu Holocene environmental Physical lithostratigraphy, change stable isotopes Hope Bay Boeckella Holocene climate change Physical lithostratigraphy, geochronology, geomorphology Boeckella Tephrochronology Geochronology (tephra) Physical lithostratigraphy James Ross Island Hidden Holocene climate change Physical lithostratigraphy, geochronology, geomorphology Boulder, Terrapin, Deglaciation, lake Physical lithostratigraphy, Keyhole development, MHH diatoms, geochemisty, zoological indicators

Reference Tatur and delValle 1986; Tatur et al. 1991 Zhao 1990 Mäusbacher et al. 1989 Schmidt et al. 1990 Martinez-Macchiavello et al. 1996 Zhao 1990; Zhao 1997 Yang and Harwood 1997 Ye and Cuihua 1997 Shen et al. 1998 Tatur et al. 1999a Tatur et al. 1999b Sun et al. 2000 Xiaomei et al. 2002

Zale and Karlen 1989 Björck et al. 1991c

Zale and Karlen 1989 Björck et al. 1996

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Appendix 1. Stratigraphic paleolimnological studies of antarctic freshwater lakes organized by region. Notes: In applications, all studies include chronological controls and considerations of lake development and history. Sample locations are indicated in Figure 1. (concluded) Lake(s) Sombre Sombre, Heywood Sombre, Midge Sombre, Heywood Sombre Upper, Lower, Maiviken Valley Tønsberg Peninsula lakes Stream bank core

Application

Indicators Maritime and subantarctic islands Signy Island Human impacts on seal Zoological indicators, populations historical records Natural vs. human impacts Zoological indicators, on seal populations historical records Tephrochronology, Geochronology (tephra) provenance Holocene environmental Diatoms, plant change, MHH macrofossils, zoological indicators Hydrological change Stable isotopes (į18O) South Georgia Vegetation, climate Pollen, spores, diatoms, change, productivity desmids Deglaciation, climate Physical lithostratigraphy, change stable isotopes (į13C) Kerguelen Island Environmental Diatoms reconstruction

Reference Hodgson and Johnston 1997 Hodgson et al. 1998b Hodgson et al. 1998a Jones et al. 2000 Noon et al. 2003 Clapperton et al. 1989; Birnie 1990 Rosqvist et al. 1999

Larson 1974

15. PALEOLIMNOLOGY OF EXTREME COLD EXTRATERRESTRIAL ENVIRONMENTS

TERRESTRIAL

AND

PETER T. DORAN ([email protected]) Department of Earth and Environmental Sciences University of Illinois at Chicago 845 W. Taylor St. Chicago, Illinois 60607-7059, USA JOHN C. PRISCU ([email protected]) Land Resources and Environmental Sciences Montana State University Bozeman, Montana 59717, USA W. BERRY LYONS ([email protected]) Byrd Polar Research Center Ohio State University Columbus, Ohio 43210, USA ROSS D. POWELL ([email protected]) Department of Geology and Environmental Geosciences Northern Illinois University DeKalb, Illinois 60115, USA DALE T. ANDERSEN ([email protected]) SETI Institute Center for the Study of Life in the Universe 2035 Landings Drive Mountain View, California 94043, USA and ROBERT J. POREDA ([email protected]) Earth and Environmental Sciences University of Rochester Rochester, New York 14627, USA

475 R. Pienitz, M.S.V. Douglas and J.P. Smol (eds), 2004. Long-term Environmental Change in Arctic and Antarctic Lakes. Springer. Printed in the Netherlands.

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Key words: Extreme environments, Antarctica, Perennial lake ice, Ice-sealed lakes, Mars, Exobiology, Lake Vostok, Subglacial lakes, Sediments, Water column

Introduction The study of environments that exist near planetary extremes has become a research area of considerable contemporary importance. Our understanding of life’s environmental limits has been largely modified as discoveries and evidence of functional ecosystems are reported in places such as the deep ocean, hot springs, deep beneath the Antarctic ice cap, and deep within the Earth. These discoveries have led to a separate field of interdisciplinary study focused on life in extreme environments. A working definition of an extreme environment is one where physical and chemical conditions approach or exceed the perceived tolerances for life (Benchley et al. 1998). Some organisms can thrive under these conditions and not survive in more clement environments. A key impetus for studying extreme environments is to understand the origin and evolution of life itself. Life originated on our planet when a select set of environmental conditions allowed molecules to organize themselves into self-replicating entities. Defining the boundaries of those conditions is essential to our understanding of life’s origins on our planet and other extraterrestrial bodies. Liquid water is essential for all life on Earth, including life that exists in extreme environments. Therefore, the study of inland waters (limnology) is by definition an essential element in the study of life in extreme environments. In temperate regions, extreme aquatic environments can be found mostly as a result of hypersalinity (e.g., salt lakes) and heat (hot springs). In polar regions, extreme aquatic environments are created by cold hypersaline water bodies, thick permanent lake ice covers, and the combined effect of high pressure, aphotic, and oligotrophic conditions in subglacial lakes. In this chapter, we consider these polar extreme limnological environments, and in particular the historical record preserved by them. We also provide a discussion of the connection between these polar extreme environments and water bodies that may exist elsewhere in our solar system, notably Mars and Europa. Perennially ice-covered lakes A number of lakes exist in the polar regions that maintain their ice cover throughout the year. The list in Tables 1 and 2 is not meant to be exhaustive as there are undoubtedly numerous lakes not accounted for here, particularly glacier-contact lakes at high altitudes. Impact of perennial ice covers Perennial ice covers have numerous effects on the underlying water column including: (a) reduction of light penetration, (b) reduction of gas exchange between lake water and

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Table 1. Documented perennially ice-covered lakes in the Arctic.

Location

Ice (m)

lat.

1. Angiussaq Lake, Greenland1

2-2.5

77º00’N 66º08’W Barnes 1960

2. Pond at Rundfjeld, central Ellesmere Island2

5.5

78º05’N 81º04’W Blake 1989

3. Ziegler Island, Franz < 2.4 Joseph Land3

long.

81º00’N 56º02’E

Selected References

Panzenbock et al. 2000

4. Lake A, northern Ellesmere Island4

2

82º08’N 78º00’W Belzile et al. 2001; Gibson et al. 2002

5. Ward Hunt Lake, Ward Hunt Island

4

83º01’N 74º02’W Villeneuve et al. 2001; R.A. Wharton, pers. comm.

1

At least one other smaller lake nearby also reported to be perennial Pond was frozen to bed in winter. 14C date indicates lake has not thawed in > 5000 yr 3 Single unnamed lake in the centre of the island 4 Lake C nearby believed perennially ice-covered but not studied 2

the atmosphere, (c) reduction in water column mixing, and (d) alteration of sedimentation pathways (Wharton et al. 1989; Gibson et al. 2002). Although all of these effects are important in defining the modern ecosystem of the lakes, and therefore the character of sedimentary deposits, the alteration of sediment pathways is likely the most important to paleolimnology. When a perennial ice cover is present, sediments can either be blown across the lakes or get deposited on the surface. In either case, sedimentation within the lake is altered. Sediments that get trapped in the ice will be warmed by the sun and melt into the ice to a depth that is largely related to the sediment grain diameter (Hendy 2000a). The depth to which any sized sediment can sink through this ice is limited, and layers of sediment collect at the dynamic equilibrium between particles melting downward and the ice cover moving upward through ablation and ice growth. This process forms sediment layers and associated lenses or inclusions of liquid water in the ice cover at approximately 2 m beneath the surface (Fritsen et al. 1998; Priscu et al. 1998). Beyond this level in a thick ice cover, sediment has to make its way through cracks in the ice, resulting in heterogeneous ridges and mounds of sediment on the bottom (Figure 1). Additionally, different lakes and different areas of lakes have varied rates of sediment accumulation on the ice, so that this process is not consistent across all lakes (Adams et al. 1998). Arguably the most studied perennial ice-covered lakes occur in the McMurdo Dry Valleys of Antarctica (77-78ºS, 160-164ºE). These lakes occur along the valley bottoms and are mostly closed basins. Groundwater exchange in the lakes is believed to be minimal, but remains largely uninvestigated. A previous review of the paleolimnology of these lakes is given in Doran et al. (1994a), and they are also treated in Hodgson et

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al. (this volume). The McMurdo Dry Valleys comprise a cold desert ecosystem, with mean annual temperatures ranging between -14.8 to -30ºC, and annual precipitation (all as snow) usually < 100 mm/yr with as little as 7 mm recorded in a single year (Doran et al. 2002a). Dry valley lake levels and lake ice thicknesses are constantly fluctuating in Table 2. Documented perennially ice-covered lakes in the Antarctic.

Location

Ice (m)

lat.

long.

1. Bunger Hills, various1

<3

66º03’S

100º08’E Doran et al. 1996, 2000

2. Frames Mountains, 3.5-4+ MacRobertson Land, various

67º08’S

62º08’E

Pickard and Adamson 1983; Chambers et al. 1986

3. Vestfold Hills, various

2.5-5+

68º05’S

78º00’E

Pickard and Adamson 1983

4. Larsemann Hills, various2

?-2.2-?

69º03’S

76º00’E

5. Schirmacher Hills, various3

1.6- > 3

70º08’S

11º07’E

Gillieson et al. 1990; D.A. Hodgson, pers. comm.; J. Burgess, pers. comm. Bormann and Fritzsche 1995; Gore 1997

6. Beaver Lake, MacRobertson Land

3-4

70º08’S

68º03’E

Laybourn-Parry et al. 2001

7. Ablation Valley4

2.5-4.5

70º08’S

68º04’W

Heywood 1977

8. Central Dronning Maud Land, various5

2.2-3.8

71º04’S

13º04’E

Bormann and Fritzsche 1995; Loopman et al. 1986; Wand et al. 1997

9. Citadel Bastion Lake, Alexander Island

3.7

72º00’S

68º03’W

D.A. Hodgson, pers. comm.

10. McMurdo Dry Valleys6

3-19

77º00’S

163º00’E Wharton et al. 1989; Chinn 1993; Doran et al. 1994a, 2002b,c

11. Lake Wilson

4

79º08’S

159º06’E Webster et al. 1996

1

Selected References

Only ice contact lakes at edge of oasis (mostly epishelf) Lake Ferris (LH11) and small lake to the north of Lake Oskar (LH18), L22, L25 (lake names from Gillieson et al. 1990, but ice observations unpublished) 3 Ice thickness data limited. Perennially ice-covered lakes are all epishelf (at least 6 lakes) 4 3 proglacial lakes: Ablation, Moutonée, and another unnamed 5 Lakes Untersee, Obersee and Burevestniksee 6 Appoximately 20 permanently ice-covered lakes and ponds (Doran et al. 1994) 2

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Figure 1. Sand mound on the bottom of Lake Hoare in the McMurdo Dry Valleys. The leg of a sediment trap can be seen in the top of the image.

Figure 2 . Perched deltas on Crescent Stream in the McMurdo Dry Valleys. Person standing on distant delta for scale. Perennially ice-covered Lake Fryxell is in the background.

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Figure 3. The food web in McMurdo Dry Valley lakes. HNAN is heterotrophic nanoflagellates. DOC is dissolved organic carbon. The X covers organisms not found in this ecosystem.

response to changing climate, primarily summer temperature (Chinn 1993; Fountain et al. 1999; Doran et al. 2002b). The current lakes have been suggested to be remnants of much larger glacial lakes that occupied the valleys during the last glacial maximum (LGM) (Stuiver et al. 1981; Hall and Denton 2000a; Hendy 2000b). Sediments from these lakes stranded on the sides of the valleys (Figure 2) provide useful information on the paleoclimate and the history of the West Antarctic Ice Sheet (Stuiver et al. 1981; Hall and Denton 2000b). However, this sediment record is sporadic, and in Taylor Valley is near 8700 yr BP; the radiocarbon (14C age) of the youngest paleolake deposit found (Hall and Denton 2000a). This suggests that lake-levels have been mostly near or below present levels since that time, and that detailed information about lake history during the Holocene must be retrieved from the lake water or sediments. The water columns of the McMurdo Dry Valley lakes are dominated by protists (few metazoans have been observed) and the benthos lacks burrowing fauna (Figure 3). There is no sediment bioturbation in the traditional sense of the word. However, because the lakes have elevated levels of dissolved gases, conditions within the shallow zones of the lakes cause microbial mats growing on the bottom to become buoyant and start to float off the bottom as “lift-off” mats (Parker et al. 1982). In areas where this occurs, disruption of the sediment surface can be significant (Figure 4). After rising through the water column, lift-off mats can become frozen in the base of the lake ice and eventually migrate up to the surface of the ice due to ablation from the top and freezing onto the bottom of the floating ice cover. Lift-off of mats located along the shallow water perimeters of the lakes (Wharton et al. 1989) may be an important process in the carbon cycle of these lakes (e.g., Parker et al. 1982). Priscu et al. (1998)

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showed that communities seeded by surface aeolian deposition actually live within the ice cover producing organic matter within the ice that eventually reaches the lake bottom. Underwater SCUBA observations have shown that lift-off does not always occur within the shallow “lift-off zone”. We have made the further observation that liftoff is frequent in areas where there is significant sediment rainout from the ice cover. It therefore appears that depth and disturbance are requirements for benthic mat lift-off in these lakes.

Figure 4. Photo of microbial mat lifting off of the bottom of Lake Hoare. Picture top to bottom is approximately 0.5 m.

Paleolimnology from sediments in perennially ice-covered lakes Lyons et al. (1985) analysed an archived core taken from Lake Vanda (77º53’S, 161º55’E) during the Dry Valley Drilling Project (DVDP). Preservation of the core stratigraphy was poor, but they were able to identify three desiccation events represented by peaks in total salt, CaCO3, organic carbon, and biogenic silica. Lawrence and Hendy (1985) extracted 14 cores from Lake Fryxell (77º61’S, 163º18’E), the longest of which was just over 1 m. Based on two 14C dates of carbonates, they concluded that the age of the core was in excess of 21 Ka BP. However, carbon reservoir results of Doran et al. (1999) indicated that these ages can

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only be considered as absolute maxima, and the closest date to the surface of the core was at ca. 0.5 m, so that a surface correction cannot be applied. Lawrence and Hendy (1985) used carbonate phase and į18O changes to infer past salinity changes. Squyres et al. (1991) used a 1.5 m spaced 3 row by 3 column grid of short (< 40 cm) cores to show the heterogeneity of sedimentation in Lake Hoare (77º63’S, 162º85’E), and tried to define a perennially ice-covered lake facies. Doran et al. (1994b) analysed the į13C of buried organic mats and carbonates, CaCO3 and organic matter (OM) content, and diatom species shifts in a ca. 30 cm core from Lake Hoare. They showed that sediment characteristics vary markedly over the short length of this core, but carbonates are consistently calcite, unlike the sediments of Lake Fryxell which contain alternating sequences of calcite and aragonite. Spaulding et al. (1997) were able to show zonation of diatoms in modern Lake Hoare surface sediments, but were unable to use that information to draw any firm paleoenvironmental conclusions from a 30 cm core of the same lake. Doran et al. (1999) used 14C dates in a Lake Hoare sediment core to infer a sediment accumulation rate of 0.015 cm yr-1. Bishop et al. (2001) did a comparative study between a 38 cm sediment core from a deep anoxic zone in Lake Hoare and a 47.5 cm core from a shallower oxic region of the lake. They showed that organic matter į13C and į15N trends provide a more complex history for the anoxic region sediments, and that biogenic pyrite found in the core from the anoxic zone is associated with depleted į34S values and high organic C values (Figure 5).

Figure 5. Lake Hoare core profiles for isotopes. į13C and į15N trends for oxic and anoxic region cores; į34S trends for the anoxic core; and percent by weight sulphur (Wt.% S) due to pyrite from Mössbauer Spectroscopy measurements and due to insoluble S (mostly pyrite) from elemental analyzer measurements (from Bishop et al. 2001).

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Paleolimnology from perennially ice-covered water columns Due to the isolation from wind mixing and strong chemical stratification in perennially ice-covered lakes, water chemistry can provide abundant paleolimnological information, a fact that was first recognized and used by Wilson (1964). Many of these lakes have evaporated to near dryness in the past, leading to the concentration of solutes and the precipitation of various salts during times of drier and/or colder conditions. As the climate ameliorated, glacier melt flowed into the lake basins on top of the concentrated monimolimnion, leaving the meromictic conditions that are observed today. This is the case for Lake Vanda, in Wright Valley (Wilson 1964), and Lake Bonney (77º71’S, 162º41’E) (Hendy et al. 1979; Matsubaya et al. 1979) and Lake Fryxell (Lyons et al. 1998a) in Taylor Valley. There is now ample evidence to suggest that these lakes have undergone numerous water-level fluctuations in response to subtle climate changes at least throughout the Holocene (Friedman et al. 1995; Hall et al. 2001), and perhaps as far back as 300 Ka BP (Hendy 2000b). By using simple diffusion models of conservative solute species such as Cl- coupled with stable isotope measurements of į18O and įD in lake waters, the timing of the beginning of the last fill event (or termination of the last desiccation event) can be determined. These model calculations indicated that for Lake Vanda and Lake Fryxell the filling of the lakes began ca. 1000 years BP (Wilson 1964; Lyons et al. 1998a). This time is generally associated with numerous other warmer climatic proxies all along the Victoria Land coast line (e.g., Lyons et al. 1998a). In addition, Lake Wilson at 80ºS in southern Victoria Land shows a similar timing of last refill, beginning ca. 1000 years ago (Webster et al. 1996). Thus, it appears that the climatic change that occurred at this time covered the entire length of the Victoria Land coast to at least 80ºS. Matsubaya et al. (1979) estimated that the drawdown event in the east lobe of Lake Bonney ended at 2600 years BP. Contrary to the work of Hendy et al. (1977), they suggested that the west lobe of Lake Bonney is only ca. 6000 years old, and represents the dissolution of previously deposited salt with Taylor Glacier melt. Conversely, Lake Hoare, another major lake in Taylor Valley, demonstrates no signs of previous cryo-concentration based on both total dissolved solids and stable isotopic evidence. Lyons et al. (1998a) speculated that Lake Hoare was a relatively new feature, possibly produced as the last refill event began ca. 1000 years ago. As the climate warmed and the Canada Glacier advanced, it separated the Lake Hoare basin from the Lake Fryxell basin to the east (Lyons et al. 2000). Consequently, both 36Cl and į37Cl profiles have supported the notion that Lake Hoare is indeed “young” and the chemistry of the lake is due solely to modern glacier melt input (Lyons et al. 1998b, 1999). This modern age also compares very well to sediment analysis suggesting the bottom of Lake Hoare was subaerially exposed prior to ca. 1000 years BP (Doran et al. 1999). The use of density or chemical stratification of antarctic lakes as paleoclimate indicators, in a more general sense, has been described eloquently by Gibson and Burton (1996). The production of “paleoepilimnia” from changes in water balance in antarctic lakes from season-to-season and even year-to-year are important indicators of climatic variation in antarctic coastal regions from the Vestfold Hills at ca. 68º30’S to Lake Wilson at ca. 80ºS.

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Figure 6. Profiles of G3He in Lake Hoare and Lake Fryxell in the McMurdo Dry Valleys.

In some cases, more recent (< 50 yr) lake conditions can be recorded by variations in H and helium isotopes. Initial work by Hood et al. (1998) utilized this approach to determine both the short time scale physical dynamics and the longer term water balance variation in Lake Fryxell. Because 3H decays to 3He, 3H/3He ages of lake water can be measured as initially outlined by Torgersen et al. (1977). Hood et al. (1998) determined that the top 7 m of Lake Fryxell water was less than 25 years old during the 1993-1994 field season. More recent data obtained during the 1999-2000 field season estimated the top 5 m of Lake Fryxell to be less than 18 years, and the top 7 meters to be less than 50 years old. Prior to the collection of water by Hood et al. (1998), Lake Fryxell had been rising at a very rapid rate (Chinn 1993). However, in the ensuing six years between these two data sets, the lake level had actually decreased due to a substantial decrease in glacier meltwater inflow (Doran et al. 2002b). This decrease in lake-level actually allows for older, deeper water to come closer to the ice-water interface, thereby explaining our increase in 3H/3He ages compared to the earlier work of Hood et al. (1998). Helium isotopic measurements are also very useful in understanding past hydrologic events occurring within these lakes. In Figure 6 we have plotted the G3He values for 3

į3He = [(Rm/Ra)-1] x 1000

(1)

both Lake Fryxell and Lake Hoare where Rm is the measured 3He/4He in the sample and Ra is the value in water in equilibrium with the atmosphere. The Lake Fryxell profile is very similar to what previously had been published by Hood et al. (1998). The change of slope of the profile at ca. 9 m is similar to that of conservative solutes such as Cl-, and suggests a strong diffusional gradient of 4He from the sediment/water interface,

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as first suggested by Hood et al. (1998). The surface waters more closely resemble waters in equilibrium with the atmosphere with additional input of tritiogenic 3He. Perennially ice-sealed lakes The perennially ice-covered lakes discussed above (i.e., lakes whose permanent ice covers overlie a well established liquid water column) form seasonal moats and have porous ice covers that influence the hydrology, physical structure and biological activity of the underlying liquid water column. However, some antarctic lakes have sufficiently thick ice covers that summer meltwater flows over the permanent ice rather than under it, so that the lake ice grows from the surface up (while still growing at the bottom when the heat flux allows it) and accumulates over time if meltwater overflow exceeds ablation. Previous authors believed these lakes were frozen to their beds and referred to them as ice block lakes (e.g., Chinn 1993). Shallower lakes are likely frozen to their beds, and such a lake (5.45 m deep) has been reported in the Canadian High Arctic (Blake 1989). Using ground-penetrating radar (GPR), at least two of these so-called ice-

Figure 7 . Physical and chemical properties of a Lake Vida ice core extracted in October 1996. Black horizons on the stratigraphy plot represent sediment layers, gray horizons are sandy ice, and vertically banded horizons contain microbial mats. Profiles are shown for chloride (Cl-), dissolved inorganic carbon (DIC), dissolved organic carbon (DOC), chlorophyll-a (Chl-a), and primary productivity (PPR). The temperature profile plotted was taken at the time of the core extraction (from Doran et al. 2003).

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Figure 8. Photograph of ca. 1 m of water trapped on the surface of the 19 m Lake Vida ice cover during the anomalously warm 2001/2002 austral summer. Photo was taken by T. Nylen on 18 January 2002 from the Lake Vida meteorological station at the west end of the lake.

block lakes in the Dry Valleys show evidence of thick ice covers overlying saturated brine that is not allowed to communicate with the atmosphere. Lake House (77º07’S, 161º04’E) in Pearse Valley and Lake Vida (77º38’S, 161º95’E) in Victoria Valley have ice thicknesses of 11 m and 19 m, respectively. The origin, evolution, and maintenance of ice-sealed Lake Vida has been described by Doran et al. (2003). Because of the way perennially ice-sealed lakes grow over time, it is very likely that clastic sedimentation is essentially halted when the lake becomes hydrologically sealed from inflowing water. From this point until the lake becomes unsealed, we anticipate that the only sedimentation occurring is through precipitation of salts. Growing ice cover will cause salts to be deposited; thinning ice cover will cause salts to be dissolved. Therefore, we do not envision that lake bottom sediments during the sealed mode will be very useful repositories of paleoenvironmental information, with the exception of changes in salt chemistry. However, if cores of sufficient length can be obtained, the timing and extent of the shifts between sealed and unsealed modes should be detectable through parameters as simple as clastic sediment load. The ice cover in perennially ice-sealed lakes contains important paleoenvironmental information. Doran et al. (2003) identified numerous layers of sediment and microbial

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mat (Figure 7) throughout the top 16 m of the Lake Vida ice cover (see also Fritsen and Priscu 1998). These layers are believed to represent sediment deposition and associated microbial mat growth during exceptionally warm summers as turbid water collects on top of the ice cover, such as during the 2001/2002 austral summer (Figure 8). Sediment deposited during these warm summers can be trapped in the liquid water layer near the surface of the ice cover in a manner similar to what happens with windblown material in the thinner ice covers. These sediments and associated microbes then become trapped within the ice cover as the liquid water freezes in winter. We believe productivity is associated with this substrate not only during the open water events, but can also occur during colder years when liquid water does not accumulate on the surface. During the latter, sediment particles would absorb solar radiation forming liquid water inclusions, similar to those observed in thin-ice lakes (Fritsen et al. 1998; Priscu et al. 1998). Radiocarbon dating of organic sediment layers in the ice indicates that the lake has been in the sealed mode for at least 2800 14C yrs BP. Microbial mats throughout the ice cover are viable when thawed in the laboratory (Fritsen and Priscu 1998; Doran et al. 2003). Subglacial lakes Water may pond below glaciers on various scales ranging from mm to km in area. Larger accumulations of water can form permanent subglacial lakes: (1) in deep glacier bed depressions, commonly below ice domes, where the depression must be deeper than the surfaces of equipotential contours above it; and (2) if the glacier bed is flat, but where a basin on the ice surface is surrounded by a ring of thicker ice and subglacial water becomes trapped below the basin area due to equipotential gradients (Nye 1976). A further classic mechanism of subglacial lake formation is from high geothermal heat flow being concentrated in one subglacial area causing local basal melting, such as in Iceland today at subglacial volcanic vents (e.g., Björnsson 1975). Although the unequivocal presence of subglacial lake deposits in the rock record is rare, inferences of their existence under paleo-ice sheets is commonly invoked to explain subglacially produced forms that are thought to have required large volumes of water to be released catastrophically. In Antarctica there are extensive dendritic channel systems that form the Labyrinth in Upper Wright Valley (77º55’S, 160º83’E), which have been suggested to have formed under more extensive outlet glaciers during such a catastrophic release of subglacial lake water (Sugden et al. 1991; Denton et al. 1993). In addition to the fossil lakes inferred to have caused the Labyrinth, there are at least 77 existing subglacial lakes beneath the East Antarctic Ice Sheet (EAIS) (Siegert et al. 1996). Airborne radar mapping in the 1970s first identified the presence of lakes beneath the EAIS (Siegert et al. 1996); however, more recent mapping has identified at least another 15 lakes in the region of Dome C, some of which may exceed 40 km in length (Tabacco et al. 2002). Lakes in the region of Dome C seem to exist in what may be considered a “Lake District” that may be connected hydraulically. Lakes larger than about 5 km in the longest dimension generally have a surface slope of < 0.5 m km-1 indicative of ice floating on liquid making them identifiable from satellite images. However, care must be taken with interpreting every low slope area as a lake until the depth of subglacial water is known (e.g., Tikku et al. 2001). Subglacial lakes remain liquid as a result of the

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melting point depression caused by the weight of the overlying ice, insulating ice cover, and heat flux from the Earth (Kapitsa et al. 1996; Siegert and Kwok 2000; Siegert et al. 2001). Lake Vostok (78º05’S, 107ºE) (Figure 9), by far the largest of these lakes, is about 280 km long and 50 km wide and > 1000 m deep at its deepest point. The lake lies under 3750 m of ice on its south end and 4150 m on the north end within the East Antarctic Precambrian craton. The difference in ice thickness can produce barotropic flow in the lake (Wüest and Carmack 2000) and differential regions where lake water melts and freezes (accretes) to the bottom of the overlying ice sheet (Siegert et al. 2001).

a

b

Flow

Cored 3623 m

Ice Sheet Flow Accretion Ice

Lake

Vostok Station Sediment

Figure 9. (a) RADARSAT image of the glacier ice surface in the region of subglacial Lake Vostok in eastern Antarctica. The flat floating ice surface can clearly be seen which outlines the lake area ca. 3.5 km below. The lake is 280 km long and 50-60 km wide. Arrows mark general direction of glacial flow at the surface (from Siegert et al. 2001). (b) cross-section of the ice sheet at Lake Vostok showing the general configuration of glacier ice, accretion ice, bedrock, and lake. The ice melts as it flows into the lake from the left, and refreezes forming the accretion ice to the right. The depth of the water column beneath the Vostok borehole is 670 m and average thickness of sediments is thought to be ca. 300 m.

Tectonic setting of East Antarctic subglacial lakes Most of the EAIS subglacial lakes appear to be deep basins associated with ice divides. However, major questions remain as to the geological origin of the deep basins beneath regions near the ice divides. The most likely geological alternatives are (e.g., Dalziel

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1998): (1) intracratonic rifts associated with extensional processes, (2) rifts from continental collision, (3) hot spot or mantle plume-driven depressions, (4) glacial scour eroding older feature, or (5) meteor impacts. Lake Vostok lies within the Vostok subglacial highlands, which are connected to the Gamburtsev Subglacial Mountains; the origin of both the mountains and lake is currently uncertain. The most recent interpretation from new airborne geophysical surveys is one of a reactivated suture zone where a thrust block overlies an old continental margin (Studinger et al. 2001). The lake is a younger, smaller extensional basin set in a sedimentary succession overlying the thrust block, thought to have been created during a later, weak reactivation phase in the tectonics. Based on helium (He) isotope ratios in accretion ice, present-day tectonic activity within the rift basin seems unlikely (Jean-Baptiste et al. 2001), though recent seismic activity in the area has been recorded (R. Bell, pers. comm., 2001). It should be noted that the He isotope evidence is from accreted ice, which reflects only the upper portion of the lake. Deeper waters may have a different He isotope signature. Inferred lake processes Physical processes Some inferences about the total physical character of Lake Vostok have been made from geophysical remotely sensed data, refrozen lake water recovered from ice cores and from modeling. However comprehensive process modeling may only be done after the documentation of the physical characteristics of the lake upon lake entry. Freezing is thought to be occurring across the entire southern section of the lake and the accreted ice is being exported from the lake at the base of the ice sheet (Bell et al. 2001). Export of accreted ice is used to indicate that fresh water must be replenishing the lake from an upstream subglacial water catchment. The residence time of the water in the lake has been estimated to range from 5000 to 125,000 years (Kapitsa et al. 1996; Jean-Baptiste et al. 2001). Mayer and Siegert (2000) have suggested a residence time of about 100,000 years based on estimates of annual meltwater mixing within Lake Vostok. The lowest value was based on helium isotopes measured in lake water accreted onto the bottom of the ice sheet (Jean-Baptiste et al. 2001). However, the ice measured in this study may have accreted in a shallow embayment isolated from the main lake basin (Bell et al. 2001) and probably underestimates the actual residence time. Isotope records, crystal growth rates and ice flow modeling suggest that the basal glacier ice overlying Lake Vostok could be as old as 1,000,000 years (Siegert et al. 2001), marking the maximum possible age of the youngest lake water. The ice above Lake Vostok has been cored to a record depth of 3623 m, stopping ca. 120 m above the surface of the lake. The upper 3500 m of glacial ice represents a 420,000 year environmental record covering four complete ice age climate cycles; ice below 3500 m is thought to represent refrozen lake water accreted to the bottom of the glacial ice (Jouzel et al. 1999; Petit et al. 1999; Priscu et al. 1999; Siegert et al. 2000). Chemical and biological processes Abyzov et al. (1998) provided evidence that a wide range of microbes (bacteria, yeasts, fungi and microalgae) exist in the Vostok glacial ice at depths ranging from 1500 to

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2750 m. The concentration of microbes, some viable, was correlated with the density of mineral particles in the ice, implying that they were deposited in the snow mainly during glacial periods when the flux of dust and the wind speed were greatest. Recent studies of accretion ice from 3590 m (Priscu et al. 1999) and 3603 m (Karl et al. 1999) have shown the presence of microbes at densities of 102 to 104 cells ml-1, and the development of metabolic activity in the presence of liquid water. Bacterial 16S rDNA revealed low diversity in the gene population (Priscu et al. 1999; Christner et al. 2001). The phylotypes were closely related to extant members of the alpha- and betaProteobacteria and the Actinomycetes. Actinomycetes were also observed in overlying glacial ice (Abyzov et al. 1998) implying that the biological seed for the accretion ice (and presumably for Lake Vostok) may have arisen from airborne particulates (e.g., Marshall 1996) deposited on the surface of the ice sheet ca. 500,000 years BP. The microbes (Figure 10) were released into the lake and subsequently refrozen to the overlying glacial ice following downward migration and melt at the glacial grounding point. Alternatively, the microbes could be remnants from an ancient Lake Vostok that existed before permanent glacial ice cover (ca. 15 million years ago). Importantly, the microbes identified within the accreted ice represent assemblages that exist near the ice/water interface and may have recently been released from the overlying glacial ice. Presumably, a completely different assemblage exists in the deeper waters and in association with the sediments, particularly if the water column is vertically stratified (Wüest and Carmack 2000; Siegert et al. 2001). Because the Lake Vostok ecosystem receives no solar radiation, all biochemical activity within the lake must depend upon chemically mediated oxidation-reduction reactions. The supply of oxidants and reductants are presumably derived from the overlying ice sheet and may result from geothermal activity, although the latter has yet to be confirmed (Jean-Baptiste et al. 2001). The geochemistry of Lake Vostok was estimated from the overlying accretion ice by Priscu et al. (1999) and Souchez et al. (2000). These estimates indicate that the waters of Lake Vostok should contain adequate quantities of dissolved organic carbon, anions and cations to provide a system that is thermodynamically favourable to support life. Lake sediment record Although planning is underway, Lake Vostok has yet to be penetrated. Many uncorroborated ideas exist concerning the history of Lake Vostok. Perhaps the best evidence of the paleolimnology of Lake Vostok can be found in the sediments. Russian seismic data have been used to estimate that the sediment succession in the lake is approximately 300 m thick near Vostok Station (Popkov et al. 1998). The lake sediments could contain an unparalleled record of antarctic paleoenvironmental information, extending well beyond the limit of ice core records. In fact, meltwater and sediment currently entering the lake comes from the base of the ice sheet so, in essence, the sediment record should pick up where the ice cores leave off. To understand the possibilities of what the lake may contain, a brief evaluation of what is currently known about the ice sheet is useful. Glaciological models of the development of the antarctic ice sheet most commonly have ice building on highlands and progressively expanding into larger ice masses until a full continental-

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Figure 10 . Atomic Force micrograph (a) and scanning electron microscope images (b-f) of particulate matter from accretion ice located 3590 metres below surface (mbs) (a-d) and from the overlying glacial ice (e = 1577 mbs; f = 763 mbs). Accretion ice images show bacterial cells (arrows) associated with organic debris (a-d); glacial ice images show diatom fragments and associated debris (e, f).

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encompassing ice sheet forms (e.g., Huybrechts 1993; Budd et al. 1994; DeConto and Pollard 2001). The glaciological models are consistent with distal proxy records of oxygen isotopes (Zachos et al. 2001) and Mg/Ca ratios (Lear et al. 2001) derived from deep sea foraminifera. These proxy records of inferred ocean temperatures and ice volumes indicate that, prior to about 38 Myr ago, glaciers were either on mountains or highlands, or that small intermittent ice sheets grew on the continent. Antarctica then entered a phase of larger, more permanent ice sheets, although their volumes still changed temporally and they probably reached their maximum size in Pleistocene times. These scenarios are also in agreement with those inferred from direct glacial deposits in long geological drill cores through glacio-marine successions in Prydz Bay (Hambrey et al. 1991; O’Brien et al. 2000) and McMurdo Sound (Hambrey et al. 1989; Powell et al. 2000, 2002). Coastal records indicate that early phases of glaciation were relatively ‘warm and wet’, with temperate glaciers (Powell et al. 2002). The drill core records also indicate that glaciation on Antarctica became progressively colder through Oligocene to Pleistocene times going through subpolar/polythermal to full polar glaciers (Powell et al. 2002). However, even the younger periods of glaciation appear to have had a range of glacial activity, especially at ice sheet margins, which may have influenced interior ice flow (cf., Powell 1981; Hambrey et al. 1989; Webb et al. 1996; Wilson et al. 1998; Hambrey and McKelvey 2000; Miller and Mabin 2000). During periods of purported temperate ice (e.g., Hambrey and McKelvey 2000), glacial dynamics would have been altered and the interior ice sheet drawn down, dramatically changing drainage divides in some of the areas of the current subglacial lakes. It has also been shown recently that the Miocene ice sheets were being driven by Milankovich forcing in a similar way to Pleistocene glaciers (Naish et al. 2001), adding evidence to the likelihood of changes in mass balance and dynamics through time. The sedimentary record contained in subglacial lakes may contain evidence that can be used to decipher early basin histories. If the basins formed before the formation of the ice sheet, then their sediment fill may contain continental interior syntectonic sediments which may help define the timing and style of tectonism by lake and/or volcanic deposits. However, given the type of history just described, it is likely that glaciers on interior highlands early in their history were temperate and very dynamic; the type of glaciers that flow down valleys and erode low areas, such as lake basins. The Lake Vostok basin, if it existed at the time, may well have been eroded over several to tens of millions of years of dynamic glaciation. Consequently, the likelihood of having a pre-glacial record in the lake basins is probably quite low, unless the sediments were somehow protected. One mechanism for sediment protection may be the formation of a perennially ice-sealed lake prior to glacial over-ride. As the local snowline descended to near lake-level, an ice-sealed lake could have formed providing the protection of a thick lake ice cover. Whether such conditions could have existed during the temperate transitional phase from the prior “greenhouse” period is unknown. There may well be records of important ice sheet and climatic fluctuations through Lake Vostok’s history retained in the sediment record. These records would be of the interior of the ice sheet rather than of its edges, as have been recovered thus far from other sources. A complication of inferring paleoclimate from Lake Vostok sediments is that any atmospheric input to the lake is delayed by hundreds of thousands of years, and possibly up to 1,000,000 years which is the expected maximum age of basal ice (Siegert

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et al. 2001). Furthermore, the length of the delay will vary with internal ice sheet dynamics. Sedimentation will be a mix of atmospherically deposited, glacier-scoured, and authigenic sediments, all from processes on very different time scales. Separating out the signals will be difficult, but certainly performing particle-size specific analysis will play a key role as we would expect the atmospheric signal to be carried into the lake through gases and fine-grained particles. It is tantalizing to contemplate what those 300 m of sediment in Lake Vostok basin could contain, but will remain speculation until it is recovered. In addition to the climatic and geological history, the record contained in the Vostok basin sediments may also reveal information on past geochemical processes, macrobiological and microbiological communities and paleoclimate. The sediments could also contain a record of extra-terrestrial material capture. Micrometeorites within Vostok accretion ice have been observed (Priscu and Mogk, unpublished data) making it likely that the sediments contain a large number of meteorites, micrometeorites and cosmic dust (e.g., interplanetary dust particles and cometary debris). Hence, sediments in the lake basin should offer an extraordinary opportunity to measure extra-terrestrial flux over millions of years. The previously unrecognized 100,000 year periodicity in the Earth’s orbital inclination has been suggested to influence the accretion rate of extraterrestrial material. Assuming that the Earth’s orbital inclination is related to the climate record, measurements of extra-terrestrial material in the Vostok sediments could provide an unparalleled record of climate change that could only be provided in this unique setting. It is interesting to speculate about the accumulation of clathrates (air hydrates) in the sediments of Lake Vostok. Clathrates are preserved in the glacier ice entering the lake and should be stable as the ice passes the pressure melting point. Air hydrate has a equilibrium density of between 0.980 to 1.025 g cm-3 (Uchida and Hondoh 2000) and Lake Vostok is estimated to have an average water density of 1.016 ± 0.001 g cm-3 (Wüest and Carmack 2000). Therefore we might expect some hydrates to float and some to sink. Lipenkov and Istomin (2001) point out that the proposed circulation in Lake Vostok would be enough to keep hydrates with a density of 1.050 g cm-3 and up to 200 Pm in diameter in suspension in the water column. Nevertheless, the fact that it is possible to have air hydrate heavier than the surrounding water holds out the potential to have some fraction of air hydrate accumulate in the sediments and provide a sedimentary record of atmospheric gas. We would also expect natural gas hydrates to accumulate in situ if conditions are appropriate. The current location of the Vostok borehole (the hole left after extraction of the Vostok ice core) is not ideal for limnological and paleolimnological studies. The borehole is very close to the southeastern edge of the lake. If sampling is restricted to one hole, it should be over the deepest part of the lake, as is standard practice in paleolimnology. The cross-section in Figure 9 is based on recent sediment profiling and suggests that sediment redistribution (slumping) has occurred over time. This information can also be used to choose the best sampling location with respect to sediments. The closer the sampling is to the west side of the lake, the more of a subglacial melt/scour record will be obtained, whereas moving towards the east may provide more of a “whole lake” signal.

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Exopaleolimnology Most available data implies that water in its liquid form is currently rare in our solar system, with the only confirmed reserves residing on Earth. There are indications that liquid water may be present beneath thick ice covers of the Jovian moon Europa (Kargel et al. 2000; Kivelson et al. 2000) and there is considerable evidence that Mars may have had liquid water at or near the surface throughout its climatic history (Malin and Edgett 2000a,b; Baker 2001). Operationally, the search for life beyond Earth can be regarded as a search for liquid water, since all life as we know it depends on liquid water. It would be of enormous interest to both the exobiological and paleolimnological communities to retrieve samples from old lacustrine deposits on Mars. The sedimentary environments of ancient fluvial features on Mars may provide the best opportunity for discovering the first evidence of life beyond Earth. An earlier review on the topic of paleolakes on Mars was provided by Wharton et al. (1995). Evidence of fluvial features on Mars The Mariner 9 orbiter provided the first indications that the Martian surface had been modified by water, unveiling what appeared to be fluvial features (Figure 11a). Perhaps with these initial observations of aquatic environments on another planet, the discipline of exopaleolimnology was born. The case for water on Mars, past and present, has been building since then through a series of orbiters and landers. Under the present climate regime, liquid water is not stable at the surface because of the low atmospheric pressure and cold temperatures. Features attributable to periglacial processes suggest that water may now be hidden beneath the surface as deposits of massive ground ice. The observation that fluvial features on Mars were localized, possibly restricted to regions of geothermal activity, and the difficulty of constructing self-consistent CO2 greenhouse models for Mars has led to the theory that early Mars, although comparatively warmer and wetter than today, was quite cold and that the fluvial features formed in association with a cold climate regime. There are two major classes of fluvial features on Mars: valley networks and outflow channels. Valley networks Valley networks (Figure 11b) often resemble terrestrial runoff channels. Estimates by Baker et al. (1992) place the formation age of most valley systems at 3.8-3.9 Gyr. There are a few examples of younger valley networks residing in what may be Hesperian or Amazonian units (for definition of Martian epochs see Doran et al. in press). The most notable are located on the flanks of the volcano Alba Patera which provide important paleoclimatic clues to more recent fluvial activity (Gulick and Baker 1989). Most common are valleys that are short, with steep gullies on the slopes of large craters in the ancient highlands. The valley walls are usually steep talus slopes and the gullies often terminate abruptly at the flat crater floors. Another characteristic common to most valley networks is that tributaries often have blunt, theatre-headed terminations. The formation of the valleys has generated a great deal of debate with most arguments focusing on the role of fluvial erosion versus other processes such as liquid CO2, ice,

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Figure 11. Mars images showing evidence of water activity. (a) Mariner 9 mosaic of about 1/3 (ca. 120 km) of the Amazonis-Memnonia channel centred at 7ºS, 151ºW. This mosaic (Mariner 9 revolution 458, pictures 12499650, 12499650, 12499790) was among the first published evidence of past water on Mars (Sagan et al. 1973). (b-d) Viking images of (b) Dense drainage network in the Southern Highlands. The image is about 250 km across and centred at 48ºS, 98ºW (Viking image 63A09). (c) Outflow channel emerging from chaotic terrain. The image is about 140 km across and centred at 1ºS, 43ºW (Viking image P-16983). (d) "Islands" near Chryse Planitia. Image is centred at 21ºN, 31ºW (Viking image 211-4987).

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lava, the role of groundwater sapping versus surface runoff, and climatological constraints. Nanedi Vallis, an 800 km long valley with only a few short tributaries, is a good example of a valley that appears to have been cut by slow erosion of running water. Its sinuosity suggests it was not a flood channel but the lack of tributaries implies that it was fed largely by groundwater rather than runoff. Aharonson et al. (2002) suggest that this feature and many other networks and outflow channels on Mars were formed by groundwater sapping. Outflow Channels Outflow channels are enormous features generally attributed to catastrophic flooding initiating from subsurface sources (Baker 2001). They can be over 150 km wide and over 2000 km long. While their initiation points are fairly distinct, they often just fade into obscurity at their downstream ends. Source regions are associated with what has been termed “chaotic terrain” (Figure 11c), a complex topography formed as the result of the removal of subsurface material and widespread collapse of topography. Typical bedforms of outflow channels include longitudinal grooves, teardrop-shaped islands (Figure 11d) and horseshoe-shaped escarpments. Outflow channels on Mars have morphological features very similar to those of the Channeled Scablands of eastern Washington State in the U.S (Baker 2001). On Earth, approximately 18,000 years ago, Lake Missoula (47º05’N, 114º05’W), a large ice-dammed lake (2x1012 m3 and up to 600 m in depth), burst through its dam releasing a tremendous volume of water (up to 120 m deep) down the regional slopes of the Columbia Plateau. Peak drainage is estimated to have been ca. 107 m3 s-1 and complete drainage occurred within a few days leaving the surface heavily scarred with large channels and other morphologic features similar to those found associated with the outflow channels on Mars. The episodic catastrophic floods on Mars probably had discharge rates that may have ranged as much as 100 times higher than the largest flood events that occurred on Earth. According to Carr (1996) most of these events seem to have been caused by the sudden release of groundwater under high artesian pressure trapped below thick permafrost. Several plausible mechanisms were proposed by Carr (1996) to account for the sudden discharge of water to the surface including volcanic activity, faulting and meteoritic impacts. The collapsed remains of the surface at the initiation site, forming the characteristic chaotic terrain, are testament to the magnitude of these voluminous discharge events. Because of their enormous discharge volumes and concomitant latent heat energy, they could, even under current climatic conditions, have flown for vast distances before being halted after boiling off to form ice crystals and water vapour. For this reason, they do not provide a great deal of useful climatic information. Nevertheless, by estimating the volume of water necessary for producing the erosional features that are observed, a lower limit may be placed on the total inventory of water on the planet. Carr (1996) reports that this lower limit is equivalent to 40 m of water spread over the whole planet. These megafloods, which most certainly would have had an impact on the martian climate, occurred mainly in the Hesparian and Hesparian/Amazonian and possibly even in recent (10 Myr) history (Burr et al. 2002).

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Springs Emergent groundwater and perhaps glaciers (Lucchitta 1982) appear to have played an important role in shaping the landscape throughout Martian history. Recent springs may have fed streams that formed the valley networks. Images obtained by the Mars Global Surveyor (MGS) spacecraft of small-scale debris flow gullies suggest spring activity during the last several million years (Malin and Edgett 2000a). Hydrothermal convection driven by magmatic activity or impact melt may have provided a mechanism of replenishing water in a Martian aquifer. It is likely that springs have formed early in Mars’ history as a result of volcanism and meteoritic impacts. Dissolved salts are likely to be present and may enhance the persistence of liquid water environments by depressing the freezing point. On Earth, highly mineralized brines are found in subarctic Canadian Shield wells and in high latitude springs of the Canadian High Arctic (Pollard et al. 1999; Andersen et al. 2002). Brines flowing onto the surface form large icings and deposit salts. The Thermal Emission Spectrometer (TES) instrument on the Mars Global Surveyor (MGS) mission has discovered an accumulation of crystalline hematite in a sedimentary rock formation that covers an area approximately 350 by 350-750 km in the Sinus Meridiani region. One possibility is that the hematite formed by precipitation from aqueous fluids under either ambient or hydrothermal conditions. The deposits, which may have been exposed by wind, provide additional evidence for long-term stability of near-surface water on early Mars (Christiansen et al. 2001). Given the nature of the site, it would not be unexpected to find other mineral precipitates as well. Sites such as this will most assuredly be targeted for exploration as part of the search for evidence of past life. Evidence for standing water On Mars, three main types of lakes have been identified – those that formed in areas of convergent drainage by valley networks in the heavily cratered uplands; those that formed within the canyons; and those that formed at the terminus of large outflow events, most notably where the circum-Chryse and Elysium outflow channels terminate (Carr 1996). These putative lakes appear to have formed by a number of mechanisms including meteoritic impact, groundwater seepage and catastrophic flooding. Lakes associated with the valley networks in the ancient, heavily cratered uplands, may contain sedimentary records of the events that led to the evolution of life on Mars since they date to 3.8-4.0 Gyr BP. McKay and Stoker (1989) point out that, since the earliest record of life on Earth has been obscured by erosion and plate tectonics, the best repository of information about the origins of life may actually reside on Mars. The Valles Marineris canyon system may have been flooded with water throughout much of Mars’ history (Carr 1996). Box canyons such as Hebes Chasma appear to have thick sequences of fine grained sedimentary material with no visible source regions. One explanation for this is that the materials are carbonates that were deposited in standing bodies of water (Nedell et al. 1987; McKay and Nedell 1988). McKay and Davis (1991) attempted to calculate the time these lakes would exist using a climate model developed by Pollack et al. (1987). They concluded that, as long as a source of meltwater entered into the lake to offset ablation of ice at the surface, the lakes could have exist for several hundred million years or more. The same authors employed calculations that explain the

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Figure 12 . Lakes on Mars? (a) Mosaic of daytime thermal infrared images (obtained by Mars Odyssey) draped over altimetry data (obtained by Mars Global Surveyor). Whiter shades are warmer temperatures and are sun-lighting effects. An ancient river following the channel in the top left is purported to have flown into a crater lake. The arrow points to what are believed to be deltaic deposits left where the river collapsed at lake level (cf., Figure 2). This image mosaic covers an area approximately 180 kilometres on each side centred near 14ºS, 175ºE, looking toward the south. Image credit: NASA/JPL/Arizona State University, PIA04260. (b) Image obtained during Mars Global Surveyor (MGS) aerobraking manoeuvres on October 18, 1997 at 15:42 PST, centred at 5º05’S, 340º07’W shows a feature hypothesized to be a playa deposit (beyond the black arrow). (c) Image taken December 29, 1997 by MGS at 1:19 p.m. PST, centred at 65º01’S, 15º01’W showing evidence of groundwater seepage from a crater rim. It is unclear what the darker deposit at the base of the crater is, but it may have been an area of pooling and evaporite formation in association with the seeps. Images in (b) and (c) provided by Malin Space Science Systems/NASA.

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physics that allows similar perennially ice-covered lakes to flourish with microbial communities in the Antarctic. If water was episodically present on Mars, it may have been possible that life survived to more recent times. Newsom et al. (1996), using a similar approach to McKay and Davis (1991) and taking into consideration the effects of the impact melt, geothermal energy released from the uplifted impact basement and the latent heat of freezing, found that the formation of large ice-covered, impact crater lakes > 65km in diameter on Mars would persist for thousands of years even under contemporary climatic conditions. The estimate of the lifespan of lakes on Mars would be extended if the lakes shift into a perennially ice-sealed mode. Cabrol et al. (1998, 1999) have argued that impact craters such as Gusev (Figure 12a) and Gale may have provided an oasis for life from the Noachian-Hesparian boundary (2.5-3.8 Gyr BP) and lasted for up to 2 Gyr. Using high-resolution Viking images to locate possible lacustrine features, they propose that Gale experienced a number of aqueous environments that transitioned from earlier warmer and wetter, to cold and icecovered waters. In this case, the sedimentary record in these environments would be important to investigate for evidence of past life. Cabrol et al. (1999) identified 179 impact crater lakes using Viking images and attempted to document their distribution, types and ages. The main implication of the study is that the 179 lakes observed most likely represent a fraction of the total number of lakes in impact craters. With new high resolution images being returned by MGS, the case for past standing water bodies on Mars is building (e.g., Figures 12b and 12c). Although not definitively under the realm of paleolimnology, we should point out the related evidence of past oceans on the surface of Mars. Parker et al. (1989, 1993) identified two contacts near the southern boundary of the northern plains and interpreted them to be shoreline features of a previous polar ocean. Baker et al. (1991) proposed that the oceans formed as a result of volcanism which triggered catastrophic flooding releasing water into the topographically lower northern plains. Head et al. (1999) have tested the original hypothesis of Parker et al. (1989) by using high-resolution altimeter data from the Mars Orbital Laser Altimeter (MOLA). They have reported their measurements to be in agreement with the observed shorelines made previously by Parker et al. (1989). Additionally, they report that their measurements revealed two major basins within the plains and that previously mapped features associated with ground ice (polygonal cracking and lobate ejecta craters) show a high degree of correlation with these basins, indicating that water may have been present at these locations in previous times. While the MOLA data seem to support the ocean theory, Malin and Edgett (1999) reported they could not find images in support of shoreline features using the MGS camera at high resolution. It may be that the processes that normally form shorelines on Earth were not operating in a similar fashion on Mars. The lack of a large moon results in solar tides that are much weaker. If the ocean was icecovered, it is not clear how definitive the shoreline features would be.

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Summary In this chapter we have attempted to review the present state of knowledge concerning limnology and paleolimnology in extremely cold environments both on Earth and beyond. The main impact of colder temperatures is to reduce the liquid water volume in that ecosystem. For instance, we can follow the progression of a closed-basin lake as temperature decreases. The first phase of lake evolution with colder temperatures will be to have occasional summers where the ice cover remains on the lake. If the temperature continues to drop, a perennial ice cover with a hydrologically connected seasonal moat will form (i.e., there is communication in the summer between the surface stream flow and under ice water column). Further cooling will cause the ice cover to become sufficiently thick that summer melting can not create a hydrologically connected seasonal moat, so that all summer melt flows onto the surface of the ice cover, and accumulates in layers (i.e., there is now ice growth at the top and bottom of the permanent ice). When the lake shifts into this perennially ice-sealed mode, summer meltwater input is proportional to ice growth because all meltwater flowing onto the surface freezes. At any time in this evolution, if long-term summer meltwater input does not exceed annual ablation from the lake, the lake will diminish and eventually dry up. If we follow the resulting sedimentation through the above scenario, we will reduce clastic sedimentation in the lake over time, and eventually eliminate it. Increasing the thickness of a perennial ice cover decreases the ability of wind-blown sediments to reach the water column. When the lake shifts into the perennially ice-sealed mode, there may initially be some sediment trapped in the basal ice, but over time the sediment will be purged as the ice melts and refreezes due to the small changes in thermal gradient that occur. Eventually, all significant amounts of sediment will be exhausted and clastic sedimentation will cease. At this point, we could only expect to record chemical and biological sedimentation. So the best record of this type of shift in lake conditions may simply be the amount and character of clastic sedimentation over time. Certainly salt chemistry and thermodynamics become increasingly important to paleolimnologists as a lake passes through these stages. We have proposed that severe climatic deteriorations could shift a perennially icecovered lake into a perennially ice-sealed lake, and eventually the lakes would disappear if ablation exceeded meltwater input. Evidence of this sequence of events may be in the geological record in relation to severe glacial periods even in temperate regions (e.g., Hoffman et al. 1998). The formation of perennially ice-sealed lakes in periglacial regions may also provide a mechanism to protect pre-glacial lake sediments as glaciers advance. This may have been the case for a pre-glacial Lake Vostok. These extreme aquatic environments also provide interesting analogs for conditions on Mars in the distant and possibly even recent past. There is strong evidence of a lacustrine history on Mars, but we can not be certain of how warm the planet was in the past. We can, however, be reasonably certain that the planet was warm enough to have perennially ice-covered or ice-sealed lakes, as we have shown these to be cold endmember lakes. Hence, the study of the paleolimnology of these cold extreme environments may provide guidance of where to look and what to look for during future life detection missions to Mars (e.g., Doran et al. 1998). A similar argument can be made for the study of subglacial lakes like Lake Vostok, which may be analogs for subglacial lakes on Mars and the proposed global ice-covered ocean of Europa.

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Acknowledgements We thank John Gibson, Damien Gore, Ian Hawes, Dominic Hodgson and Johanna Laybourn-Parry for helping us “round out” Tables 1 and 2, and Chris McKay for advice on the stability of water on modern-day Mars. Warwick Vincent provided a very helpful review of the manuscript. Collection of previously unpublished data used in this manuscript was supported by NSF Grants OPP0096250 and OPP0085400, and NASA Exobiology Grants NAG5-9427 and NCC2-1222. References Abyzov S.S., Mitskevich I.N. and Poglazova M.N. 1998. Microflora of the deep glacier horizons of central Antarctica. Microbiology 67: 66-73. Adams E.E., Priscu J.C., Fritsen C.H., Smith S.R. and Brackman S.L. 1998. Permanent ice covers of the McMurdo Dry Valley Lakes, Antarctica: bubble formation and metamorphism. In: Priscu J.C. (ed.), Ecosystem Dynamics in a Polar Desert: The McMurdo Dry Valleys, Antarctica. Ant. Res. Series, AGU. 72: 281-296. Aharonson O., Zuber M.T., Rothman D.H., Schorghofer N. and Whipple K.X. 2002. Drainage basins and channel incision on Mars. Proc. Nat. Acad. Sci. 99: 1780-1783. Andersen D.T., Pollard W.H., McKay C.P. and Heldman J. 2002. Cold springs in permafrost on Earth and Mars. J. Geophys. Res. 107(E3): 10.1029/2000JE001436. Baker V.R. 2001. Water and the martian landscape. Science 412: 228-236. Baker V.R., Strom R.G., Gulick V.C., Kargel J.S., Komatsu G. and Kale V.S. 1991. Ancient oceans, ice sheets and the hydrological cycle on Mars. Nature 352: 589-594. Baker V.R., Carr M.H., Gulick V.C., Williams C.R. and Marley M.S. 1992. Channels and valley networks. In: Kieffer H.H., Snyder C.W. and Mathews M.S. (eds), Mars. University of Arizona Press, USA, pp. 493-522. Barnes D.F. 1960. An investigation of a perennially frozen lake. Air Force Surveys in Geophysics. Volume 129. 134 pp. Belzile C., Vincent W.F., Gibson J.A.E. and Van Hove P. 2001. Bio-optical characteristics of the snow, ice, and water column of a perennially ice-covered lake in the High Arctic. Can. J. Fish. Aquat. Sci. 58: 2405-2418. Bell R.A., Studinger M., Tikku A.A., Clarge G.K.C. and Gutner M.M. 2001. Freezing and opensystem water exchange in subglacial Lake Vostok. EOS (AGU Abstracts) 82(47): F529. Benchley J., Karl D., Jackson J., Delaney J. and Welch J. (eds) 1998. Life in Extreme Environments (LExEn) Workshop Report. Report of NSF Workshop held June 17-19, 1998, Washington, D.C. Online report: http://www2.ocean.washington.edu/lexen. Bishop J.L., Lougear A., Newton J., Doran P.T., Froeschl H., Körner W., Koeberl C. and Trautwein A.X. 2001. Mineralogical and geochemical analyses of Antarctic lake sediments: A study of reflectance and Mössbauer spectroscopy and C, N and S isotopes with applications for remote sensing on Mars. Geoch. Cosmoch. Acta 65: 2875-2897. Björnsson H. 1975. Subglacial water reservoirs, jökulhlaups and volcanic eruptions. Jökull 25: 1-12. Blake W. 1989. Inferences concerning climate change from a deeply frozen lake on Rundfjeld, Ellesmere Island, Arctic Canada. J. Paleolim. 2: 41-54. Bormann P. and Fritzsche D. 1995. The Schirmacher Oasis, Queen Maud Land, East Antarctica, and its Surroundings. Ergänzungsheft 289, Petermanns geographische Mitteilungen, Gotha: J. Perthes, 448 pp. Budd W. F., Jenssen D., Mavrakis E. and Coutts B. 1994. Modelling the Antarctic ice-sheet changes through time. Annals Glaciology 20: 291-297.

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16. EPILOGUE: PALEOLIMNOLOGICAL RESEARCH FROM ARCTIC AND ANTARCTIC REGIONS

REINHARD PIENITZ ([email protected]) Paleolimnology-Paleoecology Laboratory Centre d’études nordiques Département de Géographie Université Laval Québec, Québec G1K 7P4, Canada MARIANNE S.V. DOUGLAS ([email protected]) Paleoecological Assessment Laboratory Department of Geology University of Toronto Toronto, Ontario M5S 3B1, Canada and JOHN P. SMOL ([email protected]) Paleoecological Environmental Assessment and Research Laboratory Department of Biology Queen’s University Kingston, Ontario K7L 3N6, Canada

Key words: Arctic, Antarctica, Paleolimnology, Climate change, Environmental change, Lakes

Our understanding of long-term environmental change in arctic and antarctic lakes has advanced markedly over the last two decades. This research has been driven primarily by the growing realization that circumpolar regions are critical research areas for the study of global climatic and environmental change. Logistical improvements (e.g., increased availability of aircraft, research bases, enhanced communication technology, etc.) have also made polar research safer and more feasible. Nonetheless, high costs often preclude many research initiatives, and declines in government funding programs have resulted in the closure of some key monitoring stations that are essential for understanding many of the basic environmental, limnological, and hydrological

509 R. Pienitz, M.S.V. Douglas and J.P. Smol (eds), 2004. Long-term Environmental Change in Arctic and Antarctic Lakes. Springer. Printed in the Netherlands.

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processes operating in high latitude regions. As outlined in the preceding chapters, we have many challenges ahead of us, but we also have many exciting opportunities. Paleolimnologists occupy a highly respected research niche in the global change community, as sediments archive important records of environmental and climatic change. Lakes and ponds are abundant in many parts of the Arctic and Antarctica, and our ability to interpret paleoenvironmental information is increasing rapidly. These data are being incorporated, often seamlessly, into large, multi-disciplinary research initiatives. Lake sediments are providing, at a great diversity of spatial and temporal scales, important information that often cannot be gathered from other sources, and in other cases paleolimnological data complement interpretations from other natural archives, such as ice cores and marine sediments. Polar lake sediments have been used to address many important questions, but especially those dealing with the study of climatic change. However, there is a growing realization that other scientifically and socially relevant problems require long-term perspectives (e.g., local pollution sources, long-range transport of atmospheric contaminants, penetration of ultraviolet radiation, fluctuations in economically and ecologically important animal populations). The list of possible applications is growing steadily. Although considerable progress has been achieved in a relatively short time, the contributors to this book also identify a number of challenges. Arctic and antarctic regions cover vast areas, spanning many different types of landscapes and ecosystems. With some notable exceptions (e.g., Toolik Lake area, Alaska), we currently only have a rudimentary understanding of limnological processes for most polar regions. For example, it is widely acknowledged that climatic influences are complex and include a variety of direct and indirect effects on the physical, chemical, and biological characteristics and dynamics of lakes. As paleolimnological interpretations cannot be separated from current-day limnological processes, a better understanding of the structure and functioning of high latitude lake systems remains an important research goal. A common perception that continues to persist is that circumpolar lakes and ponds are relatively homogenous systems, and that their characteristics can easily be generalized. With the possible exception of concluding that these systems are characterized by relatively low temperatures and have long periods of ice cover, our view is that high latitudes contain lakes and ponds that equal the limnological diversity found in other lake regions. This complexity, however, is not a deterrent to paleolimnologists, but an important advantage. As the various lake and pond systems will respond differently to climate forcing, as well as other environmental changes, carefully chosen study sites and methodological approaches can address many important research questions and problems. The contributors to this volume have identified additional research needs. For example, much work remains on describing the taxonomy and ecological characteristics of bioindicators from polar regions. Many ecological studies thus far have relied on surface sediment calibration sets and then employed multivariate statistical techniques to interpret these complex data sets, building on the previous successes that these approaches have enjoyed in other geographic areas. However, in many high latitude lake sets, the measured limnological variables often provide only relatively weak environmental gradients, and there is little or no present-day ecological or physiological

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data to bolster interpretations. Moreover, benthic taxa often dominate assemblages, thus posing additional problems with, for example, relating pelagic water chemistry measurements to periphytic algae. Calibration sets are clearly important, and should be pursued. But our view is that paleolimnologists should, at times, re-focus their efforts on understanding and interpreting overall community changes, preferably from different taxonomic groups and trophic levels, coupled with physical, sedimentological, and geochemical proxy data, to obtain a more holistic understanding of how overall lake processes and biotic communities may have changed in the past. Another major issue that was echoed by most contributors is the problem of developing sufficiently accurate geochronological control for high latitude sediment cores. For example, 210Pb inventories are usually much lower than those recorded in temperate and tropical lake systems. In addition, the paucity of terrestrial plant remains in many high polar regions, coupled with the slow degradation of terrestrial litter in polar landscapes, imposes additional limitations with standard AMS radiocarbon dating techniques. As noted by Wolfe et al. (this volume), other approaches are available and emerging and this is an active area of research. Moreover, the growing number of studies describing varved lake sediment profiles from polar lakes may alleviate some of these concerns, at least in some regions. As shown by the contributions to this volume, there are many similarities between arctic and antarctic regions, but also some striking differences. With few exceptions, most researchers work in either the northern or southern hemisphere. As such, there is much to learn from our colleagues working at the opposite poles, yet surprisingly there is relatively little dialogue between these two groups. Hopefully this book will stimulate a more active exchange of information and expertise between these two research communities. The publication of this book may be especially timely, as preparations for the celebration of the International Polar Year 2007-2008 are currently underway. This event marks the 125th anniversary of the First International Polar Year (1882-1883), and the 75th anniversary of the Second Polar Year (1932-1933). The previous International Polar Years catalysed many research initiatives, resulting in significant new insights. Paleolimnology was still in its infancy during the previous polar years. This is no longer the case. Never before has there been a greater need for understanding long-term changes in high latitude lake ecosystems. We are confident that paleolimnologists have the skills, approaches, and drive to meet these new challenges. No one ever claimed that this work was going to be easy or simple; we do know that this work will be important and rewarding.

513 Glossary, Acronyms and Abbreviations Į, ȕ and Ȗ particles: Energized subatomic particles emitted during radionuclide decay. 14

C: Carbon-14; radiocarbon. See Radiocarbon dating and AMS 14C dating.

137

Cs: Cesium-137 (also spelled Caesium-137); a short-lived radioisotope (half-life: 30.1 years) of anthropogenic origin that provides useful dating horizons through correlation with the history of thermonuclear weapons testing, including a major peak at 1963.

210

Pb: Lead-210; a natural occurring isotope (half-life: 22.3 years) formed in the atmosphere from Radon-222 decay and widely used for dating sediments spanning the last ca. 150 years.

239+240

Pu: Isotopes of plutonium associated with thermonuclear weapons testing and thus having stratigraphic distributions much like Cesium-137.

A.D.: Anno Domino. AABW: See Antarctic Bottom Water. Accelerator mass spectrometry (AMS): Involves the addition of a charged particle accelerator to the conventional equipment used in mass spectrometry (the separation and counting of atoms from a sample according to their mass), resulting in increased sensitivity of the system. In radiocarbon dating (counting of the radioactive 14C isotope), AMS permits dating of milligram-size samples. Active layer: The surface zone of permafrost that undergoes seasonal thaw. The depth of thaw depends on ambient climate conditions, vegetation and snow cover, and drainage, and may vary from tens of centimetres to several metres. Akinetes: Asexual spores (resting cells) which form through thickening of the walls of a normal vegetative cell. For example, the resting stage of cyanobacteria (blue-green algae). Alanine blanks: Commonly used standard for assessing the sensitivity of AMS 14C to trace levels of radiocarbon contamination. Albedo: The fraction of incident electromagnetic radiation reflected by a surface. For example, a landscape covered in snow would have a higher albedo than one of dark ground.

514 Algae: Plant-like organisms of any of several phyla, divisions, or classes of aquatic, usually chlorophyll-containing, nonvascular organisms of polyphyletic origin. They usually include the green, yellow-green, brown, and red algae in the eukaryotes and the blue-green algae in the prokaryotes. Allerød interstadial: A time period ca. 13,150 to 12,650 yr BP near the end of the last ice age. It was characterized by a relatively warm climate in the Northern Hemisphere. Allochthonous: Materials that are derived from outside a body of water and transported to lake sediments by wind, erosion, or fluvial transport. For example, terrestrial macrofossils (cf. autochthonous). AMAP: See Arctic Monitoring and Assessment Program. Amino acid racemization: The time- and temperature-dependent conversion of amino acids from their L to D isomeric configurations, which can potentially be used to date chemically-closed systems including certain protein-rich organic matrices. AMS 14C dating: Method of absolute age determination based on the radioactive decay of the carbon-14 isotope in organic material; method can use accelerator mass spectrometry, which allows the dating of 1 to 2 mg of carbon. Method can cover the last ca. 50,000 years. AMS: Accelerator mass spectrometry. Anadromous life cycle: A life cycle exploited by some animals (e.g., salmon) that includes both freshwater and marine environments. For example, young sockeye salmon hatch in nursery lakes. At a certain stage in development (e.g., 1 to 3 years for sockeye salmon), they migrate to the ocean environment where they mature to adults. They then must ascend rivers and streams to return to their nursery lake to breed, spawn and die. Anoxic sediments: Sediment devoid of oxygen. Antarctic Bottom Water (AABW): Sea water that sinks to the ocean floor off Antarctica and flows towards the equator beneath the North Atlantic deep water. Arctic Archipelago: A group of over 50 islands lying north of the Canadian mainland with an area of approximately 1.42 x 106 km2, and divided between the Northwest Territories and Nunavut. The archipelago includes the Queen Elizabeth Islands (north of Parry Channel) and the major southern islands of Baffin (the largest island in the archipelago), Victoria, Banks, Prince of Wales, and Somerset. Arctic Front: Southern boundary of the arctic air mass.

515 Arctic haze: Haze in arctic regions which reduces horizontal and slant visibility and which may extend to a height of about 10 km. It appears blue-grey when viewed away from the Sun, and reddish-brown toward it. Term introduced by Mitchell in 1956. Arctic Monitoring and Assessment Program (AMAP): A monitoring program under the Arctic Council of circumpolar countries, based in Oslo, Norway. Athalassic waters: Inland or continental saline water bodies (cf. thalassic). Atmospheric fallout: Transfer of chemical compounds and particulate matter from the atmosphere to land and water surfaces. Includes both dry and wet deposition, the latter being more important in most instances. Autecology: Division of ecology that deals with the relations of a single species to its environment. Autochthonous: Materials originating from within a given body of water and eventually accumulating in sediments. Includes both mineral precipitates and remains of aquatic organisms, such as diatom valves (cf. allochthonous). Barotropic flow: A fluid flow initiated by density differences that are due to differences in pressure. Bathymetry: Water depths of the sea floor shown on maps (bathymetric charts) by means of contour lines (isobaths). BC: Before Christ. Benthic: Pertaining to the region at or near the bottom of a lake or sea. Benthos: The biological community living in or on aquatic sediments. Biodegradation: The breakdown of organic chemicals by biological processes, e.g., microbial action under oxidative or reducing conditions in sediment. Biogenic Silica: Amorphous silica (SiO2·nH2O) produced by some living organisms, e.g., diatoms, radiolarians, silico-flagellates, freshwater sponges. Bioturbation: Disturbance of bottom sediments by the feeding and burrowing of benthic organisms. Blue-green algae: See Cyanobacteria.

516 Bottom-freezing: Influence of ice on the integrity of sediments, either in the littoral zones of deep lakes where ice is seasonally anchored, or more widely in shallow lakes (ponds) less than ca. 2 m deep. Bow wave: See Hydraulic “bow wave”. BP: Before Present (1950 AD). Brine exsolution: Ejection of salts from solution following the freezing of water. Bryozoa: Small, sessile, colonial, filter-feeding animals. Phylum Ectoprocta; also known in older literature as Polyzoa. BSEI: Backscatter scanning electron imaging. BSi: See Biogenic silica. CA: See Correspondence analysis. Cabbling: The process by which two (or more) water masses of different density combine to form a water mass of another density, especially where that increases the density of the water mass (e.g., water masses at 0 and 8ºC, respectively, combine to form water at 4ºC). Calibration training set: A dataset of micro- or macrofossil counts from modern surface sediments together with their associated water chemistry or other environmental data that are used to generate transfer functions (see transfer functions). Canadian Network for Isotopes in Precipitation (CNIP): Network of meteorological stations in Canada where precipitation is collected for analysis of G2H and G18O; a national network affiliated with the Global Network for Isotopes in Precipitation. Canonical correspondence analysis (CCA): A mathematical method for direct gradient analysis and ordinations that assumes that species have unimodal distributions along environmental gradients and removes the ‘arch-effect’ often found in principal components analysis (PCA) ordinations. CAPE (Circumpolar Arctic Paleoenvironments): Organization within IGBP-PAGES for promoting and coordinating arctic research. (Website: http://www.ngdc.noaa.gov/paleo/cape/cape.html) Carapace: Shell that covers the body of arthropods. Consists of chitin, and may be strengthened by calcium carbonate.

517 Carbon residence time effect: Also known as the radiocarbon reservoir effect. Radiocarbon samples which obtain their carbon from a different source (or reservoir) than atmospheric carbon may yield what is termed apparent ages. One of the most commonly referenced reservoir effects concerns the ocean. The average difference between a radiocarbon date of a terrestrial sample such as a tree, and a shell from the marine environment is about 400 radiocarbon years. Carbon-14: See Radiocarbon dating and AMS 14C dating. CCA: See Canonical correspondence analysis. CDOM: See Chromophoric dissolved organic matter. Cesium-137: See 137Cs. Charcoal: As used in paleoecology, charcoal is incompletely combusted wood or other plant material produced during a forest fire. Chironomid: A member of the Chironomidae or the non-biting midges. In paleolimnology, the chitinized head capsules of their aquatic larval stages are widely used, especially in paleoclimatic reconstructions. Chitin: A resistant organic compound similar in structure to cellulose containing repeating N-acetylglucosamine units, which forms the hard skeleton in many invertebrate and foraminiferal inner tests. Chlorobornanes: Components of toxaphene with a six member ring typically chlorinated at the 2,3,5 and 6 positions along with a bridged structure between the 4 and 7 positions on the ring. Chromophoric dissolved organic matter (CDOM): A fraction of organic matter composed of yellow humic and fulvic materials that are derived from terrestrial soils and vegetation or from microbial autochthonous production. Because of its strong ultraviolet (UV)-absorbing properties, CDOM forms the natural sunscreen in lakes and rivers against harmful UV radiation. Chrysophycean cysts: Siliceous stomatocysts).

resting

stages

of

chrysophyte

algae

(see

Chrysophytes: Common term for algae in the Chrysophyceae and Synurophyceae. Chydorid: Small littoral crustaceans belonging to the order Cladocera, family Chydoridae. Cladocera: Aquatic crustaceans, sometimes referred to as water fleas.

518 Clathrates: Also known as gas hydrates, these crystalline solids form when water molecules create a cage-like structure around smaller ‘guest molecules’ under high pressure and/or low temperature. CNIP: Canadian Network for Isotopes in Precipitation. Cold-based ice: Describes glaciers that are frozen to their bed. Deformation occurs in glacier ice above the bed, and may also occur within underlying sediment. Often attributed to areas of former glacier cover in which no erosional indicators can be discerned. Colluvial: Sediments derived from hillslope processes. Conchostracans: Clam shrimps belonging to the genus Eulimnadia (Phylum Arthropoda, Subphylum Crustacea, Class Branchiopoda, Order Conchostraca). Correspondence analysis (CA): An ordination method that simultaneously ordinates samples and variables (species), and maximizes the correlation between sample and variable scores. It is widely used in ecology because it assumes a unimodal (bellshaped) response of species variables to underlying gradients. Cosmopolitan: Growing or occurring in many parts of the world; widely distributed. Crustacea: A phylum or subphylum of Arthropoda. Most crustaceans are aquatic. Cryosphere: The portions of the Earth’s surface where water is in solid form, usually as snow or ice. This includes sea ice, freshwater ice, snow, glaciers and permanently frozen ground (permafrost). Cyanobacteria: (formerly Cyanophyceae = blue-green algae). A large group of prokaryotes that possess chlorophyll-a and carry out photosynthesis with the concomitant production of oxygen. Many species fix atmospheric nitrogen. Cyanoprokaryote: See Cyanobacteria. Dansgaard-Oeschger events: Dansgaard-Oeschger events are rapid climate oscillations first identified in the G18O records in the GRIP and GISP2 ice cores from Summit, Greenland. They are followed by cooling events that may end in Heinrich Events. "Dansgaard" relation: Spatial relation between the weighted mean-annual G18O of precipitation at a site and mean annual air temperature, approximated by G18O = 0.7MAT - 13.6; sometimes used to estimate or constrain paleotemperature from paleoprecipitation G18O. DCA: See Detrended correspondence analysis.

519 DDD: 4,4’-Dichlorodiphenyldichloroethane. Degradation product of DDT in sediment formed by reductive dechlorination. DDT: 4,4’-Dichlorodiphenyltrichloroethane. Widely used chlorinated pesticide now banned in most countries. Major isomers in technical DDT were 4,4’-and 2,4’-DDT. Denaturing Gradient Gel Electrophoresis (DGGE): A method for separating DNA fragments according to their mobilities under increasingly denaturing conditions (usually increasing formamide/urea concentrations). The theory behind DGGE is that the two strands of a DNA molecule separate, or melt, when heat or a chemical denaturant is applied. Dendrochronology: A chronology based on counting annual tree-rings and matching rings from different trees/tree-logs, using the obtained tree-ring time scale to date both geological and historical events. Detrended correspondence analysis (DCA): An indirect gradient analysis method that summarizes directions of variation in a given dataset by reducing the dimensionality of data. It is a special type of correspondence analysis meant to remove the arch effect occurring when one variable has a unimodal (bell-shaped) distribution with respect to a second axis. Deuterium excess (d-excess): A derived parameter used to express deviation in the isotopic composition of a given water sample from the Global Meteoric Water Line in G18O-G2H space, defined by d-excess = G2H - 8G18O, where d-excessGMWL = +10‰. d-excess: See Deuterium excess. DGGE: See Denaturing gradient gel electrophoresis. Diagenesis: Primarily chemical changes in sediment brought about after its deposition. Leads to alteration of chemical profiles in sediment due to prevailing conditions (e.g., reducing conditions). Considered to be a relatively low-pressure, lowtemperature alteration process. The term is also used for diagenesis of biological indicators or more commonly biogeochemical indicators, such as pigments. Diamicton, Diamict: Visually unsorted deposit with a wide range of particle sizes. Tills are usually diamicts carried and deposited by a glacier. Diatom: A microscopic, single-celled planktonic or benthic alga from the class Bacillariophyceae that lives in fresh and marine waters and precipitates siliceous tests (frustules). DIC: See Dissolved inorganic carbon.

520 Dimictic lake: A thermally stratified lake in which the water column mixes twice a year, typically in spring and autumn. Dissolved inorganic carbon (DIC): The fraction of soluble carbon in natural waters that is derived from remineralized organic matter, respiration, and weathering of carbonate minerals. Dissolved inorganic carbon typically includes CO2(aq.), HCO3-, and CO32+. Dissolved organic carbon (DOC): The fraction of soluble carbon that is derived from decomposed (but non-remineralized) plant, algal and bacterial biomass. Dissolved organic carbon frequently stains natural waters brown, and plays an important role in attenuating ultraviolet (UV) radiation in the water column. DOC: See Dissolved organic carbon. Dose rate: Estimated natural radioactivity of a geological sample, required to normalize the sample’s equivalent dose in order to generate luminescence ages. EAIS: See East Antarctic Ice Sheet. East Antarctic Ice Sheet (EAIS): Portion of the Antarctic Ice Sheet which is to the east of the Transantarctic Mountains. In contrast to the West Antarctic Ice Sheet, the base of the EAIS is mostly above sea-level, and is thicker and higher. Ecotone: The geographic boundary or transition zone between two different ecozones or regional bio-zones. ED: See Equivalent dose. Edaphic: Referring to the soil. EFs: See Enrichment factors. Endemic: Native to or confined to a certain region. Endorheic: Water bodies in closed or endorheic watersheds, which contain rivers or lakes that do not drain to the oceans. This interruption of surface water flow results from the balance between inputs (precipitation and surface flows) and outputs (evaporation and seepage). Lakes in endorheic watersheds are often called “terminal” or sink lakes. Englacial: The zone within a glacier but isolated from the surface (see supraglacial) and the bottom (see subglacial) of the glacier.

521 Enrichment factors (EFs): Ratio of recent flux (ng m-2 yr-1) or concentrations (ng/g) (ca. 1970 - present) to its pre-industrial flux (ca. 1860 and older) for mercury and lead. Also used for other metals. Ephippium: Resting egg of Cladocera with a resistant chitinous wall. Ephippia can endure low temperatures, drying out, or passage through the gut of a bird. Epiglacial lake: Lake found against a glacier front or beside an ice shelf. Epilimnion: The layer of relatively warm (hence less dense), oxygen-rich water, lying above the thermocline of a thermally-stratified lake. Epilithon: The assemblage growing attached to stones or rocks. Epipelon: The assemblage living on and in muds. Epiphyton: The assemblage growing attached to plants. Epipsammon: The assemblage growing attached to sand grains. Equilibrium effects: Temperature- and mass-dependent stable isotope fractionation (or partitioning) between two substances or phases under conditions of thermodynamic equilibrium; used to describe isotope exchange between liquid water and vapour. Equivalent dose (ED): The approximation of a geological sample’s total radiation dose accumulated over time, based on TL or OSL measurements of electron trap eviction. Erythemal: The erythemal UV index is a measure of UV doses and risks for human health. The clear-sky erythemal UV index expresses the susceptibility of the Caucasian skin to sunburn (erythema). Computed with the help of world-wide satellite ozone measurements. Eukaryotic: Single-celled or multicellular organisms whose cells contain a distinct membrane-bound nucleus and organelles. Euplankton: Plankton whose life cycle takes place entirely in the water column. Eutrophic: Pertaining to highly productive, nutrient-rich conditions. Evapotranspiration: The sum of the water lost from the land by evaporation and plant transpiration (usually via their stomata; see stomata). Exoskeleton: An external skeleton, like that of insects and other arthropods.

522 Fecal pellet: A sedimentary particle excreted by zooplankton. The pellet is frequently composed of mineral sediment and organic material, and settles rapidly due to its large size. Fly-ash: Fine particulate material derived from incineration and other combustion processes comprising of spheroidal carbonaceous particles (SCPs) and inorganic ash spheres (IASs). Fm: See Fraction modern. Foraminifera: A large order of rhizopods that typically have a chambered calcareous shell formed by several united zooid. Fourier Transform (FT) Raman spectroscopy: A laser excitation method used to obtain diagnostic vibrational spectroscopic data from a large number of chemical systems. Often coupled to fibre optics, macro optical arrangements and microscopes, and able to sample through glass, to study species in all physical forms including solution. It also has the highly desirable properties of being non-destructive, noninvasive and very often highly specific. Fraction modern (Fm): Ratio between the radiocarbon activities of the sample being dated and a recognized modern standard such as National Bureau of Standards oxalic acid. In determining conventional radiocarbon ages, the modern standard reflects atmospheric radiocarbon content in 1950. Thus, samples with a fraction modern > 1.0 are enriched by anthropogenic radiocarbon associated with thermonuclear weapons testing. Frustule: The siliceous component of the diatom cell wall. FT: See Fourier Transform. Fulvic acids: A class of dissolved organic matter that remains soluble in both acid and base. Gas hydrates: See Clathrates. Geological luminescence: Light emission from silicate minerals resulting from the time-dependent storage of electrons in lattice charge defects. The release and subsequent quantification of geological luminescence is accomplished by heating or exposure to light. Geomagnetic poles: The points of emergence at the Earth’s surface of the axis of the magnetic dipole that most closely approximates the Earth’s magnetic field.

523 Geomagnetic shielding: Influence of Earth’s magnetic field on the cosmic ray flux reaching the upper atmosphere, and hence influencing radiocarbon production by the bombardment of nitrogen. Geomagnetic shielding and radiocarbon production have covaried significantly over the past 50 ka. GISP: Greenland Ice Sheet Project. See also entry under GRIP. GISP2 ice core: Greenland Ice Sheet Project 2. Primarily an American effort, which led to obtaining a 250,000 year long ice core from the Greenland summit ca. twenty miles west of GRIP between 1989 and 1994. The GISP2 site was located at 72º36’N, 38º30’W. This ice core provides an extensive paleoclimate record for the Northern Hemisphere that has helped investigators piece together past environmental changes, which could shed light on possible future scenarios. Glacio-isostatic rebound: The uplifting of land once depressed under the weight of glacial ice. See Glacio-isostatic uplift. Glacio-isostatic uplift: Glacioisostasy refers to the deformation of the Earth’s lithosphere as a result of changes in the mass of an ice sheet. The equilibrium isostatic depression may be up to 1/3 of an ice sheet’s maximum thickness. The rate of ice retreat is generally much faster than rebound of the Earth’s crust in response to changes in load. Along coastal regions, this means that low-lying areas become transgressed by the sea, only to be progressively uplifted as the lithosphere rebounds, a process known as glacio-isostatic uplift. Global Meteoric Water Line (GMWL): Approximate best-fit line in G18O-G2H space, representing the locus of isotopic compositions of amount-weighted mean annual precipitation at sites worldwide, described by the equation G2H = 8G18O +10; used as a reference line in describing variations in the isotopic composition of natural waters. Global Network for Isotopes in Precipitation (GNIP): World-wide network of stations where precipitation is collected for analysis of G2H and G18O, jointly operated by the World Meteorological Organization and the International Atomic Energy Agency. GMWL: See Global Meteoric Water Line. GNIP: See Global Network for Isotopes in Precipitation. GPR: See Ground penetrating radar. Graticule: A piece of glass engraved with a scale or pattern and inserted into the microscope eyepiece, used to measure objects.

524 GRIP ice core: Greenland Ice Core Project. Primarily a European effort, which led to obtaining a ca. 250,000 year long ice core from the Greenland summit, twenty miles east of GISP2 between 1989 and 1994. The GRIP site was located at 72º35’N, 37º38’W. This ice core provides an extensive paleoclimate record for the Northern Hemisphere that has helped investigators piece together past environmental changes, which could shed light on possible future scenarios. GRIP: Greenland Ice Core Project. See also entry under GISP2. Ground penetrating radar (GPR): A radar system specially configured to work in the near surface of the Earth in order to map geologic and hydrologic details. Grounding-line: The line defining where a floating ice sheet (ice shelf) becomes a grounded glacier. Guano: Accumulated excrement and remains of birds, bats, seals, and other animals. Gyttja: Lacustrine sediment composed primarily of organic materials in various states of decay. This type of sediment is common in lakes that do not have significant inflows of mineral sediment. HA: See Humic acids. Half-life: The length of time required for the decay of half the initial inventory of radioactivity associated with an individual radionuclide. As a guideline, the dating utility of a given radioisotope can be approximated by 8 to 10 times its half-life. HCB: See Hexachlorobenzene. Head capsules: In paleolimnology, typically refers to the chitinized head capsules of chironomid larvae, which are commonly used paleoindicators. Heinrich events: Iceberg pulses and cooling events detected in North Atlantic sediments and in ice cores. The corresponding Heinrich layers found in marine sediments are rich in ice-rafted debris and poor in foraminifera, and their approximate ages are 69,000, 52,000, 35,500, 26,500, and 14,500 years BP. Hexachlorobenzene (HCB): A by-product of synthesis of many chlorinated organics and formerly used as a fungicide. Holocene: The last 10,000 years of the Earth’s history – the time since the end of the last major glacial epoch, or ‘ice age’. Equivalent to Oxygen Isotope Stage 2. Homopycnal flow: Suspended sediment dispersal in a lake at all depths due to similar densities between the lake water and the sediment plume. This type of flow typically occurs in unstratified lakes with low suspended sediment concentrations.

525 Humic acids (HA): The base soluble, acid insoluble humic fraction of decomposed organic matter. Humic acids include a range of recalcitrant high molecular mass biomolecules associated with algal, bacterial, and terrestrial precursors. Humin: The humic fraction of decomposed organic matter that is insoluble in both acid and base. Hydraulic “bow wave”: A pressure wave created by some sediment corers as they drop through the water column. Hypersaline: Describes water with an ionic concentration greater than 50‰. The hypersaline bottom waters of Lake Sophia (Cornwallis Island, Canada) have a salinity of 58‰, compared to ca. 33‰ for sea water. Hypersaline lake: A lake with a salinity usually greater than that of sea water (> 33‰). Hypersalinity: Salinity well above concentrations typically recorded in sea water. Hypolimnion: Lowest (coldest and most dense) water layer in thermally stratified lakes, where light penetration and biological activity are minimal, and turbulence is reduced. Hypopycnal flow: A sediment-laden plume of water of lower density than the water body it enters, resulting in widespread dispersal of suspended sediment across the basin. Most common where glacial meltwater enters saline water bodies. Hypsithermal: A term generally used for the postglacial climatic optimum, the warmest Holocene period. However, the term has also been used for the warmest period of any late Cenozoic interglacial. IAS: See Inorganic Ash Spheres. ICP-MS: See Inductively-coupled plasma mass spectrometry. Inductively-coupled plasma mass spectrometry (ICP-MS): Analytical technique for the detection and quantification of multiple elements at trace concentrations by combining ionization with mass spectrometry. Parts per million to parts per trillion concentrations are attainable. The analytical process includes sample nebulization into an aerosol stream, ionization by an argon plasma source, mass to charge ratio discrimination, and quantification. Inorganic ash spheres (IAS): Formed by the fusing of the inorganic component within the fuel (hence mainly from coal-series fuels). Composed primarily of aluminosilicates with varying quantities of other elements, such as iron.

526 Interstadial: A term generally used in connection with the Quaternary glaciations, meaning a relatively warm phase between colder phases (stadials). An interstadial is usually cooler and shorter than an interglacial. However, an interstadial can be as warm as an interglacial, but in that case it must be of such short duration that the glaciers of the world were not reduced to interglacial sizes, and the regional vegetation zones did not reach the extent of the interglacial zones. Isostatic rebound: See Glacio-isostatic rebound. Isotopic enrichment: Heavy-isotope build-up. Isotopic fractionation: Variations in the relative abundances of different isotopes of an element between substances or phases; also known as isotopic partitioning. Isotopic partitioning: Variations in the relative abundances of different isotopes of an element between substances or phases; also known as isotopic fractionation. Isotopomers: Molecules of the same substance having differing combinations of the isotopes of the constituent elements, as in the naturally occurring stable water isotopomers 1H1H16O, 1H2H16O and 1H1H18O. Jökulhlaup: Icelandic term literally defined as “glacier burp” used to define sudden flood events caused by the drainage of an ice-dammed lake, a lake at the margin or within a glacier. Meltwater can be released from either water ponded on the surface (see Supraglacial) or under the glacier (see Subglacial), and often results from melting caused by volcanic activity. ka: Kilo annum, 1000 years. Katabatic winds: Winds generated when high elevation cold air (e.g., atop a glacier) flows downslope to peripheral areas, driven by gravity. KB (Kajak-Brinkhurst) corer: Gravity-type corer used for collection of surface sediment cores. Kettle lakes: See Kettles. Kettles: Depressions left behind after partially-buried ice blocks melt in glacial rift deposits (especially in outwash plains). When filled with water, they form kettle lakes. The depression is frequently bowl-shaped (also called kettle holes). Kinetic effects: Stable isotope fractionation (or partitioning) between two substances or phases during one-way reactions; used to describe isotope effects arising during gas-phase molecular diffusion of water molecules during evaporation.

527 Krummholz: German word describing the stunted growth form of trees. Individual trees that have a low shrubby to mat-like growth form due to extremely harsh environmental conditions at high elevation and high latitude treelines. Last Glacial Maximum (LGM): Defines a time coinciding with a period of low global sea-level and relative climate stability. A period from at least 22,000 to 19,000 (calendar) years before present, when ice volume was at its maximum, exceeding today’s grounded ice sheets by 52.5 x 106 km. A rapid decrease in ice volume by about 10% within a few hundred years terminated the Last Glacial Maximum at 19,000 ± 250 years. The actual maximum extent of individual ice sheets may not correspond to the LGM, e.g., the maximum extent of the Laurentide Ice Sheet is considered to have occurred prior to the LGM. Late-Wisconsinan: A sub-stage of the Wisconsinan, extending from approximately 23 to 10 ka. Lattice charge defects: Loci of electron capture over geological time formed during mineral crystallization or by exposure to natural radiation. Electrons accumulated over time in lattice charge defects are the source of luminescence used in geochronology. Lead-210: See 210Pb. LEL: See Local Evaporation Line. Lentic: Pertaining to standing waters, such as lakes; cf. lotic. LGM: See Last Glacial Maximum. LIA: See Little Ice Age. Libby half-life: The half-life of radiocarbon accepted for calculation of conventional radiocarbon ages, given as 5568 years, and named in recognition of Willard F. Libby who pioneered radiocarbon dating (1946-1949). The known half-life of radiocarbon is ca. 5730 years. Little Ice Age (LIA): Climatic anomaly recorded widely throughout northern Europe, the northern North Atlantic and eastern North America from ca. AD 1200 to 1850, characterized by cooler conditions. Climatic anomalies at around this time have also been recorded in the tropics, but with a different regional expression (wetter conditions), suggesting that the LIA represents a much larger perturbation to the hemispheric or global climate system. Littoral zone: Shallow water, nearshore region of a lake, where rooted macrophytes can grow.

528 LM: Light microscopy or light microscope. LMWL: See Local Meteoric Water Line. Lobate ejecta craters: Craters found on Mars that are indicative of a fluid or ice being present at the point of impact. Local Evaporation Line (LEL): Displacement of the isotopic composition of water undergoing evaporation in G18O-G2H space resulting from heavy-isotope (2H and 18 O) build-up. Local Meteoric Water Line (LMWL): Best-fit line in G18O-G2H space representing the locus of isotopic compositions of precipitation received over the year at a given site, usually based on the mean amount-weighted G18O and G2H values of monthly precipitation. LOI: See loss-on-ignition. Loss-on-ignition (LOI): Organic matter content of the sediment estimated by measuring weight loss in sub-samples after burning at selected temperatures (typically at 550ºC for 3 hours). Often used as a first-order estimate of organic carbon content of sediments, and widely used in lake sediment studies. Lotic: Pertaining to running waters, such as streams and rivers; cf. lentic. MAAD: See Multiple Aliquot Additive Dose. Macrofossil: A fossil that can be seen without the aid of a microscope. Macrophyte: In its broadest sense, simply means a macroscopic plant. However, as used by limnologists, typically means aquatic plants (angiosperms). MAR: See Mass accumulation rate. Marine isotope stage (MIS): The oxygen isotope records of foraminifera from deep ocean cores reveal globally synchronous shifts in į18O (the ratio of 18O/16O) that can be correlated with periods of glaciation (higher, less negative į18O values) and interglacials/interstadials (lower, more negative į18O values). Glacials are assigned positive marine isotope stage numbers, whereas interglacials are assigned negative marine isotope stage numbers. The last interglaciation (Sangamonian) occurred in MIS stage 5, while the last major glaciation (Wisconsinan) occurred through stages 2-4. Marl: Fine-grained sediment mainly consisting of clay and calcium carbonate.

529 Mass accumulation rate (MAR): The amount of material that accumulates for a specified area and time interval (e.g., grams of sediment over 1 cm2 each year (g cm-2 a-1). MCA: See Medieval Climate Anomaly. MDE: Mercury depletion event. Medieval Climate Anomaly (MCA): The period between the 9th and 13th centuries when global temperatures were thought to be about 1.0°C warmer than present in the North Atlantic sector. Recently, the idea of a global or hemispheric MCA, or ‘Medieval Warm Period’ (MWP), that was warmer than today, has been questioned. Medieval Warm Period (MWP): See Medieval Climate Anomaly. Meromictic: Describes lakes that are chemically stratified, where mixing or circulation of the entire water body is incomplete on an annual basis. A lake that mixes in the upper layers (mixolimnion), but deeper waters (monimolimnion) do not. Associated with small sheltered lakes and also isolation basins, where trapped sea water or other high salinity waters form dense bottom waters overlain by less dense freshwater. MHH: See Mid-Holocene Hypsithermal. Microbial loop: A concept in which heterotrophic and phototrophic bacteria and small eukaryotic phytoplankton are consumed by heterotrophic nanoflagellates (HNAN), which are themselves consumed by larger protozoa and then metazoa, linking the energy lost as dissolved organic matter (DOM) back (loop) to copepods and other consumers of net plankton. The microbial loop helps to explain some of the role that micro-organisms may play in the global ecosystem. Microbial mats: Microbial mats are formed by cyanobacteria, organisms which thrive in environments where extreme fluctuations in conditions occur. On the early Earth, before higher multicellular organisms evolved, photosynthetic microbial mats were widespread and microbes were the only life forms. Mid-Holocene Hypsithermal (MHH): Generally synchronous climatic interval in the mid-Holocene (ca. 6000 to 2000 yr BP) characterized by higher mean temperatures. Also called the mid-Holocene climate optimum, mid-Holocene warm period and Holocene climate optimum. Discrepancies in the exact timing of the MHH at different locations are believed to be the result of local climate variations, insufficient dating, and varying carbon residence times.

530 Milankovitch cycle: Refers to cyclical changes in the Earth/Sun geometrical relationship that effect long term climate patterns. Includes three cycles: (i) eccentricity of the Earth’s orbit, with a periodicity of ca. 92,000 years; (ii) ecliptic or tilt of the Earth’s axis varying between 21°58’ and 24°36’ and having a periodicity of ca. 40,000 years; and (iii) precession of the equinoxes (occurring in the aphelion or perihelion), with a periodicity of ca. 21,000 years. Mirex: Dodecachloro-octahydro-1,3,4-methano-cyclobuta(cd)pentalene, a persistent chlorinated pesticide now banned globally by the Stockholm persistent organic pollutants (POPs) convention. MIS: See Marine isotope stage. Moat: Ring of open water around the edge of an otherwise ice-covered lake formed as the melt season begins. MOLAR: A European Union-funded research project “Measuring and modelling the dynamic response of remote mountain lake ecosystems to environmental change: A programme of Mountain Lake Research.” Monimolimnion: The unmixed, high density bottom layer of water in a meromictic lake. Monomictic lake: A lake in which there is only one annual mixing event. Mössbauer spectroscopy: Spectroscopic technique based on the Mössbauer effect. In its most common from, Mössbauer absorption spectroscopy, a solid sample is exposed to a beam of gamma radiation, and a detector measures the intensity of the beam that is transmitted through the sample. The Mössbauer effect is a physical phenomenon discovered by Rudolf Mössbauer in 1957, and refers to the resonant and recoil-free emission and absorption of gamma rays by atoms bound in a solid. Multiple Aliquot Additive Dose (MAAD): Method of determining sample equivalent dose in luminescence dating, applied principally to fine-grained polymineral or quartz fractions such as those in waterlain sediments. The method applies incremental additive doses (b or g) to natural luminescence in separate sample aliquots to construct a dose response curve that both simulates future doses and can be extrapolated to an equivalent dose at the solar reset level. Although MAAD was developed for thermoluminescence (TL) dating, it can also be used with optically stimulated luminescence (OSL) dating. MWP: See Medieval Warm Period. NAO: North Atlantic Oscillation. Dominant pattern of atmospheric circulation variability over the North Atlantic and over the Northern Hemisphere.

531 NARP: See Nordic Arctic Research Programme. Negative thermal inertia: Collectively describes the phenomena of katabatic winds, decreased thermal and albedo gradients, and the increased frequency of low-level inversions that lead to localized cooling effects proximal to glaciers and ice caps. Cooling effects are most pronounced when the surrounding area is snow-free, and largely absent when snow cover is extensive. Neogene: Geological time period from ca. 26 to 2.5 million years ago (latest part of the Cenozoic, after the Oligocene), that immediately predates the Quaternary, and comprises the Miocene and Pliocene. Neoglacial: A term used to describe a period of glacier expansion in the mid- to late Holocene following the hypsithermal climatic optimum. Distinguished from the ‘Little Ice Age’ which is more specifically applied to the period of maximum Holocene alpine glacier advance in Europe and North America during the 16th to 19th centuries. The Little Ice Age may be considered as having occurred at the end of the Neoglacial period. Nordic Arctic Research Programme (NARP): A research programme funded by the Nordic Council of Ministers. A wide range of projects have been funded in 3 main areas: (a) natural processes – land, sea and atmosphere; (b) biological diversity and environmental threats in the Arctic, and (c) living conditions of the inhabitants of the Arctic. (Website: http://thule.oulu.fi/narp/index.htm). NORLAKE: Scientific project (1999-2003) involving researchers from Denmark, Sweden, Norway, the Faeroe Islands and Iceland, focussing on the biological structure and dynamics of North Atlantic lakes, in relation to spatial variations and temporal changes in climate and land use. Part of the NARP programme. (Website: http://thule.oulu.fi/narp/pages/projects.htm). Oligosaline: Having a low concentration of salt. Oligotrophic: A general term used to describe lakes and rivers with low nutrient concentrations and productivity. Oligotrophication: A tendency in some water bodies towards lower nutrient concentrations and less productive conditions. Optically stimulated luminescence (OSL): The fraction of time-dependent luminescence accumulated in silicate minerals that is liberated by exposure to light at specific wavelength ranges, typically between 400 and 900 nm.

532 Optically stimulated luminescence (OSL) dating: A relatively new dating method where quartz and feldspar-rich sediments, which are otherwise not datable by conventional radiocarbon methods, can be absolutely dated (± ca. 10%) within a range of ca. 100,000 to 200,000 years. Organic carbon: Carbon derived from biological processing and photosynthesis. An important fraction of sediments for sorption of organic chemicals. OSL: See Optically stimulated luminescence. Ostracoda: A microscopic (generally 0.5 to a few mm long) class of aquatic Crustacea, typically having a calcified, bivalved carapace (closable by adductor muscles) enclosing a body that is usually unsegmented. Oxygen isotope stages: A sequence of sedimentation events recognized in deep oceanic cores by the oxygen isotope (18O/16O) ratios in the benthic and planktonic foraminifera therein. Oxygen isotope stage 5e: Marine (deep-sea) oxygen isotope stage 5e corresponds to the maximum of the Eemian/Sangamon Interglacial, about 130,000-115,000 years ago. Palaearctic: Zoogeographical region that comprises Europe, northern Africa, and northern Asia. Paludification: The process of water-logging mineral soils and the eventual development of fen or bog ecosystems that form peat. Palynology: The study of pollen and spores in sediments to determine vegetation patterns and changes in vegetation and climate over time. Palynomorph: Broadly defined group of ornamented microfossils composed of sporopollenin and related polymers, highly resistant to decay, and including pollen grains, spores, dinoflagellate cysts, and acritarchs. Paraglacial: A previously glaciated landscape that typically contains abundant erodable surface material. Parthenogenetic generations: Generations of animals where the eggs develop without fertilisation. Particulate organic carbon (POC): Undecomposed organic matter in natural waters, including algal cells, zooplankton, palynomorphs and fragments of detrital plant tissues (leaves, roots, stems, seeds). PCA: See Principal components analysis.

533 PCBs: See Polychlorinated biphenyls. PCDD/Fs: See Polychlorinated dibenzo-p-dioxins and dibenzofurans. Pelagic zone: The open water region of a lake or ocean, not including the bottom and the littoral/coastal zone. Periglaciation: Cold climate geomorphic processes associated with the presence of permafrost, often occurring in areas near glaciers. Periphytic algae: Algae that live attached to substrates. Permafrost: Ground material that is below 0°C for more than a year. Permafrost may occur in soil, organic or bedrock substrates, and may contain appreciable quantities of ice. Seasonal melting may occur at the surface of permafrost (see Active layer). Persistent organic pollutants (POPs): A group of mainly chlorinated organics which are only slowly degraded in the environment, bioaccumulative and toxic. Perylene: A polyaromatic hydrocarbon formed naturally in sediments from plantderived sterols by reductive processes. Photic zone: The illuminated upper zone of the water column in lakes and oceans, where photosynthesis exceeds respiration. Also called euphotic zone. Photosynthate: Pigmented organic matter associated with photosynthesis, including chlorophylls, carotenoids, and their degradation products. Phytoplankton: Photosynthetic plankton (e.g., algae that float in the open water). Pigment: A compound which absorbs light, usually in the visible range (400-700 nm), and hence imparts colour. Plankton, Planktonic: Pertaining to organisms that live in the open water, and are carried passively by currents and waves. These include many microscopic plants (phytoplankton) and small invertebrate animals (zooplankton). POC: See Particulate organic carbon. Pollen: The male gametophyte of seed plants (angiosperms or gymnosperms) contained within a microspore wall. Polychlorinated biphenyls (PCBs): Group of chemicals now banned in most countries that were formerly used as insulating fluids in electrical transformers as well as for other minor uses that required highly stable chemicals.

534 Polychlorinated dibenzo-p-dioxins and dibenzofurans (PCDD/Fs): Persistent, bioaccumulative and toxic byproducts of combustion and chemical synthesis. Polymictic lake: A lake which is almost constantly mixing, and is not thermally stratified. Polynya: Semipermanent area of open water in sea ice. Coastal polynyas characteristically lie just beyond landfast ice, i.e., ice that is anchored to the coast. They are thought to be caused chiefly by persistent local offshore winds. Openocean polynyas are larger and longer-lasting and are believed to be caused by the upwelling of deep warmer water. Pond: General term used to refer to a small lake. In polar regions, ponds are typically defined as water bodies that freeze totally to the bottom in winter. POPs: See Persistent organic pollutants. Preglacial: Prior to the onset of a glacial period in a region’s history. Principal components analysis (PCA): A numerical technique for the analysis of a multidimensional data set based on the identification of orthogonal linear combinations of variables that are selected to capture as much of the total variance in the data as possible. Proglacial lake: Lake that is hydrologically and sedimentologically influenced by the presence of one or more glaciers within its catchment. Prokaryotes, Prokaryotic: Single-celled organisms that lack a distinct membranebound nucleus and organelles (e.g., bacteria, Cyanobacteria). Protozoans, Protozoa: Single-celled, animal-like organisms that obtain energy through the consumption of other organisms. A subset of the kingdom Protista. Radiocarbon (14C) dating: There are three isotopes of carbon occurring in nature, 12C, 13 C and 14C. The percentages of the three in the atmosphere are about 98.9%, 1.1% and 0.000000000118%, respectively, and they are all oxidized to carbon dioxide (CO2). 14C is formed in the upper atmosphere by cosmic-ray bombardment of nitrogen atoms (14N). The CO2 in the atmosphere is transformed by photosynthesis to organic matter in plants on which animals feed. Therefore, all living organic matter contains carbon isotopes in about the same ratio as in the air. When organisms die and cease to exchange carbon with the atmosphere, the proportion of the radioactive isotope 14C decreases according to its decay rate. Measurement of the residual 14C radioactivity in an organic sample therefore allows its age to be estimated. The limit of radiocarbon dating is ca. 50,000 years. Radiocarbon reservoir effect: See Carbon residence time effect.

535 Rayleigh-type distillation: Preferential partitioning of heavier water isotopomers (1H2H16O and 1H1H18O) during condensation of water vapour to form precipitation, resulting in progressive depletion of 18O and 2H in residual vapour and hence in subsequent precipitation along a vapour mass trajectory. RDA: See Redundancy analysis. Redox boundary layer: Boundary between a region of reducing conditions and oxidative conditions in sediment. Redundancy analysis (RDA): A mathematical method of gradient analysis based on a canonical form of principal components analysis (PCA) in which the axes are restricted to be linear combinations of explanatory variables. Reservoir effect: See Carbon residence time effect. Rhizopods: Group of protozoans characterized by the possession of pseudopodia (e.g. amoebae). Sandur: An Icelandic term for the outwash plain formed by glacial rivers beyond the front of a glacier or ice sheet. Sandurs are subject to periodic flooding, sediment deposition, and erosion. Due to frequent flooding, the surface tends to be free of vegetation or stabilizing soil. Sediments vary depending on proximity to channels and the type of sediment transported by the river. Sangamonian: A chronostratigraphic unit defining the pre-Wisconsinan interglacial, coinciding with marine oxygen isotope stage 5a-e (approximately 130 to 72 ka). SAR: See Single aliquot regeneration. Saturated isothermal remanent magnetisation (SIRM): . Reflects the concentration of magnetic minerals in sediment cores (measured in mA (milli Ampere) m2 kg-1). SCP: See Spheroidal carbonaceous particle. SCUBA: Self Contained Underwater Breathing Apparatus. Secular variations curves: Dated records indicating variation in sedimentary magnetic properties (inclination, declination, intensity) that can be potentially correlated between sites within a region. SEM: Scanning electron microscope.

536 Single aliquot regeneration (SAR): Method of determining sample equivalent dose in luminescence dating, applicable to either fine or coarse fractions. In SAR, an age is determined for each aliquot by matching a regenerated signal from the light reset level to the original natural level. The method is particularly useful for young sediments (< 50 ka) and shows potential for decadal resolution in sediments younger than 1000 years. Single-grain optically stimulated luminescence: The application of optically stimulated luminescence to date individual mineral grains, typically sand-size quartz. SIRM: see Saturated isothermal remanent magnetisation. Snowball Earth Hypothesis: The Snowball Earth hypothesis proposes that, during several profoundly cold periods of Earth’s history that occurred from 750 to 600 million years ago, the ocean was covered by a thick sheet of ice. This would have had a huge impact on early life on Earth. While scientists generally agree that ice sheets reached low latitudes on the planet during severe glaciations, the idea of a completely ice-covered ocean is disputed. Spheroidal carbonaceous particles (SCPs): Particulates produced by the incomplete combustion of the fuel particle or droplet and composed mainly of elemental carbon. Sponge gemmule: A round over-wintering body formed as an internal bud in freshwater sponges. Sponge spicule: Small, often needle-like structures, in the tissue of sponges. Consist of opaline silica. Spore: Resting or dispersal stage of pteridophytes or bryophytes, as well as of algae and fungi. Stable isotopes: Non-radioactive atoms of an element having varying numbers of neutrons in their nuclei, and thus characterized by differing atomic masses. Statoblast: Asexual reproductive body produced by freshwater Bryozoa. The resistant wall enables it to withstand adverse conditions. The wall is composed of chitin. Statospore: Siliceous resting stage of chrysophyte algae (also called cyst or stomatocyst). Stomata: A pore in the epidermis of plants that allows gaseous exchange between the interior of the plant and the external environment. Each stomate is surrounded by two specialized guard cells. Conifer guard cells are generally lignified and often found preserved in lake sediments. Preserved stomata can often be identified to genus.

537 Stomatocyst: Siliceous resting stage of chrysophyte algae (also commonly referred to as statospore or cyst). Subglacial: Relating to the bottom or underside of a glacier (e.g., subglacial meltwater channel - a channel at the base of a glacier). Subglacial lake: A lake located under a glacier or an ice sheet. Heat originating from the Earth’s core makes the base of ice sheets and glaciers relatively warm and, if the ice is thick enough, liquid water can be found and accumulate in geological depressions. Supra-anal plates: Special plate found on the hind part of the body of some crustaceans. Supraglacial: Relating to the surface of a glacier (e.g., supraglacial meltwater pond - a pond located on the surface of a glacier). Supraglacial lake: A lake located on the surface of a glacier or an ice sheet. These water bodies can form in hollows, or in depressions bounded by ice, moraines, or rock dams. They often form in depressions on debris-covered glaciers. Suture zone: A relatively narrow zone along colliding plate boundaries where two large continental masses are welded into one unit. In dinoflagellates (Dinophyceae), the edge of a plate in thecate species, which is sometimes raised. Syntectonic: A geologic event, process, rock or feature formed during a period of tectonism. Taiga: A Russian term for the belt of coniferous forest adjacent to the tundra, more or less synonymous with the Boreal forest vegetation zone. Talik: An unfrozen zone in permafrost located under a lake or deep pond. Heat release from the overlying water prevents formation of ground frost during winter. TDS: See Total dissolved solids. Teleconnection: Relationship between weather or climate (e.g., precipitation) patterns at two or more widely separated locations. Telmatic vegetation: Vegetation on lake margins and wet ground, such as reeds and sedges. Temperature gradient gel electrophoresis (TGGE): A method for separating DNA fragments according to their mobilities under increasingly denaturing conditions when heat is applied.

538 Tephra: Solid matter that has been ejected by an erupting volcano and transported through the air, such as volcanic ash (not lava). The resulting tephro-chronology is based on correlating volcanic ash beds in sediments. Testate amoebae: Amoebae (rhizopods) that are protected by an external test (shell) composed of chitin, silica or calcium carbonate. Also referred to as testaceans or thecamoebae. TGGE: See Temperature gradient gel electrophoresis. Thecamoebae: See Testate amoebae. Thermohaline circulation: The global ocean circulation driven by differences in the density of the sea water, which is mainly controlled by temperature (thermal) and salinity (haline) conditions. Thermokarsting: Subsidence or irregular collapse of ground ice in permafrost areas due to ice melt processes, resulting in a bumpy topography with depressions. Thermoluminescence (TL): The heat-induced emittance of light from some minerals which have stored energy resulting from electron displacements in the crystal lattice. The fraction of time-dependent luminescence accumulated in silicate minerals that is liberated by heating the sample. Thermoluminescence (TL) dating: A dating method for sediments that depends upon the acquisition and long-term stable storage of TL energy by crystalline minerals contained within a sedimentary unit. This energy is stored in the form of trapped electrons. Thermoluminescence is the process in which a mineral emits light while it is being heated. The amount of light emitted is related to the sample’s age. TL dating is suitable for the dating of Quaternary age ceramics and sediments deposited within the last ca. 500,000 years. Thermophilous, Thermophilic: Warmth-loving. Used in reference to organisms characteristic of relatively warm biotopes. TL: See Thermoluminescence. Total dissolved solids (TDS): Matter dissolved in water, including both truly ionic and non-ionic species. Measured by complete evaporation of a known volume of liquid. Toxaphene: A chlorinated pesticide derived from chlorination of Į-pinene. Training set: See Calibration training set.

539 Transfer function: A mathematical function that depicts relationships between organisms and environmental variables. Transfer functions permit the reconstruction of values of different environmental variables (e.g., pH, temperature) to be inferred from the composition of a fossil assemblage. Treeline: The limit of tree occurrence, beyond which is usually shrub-tundra (e.g., lowarctic vegetation) at the latitudinal treeline or alpine tundra vegetation at the altitudinal treeline. Trophic status: Pertaining to the nutrient level of a water body. Tundra: Vegetation zone that is present beyond the treeline, typically in a permafrost area. Treeless zone with sparse or patchy vegetation composed of small plants, such as cushion plants (e.g., mosses) and grasses. Characteristic of high latitude and high altitude regions. Turbidite: Graded sedimentary deposit associated with a discrete event of upslope sediment destabilization. The base of turbidites is frequently erosional. Tychoplankton: Benthic taxa resuspended in the water column and sampled accidentally as plankton (see plankton). Ultraviolet (UV) radiation: Biologically harmful radiation from the sun that reaches the Earth’s surface in wavelengths between 280 and 400 nm (nanometres, or billionths of a metre). UV: Ultraviolet. UWITEC (Umwelt- und Wissenschaftstechnik): A company specializing in the development and manufacture of made-to-measure equipment for sediment coring and sampling of rivers, lakes and groundwater. Valve: In Ostracoda, one of two halves of the carapace, which are hinged dorsally in living specimens and in articulated fossil or subfossil material. The two parts of a siliceous diatom frustule are also called valves. Varve: Cycles of sedimentation that can be definitively attributed to annual periodicities, thus enabling the development of annually-resolved sediment chronologies. Varves are most often composed of recognizable sedimentary structures (summer and winter couplets) representing a single year of deposition. Varved sediments: Sediments with annual laminations. This annual deposit may comprise paired contrasting laminations of alternately finer and coarser silt or clay, reflecting seasonal sedimentation (summer and winter) within the year.

540 Vienna Pee Dee Belemnite (VPDB): Internationally recognized standard for reporting oxygen and carbon stable isotope ratios in carbonates, organic matter, and other substances, approximating the bulk composition of dissolved inorganic carbon in the oceans. Vienna Standard Mean Ocean Water (VSMOW): Internationally recognized standard for reporting oxygen and hydrogen stable isotope ratios in natural waters and other substances. This analytic standard for stable isotope analysis is based upon the mean natural concentrations of the different isotopes of oxygen and hydrogen in deep offshore ocean water. VPDB: See Vienna Pee Dee Belemnite. VSMOW: See Vienna Standard Mean Ocean Water. Weighted-averaging: A statistical approach in which several observations are used to provide an aggregate estimate of a mean value for a variable. The importance of each observation in determining the final estimate can be adjusted by mathematical weighting so that some observations have greater importance in determining the estimated mean than others. Weighted-averaging is important in techniques such as Canonical correspondence analysis (CCA). Wisconsinan: A chronostratigraphic unit defining the period of the last major continental glaciations, coinciding with marine isotope stage 4-2 (approximately 72 to 10 ka). YD: See Younger Dryas. Younger Dryas (YD): A period of sharp cooling that affected the northern Atlantic region and many other parts of the world around 12,000 years ago (ca. 11,000 radiocarbon years BP). Characterized by a cool or cold climate in the Northern Hemisphere. Zooplankton: Portion of the plankton community represented by mainly microscopic (up to a few millimetres) animals and protists that are suspended in the water column, with limited locomotor control (e.g., Cladocera, copepods, rotifers, heterotrophic flagellates, and ciliates).

INDEX 137

108, 109, 119, 141, 166, 169, 216, 217, 219-220, 221, 222, 270, 272, 279, 283-293, 324, 399, 510 Alaskozetes 439 Alba Patera 494 albedo 2-3, 255, 270, 272, 420, 423 Alberta 63, 221, 273, 275, 276, 277, 281 alder, see Alnus Alert 230 Alexander Island 423, 436 algae 40, 67, 117-145, 157, 211, 253, 335, 399, 405, 425, 426 alkaline 253, 333, 404, 407 alkalinity 278, 325, 327, 362, 363, 400, 412 alkaliphilous 128, 325 alkali-soluble 437 Allerød 177 allochthonous 20, 53, 64, 211, 245, 255, 328, 398, 412 allogenic 389 alloxanthin 126, 334, 438 Alnus (alder) 96, 102, 107, 249, 293, 56, 359 Alona 165, 167, 337 alpine 61, 224, 242, 253, 256, 260, 322, 356, 368, 362, 385, 389, 423, 492 Alps 260 altitude 328, 399, 405 aluminium, see aluminum aluminum (Al) 215, 218 Amadjuak Lake 244 AMAP, see Arctic Monitoring and Assessment Programme Amazonian units 494, 496 Amery Oasis 425 amino acid 438 Amituk Lake 229 Ammassalik 322 amoebae 61 AMS, see accelerator mass spectrometry Amsterdamøya 227-228 analogue 109

Cs, see cesium-137 Rn, see radon-222 226 Ra, see radium-226 238 U, see uranium-238 8.2k event 340 ȕ-decay 29 Abel Bay 446 Abisko 386-387, 403 Abiskomyia 172, 361, 366 ablation 61 Ablation Lake 425 abrasion 58 Agabus 169 Acari 439 accelerator mass spectrometry (AMS) 29-31, 44, 49, 282, 284, 303, 431, 511 accretion 490, 491 accumulation rate 35, 330, 340, 386-387, 390, 412 Ace Lake 434,449 Achnanthes 296, 327, 333, 354, 359, 363 acid 449 acidification 78, 138, 166, 213, 281 327, 407-409, 412 acidity 104, 303, 333, 401-404 acid-leachable 220 acidobiontic 128 acidophilous 128, 258 acoustic 62, 72-83 Acroperus 165 Actinomycetes 456, 490 active layer 54, 64 advection mixing 434 aeolian 55, 56, 64, 68-69, 330, 333, 340, 427 aeolian facies 68-69 aeolian processes 68-69 aerosols 226 Agassiz Ice Cap 251, 252, 255 aggradation 62 agriculture 338-339, 448 air mass 435, 446 aircraft 427 akinetes 125 Alaska 25-26, 56, 90, 102, 105, 106, 222

541

542 Andøya 171 animal 510 Annak Lake 24-25, 27 Anostraca 439 anoxia 78, 334, 412 Antarctic Bottom Water 420, 421, 456 Antarctic Convergence 9, 447 Antarctic Ocean 119, 431 Antarctic Peninsula 3, 4, 33 Antarctica 3, 4, 9, 54, 55, 69, 76, 118, 120-121, 126, 130, 137, 141142, 160, 164, 168, 210, 226, 229, 419-459, 475-500, 509511 Anthropocene 2 anticyclonic 242 antimony (Sb) 219 Apatania 173 Arachnida 160 aragonite 199, 482 Araneae 173 archeology 142-143, 291 Arctic Archipelago 241-261 Arctic Circle 138, 227, 228, 322, 384 Arctic Front 107, 272, 275, 277 arctic haze 226 Arctic Monitoring and Assessment Programme (AMAP) 210, 222, 324 Arctic Ocean 2, 3, 137, 138, 200, 213, 217, 278 Arctocephalus 141-142 Argentina 449 arid 328 Arolik Lake 38, 271, 283, 285 arsenic (As) 219 Artemisia 359, 362 artesian 496 Arthropoda 163-168 Asia 95 Assulina 161 Asterionella 295 Athabasca River 195, 200-202, 222 athalassic 136, 280, 324 Atlantic Ocean 447 atomic weapons 421, 431-432 Aulacoseira 258, 331

INDEX Australia 435 authigenic 67, 76, 493 autochthonous 30, 53, 54, 328 Axel Heiberg Island 242, 244 Axel Heiberg Project 247 Bacillariophyceae, see diatoms bacteria 334, 426, 436-438, 490, 491 bacteriochlorophyll 438 bacterioplankton 426 Baffin Bay 119, 242, 251, 256 Baffin Island 29, 31, 36, 38, 89, 98, 100, 135, 160, 169, 171, 172, 242, 246, 247, 249, 250, 251, 255-266 Baird Inlet 98-99, 100, 247 Bali 128 Banks Island 95, 98, 244 Barbeau Peak 242 Barents Sea 222, 351, 370 Barentsburg 222, 224, 227-228 Barnes Ice Cap 244 barotropic 488 Barrow 219-220 Barrow Straits 119 basal 95, 102, 108 Basaltsø 328-330 base saturation 403 Bathurst Island 247 bathymetry 72, 211 Bayesian modeling 173 Bear Island 222, 320, 323, 324, 325, 337 Bear Lake 62, 72 Beaufort Sea 138, 293 Becharoff Lake 298 bedforms 496 bedrock 253, 402 beetles, see Coleoptera Belcher Islands 24, 230 Bellsund 227-228 benthic 247, 327, 335, 398, 426, 511 benzoperylene 224 benzopyrene 225 Beringia 272, 283-292 beta-carotene 436-438 beta-carboxylase 433 Betula (birch) 96, 102, 105, 107, 244,

INDEX 249, 322, 351, 354, 356, 358, 359, 362, 409 bioclimatology 242-244 biodegradation 213 biogenic facies 67-69 biogenic processes 67-69 biogenic silica 127, 196, 282, 295, 328, 329, 330, 334, 340, 433, 481 biogeochemistry 78, 121, 126-127 biogeography 119, 120, 160, 175-176, 436, 456 biosphere 421 biotransformation 214 bioturbation 69, 70, 21, 213 birch, see Betula birds 140, 426 bismuth (Bi) 219 Bjørnøja, see Bear Island blue-green algae, see Cyanobacteria boat 427 Boeckella 168, 439 bog 161, 339 Bolingen Islands 436 Bølling 284 bone 142-143, 443 Boothia Peninsula 91, 242 bore hole 493 boreal 272, 276, 277, 366 Bosmina 166, 281, 290 Botryococcus 125 Bouvet Island 9 Bow Lake 63 bow wave 212 bowhead whales 142-143 Brachionus 162 Brachysira 358 brackish 120, 442 Branchionecta 164, 439 Branchiopoda 163-164 Bratina Island 425 Braya Sø 325, 328, 333, 334, 341 Brillia 171 brine 245, 420, 497 Bristol Bay 298 British Columbia (BC) 141, 221, 289, 362

543 broken-stick model 406-407 Broknes Peninsula 441 bromine (Br) 217 bromine oxide 217 Brooks Range 102, 107, 283, 288 Bryidae 108 Bryozoa 163, 174-175 bubbles 431 bulk sediment dating 30 Bunger Hills 423, 425, 426, 431, 432, 441, 444 caddisflies, see Trichoptera cadmium (Cd) 213, 218, 219 caesium, see cesium calcite 327, 482 calcium (Ca) 280, 398, 400, 492 Caledonian Mountains 224 calibration 92 California 362 camphene 213 Canada 3-4, 5, 7, 24, 54, 62, 63, 65, 76, 89-96, 98-110, 117-145, 157, 160, 162, 163, 172, 174, 188, 210, 217, 211, 222, 224, 225, 227, 241-261, 323, 325, 362, 398, 425, 485, 487 Canada Glacier 483 Canadian Arctic Expedition of the Eastern Arctic 120 Canadian Arctic Expedition of the Western Arctic 120 Candona 199 canonical correspondence analysis (CCA) 361, 399-400 canonical variates analysis (CVA) 170 canyon 497 Cape Herschel 132, 230, 258 Cape Roberts 429 Cape Wankarema 119 CAPE, see Circumpolar Arctic Paleoenvironments capsules 162 carbon (C) 270, 361, 366, 369-371, 433, 434, 438, 482 carbon dioxide 431, 433, 444, 448, 494 carbonaceous 242

544 carbonate 43, 188, 189, 199, 200, 282, 482 carbonate bedrock 98 caribou 140 carotenes 126-127 carotenoid 126-127, 434, 436-438 Caryophyllaceae 435 cases 173 casing 92 cation 217, 400 CCA, see canonical correspondence analysis CDOM, see chromophoric dissolved organic matter cellulose 366, 367 Cenozoic 121, 242 cesium-137 (137Cs) 20-29, 73-75, 212, 271, 431-432, 440 CFCs 4 chamber corer 429 channeled scablands 496 Chaoboridae 171 Chaoborus 171 Char Lake 139, 222, 247 Chara 199, 327 Characeae 325, 328 charcoal 108, 230, 435 chemotaxonomic 426 Chenopodiaceae 435 Chernobyl (Ukraine) 21, 24, 212 China 215-216 chironomid, see Chironomidae Chironomidae 39, 67, 169-173, 261, 280, 298, 299, 337, 338, 352, 359, 361, 363-366, 372, 352, 359, 361, 363-366, 372 Chironomus 170 chirps 72 chitin 29, 39 chloride 279 chlorine 361, 434, 438, 483, 484 chlorobornane 213 chlorophyll 118, 126-127, 334, 434, 436, 438 Chlorophyta 125, 334, 337, 435 chonchostracans 427 chromophoric dissolved organic

INDEX matter (CDOM) 133-135, 395-399 chronology 73-75, 212, 429, 431-432 chryse 497 chrysene 224 Chrysophyceae, see chrysophytes Chrysophyta, see chrysophytes chrysophytes 123-125, 127, 129, 144, 247, 253, 260, 279-280, 283, 382, 409 Chrysosphaerella 125 Chuna Lake 356-358 Churchill 92 Chydorus 337 circulation 279, 370, 371, 393, 420 circumneutral 128 Circumpolar Arctic Paleoenvironments (CAPE) 4, 110 cirque lakes 385 Cladocera 163-167, 177, 261, 337, 390, 393, 409 Cladophora 337 clastic 20, 255, 398 clasts 58, 69 clay 55, 58, 61, 66 climatic optimum 334 closed basin 328, 386, 423 clustering 400 coal 224, 227, 228, 432 coal mining 221 coastal continental lakes 426 coastal lake 441, 354 coastal marine lakes 426 Cockburn Substage 251 cocoons 162, 163 cold-based glaciers 61 cold-based ice 249 Coleoptera 160, 168-169 collections 157 colluvial 55 Colobanthus 435 colour 135, 300, 381. 399 Columbia Plateau 496 Colymbetes 168 compartments 6 Compositae 435

INDEX concentration 255 condensation 367 conductivity 331-333, 334 coniferous 277, 351, 362 constant rate of supply (CRS) model 21 contaminants 209-231 convection 497 copepod, see Copepoda Copepoda 167, 168, 281 copper 213, 219 corals 5 corers 90-91, 212 cores 73 coring 90-92 Cornwallis Island 65, 71, 136, 139, 229, 247, 258 Corynocera 170, 171, 172, 362, 367 crater 494, 499 craton 488 Cretaceous 421 crevasses 58 Cristatella 163, 174 CRS model, see constant rate of supply Crustacea 163, 431 cryosphere 2, 240 Cryptophyta 126, 127, 438 cryptophytes, see Cryptophyta cultural disturbances 338-339 Cumberland Peninsula 38, 255, 260 Cumberland Sound 244, 249, 250 CVA, see canonical variates analysis Cyanobacteria 118, 125, 126, 426, 431, 433, 434, 435, 480 cyanoprokaryotes, see Cyanobacteria cyclone 93 cyclonic 242 cyclopoid 281 Cyclotella 130, 132, 294-295, 296, 327, 330-331, 359, 409, 410 Cymbella 253, 333, 356 Cyperaceae 98, 102, 359, 362 cyst 123-125, 253, 279-280, 370, 409 Dalmutlad’do 389, 400 Dansgaard-Oeschger events 422, 457 Daphnia 165, 166, 167, 337 data-loggers 170

545 dating 160, 212, 247, 249, 251, 271, 303-304. 339, 341, 372, 390, 431-432, 511 Davis Strait 242, 322 DCA, see detrended correspondence analysis DDD 213 DDT 138, 213, 220-224 dechlorination 213 deciduous 270 decomposition 387 Deep Lake 425 Defense Research Board 247 defocusing 25 deglaciation 96, 272, 286, 320, 325, 398 degradation 436-438 degree days 169 delta 65, 429 dendroecology 5, 6 Denmark 165, 172 density 433 Denticula 333 depth 165-166, 278, 279, 280, 328, 359, 361, 371, 393, 427 Derevanoi Lake 369, 370 Deschampsia 435 desiccation 483 detrended correspondence analysis (DCA) 334, 404-405 detritus 67, 337, 391, 398 deuterium 483 Devon Island 62, 72, 218, 242, 247, 258 Devon Island Ice Cap 62, 77, 251 d-excess values 191 DGGE, see direct gradient gel electrophoresis diagenesis 328, 329 Diamesa 171 diamicton 58 Diaptomus 167-168 diatom 67, 105, 118-145, 157, 166, 196, 197, 247, 249, 253, 254255, 261, 277, 278, 279, 283, 289-292, 295, 298, 327, 330331, 352, 354, 356, 358, 359, 361, 362, 372, 375, 381, 382,

546 390, 394, 399, 401-404, 405, 406, 409, 426, 435-436, 449 diatoxanthin 334 dibenzofurans 222-223 DIC, see dissolved inorganic carbon diesel 229 diffraction 434 diffusion 213, 218, 434, 483 dilution 330 dimictic 382 dinoflagellates, see Pyrrhophyta Dinophyta, see Pyrrhophyta Diptera 169-173 direct gradient gel electrophoresis (DGGE) 439 discharge 496 Discovery 119 Disko 322 dispersal 324 dissolved inorganic carbon (DIC) 276, 278, 279, 361, 369, 431 dissolved organic carbon (DOC) 30, 32, 34, 129. 133-135, 249, 276, 678, 279, 280, 281, 287, 292-293, 294, 298-300, 324, 340, 395-399, 400, 412 dissolved organic matter 399 distillation 191 diversity 324, 405, 406 DNA 438-439, 490 DOC, see dissolved organic carbon Dolgoe Ozero 362-366, 369 Dome C 425, 444, 447, 487 Don Juan Pond 434 drift ice 119 drill 427 drowning 144 dry ice 429 Dry Valley Drilling Project (DVDP) 481 Dry Valleys (Antarctica) 4, 29, 69 Dryas 98, 362 DVDP, see Dry Valley Drilling Project E/I, see evaporation/inflow ratio EAIS, see East Antarctic Ice Sheet East Antarctic Ice Sheet 441, 487

INDEX East Greenland Current 322 eccentricity 420 echo sounding 72, 456 ecozone 272 edaphic 272 Edgeøya 161 EF, see enrichment factor effective moisture 136, 175 Efstadalsvatn 340 egg 162, 163, 166, 168, 303, 431, 439 Ehrenberg, C.G. 119 Ekblaw Glacier 247 electrode 434 electromagnetic 72 electron microprobe 76-77 electrophoresis 439 elevation 354 Ellasjöen 372 Ellesmere Island 3-4, 24, 25, 65, 95, 98, 99, 100, 120, 130, 132, 137, 217, 224, 229, 230, 242, 244, 245, 247, 252, 253, 254, 258 El-Niño 454 ELSDB, see European Lake Status Data Base Elysium 497 embedded sections 77 emergence 78 Empetrum 253 endorheic 423 enrichment factor (EF) 214-220, 339 ephippia 164, 166 epiglacial 425 epilimnion 394 epilithon 128, 137 epipelic 279 epipelon 128 epiphyte 439 epiphyton 128 epipsammic 279 epipsammon 128 epishelf 423-425, 433 equilibrium line 242 equipotential 487 equivalent dose 42 Erebus 119 Ericaceae 96, 322, 356

INDEX erosion 64, 330, 338-339, 358, 389, 398, 412, 426, 494 Estonia 165 Eucalyptus 435 eukaryote 438, 439 Eunotia 258 Eurasia 90, 349-375, 381-412 Europe 95, 169, 172, 217, 222, 225, 227, 229, 322, 386, 398, 407, 412, 396, 419, 476, 494, 500 European Lake Status Data Base (ELSDB) 371 European Plain 371 Eurycercus 337 Euryhapsis 171 eustatic 441 eutrophic 328 eutrophication 139, 141-142, 143, 166 evaporation 104, 136, 188, 193-196, 199, 200, 276, 328, 331-333, 367, 420, 427, 433, 448-451 evaporation/inflow ratio (E/I) 192, 193, 202 evaporite 76 evapotranspiration 726 excavator 429 excess 210Pb 20 exhaust 230 exobiology 494-500 exposure 432 extra-terrestrial 493, 494-55 extreme environments 475-500 exuviae 170 facies 56-57, 64-69, 138, 255, 487 Faeroe Islands 165, 320, 323, 324 fairy shrimp 439 Farewell Lake 286, 288 fecal pellets 61 feedback 270, 420, 421, 457 feedback mechanisms 118, 138 feldspar 41, 437 fen 402 Fennoscandia 258, 352, 381-412 Fennoscandian Ice Sheet 375 Figurnoye Lake 441 Filinia 162 fines 60

547 Finland 130, 161, 165, 166, 169, 172, 211, 222, 227, 228, 372, 381412 fire 108 fish 140, 141-142, 164, 166, 169, 210, 303, 324, 426, 427 fjord 423, 436 Flaherty Island 24 flatworms, see Turbellaria flocs 61 flood 55, 57, 64, 137, 200-202, 224, 226, 279, 434, 499 floodplain 279 fluoranthene 224, 225 flushing 198 fluvial 55, 64-67, 69 70, 494 fluxes 220 fly ash 226, 227-231 focusing 431 fog 322 Fog Lake 38, 39, 43, 246, 257 Foged, N. 120 Fond-du-Lac 200 food chain 426 Foraminifera 439, 492 forensic science 144 forest 401 forest-tundra 277, 292, 351, 352, 356, 359, 361, 363, 397 Former Soviet Union and Mongolian Lake Status Data Base (FSUMLSDB) 371-374 Fort Chipewyan 200 Fosheim Peninsula 242-244 Fox Lake 221 Foxe Basin 242, 244 fractionation 31, 367, 433, 434 Fragilaria 132, 247, 253, 255, 295, 296, 327, 330-331, 354-356, 358, 359, 363, 398 Franz Josefs Land 119 Frazer Lake 291 freeze coring 429 freshet 64, 80 Frobisher Bay 244 frustule 123 Frustulia 358 fry 141

548 FSUMLSDB, see Former Soviet Union and Mongolian Lake Status Data Base FT RAMAN spectroscopy 438 fungi 426 fur seal 141-142 Gale 499 Gallivare 227 Gamburtsev Subglacial Mountains 489 gas bubbles 421 gas exchange 476 gaseous phase 217 gases 423, 426, 480 gastropod 333 GCM: see general circulation models gelifluction 55 gemmules 162 gene 490 general circulation models (GCM) 2 geochemistry 104, 105, 282, 330, 350, 353, 389, 391, 434, 511 geochronology 19-45, 511 geology 371, 427 geomorphology 371, 421, 422, 427, 451 geophysical 489 geophysical characteristics 72-73 geothermal 487, 490 Germany 172 GISP2, see Greenland Ice Sheet Project Two glacial 80 glacial history 244 glacial lakes 58-64 glaciation 120, 275, 356, 358, 359, 420 glacier 55, 58-64, 211, 217, 242, 251, 253, 256, 322 443, 483, 487, 490, 492 Glacier Bay 281, 287, 399 glaciofluvial 433 glacio-fluvial deposits 55 glacio-isostatic 69 glacio-isostatic rebound 385, 412 glacio-isostatic uplift 244 glacio-lacustrine sedimentation 61

INDEX glacio-marine 492 Global Meteoric Water Line (GMWL) 190-191, 194, 197 Glyptotendipes 361 GMWL, see Global Meteoric Water Line gneiss 272 gold (Au) 226 Gomphonema 333 GPR, see ground penetrating radar grain size 77, 387, 433, 477 graminoid 359, 362 granite 242, 272, 432 graticule 108 gravel 68, 69 gravity corer 212, 428 grazer 129, 436 grazing 330 Great Bear Lake 225, 276 Great Britain 366 Great Lakes 295 Great Slave Lake 221, 295 green algae, see Chlorophyta greenhouse gas 2, 420, 447, 448-451 greenhouse models 494 greenhouse period 492 Greenland 33, 95, 96-97, 119, 127, 136, 144, 160-179, 210, 215, 219, 220, 227, 229, 247, 249, 319-323, 325-342, 370, 419, 420 Greenland Ice Core Project (GRIP) 312, 420 Greenland Ice Sheet Project Two (GISP2) 319, 420 Greenland mummies 144 grey water 139 GRIP, see Greenland Ice Core Project ground ice 494 ground penetrating radar (GPR) 485 groundwater 55, 73, 277, 387, 390, 391, 393, 399, 496, 497 guano 443 Gusev 499 Gyraulus 162 Gyrinus 168 gyttja 255, 256, 391

INDEX HA, see humic acid habitat 127-128, 130, 175, 278, 281, 282, 409 half life 20, 29, 31 Hall Peninsula 249 Halozetes 439 Hannaea 136 Hanson Lake 221 hardwater effect 30, 31, 33 Hawk Lake 225, 230 Hazen Plateau 254, 255 HCB 213, 222 head capsules 169-173, 280 Heard Island 9-10 heath, see Ericaceae heavy metals 210 Hebes Chasma 497 Heinrich events 422, 457 Hekla 339 helicopter 230 helium (He) 484-485, 489 hematite 492 Hemiptera 160 heptachlorobiphenyls 224 heptachloroborane 213 herb 96, 102, 356, 359 Hesperian units 494, 496 Heterotrissocladius 170 hexachlorobiphenyls 224 Hg, see mercury hiatuses 98 Hidden Lake 443 high closure lakes 137 High Pressure Liquid Chromatography (HPLC) 334, 436-438 high-resolution 211-212 Holocene 25 Holocene Thermal Maximum (HTM) 275, 298, 412 homopycnal 59, 61 Hooker, J.D. 119 Hope Bay 431 Horseshoe Island 442 hot spot 489 hot springs 476 HPLC, see High Pressure Liquid Chromatography

549 HTM, see Holocene Thermal Maximum Hudson Bay 24, 230, 302, 134-135, 220, 291 Hudson Strait 301 humic acid 36, 249, 432 humidity 352, 393 humus 398 Hustedt, F. 120 Hydrachnidae 173 hydrate 493 Hydrobaenus 170, 363, 366 hydrocarbon 224 hydrofluoric acid 162 hydrogen 421, 484 hydrogen sulfide 435 hydrology 187-203, 359, 366, 372, 393, 401 hydrolysis 366, 369 Hydroporus 168 hydrosphere 421 hydrothermal 497 Hydrozetes 173 Hymenoptera 160 hypersaline 76, 245, 476 hypersalinity 425, 426 hypolimnion 78, 334, 382 hypopycnal flow 43 Hypsithermal 298, 340, 444-448 IAS, see inorganic ash spheres IBP, see International Biological Programme ice 54-56, 80, 245, 294, 325, 381, 384, 418, 419-420, 432, 456 ice block lake 485-486 ice calving 58 ice cap 242, 323, 419 ice core 5, 6, 90, 127, 129-133, 319320, 331, 372, 420, 421, 427, 433, 447, 451-454, 457, 489 ice covered lakes 499 ice crystals 226 ice dam 423 ice dammed lake 4, 496 ice damming 58 ice jam 137, 195, 200 ice pan 69

550 ice rafting 64 ice rafting facies 69 ice rafting processes 69 ice scour lakes 384 ice sheet 108, 176, 325, 328, 421, 425, 431, 440-448, 488, 489, 490, 492, 493 ice shelve 251, 423, 425, 444 Iceland 120, 165, 166, 171, 320, 324, 325, 330, 332, 338, 340, 370, 487 Iglutalik Lake 98 Îles Crozet 9 Îles Kerguelen 10 Ilyocypris 163, 167 imaging 433 Imitavik Lake 24-25 immigration 340 inclusions 487 indenopyrene 224 Indian Ocean 447 indifferent 128 inductively-coupled plasma mass spectrometry (ICP-MS) 44 industry 448 infiltration 194 Inglefield Land 322 inorganic ash spheres (IAS) 227-231 insect 29, 249 insolation 447 instrumental data 421 insulation 255 interflows 61 interglacial 171, 258, 427, 98 International Biological Programme (IBP) 139, 247 International Polar Year (IPY) 511 interstitial 213 Inuit 6, 140, 142-143 Inuvik 92 invertebrate 159-179, 337 ions 276 IPY, see International Polar Year iron (Fe) 213, 215, 218, 361, 366, 391, 434 iron oxides 22 isolation 177-178, 446 isostasy 421, 440-448

INDEX isostatic 120, 438, 446 isostatic rebound 324, 423, 426 isotherm 8, 9 isothermal 382, 394, 412 isotope 104, 142-143, 166, 173, 187203, 255, 276, 282, 284, 286, 288, 289-292, 293, 331, 333, 350, 354, 362, 366-361, 393, 421, 433-434, 438, 482, 483, 484, 489, 492 isotope tracers 196-202 isotopic enrichment 191-193 isotopic fractionation, see isotopic partitioning isotopic partitioning 188 Isiurqutuuk Lake 300, 303 James Ross Island 164, 439, 443 Jan Mayen Island 12, 320, 321, 322 Java 128 Jemis Lake 200-202 Jierstivaara 372 Joe Lake 106 jökulhlaup 55 Jovian 494 Jupiter 494 Jurasee Lake 442 Jurassic 427 K, see potassium Kanent’yavr 371 Kangerlussuaq 215, 322 Kap Farvel 327 Kap København Formation 156, 171 Karelia 371 Karluk Lake 166, 281, 289-292 Kazan River 294 Keewatin 302 Kekerturnak Lake 257, 260 Kerguelen Islands 120 kerogen 196 kettle 387 kettle pans 425 kinetic effects 191, 192 King George Island 423, 438, 442, 443, 446

INDEX Kirisjes Pond 438 Kiruna 227 Kiteschsee Lake 442 Khibiny Mountains 356 Kjeldahl nitrogen 280 Kobuk 288 Kodiak Island 281 Kola Peninsula 138, 213, 218, 219, 227, 229, 213, 351, 366, 367, 370, 371, 381-412 Kolyma River 195 Komsomolskaya 447 Kråkenes 37, 166, 173 Krummholz 300 Kusawa Lake 221, 224 Kovdor 371 Labaz Lake 359 laboratory methods 92-93 Labrador 102, 109, 119, 138, 133, 167, 242, 271, 272, 275, 280, 297-303 Labrador Current 322 Labrador Sea 251 Labyrinth 487 Lac Ste. Therese 224, 225 lagoon 293 Lake 31 328, 330, 340 Lake A60 286 Lake Arresjøen 224, 225 Lake Athabasca 200 Lake B12 103 Lake Boeckella 438 Lake Bonney 434, 483 Lake C2 255 Lake Chuna 396, 403 Lake East Bonney 34 Lake EC1 103 Lake Ellasjøen 222 Lake Ferris 425 Lake Fryxell 434-435, 452, 481, 483, 484 Lake GB2 102 Lake Hazen 217, 242-244, 255 Lake Hendry 300 Lake Hoare 34, 35, 431, 432, 433, 438, 452, 482, 483, 484 Lake House 486

551 Lake Imandra 381 Lake Inari 381 Lake Jamylimujaganto 371 Lake K2 300 Lake Kachishayoot 134-135, 299, 300 Lake Kuttanen 387, 389, 391 Lake Laberge 213, 221, 224, 226 Lake LB1 103 lake level 276, 281, 371, 374, 384, 389, 390-393, 386, 413, 483, 484 Lake LR1 103 Lake LT1 103 Lake Missoula 496 Lake Mitrofanovskoe 371 Lake Mývatn 165, 166, 171, 337 Lake of the Pleistocene 284 Lake Oksana 300, 427 Lake Ontario 213-214, 425 Lake Paanajärvi 371 Lake Saanajärvi 228-229, 409 Lake Saglek 301 Lake Sarsjön 391 Lake SFL6 327, 328 Lake SFL4 327, 333 Lake Soldatskoje 399, 401 Lake SS2 325, 334, 337 Lake SS6 341 Lake Tasirlaq 300 Lake Tenndammen 221, 224 Lake Tibetannus 198-200 Lake Toskaljavri 398, 400, 407 Lake Tsuolbmajavri 394, 400, 402, 407 lake type 160 Lake Untersee 425 Lake Vanda 434, 444, 481, 483, 487 Lake Vida 34, 425, 486 Lake Vostok 425, 456, 488, 489, 490, 491, 493 Lake West Bonney 34 Lake Wilson 483 Lake Yarnyshnoe-3 367, 369, 370, 372 Lallemand Fjord 443, 446, 447 Lama Lake 363 laminations 62-64, 66, 68, 71, 255, 386, 482

552 Landnam 339 Lapland 130, 169, 381, 386, 387, 390, 393, 394, 399, 402, 405, 406, 409 larch (see Larix) Larix (larch) 95, 102, 359, 363 Larsemann Hills 229, 423, 426, 437, 435, 436, 438, 439, 442, 446 larvae 169 Last Glacial Maximum (LGM) 108, 249, 283, 358, 429, 440, 447, 452, 454, 480 Laurentian Ice Sheet 244, 275, 292 Law Dome 444, 449 Leach Lake 298 leaching 303, 389 lead (Pb) 213, 214, 218-220, 277, 339 lead shields 164 lead-206 (206Pb) 213, 219-220 lead-207 (207Pb) 213, 21-220 lead-210 (210Pb) 20-29, 73-75, 212, 271, 294, 303, 304, 431-432, 511 leaves 160 LEL, see Local Evaporation Lines Lena River 195, 351-359, 367, 369 Lena River Delta 137, 359, 361, 362 367, 370, 371, 375 Lepidoptera 160 Lepidurus 163, 164, 337 LGM, see Last Glacial Maximum LIA, see Little Ice Age Libby half life 31 lift-off mats 480-481 light 127, 381, 395-399, 476 Lille Saltsø 328, 333, 337 Lindeman Lake 221 Linnévatnet 39 lipid 196, 436 liquid nitrogen 429 Little Atlin Lake 221, 224 Little Ice Age (LIA) 258, 294, 330, 394, 395, 409, 412, 449 littoral 253, 256, 381, 390 Livingston Island 423, 431 Livingstone corer 91-92, 428-429 LMWL, see Local Meteoric Water Line

INDEX Local Evaporation Lines (LEL) 191-192, 194, 196, 197, 202 Local Meteoric Water Line (LMWL) 190-191, 196 Lochnagar 213 loess 333, 504 logistics 422, 509 LOI, see loss-on-ignition Longyearbyen 222, 224, 227-228 loss-on-ignition (LOI) 104, 105, 106, 215, 218, 298, 340, 387-390 Lotic Index 136 low closure lakes 137 Lower Dumbell Lake 229, 230 luminescence age 42 luminescence dating 432 lutien 334 Lymnaea 162 Macdonald Islands 10 Mackenzie Delta 102, 105 Mackenzie Mountains 105 Mackenzie River 195, 200, 224-225, 279 Mackenzie River Delta 137, 169, 220, 225, 279 Mackenzie River District 275 Macquarie Island 10 macrofossil 30, 36, 37, 38, 169, 200, 284, 286,, 327, 356, 358, 366, 386, 391, 435 macrophyte 108, 165, 166, 279, 337 magma 497 magnesium (Mg) 400, 492 magnetic 330, 398, 432, 433 Mallomonas 124 Malmberget 227 mammals 140, 141-142, 251, 426 mandibles 163, 164 manganese (Mn) 213, 366 manganese oxides 22 Manitoba 92, 272 Marguerite Bay 442 marine 119, 244-245, 251, 370, 421, 423, 425, 429, 431, 442-443, 446 marine cores 109 marine limit 331

INDEX Marine Oxygen Stage (MOS) 249 marine-lacustrine transition 120, 137138, 441 Mariner 9 494 Marion Island 10 Mars 334, 476, 494-500 Mars Global Surveyor (MGS) 497 Mars Orbiter Laser Altimeter (MOLA) 499 mass accumulation rates 73-75 mass spectrometry (MS) 436-438 mastodon 144 Mastogloia 333 mat 432, 480, 487 MCA, see Medieval Climate Anomaly McGill University 247 McLeod Bay 295, 296 McMurdo Dry Valleys 29, 31, 33-34, 423, 425, 426, 429, 431, 432, 433-434, 435, 444, 451, 477490, 486 McMurdo Sound 434, 492 MDE, see mercury depletion event Mean Water Line (MWL) 196 Medieval Climate Anomaly (MCA) 394 megafossils 391 Meighen Island 172 Meli Lake 286 melt 483, 485 melt water 59, 60, 77, 489, 433 Melville Peninsula 242 mercury (Hg) 138, 210, 212-214, 214218, 229, 330 mercury depletion event (MDE) 212 Meretta Lake 139 Meridion 136 meromictic 245, 324, 334, 423, 434, meromixis 69, 70-71, 335, 354, 449 Mesocricotopus 170 mesohalobous 327 mesohaline 333 mesotrophic 328 metals 210, 212, 407 metamorphic 272 meteor 489 meteoric 227

553 meteorite 493, 497 methane 279 MGS, see Mass Global Surveyor microbial loop 426 microevolution165 micrometeorite 493 micropaleontological 251 Micropsectra 170, 337, 366 microstructure 435 Microtendipes 170, 361 Middendorf Lake 367, 369 midges, see Chironomidae Milankovitch 447, 457 Milankovitch cycles 420 Milankovitch forcing 251, 492 military 229 Mimosaceae 435 mineral matter 387, 389, 398 mineral particles 490 mineralogy 76-77, 434 minerogenic 429 mire 389 mite 173, 439 mixing 278, 328, 412, 477, 483 moat 55, 485, 500 model 448, 455-456, 490 modeling 270 models 446 moisture 281, 359, 367 MOLA, see Mars Orbiter Laser Altimeter molecular taxonomy 426 Mollusca 162 mollusks, see Mollusca molybdenum (Mo) 78 monimolimnion 483 monomictic 62, 382, 394, 412 moraine 252, 384, 425 morphology 426 morphometry 72, 130, 372, 382 MOS, see Marine Oxygen Stage moss 31, 33, 39, 108, 129, 160, 211, 227, 249, 335, 381, 386, 425, 426, 431, 494 Mössbauer spectroscopy 434 Mould Bay 219-220 Moutonnée Lake 423, 426 MS, see mass spectrometry

554 multiple aliquot additive dose (MAAD) 42, 43-44 multiple regression 109 mummies 144 Murman 354 museum collections 157 muskox 140 MWL, see Mean Water Line mycosporine 438 Myriophyllum 327 Myrtaceae 435 myxoxanthophyll 335 NADW, see North Atlantic Deep Water Nanedi Vallis 496 Nansen, F. 119 NAO, see North Atlantic Oscillation NARO, see Nordic Arctic Research Programme natal streams 141 National Swedish Environmental Monitoring Program 222 Navicula 333 Nedre Æra Svatn 171 Neogene 160, 174, 175 Neoglacial 134, 240, 244, 252, 255, 257-258, 260, 286, 288, 298, 331 Neorhabdocoela 162 Nettilling Lake 98, 244 New Zealand 435, 449 nickel (Ni) 218-219 Nicolay Lake 65, 66 Nikel 219, 227, 229 nitrogen (N) 139-140, 141, 170, 260, 289- 292, 328, 338, 366, 370, 400, 429, 434, 482 no closure lakes 137 nonachlorogorane 213 Nordic Arctic Research Programme (NARO) 324, 341 Nordland 222 Noril’sk 138, 213, 219 NORLAKE 324, 341 North America 212, 217, 269-306, 419 North Atlantic 319, 322, 339, 354,

INDEX 370, 375, 393, 394 North Atlantic Current 322 North Atlantic Deep Water (NADW) 322 North Atlantic Oscillation (NAO) 319, 322, 330 North Atlantic rift zone 339 North Base 139 North Pole 119 Northwest Territories (NWT) 277, 272, 278, 279-280, 281, 295, 362 Norway 161, 166, 171, 173, 211, 215, 218, 222, 224, 323, 354, 370, 373, 384, 400, 401 Norwegian Sea 200, 370 Nostoc 435 Notofagus 435 Novaja Semlja 119 nunatak 431 Nunatak Lake 229 nursery lake 141, 281 nutrients 170, 175, 245, 278, 282, 287, 289-292, 295, 327-328, 338, 356 ocean circulation 322 octachlorobiphenyls 222 oil 227, 230 okenone 334 Oligocene 492 Oligochaeta 162 oligotrophic 67, 211, 245, 256, 276, 295, 296, 324, 327, 330, 335, 381, 386, 476 Oliveridia 363, 366 Oncorhynchus 141 open drainage 200 Operation Hazen 247 optical 299 optical environment 395-399 optically stimulated luminescence (OSC) 20, 41-44, 249 optically stimulated luminescence dating 432, 441 optics 398 optima 128, 129, 170 orbital forcing 420

INDEX orbital inclination 493 orbital induced 356 organic 402 organic carbon 213, 215, 218 organic factions 29-30 organic matter 78, 328-330, 334, 340, 361, 387-390, 412, 433, 482 organochlorines 220-224 Oribatida 173 oriented lakes 56 origin of life 476 OSC, see optically stimulated luminescence Oscillatoria 325, 435 Oscillatoriaceae 435 oscillaxanthin 335 Ostracoda 163, 167, 199, 261, 276, 282, 288 ostracode, see Ostracoda Øvre Neådalsvatn 224 outflow channels 496, 497 outwash plain 385 oxidation 490 oxygen (O) 366-371, 382, 398, 421, 433, 434, 482, 483 oxygen isotopes 173 oxyhydroxide 366 Oxyria 98 ozone 4, 398 Pacific Ocean 141, 272, 275 pack rat middens 5 PACT see Paleoecological Analyses of Circumpolar Treeline PAGES, see Past Global Changes PAH, see polycylic aromatic hydrocarbons PALE 4, 282 Palearctic 169 Paleoecological Analyses of Circumpolar Treeline (PACT) 282 paleoepilimnia 483 Paleoenvrionmental Arctic Sciences (PARCS) 282 paleohydrology 118, 187-203 paleomagnetics 75 paleomagnetism 432

555 Paleozoic 242 Palmer Deep 447 paludification 396, 402 palynology 89-110 Pangnirtung 255 PAR, see photosynthetically active (available) radiation Paracladius 361, 366 paraglacial 55 Parakiefferiella 361, 363 Paraphysomonas 125 PARCS, see Paleoenvironmental Arctic Sciences Parry Channel 244 particle size 60-61, 493 particles 210, 226-231 particulate organic carbon (POC) 30, 31, 34 particulates 398 Pasmlambina 371 Past Global Changes (PAGES) 4 Patricia Bay 98 PbS 22 PCA, see principal components analysis PCB 138, 301 PCBs 212, 213, 214, 220-224 PCDD, see polychlorinated dibenzop-dioxins Peace Athabasca Delta 195, 200-202 Peace River 200, 221 Pearse Valley 486 Peary Land 322 peat 98, 144, 255, 333, 356, 359, 396398, 402 Pechora River 359 Pediastrum 125, 175, 249, 325, 327, 435 pelagic 165 penguin 431, 438, 442, 443, 446, 451 pennate diatoms 136 Penny Ice Cap 244 percussion corer 429 peridinin 438 periglacial 144 periglacial lakes 64-67 periphyton 128, 139, 253 permafrost 2, 7, 195, 270, 272, 277,

556

INDEX

305, 470 Peronia 258 persistent organic pollutants (POPs) 209-231 perylene 224 petrogenic 225 petrology 434 pH 128, 135, 138, 175, 258, 280, 303, 352, 396, 399, 400, 401-404, 412 phenanthrene 224 phenotypic 426 pheophytin 334 phosphorus (P) 139-140, 141, 170, 295, 328, 381 photography 433 photosynthesis 121, 127, 199, 425 photosynthetically active (available) radiation (PAR) 133-135 photosynthetic bacteria 435 phycobiliprotein 438 phycology 117-145 phyla 128 Phylactolaemata 163 phylotype 490 physiography 242-244, 272 physiology 510 phytoplankton 127, 324, 382, 398, 426 Picea (spruce) 95, 102, 105, 106, 292, 293, 298, 352, 356, 369, 405 pigment 126-127, 135, 261, 325, 333336, 334, 409, 436-438 pine, see Pinus Pinnularia 142, 253, 358 Pinus (pine) 95, 351, 352, 354, 356, 358, 389, 391, 402, 405 pioneer vegetation 96 Pisidium 162, 199 piston corer 428-429 planetary extremes 475-500 plankton 255, 294, 295, 296, 381, 390, 409 platform 92, 427 platinum (Pt) 381 Pleistocene 441, 492 Pleuroxus 165 Pliocene 163, 165

Plumatella 163, 174 plutonium-238 (Pu238) 431 Poaceae 98, 99, 358, 362 POC, see particulate organic carbon Podocarpus 435 Point Barrow 217 polar desert 96, 331 polar front 220 Polar Frontal Zone 9, 10 polar sunrise 217 polarized light 77 pollen 29, 89-110, 134, 160, 200, 247, 249, 255, 270, 282, 288, 293, 297-298, 327, 340, 352, 354, 356, 358, 359, 363, 372, 372, 375, 393, 435, 547 pollen transport 93 pollutants 209-231 pollution 138, 142, 295, 331, 339, 407-409, 421, 432, 439-440, 510 Pølsa 372 polychlorinated dibenzo-p-dioxins (PCDD) 222-223 polycylic aromatic hydrocarbons (PAH) 222, 224-226 Polypedilum 170 ponds 245 poplar, see Populus POPs, see persistent organic pollutants Populus (poplar) 102, 105 Porifera 161-162 post-depositional changes 212-214 Potamocypris 163, 167 Potamogeton 327 potassium (K) 41, 42, 398 Poteryanny Zub Lake 367, 369, 370, 372 precipitate 492 precipitation 136, 190, 191, 194, 196, 200, 272, 277, 282, 284, 322, 331-337, 351, 359, 367, 393, 423, 426-427, 448-451, 483, 486 predation 164, 165, 169, 281 pre-Holocene 247-251

INDEX Prince of Wales Island 10, 98, 99 principal components analysis (PCA) 352, 359, 410 production 211, 279, 297, 399 productivity 136, 165, 325, 327, 328, 329, 330, 340, 362, 381, 383, 390, 487 pro-glacial lakes 62, 500 prokaryote 118, 438, 439 Proteobacteria 490 protist 438-439 Protozoa 161, 281, 426 proximal inflow 61 Prydz Bay 443, 446 Psectrocladius 337, 361 Pseudochironomus 170 Pup Lagoon 438 pupae 169, 170 purple sulphur bacteria 334 Pyramiden 227-228 pyrene 224, 225 pyrite 434, 482 Pyrrhophyta 109, 125, 438 Qaqortoq 322 Qausuittuq 224 Qipisarqo Lake 328-330 quartz 41, 431 Quaternary 20, 160 Québec 40, 102, 103, 107, 109, 133, 134-141, 211, 220, 242, 271, 272, 275, 282, 297-303 Queen Elizabeth Islands 242-261 Queen’s Lake 102, 132, 292-293, 398 radar mapping 487 RADARSAT 488 radial dispersal 69 radiation 251 radio echo 425 radiocarbon (14C) 20, 29-41, 75, 160, 249, 271, 282, 303, 354, 386, 390, 391, 431, 432, 441-443, 446, 451, 480 radium-226 (226Ra) 20 radon-222 (222Rn) 20, 28 Raffles Sø 328-330, 331, 340 Rainbow Lake 295

557 rainfall 76, 261 raised beach 244, 421 RAMAN 438 Rauer Islands 425, 435 recycling 328 redox boundary 212 reduction 490 reflectivity 72-73 refugia 38 regressive 137 remote sensing 72 renewal 393 Renland ice core 340 reproduction 118 reservoir effect 431, 432, 444 residence time 431, 489 resolubilization 22 Resolute Bay 139, 224 resuspension 67 retene 224, 225, 226 Rhabdocoela 162, 163 rhizopod 161 Richards Island 281 Ridge B 425 rift 489 river 385 river flow 70, 136-137 river runoff 80 Rivière George 141, 211 Robinson Lake 36, 41, 43, 98 Rock Basin Lake 247, 253, 254 rock flow 59 Romulus Lake 222 Ross Ice Shelf 456 Ross Island 451 Ross Sea 423, 452 Rotifera 162, 439 rotifers, see Rotifera RUBISCO 433 runoff 188, 193-196, 363, 494, 496 Russia 137, 138, 171, 195, 211, 215, 216, 218, 219, 220 222, 270, 349-375, 381-412 Russian peat corer 429 Saanajärvi 228-229, 409 Saksunarvatn 339

558 saline 324, 334, 420, 423 salinity 69-70, 76, 136, 137-138, 160, 175, 280, 282, 331, 395, 433, 482 Salix (willow) 96, 98, 322, 362 salmon 141, 281, 289-292 salt 421, 434, 449, 476, 481, 483, 486, 492, 500 saltating load 56 sand 55, 62, 68, 253 satellite 431 satellite images 487 scales 123-127, 260 Scandinavia 77, 210, 217, 218, 220, 322 scanning electron microscope (SEM) 77 SCAR 4 scavenger 337 Schirmacher Oasis 423, 425, 436 Schuchie Lake 229 sclerotised 163 Scotland 213 Scots pine 351 SCP, see spheroidal carbonaceous particle SCUBA 481 SD, see standard deviation sea 423, 426 sea ice 2, 3, 251, 322 sea level 270, 350, 354, 420, 423, 425, 426, 440-448, 455 sea salt 408 sea surface temperature 251, 375 seabirds 222 seal 141-142, 426 sea-level 279, 285, 293 sea-level index 138 seasonality 5, 278 sea-spray 324 sediment accumulation 73, 245 sediment catcher 429 sediment focusing 28-29, 330 sedimentation 477 sedimentation rates 211 sedimentology 260, 350, 384-385, 433, 441, 511 seeds 160

INDEX seepage 399 seepage lakes 33 seismic 72, 488-489 selenium (Se) 218 SEM, see scanning electron microscope settling rates 60, 61, 80 sewage 139 shape 77 sheep 330, 338 shells 164 shield 272, 497 shoreface 279 shoreline 429 short-wave radiation 56 shrub 96, 98, 49, 322, 356, 359, 362 Siberia 7, 109, 137, 138, 161, 169, 170, 196, 215, 227, 349-375 siderite 255, 256 Sierra Nevada 362 sieves 93 Signy Island 164, 168, 229, 423, 431, 435, 439 silicate minerals 41-44 siliceous 334 silt 55, 58, 61, 62, 66, 68, 253, 256, 330, 340 Simocephalus 165 Simuliidae 171 single aliquot regeneration (SAR) 42, 43-44 Sinus Meridiani 492 Skagi Peninsula 337 Skardtjørna 25, 27, 337 slide 433 Slipper Lake 294-295, 296 slump 433 smelter 219, 227, 407-409 smoke 225 SNAP-9A 431 snow 56, 68, 127, 211, 218, 253, 322, 398, 433, 446 snow melt 391 snow pack 194 Snowball Earth 425, 456 snowline 255 sockeye salmon 141, 281, 289-292 soil 193, 245, 270, 287, 325, 327, 333,

INDEX 338-339, 369-370, 393, 396, 401, 412, 431 solar system 494 Soldatskoe Lake 356 Somerset Island 95, 98, 99, 142-143 Søndre Strøm Fjord 322, 325, 328, 331, 333, 337 South America 435, 447 South Georgia Island 10, 120, 142, 433 South Orkney Islands 142 South Sandwich Island 10, 229 South Shetland Islands 442, 444 Southern Ocean 447 Southern Oscillation 454 spade 429 species diversity 426 species richness 164-165, 405, 406 spectroscopy 434, 438 Sphagnum 108, 135, 281, 300, 356, 358 spheroidal carbonaceous particle (SCP) 215, 227-231, 339 spherules 226-231 Spiniferomonas 125 Spitsbergen 3, 25, 119, 120, 171, 323, 337 sponges, see Porifera spores 108 springs 476, 497 Square Lake 298 stability 409, 435 standard deviation (SD) 405 statistics 510 statoblast 163, 174-175 statospores, see stomatocysts Stauroforma 358 Staurosira 331 Staurosirella 132, 331 Stephanodiscus 295 steryl 438 Stevatnet Island 337 Stictochironomus 361 Stoke’s Law 60, 61 stomate 352, 354, 359, 362, 363, 375 stomatocyst 123-125, 409 Stone Age 399 Store Saltsø 328, 333, 337

559 stoves 230 stratification 67, 69, 133, 276, 278, 296, 324, 329, 334, 382-383, 394, 395, 409, 423, 426, 490 stream 171, 432 stromatolite 434 structure 333-338 stumps 105, 391 subglacial 456 subglacial lake 487-493 subsaline 331 substrate 237 Sukkertoppen 322 sulfate 361 sulfide 303, 439 sulfur (S) 255, 407, 434, 435, 482 sulfur dioxide 128 sulphate 224, 227 sulphur, see sulfur Sumatra 128 sun 420 sunrise 217 supported 210Pb 20 supra-anal plates 163 supra-glacial 425 surface sediment calibration set 128129, 169, 249, 258, 278, 331, 352-354, 359-362, 384, 394, 407, 510 susceptibility 433 Svalbard 3, 25, 39, 72, 119, 161, 171, 173, 216, 218, 221, 222, 224, 227, 228, 320, 323, 325, 330, 334, 337, 339 Svartenhuk Halvø 328 Sweden 165, 167, 188, 197, 198-200, 211, 215, 217, 222, 227, 372, 385, 386-387, 391, 401, 402403, 404, 407 Switzerland 218 syntectonic 492 Synura 125 Synurophyceae, see chrysophytes Syowa Oasis 423 Tabellaria 354 tadpole shrimp 337 talik 73

560 tar sands 200 tarn 385 Tasiilaq 322 Tasmania 449 taxonomy 510 Taylor Dome 451-452, 454 Taylor Glacier 483 Taylor Valley 452, 454, 480, 483 Taymyr Peninsula 138, 216, 359, 363 Tazimina Lake 298 tectonic 488-489, 492 telmatic 387 temperature 129, 165, 169, 170, 173, 175, 191, 200, 251, 256, 272, 276, 278, 279, 280, 281, 282, 298, 331, 338, 359, 381, 399, 409, 421, 433, 447, 449, 454, 492, 494, 510 temperature gradient gel electrophoresis (TGGE) 439 temperature inferences 135-136 tephra 339, 431, 433 tephrochronology 20 Terra Nova Bay 452 terrace 429 terrestrial sediment 70 terrigenous 433, 435 Terror 119 Tertiary 95 TES, see Thermal Emission Spectrometer testate amoebae 161 tests 161 Tête des Cirques 242 texture 77-78 TGGE, see temperature gradient gel electrophoresis Th, see thorium thecamoebians 281 thermal characteristics 393-395 Thermal Emission Spectrometer (TES) 492 thermal expansion 55 thermal inertia 256 thermal maximum 396 thermally stimulated luminescence 249 thermocline 412

INDEX thermodynamics 500 thermokarst 55, 293, 328 thermoluminescence (TL) 41 thermoluminescence dating 432 Thidriksvallavatn 330 thin sections 77 Thingvallavatn 337 thorium (Th) 42, 45, 427, 431 Thule Inuit 142-143 Tiksi Bay 359 titanium (Ti) 398 TK-20 Lake 292-293, 296 TL, see thermoluminescence TOC, see total organic carbon tolerance 128, 129 Toolik Lake 25-28, 324, 510 top-bottom approach 132, 222, 295, 296, 407 topoclimatic 243 topography 130, 242, 270, 276, 296, 496 Toronto Lake 292-293 Toskaljarvi 173, 387, 396, 398, 400, 402 total dissolved solids (TDS) 381 total organic carbon (TOC) 395-399 toxaphene 213 trace elements 78 tracers 196-202 tractors 230 training sets, see surface sediment calibration sets Transantarctic Mountains 427 transfer function, see surface sediment calibration sets transgression 177-178, 446 transgressive 137 transparency 398 tree 272 tree ring 293, 293 treeline 7, 8, 89, 90, 93, 102-108, 150, 169, 172, 272, 275, 276, 280, 352, 356, 366, 370, 389, 396, 402, 412 tributary 494 trichome 434 Trichoptera 173 triphenylene 1224

INDEX trophic levels 324 trophic status 160, 281 truck 230 Truelove Lowland 247 trunks 391 Tsulbmajavri 398, 394, 400, 402, 407 Tsuolbmajarvi 173 Tuktoyaktuk 278, 293 Tulugak Lake 246 tundra 7, 56, 95, 102, 218, 222, 277, 284, 322, 351, 352, 356, 359, 361, 363, 386, 389, 399, 409 Turbellaria 162 turbid 279 turbidity 398 turbulence 60 turnover 405 tychoplankton 128 Tyrrell Sea 303 Ugashik Lake 298 ultraviolet (UV) radiation 4, 126, 133135, 281-282, 299-300, 349, 384, 395, 398, 438, 510 underflow 60, 61, 62, 63 Ungava 102 Ungava Bay 211 Ungava Peninsula 272 United Kingdom (UK) 324, 366 unsupported 210Pb 20 Upper Wright Valley 487 U, see uranium uranium (U) 41, 42, 45, 427, 431 uranium/thorium dating 427, 431, 432 uranium-238 (U238) 20 USA 25-26, 362, 496, 510 utilidor 139-140 UWITEC corer 429 Vaccinium 253 Valles Marineris 497 valley networks 494-496 valve 123 varve 20, 24, 62-64, 75-76, 255, 261, 294, 295, 391, 511 Vatnsdalsvatn 330 vegetation 68, 242, 253, 270, 337,

561 356, 358, 361, 362, 363, 369, 370, 387, 396, 401, 412 Vestfold Hills 423, 425, 426, 435 436, 439, 441, 449, 483 Victoria Land 444, 483 Victoria Valley 34, 454, 486 Vienna Pee Dee Belemnite (VPDB) 188-189 Vienna Standard Mean Ocean Water (VSMOW) 188, 189, 367 Viking 338-339, 499 volcanic 227 volcano 322, 431, 447, 494, 496 volume 328 Vostok 420, 447, 498 VPDB, see Vienna Pee Dee Belemnite VSMOW, see Vienna Standard Mean Ocean Water Vuoskkujávri 403 WA, see weighted averaging WAIS, see West Antarctic Ice Sheet WA-PLS, see weighted averaging partial least squares Ward Hunt Ice Shelf 3-4 warm-based glaciers 61 Washington 496 water balance 188 water column 435 water level 434 Watson Lake 213, 221 Watts Lake 441 weathering 363, 389, 412, 431, 434 Weichselian 441 weighted averaging (WA) 249, 338, 352, 362 weighted averaging partial least squares (WA-PLS) 173 West Antarctic Ice Sheet (WAIS) 444, 446, 451, 455-456, 480 West Exeter Lake 194 West Greenland Current 322 wetland 118, 281, 300 whales, 142, 143 White Smoke Lake 431, 432 Whitehorse 278

562 wildlife 140-142, 210 willow, see Salix wind 68, 398 Windmill Islands 435, 443 Windmill Lake 284 Wisconsinan 244, 249, 252 WMA, see World Meteorological Association Wolverine Lake 286, 288 wood 230, 356, 362 Wood Buffalo National Park 277, 278, 279, 295 World Meteorological Association (WMA) 448 Wrangel Island 215 Wright Valley 454, 483 xanthophylls 436-438 x-radiographs 77 x-ray 433, 434

INDEX Yamba Lake 195 Yanou Lake 443 Yarnyshnoe Bay 354 Yaya Lake 220, 221, 224, 226 yeast 456 Yellowknife 92, 102, 106, 169 Yenisey River 366, 367, 369 Younger Dryas 37, 165, 178, 284286, 354, 362, 363, 433 Yukon 76, 136, 141, 213, 221, 224, 275, 278, 279, 280, 283, 284 Yukon River 225 Zackenberg Ecological Research Operation (ZERO) 324, 341 Zalutschia 361 Zapolyarny 219 zeaxanthin 334 ZERO, see Zackenberg Ecological Research Operation zinc (Zn) 213, 218 zonal index 393 zoogeography 175-176 zooplankton 281-282, 324, 426