Earth Paleoenvironments: Records preserved in Mid- and Low-Latitude Glaciers (Developments in Paleoenvironmental Research)

Earth Paleoenvironments: Records Preserved in Mid- and Low-Latitude Glaciers Developments in Paleoenvironmental Resea...

Author: L.DeWayne Cecil | Jaromy R. Green | Lonnie G. Thompson

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Earth Paleoenvironments: Records Preserved in Mid- and Low-Latitude Glaciers

Developments in Paleoenvironmental Research VOLUME 9

Earth Paleoenvironments: Records Preserved in Mid- and Low-Latitude Glaciers Edited by

L. DeWayne Cecil U.S. Geological Survey, Idaho Falls, U.S.A.

Jaromy R. Green Garden City Community College, Kansas, U.S.A. and

Lonnie G. Thompson The Ohio State University, Ohio, U.S.A.

KLUWER ACADEMIC PUBLISHERS NEW YORK, BOSTON, DORDRECHT, LONDON, MOSCOW

eBook ISBN: Print ISBN:

1-4020-2146-1 1-4020-2145-3

©2004 Springer Science + Business Media, Inc. Print ©2004 Kluwer Academic Publishers Dordrecht All rights reserved No part of this eBook may be reproduced or transmitted in any form or by any means, electronic, mechanical, recording, or otherwise, without written consent from the Publisher Created in the United States of America

Visit Springer's eBookstore at: and the Springer Global Website Online at:

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Dedication

This book is dedicated to the scientists and explorers who came before us and those that will follow us.

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Contents

Dedication

v

Contributing Authors

xi

Preface by L.D. Cecil

xv

Foreword by E.J. Steig

xix

Acknowledgments

xxiii

List of Acronyms and Abbreviations

xxv

Published and Forthcoming Titles in the Developments in Paleoenvironmental Research Book Series

xxix

Editors and Board of Advisors of Developments in Paleoenvironmental Research Book Series

xxxi

PART I: INTRODUCTION AND METHODS 1. High-Altitude, Mid and Low-Latitude Ice Core Records: Implications for our Future by L.G. Thompson 1. Introduction 2. What we’ve learned from Mid-and Low-Latitude Alpine Ice Cores 3. Important questions still to be answered 4. References 2. Methods of Mid- and Low-Latitude Glacial Record Collection, Analysis, and Interpretation by J.R. Green et al. 1. Introduction 2. Glacial Record Collection 3. Methods (Chemical and Biological) 4. Methods (Physical Characteristics) 5. Glacial Record Interpretation 6. Summary 7. References

1 3 3 5 11

17 17 17 20 28 31 33 34

viii PART II: THE CLIMATE AND ENVIRONMENTAL CHANGE RECORD OVER THE LAST 200 YEARS

37

3. The Influence of Post-Depositional Effects on Ice Core Studies: Examples from the Alps, Andes, and Altai by U. Schotterer, et al. 1. Introduction 2. Swiss Alps 3. Subtropical and Tropical Andes 4. Mongolian and Siberian Altai 5. Conclusions 6. References

39 39 42 48 53 56 57

4. Event to Decadal-Scale Glaciochemical Variability on the Inilchek Glacier, Central Tien Shan by K. Kreutz et al. 61 1. Introduction 61 2. Sample Collection and Analytical Methods 62 3. Results 65 4. Discussion 72 5. Conclusions 76 6. Acknowledgements 77 7. References 77 5. Climatic Interpretation of the Gradient in Glaciochemical Signals Across the Crest of the Himalaya by C.P. Wake et al. 1. Introduction 2. Climatological Setting 3. Methods 4. Results 5. Discussion 6. Conclusions 7. References

81 81 82 83 85 89 92 93

6. Reconstruction of European Air Pollution from Alpine Ice Cores by M. Schwikowski 1. Introduction 2. Suitable Glaciers and Ice Cores Retrieved 3. Dating and Time Period Accessible 4. Reconstructed Air Pollution Records 5. Conclusions 6. References

95 95 96 97 100 115 116

7. Glacier Research in Mainland Scandinavia by W.B. Whalley 1. Introduction

121 121

ix 2. 3. 4. 5. 6. 7. 8. 9. 10. 11.

Present Day Glaciers – An Overview Early Work on the Glaciers in Scandinavia and Historic Variations Mass Balance Measurements and Glacier Mapping Little Ice Age Glacier Extents Thermal Regimes and Coring Projects Variations in Mass Balance and Continentality Subglacial Observations Conclusions Acknowledgements References

121 124 125 127 128 130 138 138 139 139

PART III: THE CLIMATE AND ENVIRONMENTAL CHANGE RECORD OVER THE LAST 200 – 500 YEARS

144

8. Four Centuries of Climatic Variation Across the Tibetan Plateau from Ice-Core Accumulation and į18O Records by M.E. Davis et al. 1. Introduction 2. Seasonality in the Tibetan Plateau Ice Cores 3. Precipitation Sources and Influences 4. Conclusions 5. Acknowledgments 6. References

145 145 147 154 158 159 160

9. Climatic Changes over the Last 400 Years Recorded in Ice Collected from the Guliya Ice Cap, Tibetan Plateau by Y. Tandong et al. 1. Introduction 2. Calculation of Glacial Accumulation from the Guliya Ice Core 3. A Comparison between the Precipitation Variations of the Guliya Ice Cap and its Vicinity 4. ENSO Events Recorded in the Guliya Ice Core 5. Conclusions 6. References 10. Evidence of Abrupt Climate Change and the Development of an Historic Mercury Deposition Record Using Chronological Refinement of Ice Cores at Upper Fremont Glacier by P.F. Schuster et al. 1. Introduction 2. Methods 3. Results and Discussions 4. Conclusions 5. References

163 165 169 172 177 178

181 181 185 192 211 212

x 11. Variations between į18O in Recently Deposited Snow and On-Site Air Temperature, Upper Fremont Glacier, Wyoming by Naftz et al. 217 1. Introduction 217 2. Background Information and Methodology 220 3. Results and Discussion 223 4. Conclusions 231 5. References 232 Summary by L.D. Cecil

235

Index

243

Contributing Authors

Vladimir B. Aizen University of Idaho Moscow, Idaho, U.S.A. L. DeWayne Cecil (editor) U.S. Geological Survey Idaho Falls, Idaho U.S.A. [email protected] Mary E. Davis Byrd Polar Research Center The Ohio State University Columbus, Ohio U.S.A. Shawn K. Frape University of Waterloo Waterloo, Ontario, Canada Patrick Ginot The Ohio State University Columbus, Ohio U.S.A. Jaromy R. Green (editor) Garden City Community College Garden City, Kansas U.S.A. [email protected]

xii Sichang Kang Lanzhou Institute of Glaciology and Geocryology Academia Sinica Lanzhou, China Karl J. Kreutz Institute for Quaternary and Climate Studies Department of Geological Sciences University of Main Orono, Main U.S.A. Paul A. Mayewski Institute for Quaternary and Climate Studies University of Maine Orono, Maine U.S.A. Yang Meixue Key Laboratory of Ice Core and Cold Region Environment Cold and Arid Regions Environmental and Engineering Research Institute Chinese Academy of Sciences, Lanzhou 730000, China David L. Naftz U.S. Geological Survey Salt Lake City, Utah U.S.A. Ulrich Schotterer University of Bern Switzerland Paul F. Schuster U.S. Geological Survey Boulder, Colorado U.S.A. Margit Schwikowski Paul Scherrer Institute Villigen PSI, Switzerland Willibald Stichler Institute for Hydrology Neuherberg, Germany

xiii David D. Susong U.S. Geological Survey Salt Lake City, Utah U.S.A. Hans-Arno Synal Institute for Particle Physics Zurich, Switzerland Yao Tandong Key Laboratory of Ice Core and Cold Region Environment Cold and Arid Regions Environmental and Engineering Research Institute Chinese Academy of Sciences, Lanzhou 730000, China Lonnie G. Thompson (editor) Department of Geological Sciences The Ohio State University Columbus, Ohio U.S.A. [email protected] Cameron P. Wake Department of Earth Sciences University of New Hampshire Durham, New Hampshire U.S.A. W. Brian Whalley Queen’s University Belfast UK

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Preface

The earth is in a state of constant physical, chemical, and biological change on a global scale. Global environmental alterations have occurred throughout the existence of the earth and will invariably happen in the next millenium and beyond. Global change can have immediate as well as future consequences that could affect all life on earth. As a result, the importance of understanding current and potential global environmental change has radically increased. Numerous global environmental change studies are currently underway. From monitoring ongoing natural events such as earthquakes and volcanoes to delineating potential anthropogenic effects from industrial chemical fallout from the atmosphere, all studies focus on understanding the immediate and potential environmental change and monetary impacts associated with such events. The study of global environmental change caused by anthropogenic influences requires knowing how and when the influences occurred and what effects the environment will suffer. Once these are known, the resultant future climatic and environmental changes can be projected. Additionally, studies of natural climatic and environmental alterations require the knowledge of long-term historical changes in order to predict or understand future shifts. Knowledge of past changes can only be acquired by studying and analyzing preserved environmental records that act as archives of these changes. Preserved archives of past climatic and environmental conditions do exist in nature. For example, glaciers, ice caps, and ice sheets around the world can be repositories of climatic and environmental change. Ice cores from the polar regions have provided the scientific community with an unprecedented picture of past environmental change through chemical, isotopic, and

xvi stratigraphic data. High-resolution ice core records have also been obtained from high altitude sites in the tropics. However, weather patterns and climate changes affect high-latitude regions of the world differently than mid- to low-latitude areas. In addition, the majority of the world’s population, at least 85 percent, lives between 50° N and 50° S. Therefore, understanding potential environmental change in mid- and low-latitude regions is of prime importance and could be accomplished by utilizing ice cores collected from selected alpine areas. Research on temperate ice cores faces the challenge of several commonly held beliefs about ice cores in “warm” environments. First, that the influence of meltwater percolation – which tends to smooth glaciochemical variations in the glacier forming firn and snow– precludes the use of isotopic and chemical tracers. Second, that the high accumulation rates typical for temperate glaciers and ice sheets limit the length of the record to at most, a few centuries. Third, that the availability of other climate proxies, such as pollen and tree-ring records, makes temperate ice cores unnecessary. Research at several mid-latitude sites worldwide has shown that these common beliefs are not warranted. Glacial research has already proven that ice cores collected from mid-latitude glaciers preserve the isotopic record with surprising accuracy and, for some glaciers, represents thousands of years of record. In addition, ice cores archive not only natural variations in climate and the environment but anthropogenic influences introduced over the last two centuries as well. Such additional anthropogenic information can aid in distinguishing between natural and human additions to the environment and thus further refine the understanding of future global, environmental, and climate change. There is now a small army of diverse researchers worldwide turning to the archived environmental record in mid- and low-latitude ice cores to answer diverse questions from natural and anthropogenic influences on climate change to rates of glacial retardation and growth. With the advent of ultra-sensitive analytical methods such as accelerator mass spectrometry and the experiences of diverse research teams, glaciers worldwide, with their environmental records and markers locked in, are becoming accessible. These new scientific tools and their application to understanding our influence on global environmental processes are the focus of this book. In the field of glacial research and the associated global impacts on humans there is no set of handy formulas into which various parameters can be substituted to obtain answers for the complex problems facing the world’s population. This book was designed with that fact in mind. The papers collected here represent some of the leading research and methods development in the growing scientific field of documenting global climatic and environmental changes using records archived at mid- and low-

xvii latitude sites; historically, presently, and in the future. It is hoped that current researchers and students will find the introductory “how-to” methods section useful in their work. Additionally, with a good solid grounding in the methods utilized in bringing ice core records from remote, harsh environments to the laboratory for analyses and interpretation, students will be prepared to appreciate the significance of any glacial research they may find in the literature. L. DeWayne Cecil U.S. Geological Survey Idaho Falls, Idaho U.S.A.

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Foreword

Several compilations of papers on ice-core science have been published over the last twenty-five years, among them Robin’s, “The Climatic Record in Polar Ice Sheets” (Cambridge, 1983), and Oeschger and Langway’s definitive, “The Environmental Record in Glaciers and Ice Sheets” (Wiley, 1989). The focus of these books (and many others) is the subject of ice cores from polar regions, and for good reason. Polar ice cores have yielded spectacular results, including the discovery of abrupt climate change events during the last glacial period, and the record of atmospheric greenhouse gas concentrations that demonstrate unequivocally the human influence on the atmosphere. One of the consequences of the success of polar ice core science has been that the potential of ice cores in mid- and low-latitudes has been largely overlooked for much of the last two decades. This is not to imply that no one was doing research in this area. Indeed, Paul Mayewski’s University of New Hampshire team (Mayewski is now at the University of Maine) were successfully drilling in the Himalaya in the early 1980s, while Lonnie Thompson and his crew from Ohio State University were on Quelccaya ice cap in Peru in the early 1970s (they obtained the first tropical ice core there in 1983), and have drilled dozens of high-altitude, low-latitude cores since then. Yet it was arguably not until 1995, when W. S. Broecker retracted his former skepticism in an editorial piece in Nature praising the success of Thompson’s team on Huascarán (Peru) that the community at large became aware of the importance of these records: “For years he fought not only the cold condition of his field sites, but also the lukewarm reception by many of those in our field (including me, W. S. Broecker). “Obtaining the cores from Huascarán places Lonnie Thompson in

xx the ranks of our great explorers.” (W. S. Broecker, Nature 376, p. 212, 1995). This is the first book that has been devoted entirely to ice cores from mid- and low-latitudes. In the papers compiled here, researchers highlight the work they have done over the past few decades--in parallel with the work done on polar ice cores--at high altitudes on glaciers ranging from China and Tibet to South America and Africa. Why are these records important? Their heuristic value alone ought to be sufficient justification for the modest amount of funding (compared with polar ice-coring projects) that has been devoted to obtaining them. Cores from mid- and low-latitudes allow us to extend our look at the earth’s recent history, in the beautiful, high-resolution detail that only ice cores achieve, across virtually the entire globe: ice cores have now been obtained from all the continents except Australia. Even Africa, a place that rarely conjures up images of ice, has revealed its secrets through the ice cap on Kilimanjaro. Ice cores from these regions contain information that is more directly relevant to human history than are ice cores from the polar regions. Ice cores obtained from the Alps (Schwikowski, this volume) for example, demonstrate both the dramatic increase in air pollution in Europe over the last century and the reduction in such pollution since regulatory measures were put into effect in the 1970s. These records thus provide stark evidence both of our ability to change the composition of the earth’s atmosphere, and of our ability to do something about it. They also illustrate the dependence of human civilization on climate: cores from the Andes, in particular, show evidence that significant droughts coincided with the decline of major South American cultural groups. Second, these records are important for their scientific value. Ice cores from the Andes appear to be remarkably faithful recorders of conditions in the tropical Pacific, offering the potential to document the frequency and magnitude of El Niño events in the past. Such knowledge is in turn an essential component in determining the sensitivity of tropical climate variability - which dominates global climate variability--to increased radiative forcing from anthropogenic greenhouse gases. Ice cores from Asia similarly offer the possibility of better understanding variability in the intensity and timing of monsoon rainfall. Ice cores from both regions, and elsewhere, contribute to our understanding of atmospheric transport of pollutants such as heavy metals, radioactive isotopes, organic pollutants and sulfate and nitrate aerosols (the primary causes of acid rain) (Schuster et al., this volume). On longer timescales, the growing network of cores from the tropics contributes to our understanding of the causes of ice ages and of variability in the climate system on century to millennial timescales. It has been conventional to attribute climatic variability on these timescales to

xxi various factors - such as sea ice-albedo feedbacks - to features of the climatic system that are unique to the polar regions. In part because of information from tropical ice cores, there is increasing interest in the scientific community on the potential role of the tropics in climatic change, both in the past and in the future. It is my hope that readers will view the present volume not so much as a summary of work accomplished, but as inspiration for future work. There is much that remains to be done on ice cores that have already been drilled. Unlike the major drilling programs in Antarctica and Greenland, carried out by large teams of researchers from many different universities (more than 40 institutions worldwide are involved in the recent and ongoing deep drilling efforts in Greenland), the cores from mid- and low- latitudes have been obtained exclusively by small teams from just one or two universities. Consequently, only a handful of the great number of possible measurements that can be done on the ice - ranging from trace gas concentrations to rare cosmic-ray produced isotopes - have been completed on these cores. Fortunately, researchers have been careful to archive sections of ice at their home institutions, so material is still available. There are also many remaining sites to be drilled. This is especially true of temperate glaciers and ice caps - defined not by their latitude but by the presence of ice that reaches the melting point during the summer. (Glaciers that, due to their high altitude, remain below the melting point throughout the year - often called “polar” glaciers, even if they are at not in the polar regions - dominate the list of sites where mid- and low-latitude ice cores have been obtained). It is commonly believed that temperate glaciers are of limited use because meltwater formed during the summer percolates through the summer snow and erases or homogenizes the chemical information contained therein. Yet useful information may in some cases be preserved because the formation of impermeable ice layers at the end of the summer prevents infiltration. Great success has been demonstrated on the Upper Fremont Glacier in Wyoming, where seasonal variations in stable isotope signatures (a widely-used measure of temperature) are preserved (Naftz et al., this volume). Temperate glaciers with ice hundreds or thousands of years old exist throughout the high-precipitation regions of the Canadian and Chilean west coasts, in New Zealand, and elsewhere. Even longer records are preserved beneath rocky debris in numerous additional glaciers on the drier side of these ranges. Obtaining some of these records will be easy; others, particularly those in remote locations and where rock debris is a problem, will be difficult. In some cases, the scientific payoff will not be immediately apparent to all. Yet the achievements so far suggest the effort is worthwhile. And as Lonnie Thompson has warned us, the cost of waiting

xxii may be to lose these precious records entirely, as almost all glaciers in tropical and middle latitudes are disappearing rapidly. Eric J. Steig Quaternary Research Center University of Washington Seattle, Washington U.S.A.

Acknowledgments

Many colleagues reviewed all or part of this book for its technical and editorial content. We are grateful to the following: Dr. Clay Nichols (he reviewed the entire book for us!), Dr. Mitch Plummer, Dr. Tom W.D. Edwards, Dr. Anne Palmer, Dr. Jim LaBaugh, Dr. Kendrick Taylor, Dr. Gary Gill, Dr. Jim Wiener, Dr. Erik Rolan, Dr. Per Holmlind, Dr. Atle Nesje, Dr. Wilfred Theakstone, Mr. Travis McCling, Mr. Flint Hall, Ms. Linda Channel, Ms. Barbara Kemp, Ms. Kristi Moser-McIntire, and the following contributors to this volume who reviewed collected papers other than ones they were coauthors on: Dr. Karl Kreutz, Dr. Uli Schotterer, Dr. Vladimir B. Aizen, Dr. Eric Steig, Dr. Mary E. Davis, and Dr. Margit Schwikowski. Each of the contributors to this volume and the editors thank their respective Institutions and Agencies for their support during the compilation of this book. Special thanks are due the U.S. Department of Energy, Idaho Operations and the U.S. Geological Survey for their support. Special thanks are due Ms. Nola Hartgraves for her excellent job of formatting and collecting the papers into the volume you are reading. Thanks Nola!

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List of Acronyms and Abbreviations

AMS - accelerator mass spectrometry ASRL - accelerated sea-level rise BC - black carbon BPRC - Byrd Polar Research Center CDC - Climate Diagnostics Center CI - continentality index DOE - United States Department of Energy EC - elemental carbon ECM - electrical-conductivity measurements EDA - energy dispersion analysis ELA - equilibrium line altitude ENSO - El Niño/Southern Oscillation (The term ENSO will be used to refer to both a warm and a cold episode; El Niño will be used to specify a warm episode, and La Niña will be used to specify a cold episode.)

xxvi GCM - General Circulation Model GNIP - Global Network for Isotopes in Precipitation GOALS - Global Ocean Atmosphere Land System GOOS - Global Ocean Observing System IAEA - International Atomic Energy Agency IC - ion-exchange chromatography ICPRG - Ice Core Paleoclimatology Research Group IRI - International Research Institute for Climate Prediction LGM - Last Glacial Maximum LIA - Little Ice Age LICCRE - Laboratory of Ice Core and Cold Regions Environment, Cold and Arid Regions Environmental and Engineering Research Institute NAO - North Atlantic Oscillation NASA - National Aeronautics and Space Administration NICL - National Ice Core Laboratory NOAA - National Oceanic and Atmospheric Administration NSF - National Science Foundation NVE - Norwegian Water Resources and Energy Directorate OAR - (Office of) Oceanic and Atmospheric Research OC - organic carbon QC - quality control SEM - Scanning Electron Microscopy

xxvii SLP - sea level pressure SMOW - Standard Mean Ocean Water SOI - Southern Oscillation Index SST - sea surface temperature TC - total carbon UFG - Upper Fremont Glacier UNESCO - United Nations Educational, Scientific, and Cultural Organization USGS - United States Geological Survey VEI - Volcanic Explosivity Index V-SMOW - Vienna Standard Mean Ocean Water WDMRL - Wisconsin District Mercury Research Lab WMO - World Meteorological Organization WS - weather station

Other abbreviated units used in this volume: Bq - becquerel, which is equal to one radioactive transition per second o

C - degrees Celcius

cm - centimeter cm/s - centimeters per second CO2 - carbon dioxide g/cm3 - grams per cubic centimeter

xxviii hPa - hecto Pascals Ka - thousands of years ago km - kilometer km2 - square kilometers km3 - cubic kilometers m - meter masl - meters above sea level mg/L - milligrams per liter mil - one-thousandth of an inch mm - millimeter m/y - meters per year mȍ - microohms ng/L - nanograms per liter ppt - precipitation µg/L - micrograms per liter µm - micrometer weq - water equivalent į18O - delta oxygen-18 (see Sidebar #1, page 23, for an explanation) į2H - delta deuterium (see Sidebar #1, page 23, for an explanation)

DEVELOPMENTS IN PALEOENVIRONMENTAL RESEARCH BOOK SERIES http://www.wkap.nl/prod/s/DPER 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, 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, forthcoming

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, forthcoming

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

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.

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 Executive Director, IGBP-PAGES PAGES International Project Office Sulgeneckstrasse 38 3007, Bern, Switzerland e-mail: [email protected]

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PART I: INTRODUCTION AND METHODS

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HIGH ALTITUDE, MID- AND LOW-LATITUDE ICE CORE RECORDS: IMPLICATIONS FOR OUR FUTURE

L. G. Thompson

1.

INTRODUCTION

The 20th Century has seen the acceleration of unprecedented global and regional-scale climatic and environmental changes to which humans are vulnerable, and by which we will become increasingly more affected in the coming centuries. One-half of the Earth’s surface area lies in the tropics between 30o N and 30o S, and this area supports about 70 percent of the global population. Thus, temporal and spatial variations in the occurrence and intensity of coupled ocean-atmosphere phenomena such as El Niño and the Monsoons, which are most strongly expressed in the tropics and subtropics, are of world-wide significance. Unfortunately, meteorological observations in these regions are scarce and of short duration, particularly from high elevation sites. However, ice core records are available from lowlatitude, high-altitude glaciers, and when they are combined with highresolution proxy histories such as those from tree rings, lacustrine and marine cores, corals, etc., they provide an unprecedented view of the Earth’s climatic history over several millennia. This paper provides an overview of these unique glacier archives of past climate and environmental changes on millennial to decadal time scales. Unfortunately, these glacier archives of our past climate and environmental history are at risk. This is illustrated by the recent history of one tropical glacier in particular, the Quelccaya ice cap, 3 L. D. Cecil et al. (eds.), Earth Paleoenvironments: Records Preserved in Mid- and Low-Latitude Glaciers, 3-15. © 2004 Kluwer Academic Publishers. Printed in the Netherlands.

4

High Altitude, Mid- and Low-Latitude Ice Core Records

as shown by a photo of its margin in 1977 (Figure 1a) compared with one taken in 2002 (Figure 1b).

Figure -1. The The margin of the Quelccaya ice cap in (1) 1977 and in (b) 2002

Over the last 25 years the principle objective of the Ice Core Paleoclimatology Research Group (ICPRG) at the Byrd Polar Research Center (BPRC) at the Ohio State University has been the acquisition and analysis of a global array of ice cores that can provide high-resolution climatic and environmental histories which contribute to our understanding of the complex interactions within the Earth’s coupled climate system. With the help of new light-weight drilling equipment, we have achieved one of our main scientific objectives by expanding our research from the polar regions to remote ice fields on some of the highest tropical and subtropical mountains. Ice core records from mountains in Africa, South America, and China make it possible to study processes in the subtropical and tropical latitudes where human activities are concentrated. The 15 sites from where the ICPRG has retrieved high-altitude ice cores are shown in Figure 2. We utilize an ever-expanding ice core database of multiple proxy information (i.e. stable isotopes of oxygen, or (δ18 O) and hydrogen, insoluble dust, major and minor ion chemistry, precipitation reconstruction) that spans the globe in spatial coverage and is of the highest possible temporal resolution.

L. G. Thompson

5

Some of the accomplishments and challenges of our ice core research, as well as those of our colleagues in the USGS and other American institutes, along with those of Chinese, Russian, French and Swiss scientists, are highlighted in this volume.

Figure -2. Locations of sites from where ice cores have been taken by the Ice Core Paleoclimate Research Group.

The records contained within the Earth’s alpine ice caps and glaciers provide a wealth of data that contribute to a broad spectrum of critical scientific questions. These range from the reconstruction of high-resolution climate histories to help explore the oscillatory nature of the climate system, to the timing, duration, and severity of abrupt climate events, to the relative magnitude of 20th Century global climate change and its impact on the cryosphere. The information from these ice core studies complements other proxy records that compose the Earth’s climate history, which is the ultimate yardstick by which the significance of present and projected anthropogenic effects will be assessed.

2.

WHAT WE’VE LEARNED FROM MID-AND LOW-LATITUDE ALPINE ICE CORES

The first program to drill a low-latitude mountain core to bedrock was carried out on the Quelccaya ice cap in Southern Peru (14oS, 71oW) in 1983. In 2002, a subarctic alpine site, Bona-Churchill in the Wrangell Mountains in southeast Alaska, was drilled and 623 meters of ice core were recovered, with one core from surface to bedrock measuring 460 meters in length. In

6

High Altitude, Mid- and Low-Latitude Ice Core Records

between these two programs, we have recovered cores from the Tibetan Plateau (Dunde, Guliya, Dasuopu, Puruogangri), from the Andes (Huascarán and Sajama) and from Kilimanjaro in East Africa. With the exception of Puruogangri and Bona-Churchill, all the cores have been analyzed and their overall climate records have been published. Under the Puruogangri project, two ice cores that were obtained during the collaborative expedition conducted by the Laboratory of Ice Core and Cold Regions Environment, Cold and Arid Regions Environmental and Engineering Research Institute (LICCRE, formally the Lanzhou Institute of Glaciology and Geocryology) and BPRC are being analyzed for physical and chemical parameters.

Figure -3. (a) The margin of the Puruogangri ice cap, central Tibetan Plateau demonstrating the distinct seasonal dust layering. (b) An ice core from the puruogangri ice cap.

L. G. Thompson

7

Figure 3a illustrates the distinct annual dust layers recorded in this ice cap located in the center of Tibet, where each spring aeolian sand leaves an identifiable seasonal horizon that can be seen at the margin. At the margin of this ice cap, over 2000 years can be discerned by counting these layers which average about 50 centimeters in thickness. Figure 3b illustrates one of the cores recovered from the summit of the ice cap in which these annual layers can be seen. The marked enrichment in 18O in the ice over the most recent half century at this location (Figure 4), is consistent with findings from glaciers on the northeastern and southern sides of the Plateau (Thompson et al., 1989; Thompson et al., 2000b).

Figure -4. į 18 O record (by depth) from Core 1 from the Puruogangri ice cap, shown as halfmeter averages.

8

High Altitude, Mid- and Low-Latitude Ice Core Records

Low-latitude, high-altitude ice core records have revealed the nature of climate variability over both glacial and interglacial time scales, specifically over the last 25 thousand years since the Last Glacial Maximum (LGM). Two records from the South American Andes (Huascarán in Northern Peru at 9oS, 78oW and Sajama in Bolivia at 18oS, 69oW) and one fom the Tibetan Plateau (Guliya at 35oN, 81oE) extend to or past the LGM and confirm, along with other climate proxy records (eg. Stute et al., 1995; Guilderson et al., 1994), that the LGM was much colder in the tropics and subtropics than previously believed (Thompson et al., 1995; Thompson et al., 1998). The Guliya record covers over 700,000 years and is the oldest low- latitude, high-altitude record recovered as of this writing (Thompson et al., 1997). Although the LGM period was consistently colder, it was not consistently drier through the lower latitudes as it was in the polar regions. For example, the effective moisture along the axis of the Andes Mountains during the end of the last glacial stage was variable, being much drier in the north than in the Altiplano region in the central part of the range (Thompson et al., 1995; Thompson et al., 1998; Davis, 2002). In another example in Western China, the Guliya ice cap is partly affected by the variability and strength of the Southwest Indian Monsoon system, which was much weaker during the last glacial stage than during the Holocene. However, this region of the Tibetan Plateau also receives (and received) moisture generated from the cyclonic activity carried over Eurasia by the prevailing wintertime westerlies. Not only were lake levels in the Western Kunlun Shan higher than tropical lakes during the LGM (Li and Shi, 1992), but the dust concentrations in the Guliya ice core record were consistent with those of the Early Holocene when the summer Asian Monsoons became stronger suggesting that local sources of aerosols were inhibited during this cold period by higher precipitation and soil moisture levels (Davis, 2002). Tropical and subtropical ice core records during the Holocene show evidence of major climatic disruptions, specifically droughts. Major dust events, beginning between 4.2 and 4.5 ka and lasting several hundred years, are observed in the Huascarán and Kilimanjaro ice cores (Thompson, 2000b; Thompson et al., 2002a, respectively), and the timing and character of the dust spike is similar to one seen in a marine core record from the Gulf of Oman (Cullen et al., 2000) and a speleothem δ13 C record from a cave in Israel (Bar-Matthews et al., 1997). This dry period is also documented in several other proxy climate records throughout Asia and Northern Africa (see contributions in Dalfes et al., 1994). More recently, a historically documented drought in India in the 1790’s, which was associated with monsoon failures and a succession of severe El Niños, was recorded in the insoluble and soluble aerosol concentration records in the Dasuopu ice core (Thompson et al., 2000b). Another recorded

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Asian Monsoon failure in the late 1870’s (Lamb, 1982) is noticeable in the Dasuopu dust flux record, which is a parameter that incorporates both the dust concentration and the annual accumulation rate of ice on the glacier surface.

Figure -5. Composite records of decadal averages of į18 O from ice cores from (a) the South American Andes (Huascaran, Quelccaya, Sajama) and (b) the Tibetan Platteau (Dunde, Guliya and Dasuopu) from 1000 A.D. to the present. All six ice-core records are combined (c) to give a total view of variations in į18 O over the last millennium in the tropics, which is compared with the Northern Hemisphere reconstructed temperature record (d).

High-resolution records of Late Holocene variations in temperature are available from low latitude alpine ice cores. Composites of the δ18 O profiles of the South American cores (Huascarán, Quelccaya, and Sajama) and three of the Tibetan Plateau cores (Dunde, Guliya and Dasuopu) show similar trends in decadal averages over the last millennium (Thompson et al., 2003)(Figure 5). When all six of the records from these mountain glaciers are combined, the resulting composite is similar to the Northern Hemisphere temperature records of Mann et al. (1998) and Jones et al. (1998) covering the last 1000 years. As in polar ice cores, the dominant factor controlling mean δ18 O values in Andean snowfall on decadal, centennial and millennial timescales must be temperature, while on seasonal to annual time scales both temperature and precipitation influence the local δ18 O signal (Vuille et al., 2003). Not only do these comparisons argue for the important role of

10

High Altitude, Mid- and Low-Latitude Ice Core Records

temperature in the composition of oxygen isotopic ratios in glacier ice, but they also demonstrate abrupt warming from the late 19th Century through the 20th Century. Indeed, the 20th Century was the warmest period in the last 1000 years, which also encompasses the time of the “Medieval Warming”. The recent warming is recorded in tropical alpine glaciers in other ways, both within the ice core records and by the rapid retreat of many of the ice fields. In the Andes, on the Tibetan Plateau and in the East Africa Rift Valley region this climate change has left its mark. For example, the many ice fields on Kilimanjaro covered an area of 12.1 km2 in 1912, but today only 2.6 km2 remains. If the current rate of retreat continues, the perennial ice on this mountain will likely disappear within the next 20 years (Thompson et al., 2002a). The lower elevation ice caps in the Andes are experiencing damage to their seasonal δ18 O signals from the lifting of the 0o C isotherm (Davis et al., 1995). For example, not only is the seasonal isotope signal on the Quelccaya ice cap at 14o S in Southern Peru being smoothed out as meltwater percolates through the upper layers of the snow (Thompson et al., 1993), but the ice margins are undergoing rapid and accelerating retreat.

Figure -6. The history of the retreat of the Quelccaya outlet glaciewr, Qori Kalis, from 1963 to 2002.

Figure 6 documents the retreat of Quelccaya’s largest outlet glacier, Qori Kalis, which has been studied by terrestrial photogrammetry since 1978 (Thompson et. al. 2000a). The rate of this retreat from 1983 to 1991 (14 m/yr) was almost three times that between 1963 and 1983 (5 m/yr), and in

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the 2000/2001 year reached 205 m/yr. The sequence of ten maps documents the rapid and accelerating retreat whereby the glacier front is now retreating about 40 times faster than it did in the initial measurement period from 1963 to 1978.

3.

IMPORTANT QUESTIONS STILL TO BE ANSWERED

Seasonal and annual resolution of chemical and physical parameters in ice core records from the Andes Mountains have allowed reconstruction of the variability of the ENSO phenomenon over several hundred years (Thompson et al., 1984; 1992; Henderson, 1996; Henderson et al., 1999). Because the effects of El Niño and La Niña events are spatially variable, ice core records from the northernmost (Columbia) and southernmost (Patagonia) reaches of the Andes Mountains will help further resolve the frequency and intensity of ENSO, along with temperature variations long before human documentation. This will aid in placing the modern climate changes and the modern ENSO into a more comprehensive perspective. The opportunity to study ENSO teleconnections in the most recent centuries between the tropics and the northern subarctic latitudes will be provided by the completed high-resolution record from the Bona- Churchill col in the St. Elias Range of southeast Alaska. Variability of the South Asian Monsoon is also of vital importance for a large percentage of the world’s population that lives in the affected areas. The ICPRG has drilled four cores across the Tibetan Plateau that have yielded millennial-scale histories of monsoon variability across this large region and information on the interaction between the Monsoon system and the prevailing westerlies that are traced back to the Atlantic Ocean. Although marine cores from the Arabian Sea show that the intensity of the South Asian Monsoon has increased over the last four centuries (Anderson et al., 2002), the Dasuopu record from the Himalayas demonstrates that since the early 19th Century the amount of precipitation falling on this region has decreased (Thompson et al., 2000b). However, the Dunde record from the north side of the Plateau shows an accumulation history that is opposite to that in the Himalayas (Davis and Thompson, this volume). Like ENSO, therefore, the South Asian Monsoon systems have varying geographical effects. Retrieval of ice core records from the west side of the Himalayas, which is more directly affected by the SW Indian Monsoon than is the east side where Dasuopu is located, will provide a more comprehensive overview of the precipitation and temperature histories of the Himalayas as a whole. The glaciers on these mountains are vital sources of stream water for the

12

High Altitude, Mid- and Low-Latitude Ice Core Records

populations of Nepal and India during the dry seasons, and their recent disappearance should be a source of great concern for these countries. Meteorological data from around the world suggest that the Earth’s globally averaged temperature has increased 0.6oC since 1950. The El Niño year of 1998 saw the highest globally averaged temperatures on record, while 2001 (a La Niña year) was the second warmest, and 2002 (a non-El Niño year) exceeded the 2001 average temperature. The marked warmth of the last two decades has contributed to the widespread melting of low-latitude, highaltitude glaciers. During this time, the ICPRG has been monitoring the accelerating retreat of this tropical ice in conjunction with its global ice core drilling and climate reconstruction program. Some of the clearest evidence for major climate warming underway today comes from the tropical glaciers, recorded in both the ice core į18O records and in the drastic retreats of both total area and total volume. The rapid retreat causes concern for two reasons. First, these glaciers are the world’s “water towers”, and their loss threatens water resources necessary for hydroelectric production, crop irrigation and municipal water supplies for many nations. The ice fields constitute a “bank account” that is drawn upon during dry periods to supply populations downstream. The current melting is cashing in on that account, which was built over thousands of years but is not currently being replenished. As Figure 7 illustrates, all the mountain glaciers in the tropical latitudes are currently retreating, as are most glaciers in middle and subpolar latitudes. The land between 30oN and 30oS, which constitutes 50 percent of the global surface area, is home to 70 percent of the world’s population and 80 percent of the world’s births. However, only 20 percent of the global agricultural production takes place in these climatically sensitive regions.

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Figure -7. Map demonstrating the current condition of the Earth’s cryosphere. Dark shading depicts regions where glacier retreat is underway, while lighter shading depicts where glacier advance is occurring. Shading over land between 30oN and 30oS indicates the tropical regions where most human activity is currently concentrated

The second concern that is brought about by the disappearance of these ice fields is that they contain paleoclimatic histories that are unattainable elsewhere and, as they melt, the records preserved therein are forever lost. These records are needed to discern how climate has changed in the past in these regions and to assist in predicting future changes. For example, climate records from low-latitude ice cores give us a view of widespread abrupt climate events in the tropics that occurred about 4,200 years (discussed above) and about 5,200 and B.P. (Thompson et al., 2002). These climatic “excursions” may have been catastrophic for early civilizations in Europe, Northern Africa, around the Mediterranean and in the Middle East. However, the geographic scale of these changes, their causes, the thresholds that triggered them, and the role of the tropical hydrological system are still mysteries. The multi-proxy analyses of ice cores is proving to be invaluable in the determination of long-term changes in the magnitude and the frequency of ENSO and monsoon variability over the past approximately 20,000 years, and how these systems may be related. However, high priority needs to be given to more precise dating of paleoclimate records, those from ice cores as well as other sources. In addition, paleoclimate data and climate models must be integrated to better understand past processes in climate change. The mechanisms responsible for the current global warming remain a topic of much debate, but the scientific evidence verifies that the Earth’s globally averaged surface temperature is indeed increasing, although at

14

High Altitude, Mid- and Low-Latitude Ice Core Records

varying rates. For example, over the extensive, elevated area of the Tibetan Plateau, the warming trend is amplified and is accelerating with increasing altitude (Thompson et al., 2000a). Although it is important to place this current warming and tropical glacier retreat into a long-term perspective, it is nevertheless undeniable that global water resources are at risk, and mountain glaciers and their unique climate histories are disappearing at an ever increasing rate. In order to preserve these records that are essential for examining how climate changes, we must accelerate the rate at which ice cores are being recovered and focus on those ice fields that are at the greatest risk. Thus, the loss of tropical mountain glaciers and the climate histories they contain presents an urgency to recover these archives.

4.

REFERENCES

Anderson, D.M., Overpeck, J.T., and Gupta, A.K., 2002, Increase in the Asian Southwest Monsoon during the past four centuries. Science 297, 596-599. Bar-Matthews, M., Ayalon, A. Kaufman, A., Wasserburg, G.J., 1999, The Eastern Mediterranean paleoclimate as a reflection of regional events: Soreq Cave, Israel. Earth and Planetary Letters 166, 85-95. Cullen, H.M., deMenocal, P.B., Hemming, S., Hemming, G., Brown, G.H., Guilderson, T., Sirocko, F., 2000, Climate change and the collapse of the Akkadian empire: Evidence from the deep sea. Geology 28, 379-382. Dalfes, H.N., Kukla, G., and Weiss, H., Eds., 1994, Third Millennium BC Climate Change and Old World Collapse. Springer, Berlin, 723 pp. Davis, M.E., Thompson, L.G., Mosley-Thompson, E., Lin, P.N., Mikhalenko, V.N., and Dai, J., 1995, Recent ice-core climate records from the Cordillera Blanca, Peru. Annals of Glaciology 21, 225-230. Davis, M.E., 2002, Climatic interpretations of eolian dust records from low-latitude, highaltitude ice cores. PhD Thesis, The Ohio State University. Davis, M.E. and Thompson, L.G., 2003, Four centuries of climatic variation across the Tibetan Plateau from ice-core accumulation and δ18 O records. In: Earth Paleoenvironments: Records Preserved in Mid and Low Latitude Glaciers (L.D. Cecil, J.R. Green, L.G. Thompson, eds.) Kluwer, New York. Guilderson, T.P., Fairbanks, R.G., and Rubenstone, J.L., 1994, Tropical temperature variations since 22,000 years ago: modulating inter-hemispheric climate change. Science 263, 663-665 Henderson, K.A., 1996, The El Niño-Southern Oscillation and other modes of interannual tropical climate variability as recorded in ice cores from the Nevado Huascarán col, Peru. M.S. Thesis, The Ohio State University. Henderson, K.A., Thompson, L.G., and Lin, P.N., 1999, Recording of El Niño in ice core δ O 18 records from Nevado Huascarán, Peru. Journal of Geophysical Research, D 104, 31,053-31,065. Jones, P.D., Briffa, K.R., Barnett, T.P., and Tett, S.F.B., 1998, High-resolution palaeoclimatic records for the last millennium: interpretation, integration and comparison with general circulation model control-run temperatures. Holocene 8, 455-471.

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Lamb. H.H., 1982, Climate History and the Modern World. Methuen, London, 387 pp. Li, S. and Y. Shi. (1992) Glacial and lake fluctuations in the area of the west Kunlun Mountains during the last 45,000 years. Annals of Glaciology 16, 79-84. Mann, M.E., Bradley, R.S., and Hughes, M.K., 1998, Global-scale temperature patterns and climate forcing over the past six centuries. Nature 392, 779-787. Stute, M., Forster, M., Frischkorn, H., Serejo, A., Clark, J.F., Schlosser, P., Broecker, W.S., and Bonani, G., 1995, Cooling of tropical Brazil (5oC) during the last glacial maximum. Science 269, 379-383. Thompson, L.G., Mosley-Thompson, E., and Arnao, B.M., 1984, El Niño-Southern Oscillation events recorded in the stratigraphy of the tropical Quelccaya ice cap, Peru. Science, 226, 50-52. Thompson, L.G. and 9 others, 1989, Holocene-Late Pleistocene climatic ice core records from Qinghai-Tibetan Plateau. Science 246, 474-477. Thompson, L.G., Mosley-Thompson, E., and Thompson, P.A., 1992, Reconstructing interannual climate variability from tropical and subtropical ice-core records In: El Niño: Historical and Paleoclimatic Aspects of the Southern Oscillation (H.F. Diaz and V. Markgraf, eds.) Cambridge University Press, Cambridge, pp. 295-322. Thompson, L.G. and 6 others. 1993, “Recent warming”: ice core evidence from tropical ice cores with emphasis upon Central Asia. Global and Planetary Change 7, 145-146. Thompson, L.G. and 7 others, 1995, Late Glacial Stage and Holocene tropical ice core records from Huascarán, Peru. Science 269, 47-50. Thompson, L.G. and 9 others, 1997, Tropical climate instability: the last glacial cycle from a Qinghai-Tibetan ice core. Science 276, 1821-1825. Thompson, L.G. and 10 others, 1998, A 25,000 year tropical climate history from Bolivian ice cores. Science 282, 1858-1864. Thompson, L.G., 2000a, Ice-core evidence for climate change in the Tropics: implications for our future. Quaternary Science Reviews 19, 19-36. Thompson, L.G., Yao, T., Mosley-Thompson, E., Davis. M.E., Henderson, K.A., and Lin, P.N., 2000b, A high-resolution millennial record of the South Asian Monsoon from Himalayan ice cores. Science 289, 1916-1919. Thompson, L.G. and 10 others, 2002, Kilimanjaro ice core records: Evidence of Holocene climate change in tropical Africa. Science 298, 589-593. Thompson, L.G., Mosley-Thompson, E., Davis, M.E., Lin, P.N., Henderson, K., and Mashiotta, T.A., 2003, Tropical glacier and ice core evidence of climate change on annual to millennial time scales. Climatic Change 59, 137-155. Vuille, M, Bradley, R.S., Healy, R., Werner, M., Hardy, D.R., Thompson, L.G., and Keiming, F., 2003, Modeling δ18O in precipitation over the tropical Americas, Part II: Simulation of the stable isotope signal in Andean ice cores. Journal of Geophysical Research, 108, 10.1029/2001JD002039.

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METHODS OF MID- AND LOW-LATITUDE GLACIAL RECORD COLLECTION, ANALYSIS, AND INTERPRETATION Jaromy R. Green, L. DeWayne Cecil, and Shaun K. Frape

1.

INTRODUCTION

The credibility of collecting, analyzing, interpreting and “dating” midand low-latitude glacial ice samples depends primarily on the methods used. Due to the wide variety of characteristics among mid- and low-latitude glaciers (such as location, altitude, precipitation, size, movement, and archived history), the methods used to study such glaciers vary from site to site and must be chosen with care. Of benefit to many of today’s researchers is the ability to build upon the knowledge of past studies of mid- and lowlatitude glaciers. In the 1970s and 80s, research was performed on such places as the Quelccaya Ice Cap in the Andes mountains of Peru (Thompson, 1984, 1985, 1986), the Dunde Ice Cap within the Tibetan Plateau (Thompson, 1988), and the Colle Gnifetti in the Alps (Oeschger et al., 1977). Past and present studies of these glacial sites allow current day researchers to more correctly identify and use those methods that are logistically, scientifically, and statistically sound to research additional midand low-latitude glacial sites. The following chapter details some of the methods used in current research.

2.

GLACIAL RECORD COLLECTION

2.1

Selecting a Glacial Site

Collection of glacial samples typically begins with the first and most important method of mid- and low-latitude glacial studies, which is the selection of a suitable glacial site. Initial reconnaissance of a desirable glacial site includes collecting surface snow samples for several consecutive years to determine if the mid- and low-latitude glacier meets the criteria to effectively archive geochemical and isotopic information in the ice. Some of these criteria are location, altitude, thickness, and topography. 17 L. D. Cecil et al. (eds.), Earth Paleoenvironments: Records Preserved in Mid- and Low-Latitude Glaciers, 17-36. © 2004 Kluwer Academic Publishers. Printed in the Netherlands.

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Methods of Mid- and Low-Latitude Glacial Record Collection

The first of these conditions, location, is perhaps the most important. A glacier that is often accessed and traversed tends to suffer from human contamination that could destroy the archived information stored in the ice. In contrast, remote glacial systems that are impacted minimally by human activity provide a more accurate record of environmental changes. In addition to minimal human impact, the preservation of environmental signals in ice at mid- and low-latitudes can only be accomplished if the glacier is at sufficient altitude to minimize meltwater percolation. Percolation, the process of meltwater infiltrating down through the glacier, can dampen or even wipe out the chemical and isotopic signals stored in the ice. Sufficiently high altitude can minimize or prevent problems with percolation, thus preserving the environmental signals in the ice. The preservation of signals in glacial ice is of greatest use when environmental signals, both natural and anthropogenic, can be observed in the ice. Generally, for mid- and low-latitude glaciers, “seeing” natural environmental changes coincides with ice thicknesses of 100 m or more; the greater the ice thickness, the larger the environmental picture of the past that can be reconstructed from the archived data in the ice. Anthropogenic changes to the environment, on the other hand, are typically observed in ice thicknesses of less than 100 meters (m). Slow glacial movement preserves environmental signals in glacial ice much better than fast movement. The ideal condition for slow glacial movement, low-angle topography of the underlying bedrock, prevents ice mixing that occurs with fast moving glaciers underlain by steep topography. The less that the ice is mixed, the easier it is to view stratigraphic layers, measure the chemistry, determine concentrations of isotopes, and to ascertain time markers in the ice. This leads to a more accurate chronology of the ice core, which ultimately increases the confidence that can be placed in reconstructed evidence of environmental change as obtained from the ice core.

2.2

Sampling

Once it is determined that the mid- or low-latitude glacier meets the requisite criteria, then an expedition to obtain deeper cores from the glacier usually follows. Due to the remoteness and high altitudes of most ice-coring sites, the equipment used to penetrate or drill down into the surface of the glacier must be specialized, rugged, and portable. For shallow cores, on the order of 20 to 30 m, lightweight hand-operated augers can be used relatively easily. However, deeper cores ranging from 30 m to thousands of meters require other types of drills. Choosing the correct method of drilling at a

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mid- or low-latitude glacial site can literally mean the difference between success and failure. For warmer mid- or low-latitude glaciers, with ice that is at or just below the freezing point, thermal drills can be employed that melt down through the glacial ice. Such thermal drills can be operated by generator or by solar power and are fairly compact in size. One example of a warm glacier is the Upper Fremont Glacier, Wyoming, U.S.A, where a thermal drill was used in 1991 and again in 1998 to retrieve surface-to-bedrock ice cores (see Naftz et al., this volume). If the glacier is colder, with ice that is well below freezing, then mechanical drills must be used to drill down into the glacier, even though such drills are often large and difficult to transport to the selected site. Like thermal drills, mechanical drills can be operated by solar or generator power. Many mid- or low-latitude glaciers greater than 5000 meters above sea level are cold glaciers simply because of their elevation.

2.3

Transport and Storage

Once the processes of drilling and collection of ice from a glacier are finished, the ice must be transported away from the site. Transportation of ice from extremely remote sites (such as glaciers located in Kyrghyzstan, Nepal, Peru, or Tibet) can only be successful if the logistics are worked out ahead of time: glacial ice provides little information if it melts while it is being transported. To this end, insulated boxes, dry ice, refrigerated trucks, and refrigerated airplanes are arranged before the expedition commences so that the ice can be quickly and efficiently relocated to a storage facility. Several universities in the United States have storage space for ice cores, including the University of Maine, the University of New Hampshire, and The Ohio State University; the location of storage depends on who sponsors the glacial expedition. In addition to ice-core storage at universities, one of the foremost ice-core laboratories where ice cores from glaciers and ice sheets around the world are stored, is the U.S. National Ice Core Laboratory (NICL), located in Lakewood, Colorado, U.S.A. All ice cores are stored at a constant temperature of –30 °C to preserve the physical and chemical information stored in the ice.

2.4

Processing

Before analyses of glacial ice can begin, the ice must be processed. This refers to the steps taken to prepare the ice-core sections for the various types of analyses. Glacial ice is always processed in a cold environment to maintain the frozen state of the ice. For example, NICL maintains a walk-in

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Methods of Mid- and Low-Latitude Glacial Record Collection

processing room at –24 °C. Plastic, powderless gloves are worn at all times when handling or processing ice. Horizontal and vertical stainless steel bandsaws are used to cut 1-m length ice cores into sections as small as 4 cm. These individual sections of ice are then packaged, sealed, carefully labeled, and shipped frozen to various destinations for further processing, and/or analyses. Once the ice has arrived at a destination, additional processing steps are typically performed, depending on the type of analyses to be done; such steps are designed to minimize contamination. All processing begins with wearing powderless gloves and a facemask at all times while handling the sections of ice to prevent contamination by touch or breath. Next, in order to remove any accidental contamination on the outside of the ice, the sections are typically scraped with a stainless steel microtome. However, according to a new procedure that has been developed by researchers at the University of New Hampshire that eliminates the need to scrape the ice samples, a section of ice is placed on top of a heated band of metal. This process melts out the uncontaminated center section of the core directly into a container and the outer “rind” of the core is discarded. This new procedure allows the remaining ice to be melted immediately. In contrast, a scraped sample is rinsed with greater than 18 Mohm deionized water, allowed to melt in a closed container for a short period of time, swished around in its own meltwater to remove any potential contamination still present (meltwater is poured off), and finally the remaining ice is covered again and allowed to fully melt. Chemical processing of the melted glacial samples is often a required additional step, depending on the type of analysis to be performed. When performing accelerator mass spectrometry (AMS) measurements on glacial samples, for example, there is often only a very small amount of the isotope of interest in the sample. In order to have enough mass to create a target for analysis, a chemist must add stable elemental carrier to the sample (Cecil et al., 1999)

3.

METHODS (CHEMICAL AND BIOLOGICAL)

A significant amount of information can be stored in mid- and lowlatitude glaciers. Due to complications at these sites, any single piece of information gleaned from a glacier may not represent an accurate picture of the climate and/or environment of the past. For example, meltwater percolation can dampen isotopic and chemical signals stored in the ice, making it difficult to reconstruct temperature and deposition records. Subsequently, every method of analysis, whether chemical, biological, or

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geophysical, is used in conjunction with all the other methods to correctly identify and apply the various types of information stored in a glacier. The following sections describe many chemical and some biological methods that have been or are being developed to look at and understand the record archived in mid- and low-latitude glaciers.

3.1

Oxygen-18/Oxygen-16 Ratios

The ratios of naturally occurring oxygen isotopes in precipitation samples can provide significant information relative to temperature, altitude, storm track, distance from source water, and evaporation. The ratio of the two isotopes used, oxygen-18 to oxygen-16, is determined by using some type of mass spectrometer, usually either gas or solid source, to analyze the desired sample. The resultant ratio is divided by a standard oxygen ratio maintained by the International Atomic Energy Agency IIAEA), multiplied by 1,000, and abbreviated as į18O. The į18O value can be positive or negative with respect to the standard and is reported in units of permil (or per thousand). See Sidebar 1, page 23. The values of į18O in glacial ice vary according to season. Ice representative of winter has a more negative į18O value than does ice representative of the summer season. An ice core with a well-preserved į18O signal should show the seasonal į18O oscillations (summer to winter), with more negative į18O values representing cooler air temperatures. As a result of the measurable changes in the į18O values, past changes in air temperatures can be reconstructed for the time period spanned by the icecore record (see Naftz et al., this volume).

3.2

Electrical Conductivity (Acidity)

The ultimate purpose of direct current electrical-conductivity measurements (ECM) performed on glacial ice is to assist in the determination of ice-core chronology. ECM accomplishes this by measuring the acidity in the ice. Once the acidity is known, the seasonal/summer dust layers (for layer counting) can be identified. Additionally, volcanic events in the ice that act as time-markers can be identified (see Schuster et al., this volume). The ECM technique uses a pair of electrodes (greater than 2,000 volts of potential difference) spaced 1 centimeter (cm) apart. The electrodes are moved down the continuous ice core at a constant velocity of 5 centimeters per second (cm/s) and the resultant current through the ice core is measured every millimeter (mm), providing a high-resolution profile of the acidity in the core. The ECM profile easily identifies singular volcanic events as

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Methods of Mid- and Low-Latitude Glacial Record Collection

spikes in the log, allowing the chronology of the mid- or low-latitude ice core to be refined (Schuster et al., 2000 and Schuster et al., this volume).

3.3

Major-Ion Chemistry

Major ions are considered to be chloride (Cl-), nitrate (NO3), sulfate (SO4), sodium (Na), magnesium (Mg), calcium (Ca), ammonium (NH4) and potassium (K). Determination of the concentrations of Cl-, NO3, and SO4 is accomplished by ion-exchange chromatography (IC). In the past, concentrations of Na, Mg, and Ca were determined by inductively-coupled plasma emission spectroscopy. But now, due to recent advances in lowlevel detections, concentrations of these three major ions are typically determined by IC. Concentrations of NH4 and K are also determined by IC (see Schwikowski et al., this volume). Measurements of major ions are conducted in order to see how the chemistry of precipitation at glacial sites changes; the chemistry can change due to specific events that are natural or human-induced. Natural events such as volcanic eruptions and large forest fires sharply increase the concentrations of Cl-, NO3, and SO4 in the atmosphere. These constituents spread around the globe and are deposited and preserved at glacial sites. Anthropogenic effects (biomass burning, acid rain, etc.) can also increase concentrations of these same constituents in the atmosphere. These major ions are then deposited on glaciers where they become time markers and can aid in refining the chronology of ice cores.

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Sidebar 1. Delta Notation for Reporting Stable Isotope Data The absolute measurement of isotopic ratios is a difficult analytical task and, as a result, relative isotopic ratios are measured as a matter of convention (Toran, 1982). For example, the oxygen-18/oxygen-16 (18O/16O) ratio (R in equation 1) of a sample is compared to 18O/16O of a laboratory prepared standard using the following equation: į18O = (Rsample/Rstandard – 1) x 1,000,

(1)

where, Rsample = 18O/16O in the sample, Rstandard = 18O/16O in the standard, and į18O = relative difference in concentration, in units of parts per thousand (permil). Delta 18O (į18O in equation 1) is referred to as delta notation and is the value reported by isotopic laboratories for stable isotope analysis. In equation 1, R is generally used to refer to the ratio of the heavy (or less abundant) to the light isotope (e.g. 18O/16O, 2H/1H). Delta 2H (į2H) is analogous to į18O where the ratio 2H/1H replaces 18O/16O in Rsample and Rstandard. The standard used for determining į18O and į2H in water originally was Standard Mean Ocean Water (SMOW) as defined by Craig (1961). The standard used in the chapters in this book is the Vienna Standard Mean Ocean Water (V-SMOW) that has been prepared by the IAEA. If the į18O or į2H in a water (or melted ice/snow) sample contains more of the heavier isotopes (18O or 2H) than the reference or standard material (the delta value of the standard by convention is zero), they have positive permil values and are referred to as heavier than the reference material, or as being enriched in the heavier isotope. Conversely, if the samples contain more of the lighter isotopes (16O or 1H) than the reference material, they have negative permil values and are referred to as lighter than the reference material, or as being depleted in the heavier isotope. For example, a į18O value of -17.2 can be referred to as lighter than V-SMOW or depleted in 18O relative to V-SMOW. Once the reference material has been specified, it is assumed by convention that all values are reported relative to it unless otherwise indicated. Because V-SMOW reflects the average isotopic composition of the ocean, and because of the nature of isotope fractionation processes, ҏį18O and į2H values of precipitation are always negative. The same terminology for discussions of į18O and į2H relative to V-SMOW can be applied to different samples of precipitation that have different values. For example, if two samples of precipitation have į2H values of -132.8 and -149.4, then the sample with the value of -149.4 can be referred to as lighter than the sample with a value of -132.8. In a similar fashion, the sample with the value of -150.5 is depleted in the heavier isotope relative to the other sample. Armed with this information, researchers can determine source areas for precipitation (snow) that has become glacial ice (see Kreutz et al., this volume for an example). This information in turn can be used to better understand local and regional climatic and geochemical fallout patterns.

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Methods of Mid- and Low-Latitude Glacial Record Collection

The other major ions (Na, K, Mg, NH4, and Ca) arrive at glacial sites in increased concentrations by way of: 1) seasonal dust loading to the surface of the glacier, and 2) volcanic eruptions, large forest fires, and other natural events. Such increased concentrations allow for more accurate counting of annual dust layers as well as assisting in identifying singular time markers in the ice in order to refine the chronology of the ice core. Additionally, these major ions also may indicate paleoclimate change, since colder periods, such as the Little Ice Age (LIA, see Sidebar 2, page 25), are typically windier and dryer, leading toward increased deposition of dust.

3.4

Heavy Metals

Heavy metals such as mercury (Hg), lead (Pb), plutonium (Pu), and uranium (U) naturally exist in the Earth’s environment in small concentrations. With the advent of modern-day industrialization as well as nuclear practices, however, these metals have been released in concentrations larger than the natural background levels. Because these metals can accumulate in the environment through biological uptake processes and can be potential health hazards, it is vital to understand the processes by which they are transported through and deposited in mid- and low-latitude environments as well as to identify quantities of these metals that may have been introduced into the environment. Studies of Hg concentrations in mid- and low-latitude ice were not performed until recently. Schuster et al. (2002) selected samples from the 1991 ice core collected from the Upper Fremont Glacier for Hg analysis. Chemical processing was performed on the samples to assure that all of the Hg species were oxidized. The Hg concentrations were measured by dual amalgamation cold vapor atomic fluorescence spectrometry (USEPA Method 1631, 1999) with a method detection limit of 0.025 ng/L (Schuster et al., 2002). Measurements for Pb, radioactive Pu and U (the other heavy metals potentially archived in mid- and low-latitude glacial ice) have not yet begun.

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Sidebar 2. The Little Ice Age Scientists have postulated that in recent history there was a global climatic change known as the Little Ice Age (LIA). During the period from about 1450 A.D. to 1850 A.D. air temperatures around the world were much cooler than they are today, especially in the Northern Hemisphere. There is substantial evidence that a fundamental change in air temperatures occurred during this period (see Naftz et al., Schuster et al., Thompson, and Whalley in this volume for examples). Inferential evidence abounds in the recent glacial record, in tree rings, and in human records such as documentation of an increased incidence of landslides, avalanches, and floods associated with increased precipitation and cooler air temperatures. Additionally, tax records during this period reflect the human consequences of glacier advances and the effect on surrounding landscapes (Whalley, this volume). There are significant disagreements in the scientific community on when the cooling trend started and why. There is evidence that the cooling trend started at different times in different parts of the world and lasted for centuries. Most of the LIA occurred before the Industrial Revolution in the 1800s and therefore, scientists believe that this climatic change was not due to the burning of fossil fuels but that it had natural causes. Among the postulated causes put forth by scientists, the two most often mentioned with some supporting evidence are a very slight decrease in the sun’s output of energy (about a one-quarter of one percent decrease) reaching the surface of the earth during this time and increased volcanic activity injecting a veil of sun-blocking aerosols into the atmosphere and blocking the sun’s rays (Schuster et al., this volume). The debate continues today over the importance of these two causes in climate change and some scientists add the possibility of a shift in the currents in the world’s oceans as a major driver resulting in the LIA. It seems probable that the earth’s climate was influenced by several factors that lead to this global cooling period.

3.5

Cosmogenic/Anthropogenic Isotopes

A variety of natural and anthropogenic radioisotopes exist in the environment. Radioisotopes that are created in the earth’s upper atmosphere are termed cosmogenic. Upon transfer to the hydrologic or geologic environment, these radioisotopes can potentially be used as tracers of natural and anthropogenic processes. Some cosmogenic radioisotopes that are of interest are beryllium-10 (10Be), carbon-14 (14C), silicon-32 (32Si), sulfur-35 (35S), chlorine-36 (36Cl), and iodine-129 (129I).

26

Methods of Mid- and Low-Latitude Glacial Record Collection

Cosmogenic radioisotopes are naturally created in the upper atmosphere in small quantities by the process of primary or secondary cosmic rays interacting with atmospheric atoms. Cosmic rays include, but are not limited to, high-energy protons, neutrons, muons, alpha particles, gamma rays, and neutrinos. The cosmogenic radioisotope formed as a result of this interaction depends on the energy and type of the cosmic particle and the element in the atmosphere that is interacted with. One example is the formation of the radioisotope 36Cl. The dominant production mechanism of 36Cl is by the thermal neutron capture of chlorine35. Alternatively, argon-39 can capture a thermal neutron as well and then release an alpha particle to form 36Cl. These production mechanisms follow the reactions below. 35

Cl(n,Ȗ)36Cl

39

Ar(n,Į)36Cl

(1)

Radioisotopes of anthropogenic origin can be transferred to the atmosphere as a result of nuclear activities. Some of these events include nuclear-weapons testing in the 1950s and 1960s, nuclear accidents such as Chernobyl, and emissions from nuclear power plants. Radioisotopes produced from these activities include 36Cl, 129I, cesium-137 (137Cs), U, Pu, and tritium (3H). Cesium-137 is a radioisotope that is strictly anthropogenic (fission product) and is not created in the upper atmosphere by the collision of cosmic rays with atmospheric particles. Natural and anthropogenic radioisotopes in the upper atmosphere can circulate for days, months, or years. Chlorine-36, as an example, has a residence time in the upper atmosphere of about 2 years (Synal et al, 1997). Radioisotopes are then transferred, by atmospheric circulation processes, to the lower atmosphere where they have a residence time of a few weeks. Wet and dry deposition then proceed to wash the isotopes out of the atmosphere and into the hydrologic and geologic environment. Deposition of cosmogenic radioisotopes on the surface of the earth is not uniform. The Earth’s magnetic field, atmospheric dynamics, and precipitation rate all have significant effects on where concentrations of radioisotopes will be deposited. For example, 36Cl has been shown to have larger deposition in mid- and low-latitudes than in low-or high-latitude regions of the earth (Bentley et al., 1986). Local and regional atmospheric circulation patterns, elevation above sea level, and distance from oceans also affect deposition of cosmogenic radioisotopes. Once isotopes are deposited in the hydrologic or geologic environment, they can be re-located, re-suspended, or archived. When re-suspension and

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re-location of cosmogenic radioisotopes are minimal in select mid- and lowlatitude glacial environments, such glaciers often represent excellent archives of cosmogenic radioisotope deposition (Cecil et al., this volume) Direct counting methods have allowed concentrations of some of the isotopes (3H, 14C, 35S, 137Cs, and 210Pb) to be determined for half a century or more. But it wasn’t until the advent of AMS in the late 1970s that extremely small concentrations of isotopes with long half-lives could be detected and measured with accuracy. Some of these cosmogenic isotopes that are routinely measured now include 10Be, 14C, 36Cl, and 129I. Many additional isotopes are of interest, such as 32Si, Pu and U isotopes, but the necessary AMS techniques to measure small concentrations of these isotopes in environmental samples have not been refined. For a detailed discussion of a typical AMS process see (Sharma et al. 2000).

3.6

Carbon-14

Carbon-14 is an isotope of carbon (half-life of 5,730 years) that is produced naturally in the upper atmosphere of the earth at a constant rate. Every biological organism, in the process of living, intakes a constant amount of 14C until it dies. At death, the 14C present in the organism begins to decay according to the equation A = A0e-Ȝt

(2)

where A is the current specific activity in the sample due to 14C decay, Ȝ is the decay constant for 14C, A0 is the 14C specific activity of the sample at the time of death, and t is the time since death. Occasionally, organic material can become incorporated into glacial ice. But to apply the method of 14C dating to such organic matter requires the assujption that: 1) the organic material was incorporated into the snow and ice at the actual time of death; and that 2) the initial concentration of 14C in the plant or animal material (A0) is well known and is independent of time, geographic location of the sample, and species of plant or animal. Carbon14 dating of organic matter in glacial ice can provide additional timemarker(s) that aid in refining the chronology of the ice core. A new method of measuring 14C concentrations in samples by AMS has been developed in recent years (Currie et al., 1985). The method, as with many other AMS methods, requires much smaller sample size and significantly less counting time than conventional counting methods. Now, laboratories around the world actively use AMS for 14C measurements. One example of recent and ongoing research into using AMS for 14C analysis is by the Paul Scherrer Institute, located in Zurich, Switzerland, that has

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Methods of Mid- and Low-Latitude Glacial Record Collection

developed an AMS setup for 14C analysis that is small enough to fit onto a laboratory bench.

3.7

Microbiology

The Earth’s atmosphere, replete with microbiological organisms, transports such organisms to all parts of the world, including glacial sites, by way of atmospheric circulation. Microbial cells deposited in polar environments have been shown to survive (respire) in a preserved state for extended time periods and to even colonize microhabitats within layers of snow (Karl, et al., 1999). Few microbiological studies, however, have been conducted on mid- and low-latitude glacial ice. Microbial cell populations in glacial ice may actually preserve a record of atmospheric circulation patterns, land use, and biogeographical conditions near the deposition site. This could occur if established populations within the ice adapted to local ice geochemistry that was reflective of atmospheric conditions during or close to the time of deposition. Adaptations by the dominant microbes could favor aerobic conditions if the ice retained a large capacity for dissolved oxygen. Alternatively, if sufficient organic material accumulates, then respiration could cause a depletion of oxygen, which would then favor anaerobic conditions for microbial cells to flourish in (Pedersen, 1993; Phelps at al, 1994). Additional factors, beyond aerobic and anaerobic respiration, can affect growth and development of the microbiological communities. Two of these factors are the intensity of high-altitude sunlight, which allows for radiationtolerant species to prevail, and the deposition of heavy metals, which may promote the development of metal-resistant populations. Conventional culture methods and molecular techniques based on polymerase chain reaction amplification of nucleic acids are the tools that microbiologists use to analyze ice cores for microbiological communities (Karl et al., 1999).

4.

METHODS (PHYSICAL CHARACTERISTICS)

The physical attributes of a glacier that can be studied usually begin with surface techniques. These techniques often require the collection of data representative of large volumes. Repeated surface measurements describe large-scale temporal variations in glacial structure and behavior as well as provide bulk estimates of various ice properties. The other way to study the physical attributes of a glacier is by borehole applications. In contrast to surface studies, borehole investigations focus on smaller-scale, sub-glacial

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parameters that can provide critical snapshots for the glacier as a whole. By looking at various geophysical measurements, whether surface or subsurface, the factors that influence, change, or control the glacial system can be better understood.

4.1

Radio-Echo Sounding

The technique of using radio-echo sounding to determine ice depth in polar regions has been in use for nearly half a century, but has only recently been used on warmer, mid- and low-latitude ice (0°C) due to the required modifications to the apparatus before accurate measurements could be made. A radio-echo sounding system consists of a transmitter that sends out a radio wave and a receiver that acknowledges the return of the same wave. Measurements of the elapsed time it took the wave to travel from the transmitter to the receiver are performed at various locations on the surface of the glacier by moving the apparatus. The elapsed time is then used to determine ice thickness (Welch, 2000). The technique of radio-echo sounding is also critical in mapping bedrock topography beneath the ice, calculating ice volume, and selecting drill sites.

4.2

Acoustic Televiewer Logging

One way to study the sub-surface area of a glacier is through the use of an acoustic televiewer. An acoustic televiewer is designed such that it can only be used in a borehole that is filled with fluid, thus allowing the acoustic pulses to be transmitted. The actual apparatus is a thin cylindrical instrument that is lowered into a fluid-filled borehole. The instrument is equipped with an acoustic transducer that rotates while in the borehole and emits a certain number of pulses per revolution. Pulses are transmitted through the borehole fluid, reflect off the fluid-formation interface, and return to the tool where acoustic amplitude and transit time are recorded (Zemanek et al., 1970; Morin et al., 2000). The resultant amplitude data are then converted into brightness (gray scale) or color and the resulting image appears as a planar representation of a cylindrical surface. Threedimensional cylindrical projections can be constructed by stacking the polar views. Simply put, this geophysical logging tool/technique generates a magnetically oriented image of the borehole wall.

30

4.3

Methods of Mid- and Low-Latitude Glacial Record Collection

Additional Geophysical Techniques

Other geophysical techniques are often used to characterize glaciers. These include water-level variations and video recordings in boreholes that aid in helping to characterize water movement both within and underneath glaciers (Harper and Humphrey, 1995; Fountain, 1994), borehole inclination logs and gravity measurements that help to better understand englacial deformation processes and mass balances (Hooke et al., 1992), and cross-hole electrical resistivity experiments that attempt to image drainage features (Hubbard et al., 1998).

4.4

Energy-Balance Monitoring Methods

Due to temperate locations, many mid- and low-latitude glaciers lose significant amounts of snow cover each year. Identifying the historical amounts of snow cover that remain on a mid- or low-latitude glacier is necessary to properly interpret high-resolution paleoenvironmental ice-core records. The energy balance of the snow cover (ǻQ) can generally be described by the equation: ǻQ = Snet + H + LȣE + G + M

(3)

where Snet, H, LȣE, G, and M are the net radiative, sensible, latent, conductive, and advective energy fluxes, respectively. If the value for ǻQ is negative, then the snow cover cools, preventing the melting process. Conversely, if the value for ǻQ is positive, the entire snow cover warms until it reaches a temperature of 0.0 °C, after which significant melting can occur (Marks et al., 1999). The number of models that have been developed to simulate the energy and mass balance of the seasonal snow are many and varied. One example is the model ISNOBAL, which can simulate the annual accumulation and melt of the seasonal snow cover to provide an estimate of the snow cover that remains on a glacier at the end of the melt season. In order to study the snow-cover energy and mass balance of a glacial surface, detailed measurement and monitoring of the surface climate must occur. These measurements include determination of solar and thermal radiation, air and snow temperature, relative humidity, and wind speed.

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GLACIAL RECORD INTERPRETATION

Interpretation of the records stored in mid- and low-latitude glaciers is dependant upon merging all of the data and knowledge gained from the various analytical tools. In other words, one analytical technique is not enough to interpret a paleoclimatic or paleoenvironmental record. Instead, multiple analyses provide information that can be used in conjunction to provide a more accurate basis from which to draw conclusions, or from which to interpret the records stored in the ice. Understanding variances from glacier to glacier, identifying atmospheric dynamics on local, regional, and global scales, and identifying global environmental changes are some aspects of interpreting glacial records.

5.1

Inter-Glacial Comparison of Data

A single, isolated glacial record (for example, the Upper Fremont Glacier in southern North America) is not representative of mid- and low-latitude areas worldwide. Using the combined records from mid- and low-latitude glaciers around the world is crucial to piecing together the various aspects of local, regional, and global changes that occur. However, comparing data between glacial sites that are thousands of miles apart is not always simple due to the vast differences in such things as location, prevailing wind patterns, elevation, accumulation, and ablation. These differences, in large part, can be negated if the average background precipitation flux is factored into the equation when calculating fluxes for each glacial site. The flux values can then be compared with other mid- and low-latitude sites with greater confidence.

5.2

Atmospheric Dynamics

The dynamic atmosphere is of great concern to everyone who lives on the earth. Understanding how the past atmospheric signals have changed, or shifted, can provide valuable insight toward interpreting or predicting future changes in the atmosphere and what those changes could mean to people living at mid- and low-latitudes (see Thompson, this volume). While significant studies of atmospheric dynamics of the last 100,000 years have been performed on recovered Greenland ice-cores (Mayeski et al., 1997; Meeker et al., 1997), these records represent at best only proxy records of the area of the Earth where the majority of the population live—mid- and low-latitudes. To the people who live at mid- and low-latitudes, understanding the short-term (less than 1,000 years) atmospheric dynamics

32

Methods of Mid- and Low-Latitude Glacial Record Collection

that control seasonal shifts as well as precipitation events in these areas is of paramount concern. Mid- and low-latitude ice cores, of course, can preserve information about the atmosphere and the associated dynamics at the time of deposition. In particular, such cores can preserve records of natural and anthropogenic atmospheric content, storm trajectories and wind patterns, aerosol and contaminant transport and deposition, and particulate loading. These cores can span the time period from present day to a few thousand years ago, depending on the glacial site. In recent years, atmospheric studies based on mid- and low-latitude glacial-ice research in Central Asia, Europe, and North and South America have yielded a wealth of information that is directly applicable to local, regional, and global scale studies. This volume presents a few examples of this research.

5.3

Global Environmental Change

Mid- and low-latitude glaciers are becoming recognized as valuable tools for reconstructing records of global changes to the environment (Cecil et el., 2000; Cecil and Vogt, 1997; Naftz, 1993; Schuster et al., 2000; Steig, 1999; Thompson et al., 1995, 1998). As an example, the end of the LIA is clearly preserved in the 18O isotopic record in mid- and low-latitude ice cores (Naftz et al., 1996; Thompson et al. 1986). Also preserved in the glacial ice is a continuing warming trend (Naftz et al, this volume) that parallels warming trends in alpine and high-latitude areas worldwide (Hileman, 1999; Haeberli and Beniston, 1998; Mikhalenko, 1997; and Sin’kevich, 1991). Results from analyses of mid- and low-latitude glacial ice often must be interpreted before they can be understood. Interpretation is usually based on measuring as many different aspects of the ice core as possible (physical, chemical, biological, etc.) and then using all of the information together to ascertain what (if any) changes have occurred to the environment and how those changes affect the present and future.

5.4

Glacial Record “Dating”

Any information stored in a glacial ice core is virtually useless unless the core can be dated. “Dating” in the hydrological sense is usually accomplished through the use of environmental tracers that aid in delineating flowpaths, determining water ages, allowing groundwater recharge rates to be determined, and so forth. Glacial systems, while much different from ordinary hydrologic systems, also contain environmental tracers that are preserved over time. These tracers aid in establishing chronologies for the

J.R. Green, et al.

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ice cores. An established chronology transforms otherwise meaningless glacial data into useful information about paleoclimatic and environmental changes. Environmental tracers in glacial ice are typically called event markers. Sometimes events occur, whether natural or anthropogenic, that last for a given amount of time or that are out of the ordinary. Such events that end up being preserved in glacial ice are known as event markers because they represent a specific event that happened during a specific period of time, whether a few days or a few years. Perhaps the easiest event markers to see in glacial ice are the annual stratigraphic layers that accumulate as a result of dust loading to the surface of the glacier. In the case of polar or Greenland ice cores, these layers can be counted with relative ease. Mid- or lowlatitude ice cores, in contrast, may or may not preserve annual stratigraphic layers consistently as a result of meltwater percolation that can dampen, smear, or remove the layers during any given year. Establishing a chronology (or method of dating) for mid- and low-latitude ice cores by this method is therefore ineffective and unreliable. As a necessary, alternative method of dating, other event markers preserved in mid- and low-latitude ice cores must be used to properly reconstruct the chronology of the core. Event markers can include radioactive tracers from nuclear-weapons tests, radiocarbon from insects, natural events such as volcanic eruptions, large forest fires, droughts, El Nino processes, and additional anthropogenic events (see Schuster,et al., this volume).

5.5

Collaborative Records

Information gleaned from mid- and low-latitude ice core research does not stand alone. Many other environmental records exist that support and confirm the records stored in ice. These include instrumental data, corals, tree rings, varves (lakes), ocean sediments, etc. When all of these resources are used in conjunction with ice core results does a more accurate picture of the changing environment present itself.

6.

SUMMARY

Because the majority of the word’s population live at mid- and lowlatitudes, it is vital to understand how the environment is changing, on local, regional, and global scales. Mid- and low-latitude glaciers provide a unique opportunity to look at how the environment has changed in the past, how it is changing today, and to project possible changes in the future. The study

34

Methods of Mid- and Low-Latitude Glacial Record Collection

of mid- and low-latitude glacial ice is not simple, but the process of collecting, analyzing, and interpreting such ice has been, and continues to be, refined. New techniques, lower analytical detection levels, and expanded studies have provided a solid foundation on which to base the wealth of data recorded in mid- and low-latitude glaciers, which in turn facilitates application of the information obtained to societal problems and resource evaluation.

7.

REFERENCES

Bentley, H.W., Phillips, F.W., and Davis, S.N., 1986, Chlorine-36 in the terrestrial environment, in: Handbook of Environmental Isotopes, Volume 2 (P. Fritz and J-C. Fontes, eds.), Elsevier, New York, pp. 422-480. Cecil, L.D., Naftz, D.L., and Green, J.R., 2000, Global ice-core research: understanding and applying environmental records of the past, U.S. Geological Survey Fact Sheet FS-003-00. Cecil, L.D., Green, J.R., Vogt, S., Grape, S.K., Davis, S.N., Cottrell, G.L., and Sharma, P., 1999, Chlorine-36 in water, snow, and mid- and low-latitude glacial ice of North America: Meteoric and weapons-tests production in the vicinity of the Idaho National Engineering and Environmental Laboratory, Idaho. U.S. Geological Survey Water Resources Investigations Report 99-4037, 27p. Cecil, L.D. and S. Vogt, 1997, Identification of bomb-produced 36Cl in mid- and low-latitude glacial ice of North America, Nuclear Instruments and Methods in Physics Research B 123:287-289. Craig, Harmon, 1961, Isotopic variations in meteoric waters. Science, 133:1,702-1,703. Currie, L.A., Klouda, G.A., Elmore, D., and Gove, H.E., 1985, Radiocarbon dating of microgram samples: Accelerator mass spectrometry and electromagnetic isotope separation, Nuclear Instruments and Methods in Physics Research B 12:396-401. Fountain, A.G., 1994, Borehole water-level variations and imperfections for the subglacial hydraulics of South Cascade Glacier, Washington State, U.S.A., Journal of Glaciology 40:293-304. Haeberli, W. and Beniston, M., 1998, Climate change and its impacts on glaciers and permafrost in the Alps, Ambio 27:258-265. Harper, J.T. and Humphrey, N.F., 1995, Borehole video analysis of a temperature glacier’s englacial and subglacial structure: Implications for glacier flow models, Geology 23:901904. Hooke, R.LeB., Pohjola, V.A., Jansson, P., and Kohler, J., 1992, Intra-seasonal changes in deformation profiles revealed by borehole studies, Storglaciären, Sweden, Journal of Glaciology 38:348-358. Hubbard, B., Binley, A., Slater, L., Middleton, R., and Kulessa, B., 1998, Inter-borehole electrical resistivity imaging of englacial drainage, Journal of Glaciology 44:429-434. Karl, D.M., Bird, D.F., Bjorkman, K., Houlihan, T., Shakelford, R., and Tupas, L., 1999, Microorganisms in the accreted ice of Lake Vostok, Antarctica, Science 286:2144-2147. Marks, D., Domingo, J., Susong, D.D., Link, T., and Garen, D., 1999, A spatially distributed energy balance snowmelt model, Hydrological Processes 13, 1935-1959.

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Mayeski, P.A., Meeker, L.D., Twickler, M.S., Whitlow, S., Yang, Q., Lyons, W.B., and Prentice, M., 1997, Major features and forcing of high-latitude northern hemisphere atmospheric circulation using a 110,000-year-long glaciochemical series, Journal of Geophysical Research 102(C12):26,345-26,366. Meeker, L.D., Mayewski, P.A., Twickler, M.S., Whitlow, S.I., and Meese, D., 1997, A 110,000-year history of change in continental biogenic emissions and related atmospheric circulation inferred from the Greenland Ice Sheet Project Ice Core, Journal of Geophysical Research 102(C12):26,489-26,504. Mikhalenko, V.N., 1997, Changes in Eurasian glaciation during the past century: Glacier mass balance and ice-core evidence, Annals of Glaciology 24:283-287. Morin, R.H., Descamps, G.E., and Cecil, L.D., 2000, Acoustic televiewer logging in glacier boreholes. Journal of Glaciology, vol 46, no. 155, 695-699. Naftz, D.L., 1993, Doctoral Thesis, Colorado School of Mines, 204 p. (unpublished). Naftz, D.L., Klusman, R.W., Michel, R.L., Schuster, P.F., Reddy, M.M., Taylor, H.E., Yanosky, T.M., and McConnaughey, E.A., 1996, Little Ice Age evidence from a southcentral North American ice core, U.S.A., Arctic and Alpine Research 28:35-41. Oeschger, H., Schotterer, U., Haeberli, W., and Röthlisberger, H., 1977, First results from alpine core drilling projects, Z. Gletscher. Glazial. 13(H1/2): 193-208. Pedersen, K., 1993, The deep subterranean biosphere, Earth-Science Reviews 34:243-260. Phelps, T.J., Murphy, E.M., Pfiffner, S.M., and White, D.C., 1994, Comparison between geochemical and biological estimates of subsurface microbial activities, Microbial Ecology 28: 335-349. Schuster, P.F., Krabbenhoft, D.P., Naftz, D.L., Cecil, L.D., Olson, M.L., Dewild, J.F., Susong, D.D., Green, J.R., and Abbott, M.L., 2002, Atmospheric mercury deposition during the last 270 years: a glacial ice core record of natural and anthropogenic sources. Environmental Science and Technolog, 36, 2303-2310. Schuster, P.F., White, D.E., Naftz, D.L., and Cecil, L.D., 2000, Chronological refinement of an ice core record at Upper Fremont Glacier in south central North America, Journal of Geophysical Research 105:4657-4666. Sharma, P., Bourgeois, M., Elmore, D., Granger, D., Lipschutz, M.E., Ma, X., Miller, T., Mueller, K., Rickey, F., Simms, P., and Vogt, S., 2000, PRIME lab AMS performance, upgrades and research applications, Nuclear Instruments and Methods in Physics Research B, 172:122-123. Sin’kevich, S.A., 1991, Climate warming in the Twentieth century as reflected in Svalbard ice cores: Glaciers-Ocean-Atmosphere: Interactions, International Association of Hydrological Sciences Publication No. 208, 257-267. Steig, D.J., 1999, TITLE, Eos, Trans., Amer. Geo. Union 80:S143. Synal, H-A., Beer, J., Bonani, G., Suter, M., and Woelfli, W., 1991, Atmospheric transport of bomb-produced 36Cl, Nuclear Instruments and Methods in Physics Research B 52:483488. Thompson, L. G., Davis, M. E., Mosely-Thompson, E., Sowers, T. A., Henderson, K. A., Zagorodnov, V. S., Lin, P.-N., Mikhalenko, V. N., Campen, R. K., Bolzan, J. F., Cole-Dai, J., and Francou, B., 1998, A 25,000-year tropical climate history from Bolivian ice cores, Science 282:1858-1864. Thompson, L.G., Xiaoling, W., Mosley-Thomson, E., and Zichu, X., 1988, Climatic records from the Dunde Ice Cap, China, Annals of Glaciology 10:80-84.

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Methods of Mid- and Low-Latitude Glacial Record Collection

Thompson, L.G., Mosley-Thompson, E., Bolzon, J.F., and Koci, B.R., 1986, The little Ice Age as reported in the stratigraphy of the tropical Quelccaya Ice Cap, Science 234:361364. Thompson, L.G., Mosley-Thompson, E., Bolzon, J.F., and Koci, B.R., 1985, A 1500-year record of tropical precipitation in ice cores from the Quelccaya Ice Cap, Peru. Science 229(4714):971-973. Thompson, L.G., Mosley-Thompson, E., and Arnao, B.M., 1984, El Nino-Southern oscillation events recorded in the stratigraphy of the tropical Quelccaya Ice Cap, Peru, Science 226:50-53. Toran, Laura, 1982, Isotopes in ground-water investigations. Ground Water, 20(6):740-745. USEPA Method 1631 Revision B, 1999, Mercury in water by oxidation, purge and trap, and cold vapor atomic fluorescence spectrometry, U.S. Environmental Protection Agency, Office of Water, Office of Science and Technology, Engineering and Analysis Division (4303), Washington, D.C. Welch, B.C., 2000, How does radio-echo sounding work?, University of Wyoming research webpage, (April 29, 2002); http://research.gg.uwyo.edu/iceradar/radworks.html. Zemanek, J., Glenn, E.E., Norton, L.J., and Caldwell, R.L., 1970, Formation evaluation by inspection with the borehole televiewer, Geophysics 35:254-269.

PART II: THE CLIMATE AND ENVIRONMENTAL CHANGE RECORD OVER THE LAST 200 YEARS

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THE INFLUENCE OF POST-DEPOSITIONAL EFFECTS ON ICE CORE STUDIES: EXAMPLES FROM THE ALPS, ANDES, AND ALTAI Ulrich Schotterer, Willibald Stichler, and Patrick Ginot

1.

INTRODUCTION

Glaciers and ice sheets preserve paleo-precipitation in its most direct form. However, the stored sequence of individual precipitation events and their imbedded isotopic and chemical information is influenced after deposition by various processes like wind drift and erosion, melt, sublimation, and diffusion of water vapor. The most important changes occur during the surface snow to firn transition. For instance, smoothing of a stable isotope record essentially stops when the critical density of 55 grams per cubic centimeter (g/cm3) is reached. In solid ice the thinning of the annual layers with increasing depth becomes dominant; however, this process is less important for alpine glaciers than for polar ice sheets because of the limited ice thickness and the resulting shorter time scales involved. Moreover, models are able to account for the influence of diffusion and the thinning of annual layers and to reconstruct the original seasonal variability (Johnsen 1977, 2000). In general, there is a better and more quantitative knowledge about changes after deposition on a glacier surface for stable isotopes than for chemical constituents simply because the input data and driving forces are better known. This is mainly a result of both GNIP, the Global Network for Isotopes in Precipitation (IAEA/WMO, 2001) and the advanced incorporation of water isotopes in atmospheric general circulation models (e.g. Hoffmann et al. 2000). The influence of snow drift and diffusion was first reported from isotope studies on polar ice sheets where dry snow and low accumulation caused inverse altitude effects or irregularities in the stable isotope/temperature relation (e.g. Lorius et al. 1969, Dansgaard et al. 1973). Changes in the įD/į18O relation due to sublimation have been reported recently (Satake and Kawada 1997; Stichler et al. 2001). In contrast to central polar ice sheets, melt can play an important role on mid and low latitude glaciers and ice 39 L. D. Cecil et al. (eds.), Earth Paleoenvironments: Records Preserved in Mid- and Low-Latitude Glaciers, 39-59. © 2004 Kluwer Academic Publishers. Printed in the Netherlands.

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The Influence of Post-Depositional Effects on Ice Core Studies

sheets, many of which are temperate or close to temperate. Evaporation, melt, and rain cause a re-distribution and/or a washout of isotopic and chemical tracers. Regarding the latter, the bulk of the solute in a seasonal snow cover is already removed in the early meltwater fraction (e.g. Davies et al. 1987). For stable isotopes, a fractionation takes place during phase transitions. For instance, meltwater has depleted isotope ratios as compared to the remaining snow cover, which is correspondingly enriched. Coldchamber experiments, theoretical considerations and field studies have demonstrated this effect (e.g. Buason 1972; Hermann et al. 1981; Stichler and Schotterer 2000). To avoid the influence of percolating meltwater through several annual layers, cold glaciers with temperatures well below 0°C throughout the year (in general temperatures at 10m depths are representative) are better suited for ice core studies. For mid and low latitude glaciers, appropriate conditions only exist either at high altitudes and/or high latitudes. For this reason, drilling sites necessarily are often situated in saddle or summit regions. Unfortunately, such sites are frequently exposed to extreme meteorological conditions. Unconsolidated dry winter snow depleted in stable isotopes is easily blown away and accumulates at lower altitudes where accumulation rates increase. Consequently, the snow layers at higher elevations are more enriched in stable isotopes. This may lead to an inverse altitude effect and complicates the isotope thermometry. In the dry regions of the tropical and subtropical Andes, ice core drilling sites are additionally exposed to high solar radiation that favors sublimation. A glacier may undergo substantial loss in both mass and isotopic information under such conditions. In addition, the concentration patterns of the remaining chemical tracers may be changed severely (Stichler et al. 2001; Ginot et al. 2001).

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Figure -1. Examples of post-depositional effects. Wind (left, at Colle Gnifetti, Swiss Alps), sublimation and melt (right, Cerro Tapado, Chilean Andes). Effects of wind and melt are also visible in a firn core from Colle Gnifetti in the middle. See text for an explanation.

Some of these post-depositional effects are illustrated in Figure 1. On the left, strong winds erode the smooth surface after a winter snowfall in the Swiss Alps. During long dry and stormy periods the wind shapes and hardens the surface. The wind formed sastrugi (snow dunes) may be removed as a whole or are covered by the next snowfall. Accumulation rates on the order of 80-100 cm snow (or 30 cm water equivalent [weq]) per year cannot ensure that a loss of 30 cm of snow, as in this example from Colle Gnifetti, may be sufficiently low to keep the signal of seasonal variability. Penitents are the frequent visible sign of advanced loss in snow cover by sublimation. In the dry regions of the subtropical Andes, a distinct enrichment in stable isotopes has been observed on the surface of penitents depending on their exposition to wind and sun (Peña, 1989). The meter-high snow pyramids surrounding a high-altitude camp near the drilling site at Cerro Tapado in Chile (on the right in Figure 1) are remnants of a long dry period following the 1997/98 El Niño. Near the margins of the snow field additional heating of the dark surface leads to melting. Since the meltwater cannot percolate through the frozen ground it floods the penitents and forms small frozen ponds. Due to the additional cooling effect of sublimation, even at noon, firn temperatures remain below 0°C at a depth of a few centimeters. Post-depositional effects may be recognized in the structure of the firn cores. The example from the Alps in the middle of Figure 1 shows clear signs of melt and wind influence. The latter is indicated by a sloping melt

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The Influence of Post-Depositional Effects on Ice Core Studies

horizon (dashed line). Obviously a sastrugi started to melt during a warm period and was covered by a new snowfall. Awareness of such complications, as well as the short periods of time covered by alpine glaciers, and the difficult logistical challenge of successfully recovering ice cores has certainly retarded the development of this research in comparison to that in polar regions. Nevertheless, during the last two decades, an increasing number of results from mid and low latitude glaciers has documented their importance to supplement paleo-climatic and paleo-environmental information obtained from polar ice cores. The reconstruction of atmospheric pollutants in Europe or information on the variability of paleo-precipitation in the Himalayas or the Andes from glaciochemical and isotope ice-core records has demonstrated the increasing value of such studies (e.g. Wagenbach et al. 1988; Schwikowski et al. 1999, Thompson 2000, Schwikowski, this volume). Importantly, regardless of the time scale, the specific record, or the drilling site being considered, a critical assessment of the influence of post-depositional processes on the integrity of the original precipitation sequence is needed to evaluate its continuity and how the isotopic and/or chemical information may have been modified. The following examples from our own experience span a broad spectrum of climatic regimes focused on isotope records from the surface to the firn-ice transition. Where possible, we employ direct isotope-in-precipitation and other meteorological data to quantitatively assess post-depositional effects. A number of ion-chemistry records are also discussed, demonstrating effects on individual concentration profiles. Most of our examples are from the Swiss Alps, supplemented by results from the tropical and subtropical Andes and the Russian and Mongolian Altai.

2.

SWISS ALPS

Most of the glaciers in Switzerland are temperate. Depending on slope and exposure, sufficient cold firn and ice exists only at summit regions above 3800-4000 m (Suter 1995). The first attempts to test and apply nuclear dating techniques for environmental studies started in the 1950s in temperate firn at Jungfraujoch (e.g. Oeschger et al. 1962, Ambach et al. 1969, Schotterer et al. 1977). The first ice core data from the Jungfraujoch Saddle revealed that this site is too wind-exposed for a reliable reconstruction of environmental parameters. Changes in net accumulation of more than a factor of two over a horizontal distance of 200 m demonstrated how crucial the site selection in saddle regions can be. However, the easy access by train, a fully equipped research station, plus the availability of meteorological data and nuclear fallout and pollutant monitoring have greatly facilitated field

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studies related to ice core research. Stable isotopes in precipitation have been sampled on a monthly basis since 1983. Although no official data on the amount of precipitation are published, due to excessive snow drifting, the total monthly net weight of samples collected after a precipitation event may be a fairly good indication of the seasonal precipitation distribution under field conditions. This data set affords a unique possibility to compare stable isotope data from ice cores recovered from neighboring sites and to assess quantitatively the influence of post-depositional effects on δ-values under present-day conditions.

Figure -2. Map section of the Jungfraujoch region with drilling sites and the research station where the precipitation is sampled (P) together with an aerial view of Colle Gnifetti and the uppermost part of Grenzgletscher.

The location of the drilling sites in the Jungfraujoch and Colle Gnifetti region are shown in Figure 2. Fiescherhorn Plateau (3900 m), a flat smoothsloped glacier 6 km distant from Jungfraujoch research station was drilled three times, in 1987, 2000 and 2002. Bedrock under the deepest part of the glacier at 150m depths has been reached in December 2002. Borehole temperatures in the upper part of the firn zone vary between -4°C and - 6°C. Surface melting during warm periods produces ice layers up to 10 cm thick. Due to the high annual net accumulation of 1-2 m weq, negligible percolation of melt seems to occur through underlying annual layers (Schotterer et al. 1997, 2002a). Several ice cores have been obtained from Colle Gnifetti (4500 m, Swiss/Italian border) since 1976, several of which reached bedrock (at 66 m and 124 m in 1982). The saddle, with steep crests dropping down more than 2000 m to the southeast, is also wind-exposed. It lies at the border between the cold infiltration and the infiltrationrecrystallisation zones. The glacier is frozen to bedrock, as indicated by bore

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hole temperature of -14°C. The mean annual net accumulation at the two 1982 drilling sites, approximately 150 m apart, is 32 cm and 22 cm weq respectively (e.g. Oeschger et al. 1977; Alean et al. 1983; Schotterer et al. 1985). The net accumulation increases further downstream Grenzgletscher by a factor of 6-8 at the 1994 drilling site at 4200 m altitude. Bore hole temperature variations in the firn layer from -0.4°C to -3°C indicate the possibility of percolating meltwater at this site. Below the firn/ice transition the temperature decreases to -9°C (Suter 1995)

Figure -3 į18O profiles from drilling sites in the Swiss Alps. The mean į -values of the ice cores are given together with their difference to an extrapolated precipitation value at the respective altitude (expressed by the term į precip/ice core).

The influence of post-depositional effects on the į18O variability in firn and ice cores is summarized in Figure 3. The selected records represent temperate and cold firn and cover the same interval of time. They are accompanied by an extreme example from Plaine Morte, a flat and sheltered glacier at the equilibrium line where the average mean annual net accumulation amounts to 5-10 cm weq of superimposed ice. Tritium peaks from nuclear weapon tests were used to date the records, which have been adjusted to the same vertical axis for better comparison. Two facts are obvious: the “inverse altitude effect” between Colle Gnifetti and Fiescherhorn, Jungfraujoch Saddle and Plaine Morte, and the markedly different degree of smoothing of the į-values. Only at Fiescherhorn is the į18O variability comparable to what is observed from monthly composites of

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precipitation. On Colle Gnifetti a few cycles are preserved, which probably indicate seasonal variations, whereas no trace of seasonal variability is apparent in the other two records (though acknowledging the coarse sampling resolution for Plaine Morte). The offset in mean ice core į18O from δ-values in precipitation (extrapolated for the respective altitude) is substantial. The 3.5‰ in į18O at Colle Gnifetti or Jungfraujoch Saddle, for example, are comparable to the summer/winter difference (3-5‰) in Swiss precipitation. The surprisingly small offset in superimposed ice may be explained by the absence of erosive loss, whereas Jungfraujoch Saddle is smoothed by the combined effect of wind and melt. Conservation of winter snow is responsible for the apparent full seasonal cycling of the į-values on Fiescherhorn Plateau. Probably the fresh snow contains more moisture as compared to Colle Gnifetti, promoting formation of a frozen surface crust that provides shelter against wind scour. A similar mechanism might account for an 8 m increase in altitude of the Mönch summit (Figure 1) between 1986 and 1993 (according to new measurements from the Swiss Topographic Institute), reflecting build-up of winter snow from higher temperatures and warmer precipitation events. In 1998 and 2000 two, shallow cores were drilled on the temperate Jungfraufirn and the cold Fiescherhorn Plateau, respectively. The į-records of the cores display differing reflections of the isotope variability in precipitation collected at the Jungfraujoch research station (Schotterer et al. 2002a). In Figure 4, the precipitation data from 2001 back to 1993 are plotted twice, versus time and versus accumulated weight. The accumulated weight accounts for the monthly precipitation distribution and compares better with the net accumulation given in weq for the ice core from Fiescherhorn Plateau. Arrows indicate the į-peaks in corresponding summer precipitation. As expected, in the temperate Jungfraufirn summer, į-values are sometimes missing, especially in the years 1993 and 1994. This is attributable to enhanced melting of summer layers and to occasional rain from July through September. This rain is part of the monthly composite sample at the research station, but at the glacier site, it percolates through the firn layer and is thus lost from the record. In contrast, all seasonal cycles are preserved in the record from Fiescherhorn Plateau, including much of the short-term į-variability as recorded in the monthly precipitation data. However, the distribution of this pattern with depth (or the amount in accumulated precipitation related to this pattern) is different. For example, going back in time, in the transition from winter 1999 to summer 1998 much more is accumulated (relative to the precipitation record), winter 1998 is not as pronounced as in the precipitation record, and in the transition from summer 1997 to winter 1997 the net accumulation on Fiescherhorn Plateau is again higher.

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The Influence of Post-Depositional Effects on Ice Core Studies

Figure -4. Comparison of of į18O records recovered from temperate firn (left) and cold firn (right) with į18O in composites of monthly precipitation collected at the Jungfraujoch research station.

Similar results are obtained when the variability in net accumulation on Fiescherhorn is compared to the variability in precipitation amount recorded on other high-altitude meteorological stations: The annual deviation from the 1971-1999 mean in net accumulation and precipitation amount are not correlated (Schotterer and Stichler 2002b). Despite the close agreement of the long-term δ-values on Fiescherhorn Plateau and the extrapolated įvalues in precipitation, the difference in seasonal and annual distribution of net accumulation also results in different weighting of the isotopic expression of climate and climate variability (the stable isotope/temperature relation, for example). The influence of post-depositional kinetic isotopic fractionation on the deuterium excess and the related climatic information during evaporation and/or sublimation may be evaluated on a įD-į18O diagram. On the left side in Figure 5, the į-values from the Jungfraufirn core and the monthly precipitation are compared. Neither the slope of the regression lines nor the deuterium excess values (i.e., with slope constrained to 8) differ markedly for the period between 1998 and 1993. Separate consideration of the two

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core sections, which were influenced by differing climatic conditions, yields the same result, i.e., that comparable d-excess values existed in directly sampled precipitation and net accumulation on the glacier during both periods (right side in Figure 5), verifying that the record has not been distorted by sublimation or evaporation effects and may hence be well-suited for stable isotope/climate studies. Closer consideration of the data reveals that caution is still warranted, however, since it is clear that the approximate 3‰ shift in mean δ18O values between the 1993-96 and 1996-99 intervals in the firn core, as opposed to the occurrence of very similar mean values in annual precipitation, reflects loss of summer precipitation rather than a change in temperature.

Figure -5. įD-į18O diagrams for two sections of the Jungfraujoch Firn core separated according to the different influence by evaporation, melt, and rain (left) and the comparison with the respective precipitation data (right).

Additional complications can also occur when percolating meltwater does not remove the seasonal cycle in stable isotopes, yet chemical records seem to be disturbed. In a 10 m-long firn section of an ice core drilled on Grenzgletscher at 4200 m altitude, for example, this post-depositional effect was attributed to the inflow of surface meltwater via a system of crevasses

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The Influence of Post-Depositional Effects on Ice Core Studies

that opened temporarily (Eichler et al. 2001). Figure 6 shows some selected records from this study, revealing that major ion species were leached with varying efficiencies. The known seasonal pattern for ammonium and chlorine, for instance, is more or less well preserved whereas sulfate and sodium (among others) are leached in a very specific elution sequence. The dashed line in the chlorine/sodium record represents the sea salt ratio, indicating preserved stable chlorine content and the loss of sodium.

Figure -6. Section of the Grenzgletscher core (between 1989 and 1985) influenced by meltwater. Although the į18O does not display a major disturbance, the selected records from ammonium, sulfate, and the chlorine/sodium ratio exhibit widely differing degrees of perturbation.

The elution sequence is explained by ion re-arrangement during snow metamorphism. Because no distinct enrichment could be found in the chemical records further downcore, it was assumed that the draining meltwater did not re-freeze but was discharged completely to the glacier groundwater table.

3.

SUBTROPICAL AND TROPICAL ANDES

Glaciers in the Andes are particularly important natural archives of present and past climatic and environmental changes because of the N-S orientation of this topographic barrier and its influence on the atmospheric

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circulation of the Southern Hemisphere. Between the equator and 30°S, the seasonality and the amount of precipitation changes drastically. Large differences in amount east and west of the Andean divide occur as well as a change from tropical summer precipitation to extra-tropical winter precipitation. Several ice cores down to bedrock have been recovered from this area in the last decade. From Huascaran in Peru (Thompson et al. 1995) and from Sajama (Thompson et al. 1998) and Illimani in Bolivia (Ramirez et al. 2003) climatic information has been reported as far back as to the last glacial maximum (about 15,000 years). Work on Chimborazo (Ecuador) and Tapado (Chile) is in progress. Most of this information is based on records of δ18O, dust, and some major ions. Due to the extreme climate conditions at the drilling sites and their assumed changes in the past, it is essential to assess the influence of post-depositional processes today on the recovered ice-core information. For Quelccaya in Peru it was already reported that seasonal changes in evaporation could remarkably amplify seasonal changes in δ18O (Grootes et al., 1989). The French–Swiss drilling sites on Cerrro Tapado, Illimani, and Chimborazo were drilled twice in a one-year interval. The uppermost 10-15 years should allow assessment of possible post-depositional effects over a longer period of time. Additional field experiments during drilling were carried out to study short-term changes in relation to actual meteorological conditions (Stichler et al. 2001; Ginot et al. 2001). The results from Cerro Tapado and Chimborazo document to what extent ice-core information may be altered by melt and sublimation (Schotterer et al. 2003, Ginot et al 2002). The summit glacier of Chimborazo (6250 m), situated on the equator, is cold, (core hole temperature of -4°C at bedrock) but surface melting cannot be excluded. This happened between the two drilling campaigns when a volcanic eruption covered the glacier with a dark ash-layer, sharply altering the albedo. The combined effect of sublimation and melt removed large parts of the annual accumulation. Accumulation on Cerro Tapado (5550 m) at 30°S is also strongly influenced by sublimation. This glacier is situated at the border of a dry axis that divides tropical and extra-tropical precipitation belts. At the summit drilling site the bore hole temperature of -12.5°C indicates that the glacier is frozen to bedrock. Ice cores were recovered during the 1997/98 El Niño period and one year thereafter. From meteorological records at the site it could be concluded that El Niño was followed by an extreme dry period. The snow at the surface was exposed to intense sublimation over several months. In Figure 7, isotopic and chemical results are combined from an experiment that investigated the influence of such sublimation on the accumulated snow. The isotope data plotted on the left are derived from pit samples containing the precipitation left since the El Niño event and from thin slices of hardened

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The Influence of Post-Depositional Effects on Ice Core Studies

surface snow formed by the sublimation/condensation process. The slices were removed twice a day to study changes at the surface with exposure time.

Figure -7. Sublimation experiment on Cerro Tapado. The data presented are from a highresolution snow pit (b), from thin slices of the snow surface removed twice a day (a, d, e, f). They are accompanied in (c) by samples from the first 7 cm below surface.

į18O increased by 3.5‰ within 3 days, while deuterium excess declined by 10‰ (Figure 7a). A įD-ҏį18O plot also confirms this clear sign of isotopic enrichment. In (Figure 7b) all pit samples below 7 cm are plotted and in

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(Figure 7c) the samples from the upper 7 cm together with the slices of surface snow as plotted. The slope of the pit samples matches that of the meteoric water line, while the slope of the surface samples corresponds to an evaporation line. The fact that enriched δ-values do not penetrate substantially deeper than several centimeters is consistent with evaporation pan data and modeling which show that the mass loss due to sublimation during the experiment was in the order of 2 mm weq (or 0.5 cm snow) per day. Continuous removal of the surface snow enriched in isotopes at such high rates during the experiment ensured that į-values in the deeper firn layers remained practically unchanged. However, it is obvious that persistence of excessive sublimation rates (higher than precipitation) could lead to substantial loss of climate information from intervals within an ice core. Today the annual loss by sublimation on Cerro Tapado is of the same order as the annual net accumulation (30 cm weq). Most major ions varied significantly at the surface during the experimental exposure time (Fig. 7, d, e, f). For instance, chloride concentration more than doubled and calcium increased by a factor of 5. A normalized concentration-sublimation factor allows the comparison between the different ionic species. Three groups of ions can be distinguished according to this factor. The highest enrichment is observed for species irreversibly trapped in the snow matrix and originating from wet and dry deposition as well (d), followed by a group where the enrichment in the surface layer is in proportion to the water loss by sublimation (e). Species that are present in a volatile form are released from the snow (f). Chloride has turned out to be the best quantitative chemical indicator for sublimation on Cerro Tapado, leading to efforts to reconstruct the sublimation history (Ginot et al. 2001). Documented volcanic eruptions serve as important dating tools for both polar and alpine ice cores because they may deposit a specific chemical matrix on a glacier. However, if ash layers change the surface albedo and cause surface melting, this matrix and the accumulated chemical species in the underlying firn may be also changed. In Figure 8, some records from Chimborazo before and after the volcanic eruption are plotted together, using a depth scale related to the surface in 2000. The 1999 accumulation with the depleted δ18O- values is nearly completely removed in 2000 by melt and sublimation. Black and gray stars in the upper part of the į18O-plot mark the respective horizons. The loss corresponds to approximately 70 cm weq. Below a depth of 1.2 m weq the į18O profiles are identical for both cores. Even the individual wiggles match perfectly although the two drilling sites are approximately 100 m apart. Neither the weighted mean δ-values differ, nor do the values for deuterium excess. This is a clear indication that below 1.2 m weq no further loss and/or disturbance of the accumulated isotopic

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information occurred. Any melt infiltrated without refreezing. Indeed, part of the drained meltwater evidently reached a depth of 24 m in the saddle between the two Chimborazo summits, leading to suspension of drilling because the cores were soaked with water. For chemical species, the melt process causes a re-distribution within the individual concentration profiles. On the right side of Figure 8, profiles are plotted for ammonium, sulfate, and the chlorine/sodium ratio.

Figure -8. Comparison of the the į18O record and concentration profiles of chemical species from Chimborazo before and after a volcanic eruption. Dark ash layers changed the albedo and caused intensive surface melting.

Ammonium behaves relatively conservatively. This may be explained (according to the wash-out event on Grenzgletscher, Eichler et al. 2001) by the high solubility and the position of this ion in the matrix of the snow grains. Sulfate, and especially cations like sodium or calcium, are situated more at the surface of the snow crystals and may therefore be washed-out more readily. However, since refreezing can be excluded according to the isotope balance on Chimborazo, the double-peak of sulfate (most probably originating from the volcanic eruption) needs further explanation. This is beyond the scope of this volume.

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MONGOLIAN AND SIBERIAN ALTAI

Similar to the Andes, the climate of the Altai region is also characterized by climatic extremes, in this case due to the exceedingly high contrast in seasonal temperature. Mongolia is perhaps the most continental region in the world. Temperatures in January often fall below -40°C, whilst the summer is short and hot with average July temperatures around 20°C. Precipitation is highly variable, most of it occurring during summer, with average values no greater than 15 cm per year. The glaciated area covers approximately 500 km2 and comprises small glaciers and ice caps in the Altai region of western Mongolia. Several shallow cores were drilled at the summit of the Tsast Ula ice cap (4200 m) in June 1991. Bore hole temperatures are stable at -18°C from 5 m downwards. Only a very few, thin ice layers are observed although the summer temperature at the summit sometimes approaches 0°C. This indicates that sublimation exceeds melting under these dry conditions. Meteorological data for the last 30 years are obtained from the station Khovd (1500 m altitude) at a distance of about 100 km. The yearly mean precipitation from 1970-1990 (the time period covered by the ice core) is 12 cm and varies between 6 and 21 cm. Summer precipitation dominates, constituting up to 80 percent of the annual total. The mean annual net accumulation on Tsast Ula is 25 cm weq, which points to an increase with altitude in the amount of precipitation. From the seasonal shift in δ-values and the relation of high to low values in deuterium excess, it is concluded that at least a part of the winter precipitation is normally preserved (Schotterer et al. 1997). Belukha (4500 m), the highest summit of the entire Altai mountain range, is situated in southern Siberia at the border to Kazakhstan. The glacier between the two Belukha summits was drilled at 4060 m altitude to bedrock (140 m) in 2001. Ice lenses up to 30 cm thick indicate that intensive melting can occur during summer; yet the bore hole temperature of -17.2°C suggest that the glacier is frozen to bedrock (Olivier et al. 2003). Prior to this deep drilling, an exploratory study recovered pit samples and a shallow core from Belukha west plateau at 3900 m. Pit samples were also taken near the equilibrium line of the Ak-tru glacier (3150 m) in order to evaluate the influence of post-depositional effects on the isotope records. For the Belukha area the meteorological station at Ak-kem, situated at 2000 m altitude in vicinity of the Belukha north face, reports a mean annual precipitation of 56 cm. The Siberian drilling site lies about 300 km west of the Mongolian site in the same mountain range. The nearly 5 times higher precipitation amount indicates a strong windward-lee effect. In contrast to Tsast Ula, melt has an important post-depositional effect on the glaciers of the Siberian Altai due to

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the higher air humidity. At Ak-kem station, the relative humidity during the summer months is around 70 percent. Summer dominates winter precipitation in both regions and winter temperatures are comparable (Figure 9). It is therefore of interest to examine the influence of post-depositional effects on the isotope records and to what extent the seasonality is preserved. Fortunately, at least one full annual cycle of isotopes in precipitation is available from Ulan Bator in the GNIP database (IAEA/WMO, 2001) for a comparison with the ice core data from Tsast Ula (Figure 9). Despite the restrictions (a distance of nearly 1500 km, different length and period of time), some qualitative conclusions are possible. The same slope in the δD-δ18O diagram as in Ulan Bator and the high deuterium excess exclude a major influence of sublimation on the isotope record at the Tsast Ula drilling site. Despite an altitude difference of nearly 3000 m, the precipitation in winter months in Ulan Bator is more depleted in δ-values. This is an indication that some precipitation may be lost by wind scour on Tsast Ula. However, the mean δ18O values (-15.8‰ and - 10.7‰ for Tsast Ula and Ulan Bator, respectively) translate to a reasonable lapse rate of -0.17‰ in δ18O per 100 m gained in altitude. The isotope record of the Tsast Ula ice core may therefore also represent a reasonable climate record.

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Figure -9. Long-term seasonal temperature and precipitation distribution near the two ice core drilling sites in the Mongolian and the Siberian Altai (left) and a δD-δ18O plot of precipitation and ice-core data (right).

For the Siberian Altai, the situation is more complex. In Figure 10, the isotope records from the exploratory pit studies prior to the deep drilling are plotted. On the left, the δ18O from the two pits are plotted against depth. Although the obtained Ak-tru record near the equilibrium line is smoothed by the intensive melt process, the mean δ18O value is more than 7‰ lower than the respective record from the nearly 800m higher Belukha plateau. At this site the additional wind scour removes a considerable part of the annual precipitation. Although some seasonality seems to be preserved, the average δ18O values at the Belukha saddle are comparable to the Belukha plateau (Olivier et al. 2003). Moreover, the isotope record of the remaining net accumulation in the recovered ice cores might be influenced sometimes by kinetic fractionation during post-depositional alteration. Despite the high deuterium excess, the slopes in the δD-δ18O diagram on the right side in Figure 10 at least indicate such a possibility.

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Figure -10. Isotope records from snow-pit studies in the Siberian Altai. Accumulation is influenced by wind scour as indicated by an inverse altitude effect with more enriched δvalues at the higher glacier site (left) and melt processes that additionally may cause isotopic enrichment (right).

5.

CONCLUSIONS

Post-depositional processes remove, redistribute, and change the isotopic and chemical information about climate and environment that arrives with the snow flakes falling on a glacier's surface. In nature, it is difficult to evaluate the influence of a single process because they interact with varying intensities. Sublimation hardens a fresh snow cover and may counteract wind scour. Refreezing of melt has the same effect, but sublimation and melting can also change the accumulated information. From the examples considered, it must be concluded that wind scour is probably the most important process at high-altitude drilling sites because it may disturb or even effectively remove seasonal cycles. Careful site selection and exploratory process studies prior to deep drilling may prevent unfortunate surprises. Saddle sites with channeling of winds and steep crests are obviously less well-suited than more open summit sites. If a cold glacier site is excessively wind-exposed, a more sheltered temperate glacier nearby might offer a superior record, at least for acquisition of a stable-isotope record. Chemical species are affected by post-depositional processes in a much more complex manner than stable isotopes. The latter are fundamental

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constituents of water molecules and interpretation of isotope records benefits from quantitative, physically based understanding of their behavior during phase transitions. In spite of this knowledge, no fool-proof strategy exists to anticipate or avoid such complications. Glaciers are dynamic and open systems for precipitation. Wind, sun, cloudiness, precipitation rate, temperature distribution and a myriad of other factors change as climate changes, and so does the information on climate and environment left from the precipitation that accumulates on a glacier.

6.

REFERENCES

Alean, J., Haeberli, W., and Schädler, B., 1983, Snow accumulation, firn temperature and solar radiation in the area of the Colle Gnifetti core drilling site (Monte Rosa, Swiss Alps): distribution pattern and interrelationships. Zeitschrift für Gletscherkunde und Glazialgeologie 19 (2) 131-147. Ambach, W., Eisner, H., and Sauzay. G., 1969, Tritium profiles in two firn cores from alpine glaciers and tritium content in precipitation in the alpine area. Arch. Met. Geophys. Bioklim. Serie B, 17, 93-104. Buason, Th., 1972, Evaluations of isotope fractionation between ice and water in a melting snow column with continuous rain and percolation. J. of Glaciology Vol 11, 387-400. Dansgaard, W., Johnsen, S., Clausen, H.B., and Gunderstrup., H.G., 1973, Stable isotope glaciology, Medd. Groenland 197, 1-53. Davies, T.D., Brimblecombe P., Tranter, M., Tsiouris, S.. Vincent C.E., Abrahams P., and Blackwood I.L., 1987, The Removal of Soluble Ions from Melting Snowpacks, Proceedings of the NATO Advanced Study Institute on Seasonal Snowcover Physics, Chemistry, Hydrology, Les Arcs, France, July 13-25,1986. Edited by H.G. Jones and W.J. Orville-Thomas, pp 337-392. Eichler, A., Schwikowski, M., Gäggeler, H.W., Furrer, V., Synal, H.A., Beer, J., Saurer, M., and Funk, M., 2001, Glaciochemical dating of an ice core from upper Grenzgletscher (4200m a.s.l.), J. of Glaciology Vol. 46 No. 154 307-315. Eichler, A., Schwikowski, M., and Gäggeler, H.W., 2001, Meltwater-induced relocation of chemical species in Alpine firn, Tellus 55B, 192 203. Ginot, P., Kull, Ch., Schwikowski, M., Schotterer U., Pouyaud B., and Gäggeler, H.W., 2001, Effects of post-depositional processes on snow composition of a subtropical glacier (Cerro Tapado, Chilean Andes), J. Geophys. Res. Vol.108, 32375-32386. Ginot, P., Schwikowski, M., Schotterer, U., Stichler, W., Gäggeler, H. W., Francou, B., Gallaire, R., and Pouyaud, B., 2002, Climate variability reconstruction from Andean glaciochemical records, Annals of Glaciology 35. 443-450. Grootes, P..M., M. Stuiver, M., Thompson, L.G, and Mosley-Thompson, E., 1989, Oxygen Isotope Changes in Tropical Ice, Quelccaya, Peru. J. Geophys. Res. Vol. 94, 1187-1194. Hermann, A., Lehrer, M., and Stichler, W., 1981, Isotope Input into Runoff Systems from Melting Snow Covers, Nordic Hydrology, 12: 308-318. Hoffmann, G., Jouzel, J., Masson, V., 2000, Stable water isotopes in atmosphric general circulation models, Hydrol. Proc. 14, 1385-1406. IAEA/WMO, 2001, Global Network of Isotopes in Precipitation, the GNIP Database accessible at: http://isohis.iaea.org.

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Johnsen, S.J., 1977, Stable Isotope Homogenisation of Polar Firn and Snow, in: Isotopes and Impurities in Snow and Ice, IAHS Publ. 118, 210-219. Johnsen, S. J., Clausen, H. B., Cuffey, K. M., Hoffmann, G., Schwander J., and Creyts T., 2000, Diffusion of stable isotopes in polar firn and ice; the isotope effect in firn diffusion, in: Physcics of Ice Core Records (T. Hondoh, ed.), Hokkaido University Press, Sapporo, 121-140. Lorius, C., Merlivat, L., Hagemann, R., 1969, Variation in the mean deuterium content of precipitation in Antarctica, J. Geophys. Res. 74, 7027-7031. Oeschger, H., Renaud, A., and Schumacher, E., 1962 Essai de datation par le Tritium des couches de néve du Jungfraufirn et détermination de l’accumulation annuelle, Bull. Soc. Vaud. Sc. Vol. 68, No 306 (Lausanne, Suisse), 49-56. Oeschger, H., Schotterer, U., Stauffer, B., Haeberli, W., and Röthlisberger, H., 1977, First results from Alpine core drilling projects, Zeitschrift für Gletscherkunde und Glazialgeologie, 13 (1/2), 193-208. 1977. Olivier, S., Schwikowski, M., Brutsch, S., Eyrikh, S., Gaggeler, H.W., Luthi, M., Papina, T., Saurer, M., Schotterer, U., Tobler, L., and Vogel, E., 2003, Glaciochemical investigations of an ice core from Belukha glcier, Siberian Altai, Geophys. Res. Lett. (in press). Peña, H., 1989, Mediciones de 18O y 2H en „penitentes“ de nieve, in Proceedings of a meeting on Estudios de Hidrologia Isotopica en America Latina, Mexico City, 1987, IAEATECDOC-502, Vienna, 143-154. Ramirez E., Hoffmann, G., Taupin, J.D., Francou, B., Ribstein, P., Cuillon, N., Landais, A., Petit, J.R., Pouyaud, B., Schotterer, U., and Stievenard, M., 2003, A new Andean deep ice core from the Illimani (6350m), Bolivia, EPSL, 212, 337-350. Satake, H., and Kawada, K., 1997, The quantitative evaluation of sublimation and the estimation of original hydrogen and oxygen isotope ratios of a firn core at East Queen Maud Land, Antarctica, Bulletin of Glacier Research 15 93-97. Stichler, W., Schotterer U., Fröhlich K., Ginot P., Kull C., Gäggeler H. W., and Pouyaud, B., 2001, The influence of sublimation on stable isotope records recovered from high altitude glaciers in the tropical Andes, J. Geophys. Res., Vol. 106, 22613-22620. Stichler, W. and Schotterer, U., 2000, From accumulation to discharge: modification of stable isotopes during glacial and postglacial processes, Hydrol. Process. 14, 1423-1438. Schotterer, U., Finkel, R., Oeschger, H., Siegenthaler, U., Bart, G., Gäggeler, H., and Von Gunten, H.R., 1977, Isotope measurements on firn and ice cores from Alpine glaciers, in: Isotopes and Impurities in Snow and Ice, IAHS Publ. No. 118, 232-236. Schotterer, U., Oeschger, H., Wagenbach, D., and Münnich, K.O., 1985, Information on paleo-precipitation on a high-altitude glacier, Monte Rosa, Switzerland, Zeitschrift für Gletscherkunde und Glazialgeologie 21, 379-388. Schotterer, U., Fröhlich, K., Gäggeler, H.W., Sandjordj, S., and Stichler, W., 1997, Isotope records from Mongolian and Alpine ice cores as climate indicators, Climatic Change Vol. 36, 3-4, 519-530. Schotterer, U., Stichler, W., Graf, W., Bürki, H.U., Gourcy, L, Ginot, P., and Huber, T., 2002a, Stable isotopes in alpine ice cores: do they record climate variability? in: Proceedings of an International Symposium on the Study of Environmental Change using Isotope Techniques, IAEA Vienna, 23-27 April 2001, IAEA, Vienna, 292-300. Schotterer, U., and Stichler, W., 2002b, Extending Isotope in Precipitation Data Beyond Direct Measurements: The Perspective from Glacier Ice-Core Measurements in Switzerland, in: Stable Isotopes (Edwards, T.D., Kull, Ch., Alverson, K., eds.), PAGES NEWS, Vol. 10-2, 6-7.

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Schotterer, U., Grosjean, M., Stichler W., Ginot, P., Kull, C., Bonnaveira, H., Francou, B., Gäggeler, H. W., Gallaire, R., Hoffmann, G., Pouyaud, B., Ramirez, E., Schwikowski, M., and Taupin, J. D., 2003, Glaciers and climate in the Andes between the Equator and 30°S: What is recorded under extreme environmental conditions?, Climatic Change 59, 157-175. Schwikowski, M., Brütsch S., Gäggeler H.W., and Schotterer, U., 1999, A high-resolution air chemistry record from an Alpine ice core: Fiescherhorn glacier, Swiss Alps, J. Geophys. Res. 104 13709-13719. Suter, S., 1995, Die Verbreitung kalter Firn- und Eisregionen im Alpengebiet. Diploma Thesis, ETH Zürich. Thompson, L.G., Mosley-Thompson E., Davis M.E., Lin P-N., Henderson K.A., Cole-Dai J., Bolzan J. F., and Liu, K., 1995, Late Glacial Stage and Holocene Tropical Ice Core Records from Huascaran, Peru. Science, 269, 47-50. Thompson L.G., Davis M.E., Mosley-Thompson E., Sowers T.A., Henderson K.A., Zagoradnov V., Lin P.-N., Mikhalenko V.M., Campen R.K., Bolzan J.F., Cole-Dai J., and Francou, B., 1998, A 25000-year tropical climate history from Bolivian ice cores, Science, 282, 1858-1864. Thompson L. G., 2000, Ice core evidence for climate change in the Tropics: implications for our future, Quarternary Science Reviews 19, 19-35. Wagenbach, D., Münnich, K.O., Schotterer U., and Oeschger, H., 1988, The anthropogenic impact on snow chemistry at Colle Gnifetti, Swiss Alps, Annals of Glaciology, 10, 183187.

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EVENT TO DECADAL-SCALE GLACIOCHEMICAL VARIABILITY ON THE INILCHEK GLACIER, CENTRAL TIEN SHAN

Karl J. Kreutz, Cameron P. Wake, Vladimir B. Aizen, L. DeWayne Cecil, Jaromy R. Green, and Hans-Arno Synal

1.

INTRODUCTION

Glaciochemical records developed from mid- and low-latitude Asian ice cores provide a unique archive of past atmospheric conditions, and can be used for high-resolution reconstructions of climatic and environmental variability (e.g., Mayewski et al., 1984; Kang et al., 2000a; 2000b; 2002a; Qin et al., 2000; Hou et al., 1999; 2002a; Yao et al., 1996; 1997; Thompson et al., 1995; 1997; 2000). To realize the full potential of chemical signals preserved in ice cores, detailed modern proxy calibration studies must be undertaken to understand the effects of local deposition noise and the relationships between meteorological conditions and time-series chemical variability. Previous work in the Tien Shan mountains of Central Asia (Figure 1) has demonstrated the usefulness of stable water isotopes and soluble ions for investigating temperature, moisture flux, atmospheric circulation, and dust loading on different timescales. On a seasonal basis, Yao et al. (1999) demonstrated a good correlation between į18O ratios and site temperature. Alternatively, Aizen et al. (1996) interpreted fresh snow event isotope data in terms of moisture source and transport pathway. For soluble ions, the strong influence of dust derived from surrounding arid regions has been noted in snow, firn core (Wake et al., 1992; Williams et al., 1992; Kattelmann et al., 1995; Kreutz and Sholkovitz, 2000; Kreutz et al., 2001), and aerosol (Sun et al., 1998) studies. Here we present new fresh snow and snowpit results from the Inilchek Glacier, Central Tien Shan (Figure 2) collected during July/August 2000. During the 2000 field season, two deep ice cores were also recovered, and are being used to develop highresolution stable isotope and soluble ion records of the past 200-500 years. Our goal in this chapter is to assess local-scale spatial chemical variability in 61 L. D. Cecil et al. (eds.), Earth Paleoenvironments: Records Preserved in Mid- and Low-Latitude Glaciers, 61-79. © 2004 Kluwer Academic Publishers. Printed in the Netherlands.

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the Inilchek basin and the relation between time series chemical variability and regional meteorological parameters. Such knowledge is critical for the proper interpretation of high-resolution records developed from deep ice cores, particularly at sites such as the Inilchek where relatively high accumulation rates may allow reconstructions on seasonal or sub-seasonal timescales.

Figure -1. Location map for the Inilchek Glacier, Central Tien Shan Mountains, Kyrgyzstan

2.

SAMPLE COLLECTION AND ANALYTICAL METHODS

Snow samples were collected from each fresh snow event that occurred during the 2000-field expedition, and also from four, 4-m snowpits and one 15-m crevasse wall (Figure 2). In addition, two deep cores (Core 1, 167.05

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m, 5100 meters above sea level (masl); and Core 2, 160.48 m, 5120 masl) were drilled using a solar powered ECLIPSE electromechanical auger. Fresh snow samples were collected along an elevational transect, with five samples collected at each of eight stakes immediately following every fresh snow event. Stake elevations range from 5087 m at stake 1 to 5250 m at stake 36. Snowpits were dug to 4 m depth at four stake locations (Figure 2), and sampled continuously in 5 cm intervals. The upper 100 m of Core 1 was processed in a dedicated science trench, using a polycarbonate lathe and saw system to remove the outer portion of the core and cut continuous 4-cm samples. During fresh snow sampling, snowpit sampling, and ice core processing, workers wore clean suits and polyethylene gloves to prevent contamination. All samples were collected into pre-cleaned polyethylene bottles, and returned frozen to the U.S.A. for analysis. Sampling of the crevasse wall was done without clean gear, and thus major ion analyses were not performed on that sample set. In addition to the ice core samples processed from Core 1, core chips from each Core 1 and Core 2 drill run (typically between 0.5 m and 1 m length) were collected into plastic bags and shaken to homogenize the chips. Two, 20-ml plastic vials were filled with chips from each drill run, and 10 vials were filled from selected runs to ensure that samples were representative of the entire run. Because core chips were not collected using clean sampling protocols, soluble ion measurements were not made.

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Figure -2. Location map showing sampling locations within the Inilchek Glacier basin during the 2000 field expedition. All sites were surveyed with a high-recision Trimble GPS unit. Elevation contours are given at the left side of the figure and contour lines run roughly eastwest (not shown). Core 1 is located south of Core 2. The crevasse wall sampled during 2000 is located approximately 200 m south of the southernmost snowpit. Also shown is the location of the 14 m firn core drilled during the 1998 field season (Kreutz and Sholkovitz, 2000; Kreutz et al., 2001)

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Samples were melted and aliquoted for major ion and stable isotope analyses, and refrozen as necessary for transport. Analysis for major ions at the University of New Hampshire was performed via suppressed ion chromatography using Dionex instruments (e.g., Buck et al., 1992). Samples from the July 30 and July 31 fresh snow events were not run due to low sample volume. Prior to major ion analyses, all snowpit samples were filtered through pre-cleaned Millipore 0.22 įm Durapore filters. Filter blanks for all ion species were <0.06 įmol/L, except Na+ which was 0.51 įmol/L, and well below the concentrations observed in samples. Thirty samples were run in duplicate, with the precision better than 5% for all ions except K+, which had a precision of 9%. Fresh snow and snowpit stable hydrogen isotope ratios (įD) were measured at the University of Maine Stable Isotope Laboratory via Cr reduction with a Eurovector elemental analyzer coupled to a Micromass Isoprime mass spectrometer (Morrison et al., 2001). Core chip įD ratios were measured via U reduction on a Finnigan Delta plus mass spectrometer at the University of Pennsylvania. Each sample was run in duplicate, with a resultant precision of ±0.5‰ based on sample and internal standard replicates. All data are reported in standard delta (į) notation vs. standard mean ocean water (SMOW).

3.

RESULTS

3.1

Fresh snow spatial variability

Figure 3 shows the distribution of fresh snow įD values during 5 of the 8 fresh snow events. Note that data from only 5 of the 8 fresh snow events is shown for clarity, as values for the remaining 3 events (July 21, July 24, July 31) fall between the 5 events shown. For each of the 8 events, there is no discernable trend in įD with elevation, and in most cases the mean value of the 5 samples collected at each stake is within ±2 sd. of the event mean. The typical range of mean įD values within each fresh snow event is ~5‰. Mean įD values among the 8 fresh snow events, however, are significantly different (Figure 4).

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Figure -3. Spatial variability of fresh snow įD values collected across an elevation transect on the Inilchek Glacier. Horizontal dashed lines represent mean values for each event, and error bars represent ±1 standard deviation for the five samples collected at each stake.

Between July 15 and July 25, there is a general decreasing trend in mean event įD values, followed by a rise to ~-35‰ for the July 30 and July 31 events. The largest change occurs between July 31 and August 3, with the August 3 fresh snow event having įD values ~70‰ lower than those of July 31. The large įD difference between July 31 and August 3 does not appear to correspond to a significant change in the 2-m air temperature at the site (Figure 4). We have not presented wind speed data from the AWS, as it is likely only of local significance. Published įD values from high elevation summer snowfall in Asia for comparison are sparse, and are limited to Hailougou Glacier (2940 m), southeast Tibet (Aizen et al., 1996), and the Xixibangma (5680-7000 m) and Everest (6500 m) massifs (Tian et al., 2001; Kang et al., 2002b). On Hailougou Glacier, įD values varied from –180‰ to –45‰ during summer 1990 (Aizen et al., 1996). Summer fresh snow isotope values are therefore broadly consistent between the Tien Shan and Gongga Massif. Mean summer įD values in Xixibangma precipitation (135.7 ‰), and Everest fresh snow įD ranges (-60 to –245 ‰) are lower than those observed in Inilchek Glacier samples, likely due to the higher elevation of the Xixibangma and Everst sampling locations. We refrain from using the global meteoric water line (Dansgaard, 1964) to calculate į18O values from existing Inilchek įD data for comparison with published į18O data,

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given that the local meteoric water line in Inilchek samples is significantly different from the global value (Kreutz et al., manuscript in preparation).

Figure -4. (upper panel) Time-series variability of mean įD values for each fresh snow event. Vertical error bars represent ±1 standard deviation of all samples collected for a particular fresh snow event. Horizontal error bars represent the duration of each fresh snow event. (lower panel) 2 m surface temperature collected by an automatic weather station during the 2000 field season.

On an event basis, soluble ion data from fresh snow samples do not show any systematic change in concentration with elevation (not shown). As with įD, however, there are large inter-event changes in concentration that are evident in all species (Figure 5). The time-series variability for several species pairs is almost identical (i.e., Na+, Cl-, and K+; Ca2+ and Mg2+; and SO42- and NO3-), and thus only 4 species are plotted (Figure 5) for simplicity. The two largest changes occur in Ca2+ during the July 21 event (a threefold increase in concentration), and in NH4+ during the August 3 event (a twofold decrease in concentration). The high dust concentrations during the July 21 event are associated with the highest nighttime temperatures recorded at the site during the sampling period (Figure 4). Like NH4+, Ca2+ concentrations

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are lowest during the August 3 event. Few fresh snow ion data exist in the Tien Shan for comparison. With the exception of NH4+, ion concentrations in Inilchek fresh snow samples collected during 2000 are lower than those previously reported in the Tien Shan (Wake et al., 1992; Williams et al., 1992; Sun et al., 1998). One possibility is the short summer sampling period represented by the Inilchek 2000 samples, which does not take into account the large dust events commonly observed in winter and spring. Alternatively, the lower NH4+ concentrations may reflect a decreased proportion of dry deposition at the Inilchek site. However, we have no a priori reason to suspect that the ratio between wet and dry chemical deposition is different within the Tien Shan range.

Figure -5. Time-series variability of mean soluble ion values for each fresh snow event. Error bars represent ±1 standard deviation for all samples collected during a particular fresh snow event. Note that ion data is not available for the July 29 and July31 fresh snow events.

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Fresh snow isotope and ion relationships

Figure -6. Scatter plots of įD and major ions in fresh snow events. Note that ion data is not available for the July 29 and July 30 fresh snow event.

To investigate the relationships among isotope and soluble ion data in fresh snow, data for each event is shown using scatter plots (Figure 6). Based on this comparison, three groups of data are identified: 1) The August 3 event, with low įD, Ca2+, Mg2+, and NH4+ values; 2) the July 21 event, with median įD values and high Ca2+, Mg2+, and NO3- values; and 3) the remaining four events with intermediate isotope ratios and major ion

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concentrations. The relationship between įD and major ions during the August 3 event is unique, with obviously low concentrations of primary dust species (Ca2+, Mg2+), and NH4+. The remainder of the ion species display concentrations during the August 3 event that have more scatter, and also are close to the mean of all fresh snow events. Therefore, it appears that the August 3 event has a unique chemical signature. We note that there is no obvious change in site temperature prior to or during the August 3 event (Figure 4). The July 21 event also has a unique chemical signature, with high primary dust species (Ca2+, Mg2+), and NO3-. Isotope values for the July 21 event are not significantly different from other events. For the remainder of events, there does not appear to be any systematic relationship between įD and ion concentration. Likewise, nothing definitive can be seen in Na+ or Cl- concentrations – the concentration range for each event appears to not be associated with a particular range or įD values.

3.3

Snowpit and ice core spatial variability

Figure -7 Box and whisker plots of snowpit, crevasse wall, and ice-core įD data. Pit 3 is not shown because it does not cover a full annual cycle. Upper, middle, and lower lines represent the upper quantile, median, and lower quartile. Whiskers represent ±2 standard deviation, with outliers plotted as *symbols. Data are arranged according to increasing elevation in the basin (i.e., the crevasse wall is at the lowest elevation, and Pit 4 at the highest elevation)

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The distribution of įD values in snowpits, the 15-m crevasse wall, and the upper 18 m of Core 1 are shown in Figure 7. Pit 3 is not included because it does not contain a complete accumulation year. Median values for the three snowpits and the crevasse wall are within the upper and lower quartile boundaries for each location. The median value for Core 1 is slightly below the other sites, and may be related either to the longer (~7 years) record contained in the Core 1 dataset vs. the snowpits (~2 years), or the different time periods represented. Outliers in each case are restricted to the lower part of the distribution, possibly reflecting strongly depleted įD values during severe winter seasons. As with the fresh snow samples, there is no significant trend in mean įD with increasing elevation across the basin.

Figure -8. Box and whisker plots of major ion data from the four Inilchek 2000 snowpits. Upper, middle, and lower lines represent the upper quartile, median, and lower quartile. Whiskers represent ± standard deviation, with outliers plotted as * symbols.

Major-ion data in Pits 1, 2, and 4 are shown in Figure 8. Overall, all three pits show similar major-ion concentration distributions. Outliers are

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confined towards high events, reflecting positively skewed distributions characteristic of major-ion data in Asia (Wake et al., 1992). While there is minor variability in ion concentrations among the snowpits, there is no apparent trend with elevation (Figure 8). Ice core chip įD values from the two deep cores are plotted in Figure 9. Mean įD values for the two cores are nearly identical (Core 1 = -112.7‰; Core 2 = -112.2‰), as is the down-core variance (Core 1 = 38.6‰; Core 2 = 41.2‰). Moreover, no long-term trend in įD with depth is apparent. Given that there are likely slight changes in both spatial and temporal accumulation rates between the two cores, and that different sampling intervals are represented in the two cores (i.e., each core run is not the same length), the įD profiles for each core are reasonably similar. There does not appear to be any indication of ice flow effects in lower portions of the core (i.e., no large breaks or discontinuities in the įD records). Given that the core depths represent approximately 60% of the total ice depth (depth to bedrock measured via radio echo sounding as 285-300 m at Core 1 and 250-160 m at Core 2), we expect that the cores are free of any flow deformation. Overall, the similarity of mean įD values and variance between the two cores indicates that the primary signal recorded in the cores is related to climate variability, and not to surface spatial chemical variability or postdepositional flow effects.

4.

DISCUSSION

The spatial variability of įD in fresh snow and snowpit samples is valuable for evaluating snow formation and deposition processes. Given typical temperature lapse rates in mountain regions (~5-8°C km-1; Barry, 1992), a corresponding decrease in precipitation isotopic composition with increasing altitude is often assumed based on isotope/temperature relationships developed from global datasets (e.g., Dansgaard, 1964; Rozanski et al., 1992; Friedman et al., 1992; Rowley et al., 2001; Hou et al., 2002b). Indeed, where precipitation isotope measurements have been made spanning large elevation ranges in mountain regions, į18O lapse rates ranging from -2 to -5‰ km-1 (e.g., Siegenthaler and Oeschger, 1980; Moser and Stichler, 1970; Niewodniczanski et al., 1981; Holdsworth et al., 1991; Garzione et al., 2000; Gonfiatini et al., 2001) have been observed. Thus, when collecting samples across an elevation range in a large basin, differences in isotopic composition related to air mass uplift and rainout at progressively lower temperatures might occur. Assuming a conservative value of 40‰ km-1 (using a slope of 8 for įD/į18O), we would expect to observe a change in įD of 6.5‰ over the elevation range in the Inilchek

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basin. Several fresh snow events from 2000 have ±1 standard deviation values for replicate measurements at each stake of close to 10‰ (Figure 3). However, there is no systematic decrease in mean įD values for any of the eight fresh snow events sampled during 2000 in the Inilchek Basin that approaches the estimated value of 6.5‰. We speculate that the mechanism responsible for the lack of įD/elevation trend is snow formation at a constant cloud base elevation throughout the basin. Precipitation during the 2000 fresh snow events occurred under calm wind conditions, suggesting that a relatively stationary horizontal cloud deck may have been responsible for precipitation formation at a constant elevation, with subsequent deposition over a large elevation range. Similar results have been observed for previous fresh snow events sampled during 1991 in the Inilchek Basin, and which cover a larger elevation range (4000 – 5100 m; Aizen et al., 1996; Kreutz et al., 2001). Snowpit įD values support the cloud deck model, as there is no discernable trend in median įD values with elevation (Figure 7). Thus, over limited elevation ranges were air mass uplift is minimal, precipitation isotope variability may be related primarily to local deposition noise (e.g., slight wind redistribution, surface irregularities) and random sampling and analytical error.

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Figure -9.. įD data from ice core chips collected from the two deep cores drilled in 2000. The black dashed lines represent raw data (average įD from two replicate samples taken from each core run), and the filled regions represent a 5-point running mean above and below the overall įD record mean value.

This lack of įD/elevation trend in fresh snow and snowpit samples has important implications for interpreting the time-series įD records generated from the two deep ice cores (Figure. 9). If long-term trends in įD were observed in the core records, one possible explanation would be ice flow bringing lighter įD snow/firn/ice deposited at higher elevations in the basin. The lack of observed įD trends in the core, coupled with the lack of observed įD/elevation trends in fresh snow and snowpits, argues against this possibility. In addition, assuming that the 5-point running mean used in Figure 9 represents a 3-5 year average, the interannual įD variability is ~4050‰. This range of įD values is much higher than the inter-event įD variability observed for the fresh snow events (typically 5‰), but comparable to the inter-event įD variability (35-40‰). Therefore, we conclude that the majority of the įD variability observed in the core profiles is related to climate-related processes, rather than ice flow effects or local deposition noise. Comparison of įD and major-ion data in the fresh snow events indicates that several atmospheric parameters may be involved in determining the

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įD/relationships for each event, and for the time-series variability in each parameter. Previous work in the Tien Shan has shown that three primary dust compositions are deposited on the Inilchek Glacier – loess, calcite, and gypsum (Kreutz and Sholkovitz, 2000). Based on ion balances and trace element data, calcite is likely the dominant source of Ca2+ on the Inilchek, with gypsum contributing the remainder. Thus, the high Ca2+ levels in the July 21 fresh snow event suggest that the air mass passed over a calcite source area during transport to the Tien Shan. Conversely, the low Ca2+ values for the August 3 event imply that calcite was not an important component, while SO42- values are higher than in other events. The presence of significant SO42- concentrations suggests that the August 3 event was influenced by a gypsum source rather than calcite. In addition, NH4+ concentrations during the August 3 event are the lowest of any 2000 fresh snow event. One possible source for NH4+ in Tien Shan snow is regional agricultural activities (nitrogen-rich fertilizer use; Kreutz et al., 2001; Hou et al., 2002a). The low NH4+ concentrations in August 3 snow implies that the air mass did not travel over significant agricultural regions (e.g., the Ferghana Valley). Lastly, įD values for the August 3 event are the lowest of any event. Assuming that event-scale įD variability is not directly related to site surface air temperature (Figure 4) but rather to regional-scale moisture source and transport distance (Aizen et al., 1996), then the low įD values imply a long-traveled air mass. įD values for the July 21 event are similar to the majority of fresh snow events, which may imply a more local moisture source. We speculate that the July 21 event was caused by an air mass originating from the North Atlantic/Mediterranean region, passing across the arid regions west of the Tien Shan where extensive calcite deposits exist (Digital Soil Map of the World, 1995). We further speculate that the August 3 event originated in the north sea region, traveling southeast across the large gypsum deposits lying north of the Tien Shan. These scenarios potentially explain both the time-series glaciochemical changes as well as the isotope/ion relationships seen in each event. Confirmation of these ideas will require detailed investigations of back trajectories, and correlations between glaciochemical data and regional meteorological conditions. Fortunately, the former Soviet Union maintained a robust network of meteorological stations in the Tien Shan (e.g., Aizen et al., 1997), and thus there is a large database to use for glaciochemical proxy calibration in the region.

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Event to Decadal-Scale Glaciochemical Variability

CONCLUSIONS

Based on data collected from fresh snow, snowpit, crevasse wall, and icecore samples during the 2000 Inilchek Glacier expedition, we conclude the following: Limited spatial and elevational variability exists in fresh snow isotope and major-ion data from the Inilchek Basin. There is no apparent trend with increasing elevation for any of the parameters measured, and the intra-event variability in įD ratios and major-ion concentrations is much smaller than the inter-event variability. Based on the lack of isotope variability with elevation, we conclude that snow formation occurs from a horizontal cloud base across the Inilchek Basin. This model is supported by multi-year snowpit, crevasse wall, and ice-core data, which also do not display significant elevational variability in mean isotope or major-ion data. Isotope profiles from two deep ice cores have similar mean values and variability (±1 standard deviation), and there is no long-term trend in either įD profile. Based on estimated accumulation rates, the interannual-scale variability in the core profiles is significantly greater than the spatial variability during individual fresh snow events. Given the lack of isotope/elevation trend in fresh snow and snowpit data, we conclude that a majority of the down core isotope variability reflects climate changes rather than ice flow effects or deposition noise. Large differences in fresh snow isotope ratios and major-ion concentrations are observed among the 8 fresh snow events sampled in 2000. Correlations between įD and major ions are not consistent among the 8 events, with three different relationships identified: 1) low įD, low dust (Ca2+, Mg2+) and NH4+ concentrations during the August 3 event; 2) high dust and intermediate įD values during the July 21 event; and 3) intermediate įD and major-ion concentrations during the remaining events. We interpret these different relationships as reflecting different air mass source and transport pathway to the Tien Shan. These conclusions have implications for the interpretation of the highresolution isotope and major-ion records eventually developed from the two deep Inilchek ice cores. First, high-resolution records from the two cores can be used for detailed evaluations of time-series signal-noise ratios related to small-scale deposition processes. Second, given the high accumulation rate at the site (~1.5 m WE/yr; Kreutz et al., 2001), it may be possible to reconstruct climate and environmental variability on an event basis. The large changes in fresh snow glaciochemistry, among other events, implies that such variability might be apparent in the down-core record. Lastly, given the different isotope/ion relationships observed in the fresh snow events, it might be possible to use the down-core records to investigate

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specific moisture source and transport pathways through time. To do so will require detailed glaciochemical calibration with regional meteorological conditions available from the extensive array of Former Soviet Union stations. Thus, the high-resolution, multivariate records being developed from the Inilchek cores hold the potential for detailed climate and environmental records from the Tien Shan.

6.

ACKNOWLEDGEMENTS

We thank our colleagues (D. Susong, S. Nitikin, G. Hornock), guides, and staff for their assistance in the field during the 2000 Inilchek Expedition. We further thank D. Introne (University of Maine Stable Isotope Laboratory), E. Steig (University of Pennsylvania, now at University of Washington), and J. Morrison and A. Phillips (Micromass) for stable isotope measurements, and S. Whitlow (University of New Hampshire) for majorion measurements. B. Hall and S. Kang kindly reviewed the manuscript, and provided useful feedback for clarification and improvement. Research presented here was supported by the National Science Foundation (ATM0096323 [KJK], ATM-0000561 [CPW], ATM-9905670 [VBA]), the U.S. Department of Energy, and the U.S. Geological Survey.

7.

REFERENCES

Aizen, V.B., Aizen, E., Melack, J., Martma, T., 1996, Isotopic measurements of precipitation on central Asian glaciers (southeastern Tibet, northern Himalayas, central Tien Shan), Journal of Geophysical Research 101(D4):9185-9196. Aizen, V.B., Aizen, E.A., Melack, J.M., Dozier, J., 1992, Climate and hydrographic changes in the Tien Shan, Central Asia, Journal of Climate 10(6):1393-1404. Barry, R.G., 1992, Mountain Weather and Climate, Methuen, London. Buck, C.F., Mayewski, P.A., Spencer, M.J., Whitlow, S., Twickler, M.S., Barrett, D., 1992, Determination of major ions in snow and ice cores by ion chromatography, Journal of Chromatography 594:225-228. Dansgaard, W., 1964, Stable isotopes in precipitation, Tellus 16:436-468. Digital Soil Map of the World, FAO/UNESCO, 1995, 1:5,000,000, Version 3.5. Friedman, I., Smith, G.I., Gleason, J.D., Warden, A., Harris, J.M., 1992, Stable isotopic composition of waters in Southeastern California 1. Modern precipitation, Journal of Geophysical Research 97(D5):5795-5812. Garizone, C.N., Quade, J., DeCelles, P.G., English, N.B., 2000, Predicting paleoelevation of Tibet and the Himalaya from į 18O vs. altitude gradients in meteoric water across the Nepal Himalaya, Earth and Planetary Science Letters 183:215-229. Gonfiantini, R., Roche, M.A., Olivry, J.C., Fontes, J.C., Zuppi, G., 2001, The altitude effect on the isotopic composition of tropical rains, Chemical Geology 181:147-167.

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Holdsworth, G., Fogarasi, S., Krouse, H.R., 1991, Variation of the stable isotopes of water with altitude in the Saint Elias Mountains of Canada, Journal of Geophysical Research 96(D4):7483-7494. Hou, S., Dahe, Q., Wake, C.P., Mayewski, P.A., 1998, Abrupt decrease in recent snow accumulation at Mount Qomolangma (Everest), Himalaya, Journal of Glaciology 45(151):585-586. Hou, S., Qin, D., Zhang, D., Kang, S., Mayewski, P.A., Wake, C.P., 2002a, A 154a high resolution ammonium record from the Rongbuk Glacier, north slope of Mt. Qomolagma (Everest), Tibet Himal region, Atmospheric Environment in press. Hou, S., Masson-Delmotte, V., Qin, D., Jouzel, J., 2002b, Modern precipitation stable isotope vs. elevation gradients in the High Himalaya: Comment on “A new approach to stable isotope-based paleoaltimetry: implications for paleoaltimetry and paleohypsometry of the High Himalaya since the Late Miocene” by D.B. Rowley, R.T. Pierrehumbert, and B.S. Currie, Earth and Planetary Science Letters in press. Kang, S., Wake, C.P., Dahe, Q., Mayewski, P.A., Tandong, Y., 2000a, Monsoon and dust signals recorded in Dasuopu glacier Tibetan Plateau, Journal of Glaciology 46:222-226. Kang, S., Qin, D., Mayewski, P.A., Wake, C.P., 2000b, Recent 180 year oxalate records recovered from the Mount Everest ice core: some environmental implications, Journal of Glaciology 46:155-156. Kang, S., Mayewski, P.A., Qin, D., Yan, Y., Hou, S., Zhang, D., Ren, J., Kreutz, K., 2002a, Glaciochemical records from a Mt. Everest ice core: Relationship to atmospheric circulation over Asia, Atmospheric Environment 36(21):3351-3361. Kang, S., Kreutz, K.J., Qin, D., Mayewski, P.A., 2002b, Stable isotopic composition of precipitation over the northern slope of the central Himalayas, Journal of Glaciology in press. Kattelmann, R., Elder, K., Melack, J., Aizen, E., Aizen, V., 1995, Some surveys of snow chemistry in the Tien Shan of Kirghizstan and Kazakhstan, in: Biogeochemistry of Seasonally Snow-Covered Catchments, IAHS, Boulder, CO, pp. 185-190. Kreutz, K.J., Sholkovitz, E.R., 2000, Major element, rare earth element, and sulfur isotopic composition of a high-elevation firn core: Sources and transport of mineral dust in Central Asia, Geochemistry, Geophysics, and Geosystems 1:2000GC000082. Kreutz, K.J., Aizen, V.B., Cecil, L.D., Wake, C.P., 2001, Oxygen isotopic and soluble ionic composition of precipitation recorded in a shallow firn core, Inilchek Glacier (Central Tien Shan), Journal of Glaciology 47(159):548-554. Mayewski, P.A., Lyons, W.B., Ahmad, N., Smith, G., Pourchet, M.., 1984, Interpretation of the chemical and physical time-series retrieved from Sentik Glacier, Ladakh Himalaya, India, Journal of Glaciology 30(104):66-76. Morrison, J., Brockwell, T., Merren, T., Fourel, F., Phillips, A.M., 2001, On-line highprecision stable hydrogen isotopic analyses on nanoliter water samples, Analytical Chemistry73(15):3570-3575. Moser, H., Stichler, W., 1970, Deuterium measurements on snow samples from the Alps, in: Isotope Hydrology 1970, IAEA Symposium 129, Vienna, pp. 43-57. Niewodniczanski, J., Grabczak, J., Baranski, L., Rzepka, J., 1981, The altitude effect on the isotopic composition of snow in high mountains, Journal of Glaciology 27(95):99-111. Qin, D., Mayewski, P.A., Wake, C.P., Kang, S., Jiawen, R., Shugui, H., Yao, T., Yang, Q., Jin, Z., Mi, D., 2000, Evidence for recent climate change from ice cores in the central Himalaya, Annals of Glaciology 31:153-158.

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Rowley, D.B., Pierrehumbert, R.T., Currie, B.S., 2001, A new approach to stable isotopebased paleothermometry: implications for paleoaltimetry and paleohypsometry of the High Himalaya since the Late Miocene, Earth and Planetary Science Letters 188:253-268. Rozanski, K., Araguas-Araguas, L., Gonfiantini, R., 1992, Isotopic patterns in modern global precipitation, in: Climate Change in Continental Isotope Records (P.K. Swart, K.C. Lohmann, J. McKenzie, and S. Savin, eds.), Geophysical Monograph 78, American Geophysical Union, pp. 1-36. Siegenthaler, U., Oeshger, H., 1980, Correlation of 18O in precipitation with temperature and altitude, Nature 285:314-317. Sun, J., Qin, D., Mayewski, P.A., Dibb, J.E., Whitlow, S.I., Li, Z., Yang, Q., 1998, Soluble species in aerosol and snow and their relationship at Glacier 1, Tien Shan, China, Journal of Geophysical Research 103(D21):28021-28028. Thompson, L.G., Mosley-Thompson, E., Davis, M.E., Lin, P.N., Dai, J., Bolzan, J.F., Yao, T., 1995, A 1000-year climate ice-core record from the Guliya ice cap, China: its relationship to global climate variability, Annals of Glaciology 21:175-181. Thompson, L.G., Yao, T., Davis, M.E., Henderson, K.A., Mosley-Thompson, E., Lin, P.E., Beer, J., Synal, H.A., Cole-Dai, J., Bolzan, J.F., 1997, Tropical climate instability: The last glacial cycle from a Qinghai-Tibetan ice core, Science 276:1821-1825. Thompson, L.G., Yao, T., Mosley-Thompson, E., Davis, M.E., Henderson, K.A., Lin, P.E., 2000, A high-resolution millennial record of the South Asian monsoon from Himalayan ice cores, Science 289:1916-1919. Tian, L., Masson-Delmotte, V., Stievenard, M., Yao, T., Jouzel, J., 2001, Tibetan Plateau summer monsoon northward extent revealed by measurements of water stable isotopes, Journal of Geophysical Research 106(D22):28,081-28,088. Wake, C.P., Mayewski, P.A., Ping, W., Yang, Q., Jiankang, H., Zichu, X., 1992, Anthropogenic sulfate and Asian dust signals in snow from Tien Shan, northwest China, Annals of Glaciology 16:45-52. Williams, M.W., Tonnessen, K.A., Melack, J.M., Daqing, Y., 1992, Sources and spatial variation of the chemical composition of snow in the Tien Shan, China, Annals of Glaciology 16:25-32. Yao, T., Thompson, L.G., Mosley-Thompson, E., Zhihong, Y., Xingping, Z., Lin, P.N., 1996, Climatological significance of 18O in north Tibetan ice cores, Journal of Geophysical Research 101(D23):29,531-29,537. Yao, T., Shi, Y., Thompson, L.G., 1997, High resolution record of paleoclimate since the Little Ice Age from the Tibetan ice cores, Quaternary International 37:19-23. Yao, T., Masson, V., Jouzel, J., Stievenard, M., Weizen, S., and Keqin, J., 1999, Relationships between į 18O in precipitation and surface air temperature in the Urumqi River Basin, east Tienshan Mountains, China, Geophysical Research Letters 26(23):34733476.

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CLIMATIC INTERPRETATION OF THE GRADIENT IN GLACIOCHEMICAL SIGNALS ACROSS THE CREST OF THE HIMALAYA

Cameron P. Wake, Paul A. Mayewski, and Sichang Kang

1.

INTRODUCTION

Documenting and identifying the causes of spatial variability in snow chemistry preserved in mid- and low-latitude glaciers is critical for improving our ability to interpret longer term glaciochemical time-series developed from these sites. Unlike the central regions of the polar ice sheets, mid- and low-latitude glaciers often lie adjacent to a variety of source regions for major ions, such as deserts, salt flats, agricultural fields, and cities. Previous work in the mountains of central Asia has shown that the spatial distribution of snow chemistry is controlled predominantly by the influx of dust from arid and semi-arid regions (Williams et al., 1992; Wake et al., 1993; Wake and Mayewski, 1993), while anthropogenic emissions may dominate downwind of major urban centers (Wake et al., 1992). Conversely, snow and ice from the southern slopes of the Himalaya display relatively low concentrations and fluxes. Here we present new glaciochemical data from six glaciers in the eastern and central Himalaya; three on the southern slopes and three on the northern slopes (Figure 1). These data represent a substantial improvement in spatial coverage of glaciochemical records across the Himalaya and allow for an improved definition of the gradient in snow and ice chemistry from the southern to the northern slopes. This data set provides the basis to investigate the processes responsible for the spatial variability of glaciochemical signals and their relationship to climate variability. 81 L. D. Cecil et al. (eds.), Earth Paleoenvironments: Records Preserved in Mid- and Low-Latitude Glaciers, 81-94. © 2004 Kluwer Academic Publishers. Printed in the Netherlands.

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Climatic Interpretation of the Gradient in Glaciochemical Signals

CLIMATOLOGICAL SETTING

The main crest of the Himalaya acts as a significant climate divide separating the relatively wet tropical climate of the southern slopes with the relatively arid highland climate on the northern slope (Table 1). The southern slopes of the Himalaya are characterized by mixed forest and small scale agriculture below 4000-4500 m, with glaciers and mountains dominating above this elevation. The climate varies considerably with the large range in elevation (2000 to 8000 m). In general, the region receives summertime precipitation from orographic lifting of monsoonal air masses as they push up against the Himalaya (Inoue, 1976). During the winter, westerly depressions traveling along the southern slope of the Himalaya result in precipitation events at higher elevations (Barry, 1981). Precipitation amounts in the Khumbu Himal range from 1.0 to 1.4 m per year, with the majority of precipitation falling during the summer (Chalise et al., 1996). Conversely, the northern slopes lie in the rainshadow of the Himalaya and experience relatively arid conditions. Precipitation falls mainly in summer as part of plateau monsoon circulation. Annual precipitation amounts range from 0.4 to 0.5 m (Domrös and Peng, 1988). The highlands to the north (elevations greater than 4000 m) are characterized by grassland steppes which transition into the semi-arid and arid desert regions of the central and northern Tibetan Plateau.

Figure -1. Location map for snowpit and ice-core samples collected in the eastern Himalaya.

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Table -1. Comparison of climate on the north and south slopes of the Himalaya. The period of record is from 1967 - 1993. (ppt = precipitation)

3.

Elevation (m) Latitude (N) Longitude (E)

Tingri 4300 28.38 87.05

Nyalam 3810 28.11 85.58

Annual ppt (mm) Summer ppt (mm) Mean annual T (0C)

269 256 2.2

624 315 3.5

METHODS

Samples were collected from snowpits and firn/ice cores in the accumulation zones of Himalayan glaciers over a range of elevations (Table 2). Three of the sites are on the southern slopes of the Himalaya (Kangchung Glacier and Nangpai Gosum Glacier in the Khumbu Himal, and Chago Glacier in the Makalu region) and three are on the northern slopes (Hidden Valley in the Dhaulagiri Himal, Kangwure Glacier on the north slope of Mt. Xixabangma, and the Far East Rongbuk Glacier on the north side of Mt. Everest). All snowpit and ice-core sampling was performed employing techniques designed to minimize or eliminate contamination. Samples were collected and processed by personnel wearing clean gear (non-particulating clean suits, particle masks and clean gloves). Snowpit samples were collected directly into precleaned, watertight 60-ml HDPE containers. Firn and ice-core samples were processed in the field in a science trench excavated in the snowpack. The core samples were then placed in precleaned 60-ml HDPE containers. Samples were transported frozen from the field to the freezers at the University of New Hampshire (UNH). Major-ion (Na+, NH4+, K+, Mg2+, Ca2+, Cl-, NO3-, SO42-) concentrations were determined via ion chromatography in our laboratories at UNH using techniques described elsewhere (Wake 1993; Buck et al., 1992). We have also calculated excess cations (denoted as ǻC) which equals the sum of cations minus the sum of anions on a sample by sample basis. Annual fluxes were calculated by multiplying the weq depth of each sample by the ion concentration, and then summing by year.

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Table -2. Site location information, major-ion concentrations and annual fluxes for snow and ice samples collected from glaciers on the southern and nothern slopes of the Himalaya. SOUTH NORTH Nangpai Hidden Far East Kangchung Chago Kangwure Gosum Valley Rongbuk Elevation (masl) 5640 5800 6170 5150-5850 6140 6500 Latitude (N) 28.0 28.0 27.9 28.8 28.3 28.1 Longitude (E) 86.7 86.6 87.1 83.6 85.2 87.0 Type of samples Pit pit/core pit/core pits pit/core pit/core No. of samples 41 455 117 52 170 1010 Period of record 1995-96 1981-98 1990-96 summers 1952-91 181493,94,95 1997 Accum. (m weq/yr 1) 0.71 0.78 0.70 na 0.32 0.21 Median (µEq Kg-1) Ca2+ 0.13 0.22 0.81 3.15 8.59 8.61 NH4+ 0.56 0.98 0.84 0.57 2.89 2.07 0.05 0.06 0.11 0.39 1.16 0.77 Mg2+ Na+ 0.02 0.05 0.10 0.25 0.85 0.62 0.06 0.03 0.08 0.05 0.40 0.14 K+ ǻC NO3SO42ClCa2+ NH4+ Mg2+ Na+ K+ ǻC NO3SO42ClCa2+ NH4+ Mg2+ Na+ K+

0.14 0.79 0.43 0.24 0.22 0.09 0.08 0.28 Maximum (µEq Kg-1) 0.41 3.42 1.98 9.00 0.16 0.73 0.66 4.93 0.50 1.56

1.14 0.50 0.29 0.16

3.31 0.60 0.32 0.26

9.25 1.84 1.61 0.90

9.32 1.30 0.81 0.52

5.64 2.41 0.80 0.62 0.36

40.2 3.71 7.57 2.77 3.77

242 21.2 14.1 13.6 5.08

375 11.2 20.7 7.28 4.64

0.99 8.95 5.55 3.81 1.09 2.95 1.20 3.00 2.04 0.50 3.94 0.53 Annual Flux (nEq cm-2) 13.0 21.2 97.9 54.1 100 70.1 4.33 5.30 14.3 5.55 6.76 9.95 5.72 3.21 6380

42.1 10.2 18.7 2.77

230 16.3 36.0 10.6

375 29.3 36.7 7.38

na na na na na

594 117 50.3 49.9 11.6

325 50.4 30.9 19.2 4.74

na na na na

601 89.5 85.1 44.6

341 44.2 30.8 14.4

ǻC -3.53 76.2 49.7 22.9 NO325.7 10.8 SO4210.1 26.9 ClǻC = (sum of cations) - (sum of anions) na means not available

106 47.7 31.1 14.7

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Depth-age relationships for the four sites where ice cores were recovered (Table 2) were established through annual layer counting of the major ion and the oxygen isotope records (Wake 1992; Wake and Steivenard, 1996; Wake et al., 2001) and, for the Nangpai Gosum, Kangwure and Far East Rongbuk cores, through identification of the 1963 marker horizons resulting from fallout of radioactive debris from atmospheric nuclear-weapons testing. Analysis of physical and chemical stratigraphy in the Kanchung Glacier snowpit allowed for identification of monsoon verses winter time snow and indicated that we sampled a full year of snow accumulation. All accumulation on the glaciers in Hidden Valley was in the form of superimposed ice and we therefore only collected samples for chemical analysis from shallow snowpits consisting of monsoon snow.

4.

RESULTS

For comparison of snow chemistry between sites, the median, maximum, and annual flux values as well as the mean annual snow accumulation rate (in m weq/yr) for each site are listed in Table 2. We report median rather than mean values for the major-ion concentrations as the data is lognormally distributed; the median value therefore provides a more accurate measure of the central tendency of the data set as it is not strongly skewed by a few very high concentration levels that characterize glaciochemical data. The median ion concentration and annual ion flux data are also illustrated in Figure 2.

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Figure -2. (a) Median major-ion concentrations (µeq 1-1) and (b) mean annual fluxes (neq/cm-2 yr--1) for snow and ice samples collected from glaciers on the southern and northern slopes of the eastern Himalaya. No annual flux data is available for Hidden Valley (see text for details).

Supporting our previous findings, ion concentrations for glaciers on the northern slopes are greater than those on the southern slopes, and maximum

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concentrations are much greater. Calcium and ǻC account for the majority of the ion burden in snow deposited on the northern slopes indicating that precipitation there is alkaline. Measurements of alkalinity on frozen snow samples from the Tien Shan indicate that the ǻC term is made up primarily of CO3=/HCO3- (Williams et al., 1992). The very strong relationship between the Ca++ and ǻC at all three sites, as expressed by high correlation coefficients and slopes close to one (Table 3), suggests that the primary source of Ca++ and ǻC for glaciers on the north slope of the Himalaya is CaCO3 rich dust. For the northern slope sites, NH4+ is the second most important cation, followed by Mg2+ and Na+ while for anions NO3- is the second most abundant, followed by SO42-. On the southern slopes, NH4+ is the most abundant cation, followed by Ca2+. Anions (including ǻC) display a variable ranking among the three sites on the southern slopes. Table -3. Relationship between Ca2+ and ǻC for glaciers on the northern slopes of the Himalaya. Site r slope Hidden Valley 0.98 1.06 Kangwure 0.99 0.96 Far East Rongbuk 0.99 0.98

Snow accumulation rates are two to three times lower for the two sites on the northern slope for which there are multi-annual records (Kangwure and Far East Rongbuk) compared to the southern slope sites (Table 2). However, annual fluxes of dust related ions (i.e., Ca2+, Mg 2+, Na+, and ¨C) are still substantially greater for the northern slope sites (Figure 2) indicating that a dilution effect due to greater snow accumulation rates on the southern slopes cannot explain the higher concentrations of these dust related species on the northern slopes. Conversely, fluxes of NO3- and SO42- from northern slope sites are equivalent or up to a factor of 2 greater than those on the southern slopes. Ammonium fluxes vary by a factor of three, with the greatest flux from the Kangwure Glacier and the lowest from the Far East Rongbuk glacier, both on the northern slopes. To further investigate the relationship between ion concentration and snow accumulation, scatterplots for six species [representing dust (Ca2+ and ¨C), salt (Mg2+ and Na+), and anthropogenic sources (SO42- and NO3-)] comparing the mean annual snow accumulation rate and median ion concentration are presented in Figure 3. The scatterplots reveal two distinct groups; high ion concentration/low snow accumulation for northern slope glaciers (circles in Figure 3) and low ion concentration/high snow accumulation for southern slope glaciers (squares in Figure 3). While a cursory inspection of Figure 3 suggests there is a strong ion concentration/snow accumulation relationship, previous work has shown that

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Climatic Interpretation of the Gradient in Glaciochemical Signals

other factors (e.g., sources, transport pathways, wet versus dry deposition) need to be considered when comparing data collected from different geographic regions (e.g., Yang et al., 1996). This is especially true for our data, given the significant differences in source regions, transport pathways, and snow accumulation rates for the south slope verse north slope sampling sites.

Figure -3. Major-ion concentration versus snow accumulation for snowpit/ice-core data for glaciers on the northern (circles) and southern (squares) slopes of the Himalaya.

While there is some overlap, the length and period of record are not the same for all sites (Table 2). The longer records (e.g., 40 years from Kangwure Glacier, 184 years from the Far East Rongbuk Glacier) may be

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influenced by periods of different climate regimes compared to the shorter records. This in turn could skew the average concentrations and mean annual fluxes either higher or lower. To test this, we recalculated median and maximum ion concentrations, and annual fluxes, for the Kangwure and the Far East Rongbuk records beginning in 1981 and compared the results to the longer records. Maximum ion concentrations are quite similar for both periods of record; however there some differences in the median concentrations and annual fluxes. Median concentrations for Ca2+, Mg2+, Na+, ¨C, NO3-, and SO42- are about two times greater , and NH4+and Clconcentrations 1.3 times greater, in the longer records. The annual fluxes also show values that are about one-third greater for the Kangwure record and two-thirds greater for the Far East Rongbuk record. Even considering these differences, the median ion concentrations and annual fluxes for dust related species from the longer north slope records are still five to ten times greater than, and the NO3-,SO42- and NH4+ fluxes are similar to, the south slope records. This indicates that the Himalayan glaciochemical data presented here is relatively robust, and the conclusions drawn below are supported by comparison of either the longer or shorter term data sets.

5.

DISCUSSION

There exists a strong gradient in dust related (Ca++, Mg++, Na+, and ǻC) ion concentrations and fluxes across the Himalaya, with northern slope sites showing significantly greater values (Table 2; Figure 2). The primary source of dust for northern slope glaciers appears to be dust derived from the arid and semi-arid regions of central Asia. Previous research (e.g., Wake and Mayewski, 1993; Wake et al., 1993) revealed a strong spatial gradient in the concentration of these ions in Central Asia, with decreasing concentrations the further the glacier is from the major dust source regions - the large arid and semi-arid regions in the northern and western Tibetan Plateau.

5.1

Discussion of Dust Generation

There exists a strong gradient in dust related (Ca2+, Mg2+, Na+, and ¨C) ion concentrations and fluxes across the Himalaya, with northern slope sites showing significantly greater values (Table 2; Figure 2). The primary source of dust for northern slope glaciers appears to be dust derived from the arid and semi-arid regions of central Asia. Previous research (e.g., Wake and Mayewski, 1993; Wake et al., 1993) revealed a strong spatial gradient in the concentration of these ions in Central Asia, with decreasing concentrations the further the glacier is from the major dust source regions - the large arid

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and semi-arid regions in the northern and western Tibetan Plateau. Several arid and semi-arid regions in central Asia act as source area for dust (Goudie, 1983; Middleton et al., 1986) that are subsequently transported eastward over the Tibetan Plateau, eastern China and the Pacific ocean. There are potential local sources of dust that may be deposited in the accumulation zones of northern slope glaciers. These include the exposed rock walls of the glacier basin as well as extensive glaciofluvial deposits downstream (e.g., Williams, 1983). However, if these local deposits were the predominant sources for dust deposited on glaciers in the Himalaya, we would expect similar concentrations in snow and ice samples collected from both the northern and southern slopes of the Himalayas. This is clearly not the case, providing additional evidence that distant north and north-westerly regions are the source for the majority of dust deposited on the northern slopes of the Himalaya. Furthermore, the south-north gradient and the relative proportions of dust related species for the north slope sites (i.e., Ca2+ > Mg2+ > Na+ and ∆C > Cl-) argues for an inland desert dust source for northern slope glaciers as opposed to a modern marine source to the south. The sharp division between dust concentrations on the northern and southern slopes of the Himalaya indicate that the main crest of the Himalaya not only defines different climatic zones, but also separates two different air masses with very different dust loads. For the northern slope sites, SO42-, NO3- and NH4+ show spatial variability that is similar to Ca2+, and ¨C. For example, Kangwure glacier shows approximately twice as much Ca2+ , ¨C, SO42-, NO3- and NH4+ compared to the Far East Rongbuk (FER) glacier. The time-series of all major ions from Kangwure and FER also show similar temporal variability, with correlation coefficients among all ions greater than 0.75. This suggests that most of the major ion chemistry for glaciers on the northern slopes has either a common source (i.e., dust generated from the arid and semi arid regions of the Tibetan Plateau) or has separate sources that lie at sufficient distance upwind (i.e., dust combined with anthropogenic emissions from upwind sources such as Europe and Russia) that the major ions from different sources have become well mixed by the time they are deposited on the northern slopes of the eastern Himalaya. In contrast to the situation for dust, the strongest regional source of anthropogenic aerosols (such as SO42-, NO3- and NH4+) lie to the south of the Himalaya. For example, the 1999 Indian Ocean Experiment documented an extensive pollution layer over much of the Indian subcontinent and Indian Ocean (Lelieveld et al., 2001). Analysis of aerosol and precipitation samples over an entire year in the Khumbu Himal (Shrestha et al., 2000a; 2002) show strong seasonal variability in major ion concentrations, with SO42-, NO3- and NH4+ concentrations 5 to 10 times greater in the pre-monsoon

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(spring) season compared to the summer monsoon season. These seasonal variations can be explained by changes in precipitation regimes associated with large scale changes in atmospheric circulation. Precipitation during the summer monsoon plays an important role in cleansing the atmosphere and results in dramatic reduction of aerosol concentrations in the Himalaya. Conversely, a dearth of precipitation during the pre-monsoon season is characterized by a gradual buildup of pollutants in the Himalaya. Snow on the southern slopes of the Himalaya displays similar or lower fluxes of NO3- and SO42-, and similar fluxes of NH4+, compared to snow on the northern slopes. Given the extensive pollution that characterizes the atmosphere over India, fluxes of NO3- and SO42- from the southern slopes of the Himalaya are surprisingly similar to fluxes measured in snow from the South Pole and from pre-1900 A.D. snow from Summit, Greenland (Table 4). Table -4. Comparison of annual fluxes of NO3 and SO42- in Himalayan and polar snow. Species Nangpai Kangchung Chago SouthPole* Summit Gosum Greenland* SO4211 26 31 9 16 NO323 59 48 11 23 100 54 70 <1 7 NH4+ *South Pole and Summit, Greenland data from Whitlow et al,m 1992.

Even though precipitation samples collected during the pre-monsoon contain an anthropogenic influence (Shrestha et al., 2002), similarities in the annul fluxes of NO3- and SO42- between the Himalayas and the interior of the polar ice sheets indicates that a strong anthropogenic signal is not recorded in Himalayan snow. This is at least in part due to the limited precipitation that occurs during the pre-monsoon season. The precipitation that does fall in the summer on Himalayan glaciers is part of the same precipitation regime that serves to scavenge aerosols and thereby cleanse the air masses before they reach the Himalaya. The data suggests that the interannual to decadal variability of anthropogenically derived major ions such as SO42- and NO3on the southern slopes is controlled primarily by changes in atmospheric circulation and the amount and intensity of precipitation that the air mass experiences as it travels from the source region to the Himalaya. Any reliable interpretation of glaciochemical time-series that relate to changes in source strength must first take this into account. The distinct groupings revealed in Figure 3 reflect a combination of environmental factors that distinguish the southern slopes from the northern slopes. Climatically, this is reflected by different precipitation regimes with relatively wet conditions on the southern slopes and relatively dry conditions on the northern slopes. The source regions, transport pathways, and

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deposition processes for air masses also differ across the Himalaya. On the northern slopes, the major ion signals indicate that relatively “dirty” air masses, derived primarily from the dust producing regions to the west and north, dominate. The high ion concentration in low snow accumulation areas suggests that most of the ions are deposited via dry deposition. On the southern slopes, the absence of a dust related signal combined with an understanding of regional circulation patterns associated with the monsoon indicates air masses are derived primarily from southerly sources. These southerly-derived air masses are initially very polluted, especially during the pre-monsoon period. However, precipitation scavenging during transport of the summer monsoon plays an important role in reducing aerosol concentrations in the Himalaya (Shrestha et al., 2002). The relatively low ion concentrations and high snow accumulation rate, even for pollution derived aerosols such as NO3- and SO42-, suggests that dry deposition does not appear to be a dominant process for southern slope snowpacks.

6.

CONCLUSIONS

Samples from firn/ice cores and snowpits have been collected for analysis of major-ion concentrations (Na+, NH4+, K+, Mg2+, Ca2+, Cl-, NO3-, SO42-) from high elevation accumulation zones from six glaciers in the Himalaya; three on the north side and three on the south side of the main crest. Depthage relationships, established via annual layer counting and identification of marker horizons, allow for calculation of annual accumulation and fluxes of major ions. Confirming previous studies, concentrations and fluxes of dust related major ions (Ca2+, Mg2+, Na+, and ¨C) are 2-3 orders of magnitude greater for northern slope sites compared to the southern slope sites. The current study defines the relatively short distance over which this change occurs with the boundary defined by the main crest of the Himalaya. This crest not only defines different climatic regimes, but also different air masses. Glaciers to the north of the main crest are strongly influenced by dust derived from the arid regions in central Asia. As a result, glaciochemical records that document changes in the dust source regions and/or changes in atmospheric circulation patterns that transport dust across central Asia can and should be developed from north slope glaciers (e.g., Wake et al., 2001). While regional sources of pollution are derived primarily from the south, the spatial variation of pollution related aerosols does not display a strong gradient across the Himalaya. In fact, annual fluxes of NO3- and SO42- in Himalayan snow are comparatively low. This is in part due to the scavenging of aerosols in summer monsoon air masses by precipitation

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during transport to the Himalaya. Thus the pollution signal stored in Himalaya glaciers, especially those on the southern slope, is primarily a function of atmospheric transport and the amount/intensity of precipitation. Extracting pollution related signals from these glaciochemical records requires that changes in atmospheric circulation and precipitation be taken into account. Nitrate and SO42- deposition on the north slope glaciers are closely linked to dust deposition. Extracting anthropogenic signals from the dust dominated glaciochemical record stored in north slope glaciers will require innovative approaches, such as developing detailed records of trace elements, rare earth elements, and sulfur isotopes (e.g., Kreutz and Sholkovitz, 2000). Multi-parameter glaciochemical records provide the basis for identifying potential sources and transport pathways for major ions and thereby identifying the causes for spatial variability in major-ion records preserved in mid- and low-latitude glaciers. Improving our understanding of these processes is critical for increasing our confidence in paleoclimate reconstructions from ice-core records in these regions. The spatial variability of glaciochemical records presented in this paper will assist in the interpretation of longer-term multi-parameter ice core records that have recently been developed from glaciers in the eastern Himalaya.

7.

REFERENCES

Barry, R.G., 1981, Mountain Weather and Climate. Methuen, New York. Buck, C.F., P.A. Mayewski, M.J. Spencer, S.I. Whitlow, M.S. Twickler, D. Barrett, Determination of major ions in snow and ice cores by ion chromatography, J. Chrom., 594:225-228, 1992. Chalise, S.R., Shrestha, M., Thapa, K.B., Shrestha, B.R., and Bajracharya, B 1996 Climatological and Hydrological Atlas of Nepal, International Center for Integrated Mountain Development (ICIMOD), Kathmandu, Nepal. Domrös, M. and Peng, G., 1988, The Climate of China. Springer Verlag, Berlin. 361 pp. Goudie, A.S., 1983, Dust storms in space and time. Progr. Phys. Geogr. 7, 502-530. Inoue, J. 1976. Climate of the Khumbu Himal. Seppyo 38 (special edition), 66-73. Kreutz, K.J. and Sholkovitz, E., 2000, Major element, rare earth element, and sulfur isotopic composition of a high elevation firn core: Sources and transport of mineral dust in central Asia. Geochemistry, Geophysics, Geosystems 1, paper 2000GC000082. Lelieveld, J. and 26 others., 2001, The Indian Ocean Experiment: Widespread air pollution from south and southeast Asia. Science 291, 1031-1036. Middleton, N.J., Goudie, A.S., and Wells, G.L., 1986, The frequency and source areas of dust storms. In: W.G. Nickling (ed.) Aeolian Geomorphology, Allen & Uniwin, Boston, pp. 237-259. Shrestha, A.B., Wake, C.P. Dibb, J.E., and Whitlow, S.I., 2002, Aerosol and precipitation chemistry at a remote Himalayan site in Nepal. Aerosol Science and technology, 36, 441456.

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Shrestha, A.B., Wake, C.P., Dibb, J.E., Mayewski, P.A., Whitlow, S.I., Carmichael, G.R., and Ferm, M., 2000a, Seasonal variations in aerosol concentrations and compositions in the Nepal Himalaya. Atmos. Environ. 24, 3349-3363. Shrestha, A., Wake, C.P., Dibb J., and Mayewski, P., 2000b, Precipitation fluctuations in the Nepal Himalaya and its vicinity and relationship with some large scale climatological parameters. International J. Climatol. 21, 317-327. Wake, C.P., Mayewski, P.A., Qin D., Yang Q., Kang S., Whitlow, S., and Meeker, L.D., 2001, Changes in Atmospheric Circulation over the South-Eastern Tibetan Plateau over the last Two Centuries from a Himalayan Ice Core. PAGES News. Vol 9, No. 3, p. 14-16. Wake, C. and Stievenard, M., 1996, The amount effect and oxygen isotope ratios recorded in Himalayan snow. Proceedings of the IGBP-PAGES/PEPII Nagoya Symposium, Nov. 28Dec. 1, 1995, Nagoya, 236-241. Wake C.P., Mayewski, P.A., Li Z., Han J., and Qin, D.,1994, Modern eolian dust deposition in central Asia. Tellus 46B, 220-233. Wake, C.P. and Mayewski, P. A., 1993, The spatial variation of Asian dust and marine aerosol contributions to glaciochemical signals in central Asia. In G. Young (Ed.), International Symposium on Snow and Glacier Hydrology, Kathmandu. IAHS Pub. No. 218, 385-402. Wake, C.P., Mayewski, P.A., Xie, Z., Wang, P., and Li, Z., 1993, Regional variation of monsoon and desert dust signals recorded in Asian glaciers. Geophys. Res. Lett. 20, 14111414. Wake C.P., Mayewski, P.A., Wang P., Yang Q., Han J., and Xie Z., 1993, Anthropogenic sulfate and Asian dust signals in snow from Tien Shan, northwest China. Annal. Glaciol. 16, 45-52. Whitlow, S., Mayewski, P.A., and Dibb, J.E., 1992, A comparison of major chemical species seasonal concentration and accumulation at the South Pole and Summit, Greenland. Atmos. Environ. 26A, 2045-2054. Williams, M.W., K.A. Tonnessen, J.M. Melak and D. Yang, 1992, Sources and spatial variation of the chemical composition of snow in the Tien Shan, China, Annals Glaciol. 16, 25-32. Williams, V.S., 1983, Present and former equilibrium-line altitudes near Mount Everest, Nepal and Tibet. Arctic and Alpine research 15, 201-211. Yang, Q., P.A. Mayewski, E. Linder, S. Whitlow and M. Twickler. 1996, Chemical species spatial distribution and relationship to elevation and snow accumulation rate over the Greenland ice sheet. J. Geophys. Res 101, 18,629-18,637.

RECONSTRUCTION OF EUROPEAN AIR POLLUTION FROM ALPINE ICE CORES

Margit Schwikowski

1.

INTRODUCTION

There is a rather long tradition of glaciological research in the European Alps with, for instance, 200-m deep boreholes drilled nearly a hundred years ago in order to measure glacier thickness. However, detailed paleo ice-core studies were not performed before 1977. One of the reasons was that glaciers in the Alps were generally believed to be temperate. In a temperate glacier melt-water percolation destroys the chemical stratigraphy of water-soluble components, making such a glacier unsuitable for paleo-chemical investigations. However, in mid and low latitudes, the existence of so-called cold glaciers (temperature well below the pressure melting point) is strongly dependent on altitude due to the temperature lapse rate in the atmosphere. In the Alps, sufficiently cold firn temperatures that suggest a cold glacier are generally found above 4000 masl. in the northern part and above 4300 masl. in the southern part (Funk, 1994; Suter et al., 2001). Glaciers in the Alps are well suited for the purpose of reconstructing atmospheric concentrations of short-lived species and to document the effects of anthropogenic emissions on air pollution in Europe, since they are located in the center of the highly populated and industrialized areas of Europe. A good example of such short-lived species are atmospheric aerosol particles which have attracted attention in climatic research recently since their cooling effect on the atmosphere might partly compensate the warming by anthropogenically emitted greenhouse gases. In order to quantify the 95 L. D. Cecil et al. (eds.), Earth Paleoenvironments: Records Preserved in Mid- and Low-Latitude Glaciers, 95-119. © 2004 Kluwer Academic Publishers. Printed in the Netherlands.

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effect of anthropogenic aerosols, reliable concentration data sets from the source regions are necessary. Glaciochemical records from the European Alps offer the possibility to test the reliability of translating ice-core proxy data into paleoclimatic and paleoatmospheric records by comparing them with corresponding historical records and historical data from meteorological and air quality measurements (Preunkert et al., 2002; Schwikowski et al., 1999a, Wagenbach, 1997). This combination of suitable glacier archives with longterm, high quality instrumental records available in the vicinity is unique. Alpine ice-core data from pre-industrial times allow the establishment of natural concentration levels for various atmospheric species whose atmospheric cycles are now influenced by human activities (Harnisch et al., 1996). However, glaciers in the Alps are retreating rapidly due to global warming, which poses a threat to these valuable archives. For example, partial damage of a glaciochemical record due to melt-water percolation in the firn part of an ice core drilled at 4200 masl from the Grenzgletscher was recently observed (Eichler et al., 2001).

2.

SUITABLE GLACIERS AND ICE CORES RETRIEVED

In the Alps, only a few potential ice core drilling sites exist. For the purpose of paleo-environmental and paleo-climate studies, the main criterion for selecting a glacier archive is the absence of both melt-water percolation and temperate firn. The influence of melt-water percolation is absent or minimal only in the recrystallization, the recrystallization-infiltration, and cold infiltration zones of cold glaciers (Shumskii, 1964). In the recrystallization-infiltration zone melt-water, which has formed at the surface by solar radiation, immediately refreezes some centimeters below the surface. In order to facilitate understanding and modeling of glacier flow, geometry is preferred where fast horizontal advection of ice from steep slopes does not occur towards the drilling area. Hence, ice caps or saddles between peaks are favorable (Funk, 1994). In addition, the glacier at the drilling site should show maximal thickness in order to extend the time period covered by the ice core as far back into the past as possible. Following these criteria, ten different sites in four high-elevation areas in the Alps were identified (Funk, 1994): Fieschersattel, Aletschhorn, Jungfrau (in the Bernese Alps), Colle Gnifetti, Seserjoch, Nordend (in the Monte Rosa area), Col de Valsorey (Grand Combin), and Col du Dôme, Col Major, Col de la Brenva (in the Mont Blanc area). Up to now major drilling projects have been conducted in three of these areas. The locations of the areas are

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shown in Figure 1 and the key parameters of the resulting ice cores are listed in Table 1.

3.

DATING AND TIME PERIOD ACCESSIBLE

The main dating tool applied for Alpine ice cores is the counting of annual layers of one or more seasonally varying parameters. Parameters such as the stable isotope ratios δ18O ( Schotterer et al., in press; Eichler et al., 2000b; Schotterer et al., 1998) and δD (Preunkert et al., 2000), tritium (Schwikowski et al., 1999a), NH4+ (Eichler et al., 2000b), aluminium (de Angelis and Gaudichet, 1991), Ca2+, and ice crystal size (de Angelis and Gaudichet, 1991) were used. As an example the annual signal cycles of the tritium activity concentration and δ18O along the Fiescherhorn ice core are shown in Figure 2.

Figure 1. Topographic map of the Alps showing the locations of the three high-elevation areas: Bernese Alps (Fiescherhorn), Mont Blanc, and Monte Rosa where major drilling projects have been conducted. The Grand Combin (not marked) is located close to the Monte Rosa. Inset: Map of Europe showing geographical extension of the Alps

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Table 1. Key parameters of major ice cores (i.e. penetrating the firn-ice transition) retrieved from glaciers in the Alps Area

Glacier

Elevation (m a.s.l.)

Year

Ice core length (m)

Time period covered (years)

Monte Rosa

Colle Gnifetti

4450

1977

55 and 65

About 200a

Colle Gnifetti

4450 4470

1982

124 (bedrock), 109 Greater than 2000*b 66 (bedrock)

Colle Gnifetti

4450 4480

1995

101 (bedrock) 62 (bedrock)

about 1000*c

Grenzgletscher 4200

1994

125

1994-1937 ±2d

Col du Dôme

4250

1986

70

1986-1955±1e

Col du Dôme

4250

1994

139 (close to bedrock), 126

1994-1919±10f

Mont Blanc

* Estimation by glaciological flow modeling, see Figure 3 a Oeschger, 1977 b Schotterer et al., 1985 c Luethi and Funk, 2001 d Eichler et al., 2000b e de Angelis and Gaudichet, 1991 f Vincent et al., 1997 and Pruenkert et al., 2000

Distinct stratigraphic markers, for instance, from the atmospheric nuclear-weapons tests conducted between 1952 and 1962 (tritium (Döscher et al., 1995; Eichler et al., 2000b; Schotterer et al., 1985; Schwikowski et al., 1999a; Vincent et al., 1997; Wagenbach et al., 1988), gross-β radioactivity (de Angelis and Gaudichet, 1991), 137Cs (Eichler et al., 2000b)), the reactor accident in Chernobyl (137Cs (Eichler et al., 2000b)), Saharan dust falls (yellowish layers, Ca2+ (Döscher et al., 1995; Eichler et al., 2000b; Haeberli, 1977; Schotterer et al., 1985; Wagenbach and Geis, 1989; Wagenbach et al., 1988; Wagenbach et al., 1996)) and volcanic eruptions (SO42-, SO42- to Ca2+ ratio (Schäfer, 1995; Schwikowski et al., 1999b)) are used to reduce the error of annual layer counting. Nuclear dating based on the radioactive decay of 210Pb (Döscher et al., 1995; Eichler et al., 2000b; Gäggeler et al., 1983; von Gunten et al., 1982) as well as the distinct methane concentration trend, well known and dated from Greenland ice cores, is also applied to support establishing a time scale (Blunier, 1995; Dällenbach, 2000).

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Using these methods, precise time scales have been established for the upper parts of the ice cores, whereas the age of the deepest 20 percent is still undetermined (Figure 3). Since the global methane concentration trend before 1800 is relatively indifferent, the dating based on methane concentrations prior to this time is not very accurate and does not exclude the time scale deduced from annual layer counting (T. Blunier, personal communication). Several observations in the ice cores from Colle Gnifetti, such as the dramatic decrease of δ18O close to bedrock (Wagenbach, 1994) and the low δ18O values of O2 in air bubbles (U. Schotterer, personal communication) suggest that the basal ice might consist of Pleistocene precipitation.

Figure 2. Variations of the concentrations of tritium and of δ18O along the Fiescherhorn ice core (reproduced by permission of American Geophysical Union from Schwikowski et al., 1999.

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The Colle Gnifetti site is still believed to contain the longest paleo records in the Alps, although a detailed radar survey at the Fiescherhorn glacier showed a surprisingly large ice thickness of up to 150 m in the accumulation area (M. Funk, personal communication), which might contain a long-term paleo-record spanning several centuries. The chronologies of all ice cores are generally in agreement with modeled depth-age relationships (Haeberli et al., 1988b; Luethi and Funk, 2000; Luethi and Funk, 2001; Vincent et al., 1997; Wagner, 1996).

Figure -3. Age-depth relationship of the 124-m ice core from Colle Gnifetti. The dating points from annual layer counting of Ca2+(cross), stratigraphic markers (diamonds, triangle, and squares), and the methane concentration (circles, (Dällenbach, 2000)) as well as results from 3D-modelling (Luethi and Funk, 2000; Wagner, 1996) are depicted. Note the large uncertainty for the deepest 20 m.

4.

RECONSTRUCTED AIR POLLUTION RECORDS

Various concentration records of a number of chemical trace species and gases obtained from the different Alpine ice cores have been published. These records clearly demonstrate the impact of anthropogenic emissions on the impurity content of snow. They show a generally consistent picture of a vastly altered atmospheric composition due to industrialization. This is the

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case for major aerosol components, trace constituents, gaseous species, and radionuclides, which are discussed below. The total SO42- record shows an increasing trend beginning in 1870 and reaching a maximum between about 1970 to 1985, with declining concentrations afterwards (Preunkert et al., 2001b; Schwikowski et al., 1999b) (Figures 4 and 5). Between 1760 and 1870 a constant level was observed (Schwikowski et al., 1999b). Using Na+ and Ca2+ as sea-salt and mineral dust tracers, respectively, it was demonstrated that the increasing trend in SO42- was due to anthropogenic SO2 emissions and not variations in the contributions of the natural sources (Figure 4), since sea-salt and mineral dust SO42- records do not show a trend. The contribution to total SO42- of these two natural sources is negligible, except in the case of mineral dust during pre-industrial times (Table 2). The exSO42- concentration, which is the total SO42- corrected for the sea-salt and mineral dust contribution and which is assumed to originate exclusively from oxidation of SO2 in the atmosphere, increased dramatically from 0.047 mg L-1 in the pre-industrial (1760 to 1870) to 0.602 mg L-1, in the maximum in the industrial time period (1963-1981) (Schwikowski et al., 1999b) (Figure 4). This exSO42- record is interpreted as a direct measure of anthropogenic SO2 emissions. The record exhibits interesting details, e.g. the decrease in the exSO42- 5-year average after World Wars I and II as well as the extremely steep increase after 1950, due to the enhanced combustion of liquid fuels, replacing coal combustion. Between 1963 and 1981 anthropogenic exSO42- formed the major portion (79 percent) of the total SO42- (Table 2).

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Figure -4. Record of 5-year averages of exSO4 2- from the 1982 Colle Gnifetti ice core, obtained by applying the following formula to each sample: (exSO4 2-) = (SO4 2-) - 0.12 (Na+) - 0.175 - (nssCa2+) (concentrations in ueq L-1), along with the corresponding records of seasalt and mineral dust derived SO4 2- (reproduced with permission from Schwikowski et al.)

To identify potential source regions of SO42- in the high-alpine ice cores the data were compared with SO2 emission estimates. For the Monte Rosa area the historical development of exSO42- is similar to the evolution of SO2 emissions in Switzerland, France, and West Germany, where the emission maximum was reached in 1965-1975. This is in agreement with results from trajectory analyses studying the transport from the Po valley to high-alpine sites. They showed that transport from this major source area in Italy does not seem to cause higher-than-average concentrations of anthropogenic species at high-alpine sites (Seibert et al., 1998). In contrast, countries having an impact on the snow SO42- concentration in the Mont Blanc area in summer include regions within a 700-1000 km range with major contributions from France, Spain, and Italy. SO2 emissions from northern countries such as United Kingdom, Belgium, and West Germany have a weaker impact (Preunkert et al., 2001b). For the winter situation, the SO42record suggests a larger scale contamination with contributions from total Europe and possibly from the United States (Preunkert et al., 2001b). This is in agreement with the concept that in winter the high-alpine sites are

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decoupled from the planetary boundary air contamination and represent the free troposphere. The large enhancement of atmospheric SO42- levels above 4000 masl in summer compared to winter contrasts with the situation at the surface, with little variability between summer and winter levels (Figure 6). This trend reflects the more efficient convective upward motion of air masses from the polluted boundary layer in summer compared to winter. Table 2. Average concentrations of SO42- in mg L-1 from different sources and their percentage contribution to the total SO42- for the industrial (1963-1981) and pre-industrial (1756-1870) time period deduced from the 1982 Colle Gnifetti ice core (reproduced with permission from Schwikowski et al., 1999) Mineral dust Total SO42Sea-salt SO42ExSO42SO42Natural* industrial 0.707 0.011 0.092 0.047 (2%) (13%) (6%) pre-industrial 0.131 0.009 0.074 0.047 (7%) (57%) (36%) * Estimation of anthropogenic exSO42- is based on the conservative assumption that the preindustrial exSO42- was purely of natural origin and that this natural contribution remained constant over the entire time period.

Two other major ionic aerosol components analyzed in Alpine ice cores reflect the pollution history of Central Europe: NO3- and NH4+. nitrate is a product of the precursor gases NOx, showing an exponential increase in the ice core record from 1930 to 1965 (Döscher et al., 1995). Prior to 1930 a constant level of 0.067 ± 0.005 mg L-1 was observed, compared to 0.15 mg L-1, the maximum during the period 1965-1981. Thus, the NO3- record reflects increasing emissions from traffic. NH4+ is formed in the atmosphere via neutralization of NH3, which is the primary gaseous alkaline species in the atmosphere over Europe, neutralizing up to 70 percent of the original acidity in precipitation. It is involved in the conversion of SO2 and NOx into the aerosol phase. The high-resolution ice core record of this environmentally relevant species obtained from the 1982 Colle Gnifetti ice core showed a three-fold concentration increase between 1880 and 1980 (Döscher et al., 1996) (Figure 7), indicating that the NH3 emissions in Europe have substantially increased in the 20th century. Comparisons with estimated anthropogenic NH3 emissions from animal manure and application, and production of fertilizers in Europe suggest that the historical emissions from anthropogenic sources are slightly overestimated (Figure 7).

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Figure 5. (left). Comparison of the anthropogenic SO42- trend observed in summer samples of the Col du Dôme core (smoothed profile, first component of single spectra analysis with a 5 year time window) with SO2 emission inventories from (Lefohn et al., 1999) (thin solid line) and (Mylona, 1996) (dashed line) considering different countries within 700 km around the Alps: Top: France, Switzerland, Spain, Italy, and half of the emissions from former West Germany. Middle: France, Switzerland, Spain, and Italy. Bottom: France, Switzerland, and Italy (reproduced by permission of American Geophysical Union from S. Preunkert et al., 2001.)

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Figure -6 (Right) Atmospheric SO42- concentrations in winter and summer versus elevation using observations made in 1991-1993 at various European sites located around the Alps (Payerne on the Swiss Plateau; Schauinsland in Germany; Sonnblick in Austrian Alps; Jungfraujoch in the Swiss Alps) and inverted data from the Col du Dôme ice core using firn/air ratios (Preunkert et al., 2002) (reproduced by permission of American Geophysical Union from S. Preunkert et al., 2001).

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Figure 7. Ten-year average NH4+ concentrations for the period 1780 to 1890 (solid line). In addition, anthropogenic NH3 emissions from animal manure and application and production of fertilizers in Europe estimated for the period 1870 to 1980 are shown (dashed line, adopted from (Asman et al., 1988)) (reproduced by permission of American Geophysical Union from A. Döscher et al.)

A trend comparable to that of SO42- is observed for concentrations of carbonaceous particles such as black carbon (BC), elemental carbon (EC), water insoluble organic carbon (OC), and total carbon (TC) (Lavanchy et al., 1999) (Figure 8). These carbonaceous particles are emitted by combustion of fossil fuels and by biomass burning, and may alter the radiation balance by absorbing and scattering light, thus influencing the climate. The BC, EC, OC, and TC concentrations in the 1982 Colle Gnifetti ice core began to increase after 1890, corresponding to the start of industrialization. From preindustrial times (1755-1890) to modern times (1950-1975) BC, EC, OC, and TC concentrations increased by a factor of 3.7, 3.0, 2.5, and 2.6, respectively (Lavanchy et al., 1999).

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Figure 8. Ice core concentrations (dashed line) and 10-year averages (solid line) of organic carbon (OC), elemental carbon (EC), black carbon (BC), and total carbon (TC) for the time period 1755 to 1975 deduced from the 1982 Colle Gnifetti ice core. BC concentrations were calculated using αAPI = 9.3 m2 g-1 (reproduced by permission of American Geophysical Union from V.M.H. Lavanchy et al., 1999)

The sum of BC emissions of Germany, France, Switzerland, and Italy, calculated from fossil fuel consumption considering coal, wood, and petroleum products, correlates well with the EC concentrations in the ice core for the time period from 1890 to 1975 (R2 = 0.56). This indicates that EC concentrations in the ice core reflect the emissions of Western Europe,

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which corroborates the interpretation of the SO42- record (see above). For the time period 1890-1975, SO42- and TC concentrations correlate (R2 = 0.43), suggesting that both species have a common dominating source: combustion. Prior to 1860, elevated levels of BC were observed, indicating that emissions to the atmosphere of carbonaceous particles were already significant. A weak correlation between BC and exSO42- (R2 = 0.36) for the pre-industrial period suggests that these emissions resulted from wood combustion for domestic heating (Schwikowski et al., 1999b). Trace constituents such as heavy metals also show significantly enhanced concentrations in Alpine ice cores due to anthropogenic activities. For instance, Pb concentrations analyzed in two firn/ice cores from Colle Gnifetti show values 24 times larger in firn dated from the 1970’s than in ice dated from the 17th century, confirming the massive rise in Pb pollution in Western Europe during the last few centuries (Figure 9). A decline in the concentration is then observed during the last two decades, i.e. from 1975 to 1995. A major feature is the large increase in Pb concentrations from about the 1930s to the 1970s, and the subsequent decrease to the mid 1990s. This feature was also observed in a Mont Blanc ice core (Rosman et al., 2000) and in other European atmospheric archives such as peat bogs (Shotyk et al., 1998) and lake sediments (Moor et al., 1996; Von Gunten, 1997). It is clearly linked with the rise and fall of the use of Pb additives in European countries. The long time period (back to the 17th century) covered by this record facilitated the observation of trends prior to 1930. Very low concentrations are found before 1800 and are followed by progressively increasing values from the end of the 18th century to the late 19th century – early 20th century. Concentrations then decrease around 1930. These variations are due to prePb additive, anthropogenic emissions with major contributions from nonferrous metal productions, iron and steel manufacturing, and coal and wood combustion (Nriagu, 1978; Nriagu, 1998). From the middle of the 19th century to 1910, a significant proportion of the Pb concentrations can be attributed to emissions from coal burning, resulting in a secondary Pb peak in 1910. Afterwards, the worldwide economic recession led to a collapse of the Pb emissions, which recovered within a period of 10 years and stabilized until the end of World War II. A level of about 1000 ng L-1 Pb from mining and coal burning (Figure 9) was further elevated during the introduction and increased use of leaded gasoline after 1950. Following this, concentrations declined due to a significant reduction of Pb emissions of various Western European countries as a consequence of abatement measures.

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Figure 9. Lead paleo concentration record from the Colle Gnifetti ice core as 5-year (period 1900-1995) and ten-year averages (period 1650-1900) M. Schwikowski et al.

The observed changes in the Pb concentration are generally accompanied by respective changes in the Pb isotopic ratio 206Pb/207Pb, which decreased from about 1.18 in the 17th and 18th century to about 1.12 in the 1970s (Schwikowski et al., manuscript in preparation, Rosman et al., 2000). Other metals that are of high environmental interest due to the significant influence of human activities on their atmospheric cycles include: cobalt (Co), chromium (Cr), molybdenum (Mo), antimony (Sb), silver (Ag), gold (Au), platinum (Pt), palladium (Pd), and rhodium (Rh). Data from the Col du Dôme ice core reveal significantly higher concentrations in recent ice than in ice dated before the middle of the 19th century for all metals except Au and Pt (Van de Velde et al., 2000; Van de Velde et al., 1999). There are however differences in the timing and the amplitude of the observed increase from one metal to another as shown in Figures 10 and 11. The dominant sources

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for the enhanced concentrations of Co, Mo, and Sb in recent ice were suggested to be emissions from oil and coal combustion, whereas Cr is released from iron, steel, and ferroalloy industries (Van de Velde et al., 1999). A qualitative review showed volcanoes, mining and smelting activities, industry and waste incinerators to be possible natural and anthropogenic sources of Ag, Au, Pt, Pd, and Rh (Van de Velde et al., 2000). Aside from the long-lived methane (Blunier, 1995) historical concentration records of two more reactive gaseous species were reconstructed from Alpine ice cores, namely HCl and HF. On the average 16 percent of the Cl- and most of the F- deposited on Grenzgletscher in the Monte Rosa area in the period 1937-1994 could be related to HCl and HF emissions from anthropogenic sources (Eichler et al., 2000a). The record of non-sea-salt Cl- was found to represent the historical development of HCl emissions mainly from waste incineration on the Swiss Plateau (100-200 km distance to ice core site), which increased strongly until 1985 due to rising waste incineration and the large amounts of PVC in waste (Figure 12). Coal burning is considered a minor source of HCl emissions in Switzerland due to the opposite trend of coal consumption. As a consequence of the installation of flue gas scrubbers, a decline of HCl emissions from waste incinerators has occurred since 1985.

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Figure 10) Changes in Co, Cr, Mo, and Sb concentrations in snow and ice deposited at a high altitude site in the French-Italian Alps since the end of the 18th century. Individual data have been averaged over periods of 1-3 years (reprinted with permission from K. Van de Velde et al., 1999)

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Figure 11. Changes of mean annual concentrations of Ag, Au, Pt, Pd, and Rh in Mont Blanc snow and ice dated from the 18th century to the early 1990s (reprinted from Atmospheric Environment 34, K. Van de Velde et al., 2000)

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The F- deposition record reflects HF emissions from the aluminum industry in the Swiss Rhone valley (about 40 km distance to ice core site) (Figure 12). A steady concentration increase until 1965 is observed in the ice core record, followed by a steep decline due to the installation of waste-air purification systems after 1970. The data indicate a strong impact of emissions of HCl and HF on the local and regional precipitation chemistry. This is explained by the high water solubility of these gases, leading to a complete uptake by cloud droplets and subsequent wet deposition close to the sources. The findings are consistent with the F- record from the Col du Dôme ice core, where 86 ± 3 percent of the total F- deposited in the late 1960s was attributed to emissions from aluminum smelters. Coal burning accounted only for 8 ± 2 percent of the total F- during the same time period (Preunkert et al., 2001a). Prior to 1880, soil dust emissions dominated the atmospheric F- budget.

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Figure -12. Five-year average concentration records of non-sea- salt Cl- (A) and of F- (B) (lines) together with historical emission estimations of HCl and HF for Switzerland in the period 1935-90 (dashed lines). Diamonds represent emissions from coal burning, while stars indicate total emissions, derived from (BUWAL, 1995) (reproduced by permission of American Geophysical Union from A. Eichler et al., 2000)

Anthropogenic emissions of various radioactive isotopes are also recorded in Alpine glaciers, e.g. tritium (3H), 137Cs, and 36Cl, which contaminated the atmosphere as a consequence of nuclear-weapons testing between 1952 and 1962. The reactor accident in Chernobyl on April 1986 also produced a 137Cs horizon (Figure 13). The integrated 137Cs activity from nuclear weapons fallout deposited on the Grenzgletscher (3500 ± 800 Bq m2 ) aggress well with estimated values, e.g. 5040 Bq m-2 between 40° and 50° N (Kiefer, 1986). The 137Cs fallout from the Chernobyl accident calculated for the Grenzgletscher (800 ± 200 Bq m-2) was at the lower limit of the values determined at other Alpine sites (between 400 and 15000 Bq m-2)

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(Haeberli et al., 1988a). These differences in the deposited activities reflect the known regional heterogeneity of the Chernobyl fallout, which was strongly dependent on the precipitation pattern in the first few days of May 1986. The overall input at Grenzgletscher of bomb produced 36Cl, which was formed by the neutron activation of 35Cl of seawater, was 4.5 ± 0.8 x 1012 atoms m-2. This agrees well with estimated values of fallout between 40 and 50° N (4.6x1012 atoms m-2), obtained by using a box model of atmospheric transport of radionuclides (Sachsenhauser et al., 1997).

Figure 13. Records of the 137Cs (a) and 3H (b) and 36Cl concentration (c) vs depth of the Grenzgletscher ice core; the activities were calculated back to the date of drilling in October 1994 (reprinted from A. Eichler, et al., 2000)

Increased concentrations in Alpine ice cores were also detected for 129I, a long-lived fission product released continuously into the environment by nuclear fuel processing plants. No bomb peak, similar to the 137Cs bomb peak, was observed for 129I, thus it was concluded that the main source of the present day 129I fallout is nuclear fuel processing (Wagner et al., 1996).

5.

CONCLUSIONS

Paleo records from Alpine glaciers allow the reconstruction of European atmospheric pollution history. These records show a generally consistent picture of a vastly altered atmospheric composition due to anthropogenic

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emissions. Typically following the onset of industrialization, increased concentrations of major aerosol components (e.g. SO42-, NO3-, NH4+, black and elemental carbon), trace constituents (e.g. Pb and other heavy metals), gaseous species (including the greenhouse gas methane), and radionuclides (3H, 137Cs, 36Cl, and 129I) were observed in the ice cores. However, on a positive note, Alpine glaciers are also a reliable indicator of the progress achieved in environmental protection. Since approximately 1970, many of the pollutants mentioned here show a clear tendency to lower concentrations, which is a direct result of various air quality measures such as the use of filtering units in power plants, in incineration plants and in industry, the increased use of oils with low sulfur content, as well as the introduction of catalytic converters and lead-free gasoline.

6.

REFERENCES

Asman, W.A.H., Drukker, B., and Jannssen, A.J., 1988, Modelled historical concentrations and depositions of ammonia and ammonium in Europe, Atmospheric Environment, 22, 725-735. Blunier, T., 1995, Methanmessungen aus Arktis, Antarktis und den Walliser Alpen, Interhemisphärischer Gradient und Quellenverteilung, Thesis, Physikalisches Institut, Universität Bern. BUWAL, 1995, Vom Menschen verursachte Luftschadstoff - Emissionen in der Schweiz von 1900 bis 2010, pp. 122, Bundesamt für Umwelt, Wald und Landschaft, Bern. de Angelis, M., and Gaudichet, A., 1991, Saharan dust deposition over Mount Blanc (French Alps) during the last 30 years, Tellus B, 43 (1), 61-75. Dällenbach, A., 2000, Methan- und Lachgasmessungen aus Arktis, Antarktis und den Alpen, Thesis , Physikalisches Institut, Universität Bern, Bern. Döscher, A., Gäggeler, H.W., Schotterer, U., and Schwikowski, M., 1996, A historical record of ammonium concentration from a glacier in the Alps, Geophysical Research Letters, 23, N.20, 2741-2744. Döscher, A., Gäggeler, H.W., Schotterer, U., and Schwikowski, M., 1995, A 130 years deposition record of sulfate, nitrate and chloride from a high-alpine glacier, Water Air and Soil Pollution, 85 (2), 603-609. Eichler, A., Schwikowski, M., and Gäggeler, H.W., 2001, Meltwater-induced relocation of chemical species in Alpine firn, Tellus Series B-Chemical and Physical Meteorology, 53 (2), 192-203. Eichler, A., Schwikowski, M., and Gäggeler, H.W., 2000a, An Alpine ice-core record of anthropogenic HF and HCl emissions, Geophysical Research Letters, 27 (19), 3225-3228. Eichler, A., Schwikowski, M., Gäggeler, H.W., Furrer, V., Synal, H.W., Beer, J., Saurer, M., and Funk, M., 2000b, Glaciochemical dating of an ice core from upper Grenzgletscher (4200 m a.s.l.), Journal of Glaciology, 46 (154), 507-515. Funk, M., 1994, Possible Alpine ice-core drilling sites, an overview., in Proceedings of the ESF/EPC workshop on Greenhouse gases, isotopes and trace elements in glaciers as climate evidence of the Holocene, edited by W. Haeberli, and B. Stauffer, VAW Arbeitsheft, pp. 40-44, Zürich.

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Gäggeler, H., von Gunten, H.R., Rössler, E., Oeschger, H., and Schotterer, U., 1983, 210PbDating of cold alpine firn/ice cores from Colle Gnifetti, Switzerland, Journal of Glaciology, 29 (101), 165-177. Haeberli, W., 1997, Sahara dust in the Alps - A short review, Zeitschrift fuer Gletscherkunde und Glazialgeologie, 13 (1/2), 206-208. Haeberli, W., Gäggeler, H., Baltensperger, U., Jost, D., and Schotterer, U., 1988a, The signal from the Chernobyl accident in high-altitude firn areas of the Swiss Alps, Annals of Glaciology, 10, 48-51. Haeberli, W., Schmid, W., and Wagenbach,, D., 1988b, On the geometry, flow and age of firn and ice at the Colle Gnifetti core drilling site, Zeitschrift fuer Gletscherkunde und Glazialgeologie, 24 (1), 1-19. Harnisch, J., R. Borchers, P. Fabian, H.W. Gäggeler, and U. Schotterer, 1996, Effect of natural tetrafluoromethane, Nature, 384 (6604), 32-32. Kiefer, H., 1986, The consequences of the Chernobyl reactor accident for the dose commitment of the general public, KFK (Kernforschungszentrum Karlsruhe) Nachrichten, 18 (3), 133-134. Lavanchy, V.M.H., Gäggeler, H.W., Schotterer, U., Schwikowski, M., and Baltensperger, U., 1999, Historical record of carbonaceus particle concentrations from a European highalpine glacier (Colle Gnifetti, Switzerland), J Geophys Res., 104, 21227-21236. Lefohn, A.S., Husar,J.D., and Husar, R. B., 1999, Estimating historical anthropogenic global sulfur emission patterns for the period 1850-1990, Atmospheric Environment, 33 (21), 3435-3444. Luethi, M., and Funk, M., 2000, Dating ice cores from a high Alpine glacier with a flow model for cold firn, Annals of Glaciology, Vol 31, 2000, 31, 69-79. Luethi, M.P., and Funk, M., 2001, Modelling heat flow in a cold, high-altitude glacier: interpretation of measurements from Colle Gnifetti, Swiss Alps, Journal of Glaciology, 47 (157), 314-324. Moor, H.C., Schaller, T., and Sturm, M., 1996, Recent changes in stable lead isotope ratios in sediments of Lake Zug, Switzerland, Environmental Science and Technology, 30 (10), 2928-2933. Mylona, S., 1996, Sulphur dioxide emissions in Europe 1880-1991 and their effect on sulphur concentrations and depositions, Tellus Series B-Chemical and Physical Meteorology, 48 (5), 662-689. Nriagu, J.O., 1978, Biogeochemistry of lead in the environment, Elsevier, New York. Nriagu, J.O., 1998, Global atmospheric metal pollution, in From urban air pollution to extrasolar planets, edited by C.F. Boutron, pp. 205-226, Les Editions de Physique, Paris. Oeschger, H., 1997, First results from Alpine core drilling projects, Zeitschrift fuer Gletscherkunde und Glazialgeologie, 13 (1/2), 193-208. Preunkert, S., Legrand, M., and Wagenbach, D., 2001a, Causes of enhanced fluoride levels in Alpine ice cores over the last 75 years: Implications for the atmospheric fluoride budget, Journal of Geophysical Research-Atmospheres, 106 (D12), 12619-12632. Preunkert, S., Legrand, M., and Wagenbach, D., 2001b, Sulfate trends in a Col du Dome (French Alps) ice core: A record of anthropogenic sulfate levels in the European midtroposphere over the twentieth century, Journal of Geophysical ResearchAtmospheres, 106 (D23), 31991-32004. Preunkert, S., Wagenbach, D., and Legrand, M., 2002, Improvement and characterization of an automatic aerosol sampler for remote (glacier) sites, Atmospheric Environment, 36 (7), 1221-1232.

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Preunkert, S., Wagenbach, D., Legrand, M., and Vincent, C., 2000, Col du Dome (Mt Blanc Massif, French Alps) suitability for ice-core studies in relation with past atmospheric chemistry over Europe, Tellus B, 51 (3), 993-1012. Rosman, K.J.R., Ly, C., Van de Velde, K., and Boutron, C.F., 2000, A two century record of lead isotopes in high altitude Alpine snow and ice, Earth and Planetary Science Letters, 176 (3-4), 413-424. Sachsenhauser, H., Zerle, L., Beer, J., Masarik, J., and Nolte, E., 1997, Atmospheric transport of cosmogenic radionuclides, in Glaciers from the Alps: Climate and Environmental Archives (Proceedings of the Wengen workshop, 21-23 October 1996, Switzerland), pp. 103-114. Schäfer, J., 1995, Rekonstruktion bio-geochemischer Spurenstoffkreisläufe anhand eines alpinen Eisbohrkerns, Diploma Thesis, Universität Heidelberg. Schotterer, U., Oeschger, H., Wagenbach, D., and Münnich, K.O., 1985, Information on paleo-precipitation on a high-altitude glacier, Monte Rosa, Switzerland, Zeitschrift fuer Gletscherkunde und Glazialgeologie, 21 (85), 379-388. Schotterer, U., Schwarz, P., and Rajner, V., 1998, From pre-bomb levels to industrial times: A complete tritium record from an alpine ice core and its relevance for environmental studies, in International symposium on isotope techniques in the study of past and current environmental changes in the hydrosphere and the atmosphere, pp. 581-590, Vienna. Schotterer, U., Stichler, W., Graf, W., Bürki, H.U., Gourcy, L.,Ginot, P., and Huber, T., 2001, Stable isotopes in alpine ice cores: do they record climate variability? in Conference on isotope techniques in the study of past and current environmental changes in the hydrosphere and the atmosphere, IAEA Vienna, in press. Schwikowski, M., Barbante, C., Doring, T., Gaageler, H.W., Boutron, C., Schotterer, J., Tobler, L., VandeVelde, K., Ferrari, C., Cozzi, G., Rosman, K., and Cescon, P.,Post 17th century changes of European pb emission records in Alpine ice, manuscript in preparation. Schwikowski, M., Brütsch, S., Gäggeler, H.W., and Schotterer, U., 1999a, A high-resolution air chemistry record from an Alpine ice core: Fiescherhorn glacier, Swiss Alps, Journal of Geophysical Research Atmospheres, 104 (D11), 13709-13719. Schwikowski, M., Döscher, A., Gäggeler, H.W., and Schotterer, U., 1999b, Anthropogenic versus natural sources of atmospheric sulphate from an Alpine ice core, Tellus, 51B, 938951. Seibert, P., H. KrompKolb, A. Kasper, M. Kalina, H. Puxbaum, D.T. Jost, M. Schwikowski, and U. Baltensperger, 1998, Transport of polluted boundary layer air from the Po Valley to high-alpine sites, Atmospheric Environment, 32 (23), 3953-3965. Shotyk, W., Weiss, D., Appleby, P.G., Cheburkin, A.K., Frei, R., Gloor, M., Kramers, J.D., Reese, S., and Van der Knapp, W.O., 1998, History of atmospheric lead deposited since 12,370 14C yr BP from a peat bog, Jura Mountains, Switzerland, Science, 281, 1635-1640. Shumskii, P., 1964, Principles of Structural Glaciology. Translation D. Krauss, Dover Publications Inc. Dover, New York. Suter, S., M. Laternser, W. Haeberli, R. Frauenfelder, and M. Hoelzle, 2001, Cold firn and ice of high-altitude glaciers in the Alps: measurements and distribution modelling, Journal of Glaciology, 47 (156), 85-96. Van de Velde, K., Barbante, K., Cozzi, K.C., Moret, G., Bellomi, I., Ferrari, C., and Boutron, C., 2000, Changes in the occurrence of silver, gold, platinum, palladium and rhodium in Mont Blanc ice and snow since the 18th century, Atmospheric Environment, 34 (19), 3117-3127.

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Van de Velde, K., Ferrari, C. Barbante, C., Moret, I., Bellomi, T., Hong, S.M., and Boutron, C., 1999, A 200 year record of atmospheric cobalt, chromium, molybdenum, and antimony in high altitude alpine firn and ice, Environmental Science and Technology, 33 (20), 3495-3501. Vincent, C., Vallon, M., Pinglot, J.F., Funk, M., and Reynaud, L., 1997, Snow accumulation and ice flow at Dome du Gouter (4300m), Mont Blanc, French Alps, Journal of Glaciology, 43 (145), 513-521. Von Gunten, H.R., Rössler, E., and Gäggeler, H., 1982, Dating of ice cores from Vernagtferner (Austria) with fission products and lead-210, Zeitschrift fuer Gletscherkunde und Glazialgeologie, 18 (1), 37-45. Von Gunten, H.R., Sturm, M., Moser, R.N., 1997, 200-year record of metals in lake sediments and natural background concentrations, Environmental Science and Technology, 31, 2193-2197. Wagenbach, D., 1994, Results from the Colle Gnifetti ice-core programme, in ESF/EPL Workshop on Greenhouse gases, isotopes and trace elements in glaciers as climatic evidence of the Holocene, edited by W. Haeberli, Stauffer, B., VAW Arbeitsheft, pp. 19-22, Zürich. Wagenbach, D., 1997, High Alpine Air and Snow Chemistry, in Cloud Multi-phase Processes and High Alpine Air and Snow Chemistry, edited by S. Fuzzi and D. Wagenbach, pp. 163199, Springer, Berlin Heidelberg. Wagenbach, D., and Geis, K., 1989, The mineral dust record in a high altitude Alpine glacier (Colle Gnifetti, Swiss Alpsa), in Paleoclimatology and Paleometeorology: Modern and Past Patterns of Global Atmospheric Transport, edited by M. Leinen, and M. Sarnthein, pp. 543-564, Kluwer Academic Publishers, Dortrecht, Netherlands. Wagenbach, D., Muennich, K.O., Schotterer, U., and Oeschger, H., 1988, The anthropogenic impact on snow chemistry at Colle Gnifetti, Swiss Alps, Annals of Glaciology, 10, 183187. Wagenbach, D., Preunkert, S., Schäfer, J., Jung, W., and Tomadin, L., 1996, Northward transport of Saharan dust recorded in a deep Alpine ice core, in The Impact of Desert Dust across the Mediterranean, edited by S. Guerzoni, and R. Chester, pp. 291-300, Kluwer Academic Publishers, Dortrecht, Netherlands. Wagner, M.J.M., DittrichHannen, B., Synal, H.A., Suter, M., and Schotterer, U., 1996, Increase of I-129 in the environment, Nuclear Instruments & Methods in Physics Research Section B- Beam Interactions with Materials and Atoms, 113 (1-4), 490-494. Wagner, S., 1996, Dreidimensionale Modellierung zweier Gletscher und Deformationsanalyse von eisreichem Permafrost, Thesis, ETH Zürich, Zürich.

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GLACIER RESEARCH IN MAINLAND SCANDINAVIA

W. Brian Whalley

1.

INTRODUCTION

Glaciers in mainland Scandinavia cover a wide latitudinal range and are mostly in mountainous areas. Even in the mountains of the north, the glacier masses are generally not cold enough to provide ice cores which can be relied upon to give long term, unambiguous climatic records. Despite this ‘deficiency’, the glaciers of Scandinavia have provided valuable records of past glacial and climatic events and are an important source of data on mass balance. The time coverage of the mass balance record, over 50 years in some cases, is especially noteworthy and provides an important input to climatically-related scenarios for modelling. This review therefore summarises: 1) the basic disposition of ice masses in mainland Scandinavia; 2) the contributions to glacier mass balance records; and 3) the Little Ice Age (LIA) record of glacier change; 4) other investigations related to linkages between glaciers OR recent and possible future climatic events.

2.

PRESENT DAY GLACIERS – AN OVERVIEW

There are some 1900 glaciers in Scandinavia of which about 84 percent are located in Norway (Østrem et al., 1973; Østrem and Haakensen, 1993; Schytt, 1993). The total glaciated area is approximately 2900 km2. Glaciers 121 L. D. Cecil et al. (eds.), Earth Paleoenvironments: Records Preserved in Mid- and Low-Latitude Glaciers, 121-143. © 2004 Kluwer Academic Publishers. Printed in the Netherlands.

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in mainland Scandinavia are mainly on the western, maritime margin and have a southerly extent as far as Breifonn (59˚ 44’ N, 6˚54’ E); the farthest north glaciers are those of Seilandsjøkelen (ca 10 km2) at 70˚ 24’ N, 23 15 E. A report on the main glacier areas in north Norway is provided by Andreassen (2000) and of overall Norwegian glaciological investigations by Kjøllmoen, (2001). Although there are many small glaciers (greater than 1 km2) there are several ice caps, ten being greater than 30 km2 in surface area. The later include Jostedalsbreen, the largest in mainland Europe (Table 1). The highest summit in Norway is Galdøpiggen (2468m) in the Jotunheimen, an area of extensive mountain glaciation and Kebnekaise (2110m) in the north of Sweden (Figure 1). None of the ten largest ice masses are in Sweden, the largest wholly in the country is Stourrajekna at 11.8 km2, although Salajekna (Sulitjelmaisen; 24.5 km2) is on the border of Norway and Sweden. A major project in mainland Scandinavia in the 1960s and 1970s was the preparation of a glacier atlas. Data on all the glaciers are available in the two volumes of the Scandinavia Glacier Atlas, North (Østrem et al., 1973) and South (Østrem et al., 1988). Although using mainly pre-1960s aerial photography, and thus a little out of date, this major undertaking still provides the basis for assessing the areal extent of glaciers in Norway and Sweden as well as other attributes of the glaciers such as altitudinal range, aspect, and morphology. The volumes contain bibliographic data as well as the inventory, methods of compilation and general maps of meteorological conditions, glaciation (glacierisation) altitudes, etc. Although there have been some considerable changes since these surveys were made, they remain important compilations. The digital data are held in World Data Centre archives.

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Figure -1. Scandinavia with the main monitored glaciers. (Modified from Kjollmoen, 2000)

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Table -1. Sizes, altitudes, and locations of the ten largest ice masses in Scandinavia. These are data directly from the 1988 inventory and in some cases may be now somewhat smaller. Max elevation Lowest elevation Name Area (km2) (masl) (masl) Jostedalbreen 486 2000 350 Svartisen W 221 1580 20 Søndre Folgefonni 168 1660 490 Svartisen E 148 1550 208 Blåmannsisen 87 1560 810 Hardangerjøkulen 73 1850 1050 Snønipbreen 50 1830 890 Okstindbreen 46 1740 750 Øksfjordjøkelen 41 1170 330 Harbardsbreen 36 1950 1250

This latitudinal range in Scandinavia (over 1200 km) also means that there is considerable difference in the meteorological and climatic conditions, both north to south and west to east. A distinct gradient exists from the maritime west to the more continental east, reflected in the equilibrium line altitude (ELA) (Chorlton and Lister, 1971; Liestøl, 1967). This continentality is also reflected in the specific ablation with a gradient going from (at about the same latitude) Ålfotbreen in the west (5° 40’ E, 40 km from the west coast) to Gråsubreen in the east (8° 40’ E), some 150 km. These relationships were first explored by Liestøl (1967) who showed the effects of ocean distance on the glacier budget gradient as well as the equilibrium line and glacierisation level. The work of Liestøl provided the first significant impetus in examining the relationships between mass balance and climate. The concept of ‘glaciation level’ was investigated by Østrem (1972) in papers relating to the formation of ice-cored moraines but of wider significance with respect to continentality and the mapping of past glacier limits. Maps of the transient snowline and glaciation level are included in the Scandinavia glacier atlas.

3.

EARLY WORK ON THE GLACIERS IN SCANDINAVIA AND HISTORIC VARIATIONS

Some early scientific travelers and geologists have provided useful information about the positions of certain glaciers at the times of their visit. Comparisons have been made with extant glaciers to help in the reconstruction of LIA (approximately A.D. 1600 – 1900) glacier limits. These are discussed in a subsequent section of this chapter. Of particular

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importance were Rekstad and Oyen (1908) in Norway and, in Sweden, Hamberg (1910) and Svenonius (1910) in a major contribution on the 'Glaciers of Sweden in 1908' and by Wahlenberg and Westman (Holmlund and Schneider, 1997). Their photographs and markers near snouts have been of use in plotting the retreat of valley glaciers. For example, even before Rekstad’s visit to Nigardsbreen, James Forbes visited the glacier and both his paintings (Forbes, 1853) and Rekstad’s photographs have been useful in re-constructing the recent glacial history of this important outlet glacier of the Jostedalsbreen ice cap. An overview of the historic variations of glaciers in Scandinavia has been provided by Bogen et al (1989). This includes, for example, the variation of 7 glacier outlets of the ice caps of Jostedalsbreen and Folgefonni from 1904 to 1986. The common periods of advance and retreat and the association with mean summer temperatures is clear. More explicit links between glacier variations and climate are discussed below in the light of modeling and the use of mass-balance data. A new perspective on glacier research could be said to have started in Norway with the expeditions from the University of Cambridge under the leadership of W. Vaughn Lewis (Nye, 1997). The expeditions to Austerdalsbreen were primarily developed to investigate a variety of phenomena (including the annual layering of icefalls and the formation of ogives, velocities at depth, compressive stresses near the snout of glaciers, etc.) rather than to investigate the mass balance of Norwegian glaciers specifically. However, contributors to the program included J. F. Nye and J.W. Glen whose work on glacier flow and a variety of aspects of theoretical glaciology was to become seminal. In the present context, the work of Nye (1965) on modelling the budget history of a glacier from advance and retreat data has provided the starting point for subsequent modeling of glacier behaviour over time.

4.

MASS BALANCE MEASUREMENTS AND GLACIER MAPPING

Considerable credit must go to the pioneers of mass balance measurements in Scandinavia; Liestøl in Norway and Ahlmann and Schytt in Sweden. The late Prof. Olav Liestøl started mass balance measurements on the southern Norwegian glacier, Storbreen (Jotunheimen) in 1949 and this work (Liestøl, 1967), together with that on Storglaciären in northern Sweden (Schytt, 1959), pioneered much of what we know about mass balance measurements in general. Since then, the Hydrological Branch of the Norwegian Water Resources and Energy Directorate (NVE) has maintained

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a network of mass balance measurements, primarily concerned with hydroelectric power generation in southern Norway but also extending to other areas such as Engabreen (Svartisen). Investigations more concerned with climate change effects on glaciers have been underway more recently with work on Langfjordjøkelen in Troms (Kjøllmoen et al., 2000). The current mass balance coverage of twelve glaciers is shown in Figure 1. In Sweden there are fewer glacierized basins and less need for hydropower related observations but, as in the first observations in Norway, the mass balance investigations were carried out for purely scientific reasons, e.g. Wallén's work on the Kârsa glacier (Wallén, 1959). The Swedish observations were initiated by H. W Ahlmann in the 1930s (Ahlmann, 1948). Subsequently, the late Valter Schytt, under the direction of Ahlmann, started observations at Storglaciären, a relatively accessible glacier in northern (67°55’N) Sweden (Schytt, 1959). The mass balance observations at Storglaciären (Holmlund et al., 1996a) are significant in the development of glacier observations, not least as its mass balance has been measured annually since 1945. The mass balance methods used today as international ‘standards’ were, to a great extent, developed at Tarfala together with major contributions from the Norwegian experience under Gunnar Østrem and Randi Pytte Asvall (Østrem and Brugman, 1991). Although there have only been rather occasional and diverse observations of glaciers in North Norway (on and north of the Arctic Circle) a recent report edited by Andreassen (2000) has identified areal and some volumetric changes associated with a number of glacier systems. Of these, Engabreen an outlet of the Svartisen ice cap, is perhaps the best known with annual observations of mass balance since 1970. Other work on the mapping of Svartisen has been carried out by parties from the University of Manchester (Theakstone and Knudsen, 1984); Theakstone, 1990). Also straddling the arctic circle, and studied by parties from the United Kingdom and the University of Aarhus (Denmark), is the small ice cap and outlet glacier systems of Okstindan (Theakstone et al., 2000; Theakstone and Jacobsen, 1997; Knudsen and Theakstone, 1997). In Norway, as well as the mass balance measurements by conventional techniques, there has also been great interest in the use of glacier maps to calculate volumetric and areal changes. Østrem and Tvede (1986) present an early case for this and show that for Folgefonni between 1959 and 1981, there was a general decrease in volume. However, the greatest loss was on the western side of the ice cap and the summit dome had moved towards the eastern side where there was even a small increase in thickness. The presumption is that this shows a change in the wind pattern that brings the snow to the glacier. An analysis of wind directions suggested an increase in

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the westerly component from 29 percent in 1951-1959 to 43 percent in 19601979. Østrem and Tvede (1986) point out that repeated mapping of glacier systems would be helpful to show other such changes and Østrem and Haakensen (1999) have also presented a comparative case for conventional measurements and maps. The use of digital photogrammetry and its incorporation into digital elevation models provides a further stimulus for the volumetric recording of glacier mass changes (Andreassen, 1999). Considerable efforts are now being made to utilise this technology to provide accurate assessments of volumetric change and also to relate digital terrain models to conventional methods of glacier mass balance measurements. In Sweden there has also been some interest in the use of glacier mapping, in particular at Storglaciären (Holmlund, 1996). Several glaciers are currently being observed photogrammetrically, every 15 years: Storglaciären, Mårma, Stour Raaita and Pårteglaciären. With the use of ice radar, especially with adjunct GPS, there is increasing use of mapping volumetric changes over time. Studies of Mikkaglaciaren (Swedish Lapland, Holmlund, 1986) and Austre Okstindbreen (Theakstone and Jacobsen, 1997) are especially noteworthy. A summary of investigations in Norway can be found (in English) in (Kjøllmoen, 2001) and for Sweden in Holmlund et al. (1996a) and Holmlund and Jansson (1999).

5.

LITTLE ICE AGE GLACIER EXTENTS

The presence of LIA glaciers is almost everywhere quite near (less than 1 km) the present glacier limits. For example, Gellatly et al. (1989) used the sketches of James Forbes at Jøkulfjord to help reconstruct the limits of the plateau ice of Øksfjordsjøkelen (Troms - Finnmark). Similarly, the reports and photographs taken by climbers about the turn of the 19th Century have provided evidence of some glacier limits and phenomena, e.g. the thickness of the arm of Strupbreen (Troms) and the drainage of the ice dammed lake (Whalley, 1973). Much attention has been paid to the reconstruction of LIA limits usually by lichenometric methods such as used by Ballantyne in Lyngen in north Norway (Ballantyne, 1990), on Nigardsbeen, Jostedalsbreen (Bickerton and Matthews, 1993; Andersen and Sollid, 1971) and Storbreen (Jotunheimen) and around Okstindan (Griffey, 1977). Of special note in Norway is the work undertaken by the late Jean Grove, first by amalgamating a wealth of data on the LIA in general (Grove, 1988) but also with respect to the advances of glaciers in western Norway associated with the LIA. The glacier fluctuation and environmental data here were reconstructed with the help of tax records (Grove and Battagel, 1983) and provide a unique glimpse of glacier advances in an area relatively

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sparse in observations and at the same time, showing the human consequences of glacier advances, environmental deterioration, and geomorphic activity associated with the LIA (Grove, 1972). It has been supposed that the ’date’ of the LIA maximum in Scandinavia was about A.D. 1750 from a moraine date at Engabreen. Recent work by Winkler (2002) at Svattisen and Okstindan confirms this mid 18th Century maximum. Winkler considers that synchronicity between Svartisen and Southern Norway is due to increased SE-NW trend of moist airflow in winter parallel to the coast of Norway as the main forcing factor. However, the large variation on precipitation may have given even 1950 as the maximum in south Folgefonna, southern Norway. Evidence suggests that north Scandinavia had large differences in Neoglacial variations (Matthews, 1984) although Winkler (Winkler, 2001) views the north to south synchronicity covering glaciers as far north as Lyngen (70° N). Lack of good 'absolute' dating in north Norway still prevents a means of correlating north-south events but the investigations of Karlén (1973) show a wide variation in LIA maxima in the Kebnekaise area of north Sweden although advances between A.D. 1500 and A.D. 1640 were the most extensive and common. There is still no comprehensive overview of LIA limits and chronology for Scandinavia with the exception of that provided by Winkler (1996).

6.

THERMAL REGIMES AND CORING PROJECTS

The glaciers of mainland Scandinavia are predominantly ‘warm based’. This is mainly due to the fact that the glaciers respond to the maritime conditions, especially in the south of the land mass. However, there is good evidence that a few glaciers are cold based and/or polythermal. Some high plateaus (above ca 1300 masl) have glaciers which have been shown to be frozen to their beds (Whalley et al., 1996). Holmlund et al. (1996b) have suggested that ice radar profiling can indicate the presence of ‘cold’ ice according to dielectric properties of the snow/firn/ice. The two glaciers they investigated (Storglaciären and Mårmaglaciären) are both polythermal and thus capable of providing at least some climatic signatures. Holmlund (1998) reports a 27.5 m core containing 52 sedimentary layers from Mårmaglaciären with the oldest ice from the LIA. The ice had low ionic content, comparable to the other non-High Arctic cores. Some tephra (volcanic depris) was also found although it is not clear if this is from the A.D. 1783 Laki eruption or the A.D. 1815 Tambora eruption. Unfortunately, despite this knowledge, there has, to date, been no full coring through any ice masses which are fully ‘cold-based’. A coring

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project was undertaken by Japanese scientists on Jostedalsbreen in 1987 with NVE and the University of Bergen (Kawamura et al., 1989) as part of a survey of maritime glaciers in the Arctic Ice Coring Project. A core of 46.9m to bedrock was extracted from the summit dome (1960 m) in the accumulation area of Nigardsbreen. The temperature profile was, apart from near the surface, at 0˚C throughout the depth of the hole and water infiltrated the hole at the bed suggesting an aquifer confined by the bedrock. Although samples for oxygen isotope analysis were taken, the results have not been reported. However, He et al. (2001) have shown that there may be a source of climatic information from stable isotope data in their study of snowpit chemistry from Austre Okstindbreen. Some limited isotopic measurements of ice have been taken in North Norway (Gordon et al., 1988) but these are not from cores. They do however confirm the thermal regime of Balgesvarri (>1650 masl., Troms), one of the highest ice plateaus in north Norway. Despite the fact that summer melt does occur and that an isotopic signal may be “noisy”, there are suggestions that useful signatures from ice cores may be obtained under such conditions (Koerner, 1997). Coring from these high plateaus, where the ice thickness is for example about 80 m thick on Balgesvarri, may be useful in providing extra evidence for the recent movement of the North Atlantic Oscillation (NAO) and, in particular, the position of the Polar Front. Storglaciären (N Sweden) is polythermal, (Jansson, 1996) some 85 percent being temperate but the surface layer of the ablation area is cold (Hooke et al., 1983; Holmlund and Eriksson, 1989). Typically for polythernal glaciers, the cold portion is thickest at the terminus and towards the margins (maximum about 60 m) and decreases towards the center of the glacier the approximate position of the equilibrium line. The surface layer in the accumulation area is only cold seasonally but it is probably typical for valley glaciers in this area to the east of the mountain divide. However, It is probable that the small plateau cap on the summit of Kebnekaise is also below the pressure melting point and might provide a suitable place for coring (Holmlund, 1998).

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VARIATIONS IN MASS BALANCE AND CONTINENTALITY

Figure -2. Mass balance for selected glaciers in south Norway, from west (left to east (right). (From Kjollmoen, 2000)

Figure -3. Cumulative net balance for selected glaciers in south Norway, 1963-2000. (From Kjollmoen, 2000)

The length and detail of the mass balance measurement record in Scandinavia, the variation from maritime to continental glaciers, as well as the considerable wealth of data on historic and LIA climatic events and the availability of meteorological data make it an excellent location for testing or calibrating a variety of climate-related models. Figures 2 and 3 show

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some typical data for some Norwegian glaciers for a specific year and for 1963-2000. Nigardsbreen and Storglaciaren are two glaciers that have been intensively studied and thus merit some attention. The Kebnekajse area of Lapland plays an important part in Swedish contributions to Scandinavian glacier studies. In particular, the Kårsa glacier was the subject of studies by Ahlmann and its response to meteorological factors was studies by Wallén, see (Schytt, 1959; Wallén, 1959) for summary. Nigardsbreen, an eastern outlet glacier of Jostedalsbreen has been the subject of long-term observations (e.g. Østrem et al., 1976) so has been a fruitful way in which observations can be related to glaciological theory. Østrem et al. (1976) provide a table of the front variations of Nigardsbreen since A.D. 1710 as well as a summary of the observations to that time but which include estimates of volume change (Table 2) and repeat photographs by Olav Liestøl. Table -2. Changes in volume of Nigardsbreen (Jotunheimen) over time (Ostrem, et al 1976) Calculated for areas Estimates or the Year below 1000 masl entire glacier Total Per year Total 1748 - 1937 - 750 x 106 m3 - 4.0 x 106 m3 1748 - 1937 - 230 x 106 m3 - 16.4 x 106 m3 1937 - 1951 1937 - 1951 6 3 1951 - 1964 1951 - 1964 - 140 x 10 m - 10.8 x 106 m3 1964 - 1974 1964 - 1974 - 60 x 106 m3 6.0 x 106 m3 5.2 x 106 m3 1748 - 1974 1748 - 1974 - 1180 x 106 m3

Oerlemans (1992) used an energy balance model with the NVE glacier data base to examine the climate sensitivity of three glaciers in southern Norway (Nigardsbreen; Hellstugubreen, a continental glacier; and Ålfotbreen, which is very near the Norwegian coast). Two of Oerlemans’ general findings are important; 1) that maritime glaciers are more sensitive than continental glaciers, and 2) that the different hypsometry of maritime glaciers means that the ablation zone extends further down into a warmer climatic environment. Of the two effects, the former is more significant. Later work by Oerlemans (1997) used the available historic evidence for the snout positions of Nigardsbreen since A.D. 1748 to help calibrate a flowline model. This is based on the Nye (1965) theory relating mass balance to glacier length but uses numerical integration of the non-linear equation which describes the mass balance history (Oerlemans, 1986). The data for the topography were already available from a map by Østrem (1988) and the historical data from diverse sources including a mass balance record by NVE. The adjustment of flow-law parameter and basal topography discussed in the paper provide, together with the historical data, a good

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prediction of the behaviour of the glacier. Amongst the predictions was that the LIA mean temperature for the area should have been between 0.53 and 0.27 ˚C lower than for the reference period of A.D. 1900-1993. Four scenarios for climate warming rates were considered. As an example, 0.02 K a-1 and no increase in precipitation, Nigardsbreen would advance slightly until A.D. 2020 but that thereafter the ice volume remaining in A.D. 2100 would be only 10 percent of that in 1950. Oerlemans relates this back to the finding suggested by Nesje and Dahl (1993) that glaciers were absent from western Norway in the Hypsithermal (c. 8 - 6000 BP). Oerlemans points out that this result is due to the characteristic topography of high flat ice caps (which supply ice to outlets such as Nigardsbreen) where the surface elevations are only a few hundred metres above the present day equilibrium line. However, there is only a thermal configuration in this model and assumptions about the precipitation have to be made. Very little is known about the changes in precipitation, not least in direction of the storms providing winter accumulation. Evidence, e.g. Laumann and Reeh (1993) and Østrem and Tvede (1986) mentioned above, suggests that these have been changeable even in the last 100 years. Laumann and Reeh (1993) used a degree-day model to estimate the mass-balance-elevation relationship for three glaciers on a west-east transect (Ålfotbreen, Nigardsbreen, Hellstugubreen). Thirty-year normals of meteorological data were used in addition to local temperature and precipitation data. They concluded that if the temperature increases 2˚C then all three glaciers would lose mass unless there was an accompanying increase in precipitation. Ålfotbreen, being relatively low would be highly sensitive to an increase in temperature and for a 3˚C rise the equilibrium line altitude would be higher than the glacier itself unless there was a simultaneous 30-40 percent increase in precipitation. Hellstugubreen, the highest, would lose the least mass because precipitation would fall as snow rather than rain. A rise in the equilibrium line would have less effect for this glacier than for a similar rise on the Jostedalsbreen plateau which contributes to Nigardsbreen. The study clearly shows the coupled effects of temperature, precipitation and altitude/location/continentality. A later study (Tvede and Laumann, 1997) showed the significance of local topography as snow drift and storm direction appeared to play an important part in accounting for the behaviour in some glaciers in southern Norway. This ties in with the findings of Østrem and Tvede (1986); the glaciers in the latter study (Folgefonni) are only some 20 km north of those studied by Tvede and Laumann. The retreat of some major Swedish glaciers can be seen in Table 3. Of these, Storglaciären (67°55’N, 18°35’E) is a small glacier located on the eastern side of the Kebnekaise massif, northern Sweden (Jansson, 1996) and

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is the original site of Swedish mass balance measurements as mentioned previously. The glacier is 3.2 km long from its head at about 1730 m to the terminus at about 1120 m and has a total surface area of 3 km2. The glacier has an average thickness of 95 m, with a maximum depth of 250 m in the upper part of the ablation area. Three recent studies relating glaciers to atmospheric circulation and continentality in Swedish Lapland present complementary data from further north than the Norwegian investigations just mentioned. Pohjola and Rogers (1997) compared the 50-year balance record from Storglaciåren and analysed extreme events in the data and reconstructed the mean synoptic meteorology for those years. They explained the high correlation between air temperatures and summer ablation for the glacier by the sea level pressure conditions with a warm core ridge developed over southern Greenland to the Gulf of Bothnia and high pressure over the Barents Sea giving strong ablation years and maximum temperature anomalies. Summers with low ablation were given by a high pressure ridge over the British Isles to southern Norway but no high pressure over the Barents Sea. Further, they suggest that glacier growth is favoured by periods of enhanced maritime air flow in both summer and winter with the opposite giving net mass loss on glaciers because of drier winters and warmer summers. These close linkages have yet to be explored for other glaciers in north Scandinavia but they suggest that the area has experienced an increase in maritime influence with stronger and more persistent westerly flow in the last two decades. Table -3. Retreat of Swedish Glaciers 1965-1985. (from Bogen et al., 1989) Glacier Area (km2) Length (km) Retreat (m) Salajekna 24.5 10.0 209 Mikkajekna 7.6 4.6 325 Ruotesjekna 5.4 4.6 544 Rabots glaciår 4.2 4.1 225 Riukojietna 5.5 3.3 224 Suottasjekna 8.1 4.4 174 Vartasjekna 3.6 3.0 72 Storglaciåren 3.1 3.7 98 Ruopsojekna 3.6 3.9 137 Ô. Påssusjietna 1.8 1.9 117 U. Rietaglaciåren 2.0 2.1 80 Stuor Reitglåciaren 2.0 2.6 147 Hyllglaciåren 1.5 2.2 80 Isfallsglaciåren 1.4 2.1 163 A. O. Kasaglacl. 0.6 1.4 150 Kårsojetna 1.6 2.2 131

Remarks To 1984 To 1984 From 1963 To 1984 To 1984

From 1968

To 1984

From 1968

A study by Holmlund and Schneider (1997) showed a very good correlation between Accumulation Area Ratio (AAR) and the net balance of

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Storglaciåren. They then used the AAR as a proxy for investigating the relationship of net balance with a continentality index for four glaciers in northern Scandinavia (Engabreen, Salajekna, Pårtejekna and Storglaciåren). The ‘continentality index’ (CI) was defined as the difference between the coldest winter month and that of the warmest month the following summer. The CI gradient was calculated at 2.8˚ C per 100 km and rises from west to east. Although the authors claim only a preliminary analysis for northern Norway it does reflect the findings of Pohjola and Rogers (1997). In particular, the climate of the area is mainly controlled by low pressure systems from the west giving a maritime climate along the west of Norway. In some respects, this is hardly surprising. What remains unknown is how changes in the Polar Front position from year to year and over the long term, as also with the location of the NAO, affect mass balance and continentality. Some progress is being made on this by the use of modeling with glacier mass-balance data. Brugger (1997) used a time-dependent model of glacier flow to predict the response of Storglaciåren to various warming scenarios using both a linear temperature forcing and a temperature-precipitation forcing. Brugger suggests that, for Storglaciåren, long-term observations of behaviour may be required to detect climate warming. This would have to be shown above natural variations in mass balance from year to year, previous smallmagnitude high frequency signals, and a temperature response that might be exponential rather than linear. It remains to be seen if a similar analysis produces the same results for more maritime glaciers at about this latitude. A recent paper by Schneeberger et al. (2001) uses the mass balance data for predictive modeling of atmospheric CO2 content. Nesje et al. (2000), in a study (see below) on the linkages between the NAO and mass balance reported that for, the period 1963-1998, the maritime glaciers studied increased their mass balance and the continental glaciers showed decreases in mass. These glaciers are all in southern Norway except for Storglaciaren in N Sweden (Figure 1). Nesje et al. showed that the net balance for the maritime glaciers was more influenced by the winter balance than the summer balance but the opposite was the case for the more continental glaciers. Correlations have yet to be carried out for the other glaciers observed by NVE in Norway. The recent study of glaciers in north Norway (from just south of the Arctic circle) (Andreassen, 2000) shows considerable variation in behaviour of the glaciers examined (some from direct mapping, some for aerial photographic observations). In the Svartisen area for example, all five of the glaciers studies have shown marked retreats in the 20th century but since about 1960 there have been notable changes; Engabreen has advanced from a positive balance in 1970 onwards to its 1950 snout position. Two smaller

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glaciers in the area (Svartisheibreen and Dimdalsbreen) have also gained mass and their retreat seems to have stopped and the retreat of Trollbergdalsbreen appears to have slowed. Engeset et al. (Engeset et al., 2000), using a (precipitation) degree day model for the mass balance of Svartisen glaciers showed that the glacierized area (100 km2) draining to the Storglomvatn hydropower reservoirs is currently in approximate equilibrium with the present climate and that the period of recession until the middle 20th Century was predicted by this approach. Application of future climate scenarios for the region indicated a net loss of 5 x 109 m3 of water until A.D. 2050 as an upper limit, but close to zero by another scenario. Braithwaite and Zhang (1999) advocate use of degree-day modeling in general as it can be used with relatively simple meteorological/climatic data sets rather than the relative complexity of energy balance models. They used a general set of climate data, including that from Storglaciåren and some Norwegian glaciers to investigate the mass balance sensitivity to temperature change for some sub-polar to maritime glaciers. It would be interesting, given the success of degree-day modeling, to investigate this sensitivity in more detail for the full range of Scandinavian glaciers for which data could be assembled. Further to the northeast of Svartisen, but only about 100 km further north, at Blåmannsisen, aerial photograph interpretation showed that western outlets had been advancing or stable between 1961 and 1998 whereas the northern outlets have been retreating (Andreasson, 2000). The eastern outlets had also retreated since 1961 but mainly before 1985. This asymmetry of response is perhaps not surprising given the behaviour of Folgefonni mentioned above and the changes in weather patterns of recent years. The mass balance measurements on Folgefonni will be started again in 2003 and at Blåmannsisen it is hoped that joint Swedish and Norwegian observations will examine the behaviour of the whole of the ice cap in an integrated manner with both conventional measurements and a variety of remote sensing techniques coupled with ground penetrating radar of the snowpack to compare with traditional mass-balance techniques. Ice sheet reconstructions related to present day ice limits and implications of Scandinavian glaciers to the North Atlantic area The whole of the Scandinavian ice sheet has decayed from a once much larger body; the last glacial maximum is usually given as around 20 ka BP. For the most part, these reconstructions have been used in modelling exercises (e.g. Holmlund and Fastook, 1993; Fastook and Holmlund, 1994) and specifically with reference to the outermost limits of ice in southern Sweden. There is insufficient space here to go into any of the detail produced by glacio-geomorphological mapping in the last 100 or so years. However, of particular interest in the present overview are the few, but

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increasing in number, observations and interpretations related to the offshore conditions and climate-oceanic heat sources. Two types of investigation are evident; 1) those relating glacier change (mass balance, length, volume) to atmospheric/meteorological conditions from direct records, usually of the last 100 years, or 50 years for mass balance records, and 2) those using proxies such as sediment cores from the neighbouring oceans. The eastward location of mainland Norway to the position of the NAO is clearly a matter of considerable importance and is likely to continue to be so. Nesje et al. (2000) have addressed this issue. They used a NAO Index (difference or normalised sea-level pressures between Ponta Delgarda, Azores and Stykkisholmur, Iceland, 1865-1984) to compare NAO variations with the mass balance of 10 glaciers in south Norway. These glaciers included Ålfotbreen and Folgefonna in the west and Hellstugubreen and Gråsubreen in the east. All were south of latitude 63˚N. They found a correlation between the NAO index and the mass balance properties of the glaciers but that the effect of the NAO on the winter and net mass balances gradually decreases with increasing continentality, as might be expected. They also examined some other proxy records of climate change and suggested that Holocene glacier variations and ELA variations of maritime glaciers in western (i.e. maritime) Scandinavia reflects past NAO variations. Olsen et al. (2001) have recently provided new data and an overview of sediment ages and stratigraphic information from some 200 sites in Norway. They suggest, perhaps controversially, that there have been rapid shifts between glacial and interstadial conditions in semi-cycles of 5-7 ka between 40-45 and 10 ka BP. A new (‘yo-yo’) model is used to describe these rapid fluctuations. They suggest that, in addition to precipitation, the mountain/valley/fjord topography, glacial isostacy and relative sea level change were probably more important in the size of these fluctuations than air temperature changes. It is yet to be explored how these findings fit with those of both modelling experiments constrained by ice limits, especially on the east of the Scandinavian ice mass and the findings of e.g. Koç and Jansen (1994) and Koç et al. (1993) regarding polar front position and orbital forcing factors. A recent paper using high resolution tree-ring, lake core and marine records from several circum-North Atlantic sites in the early Holocene (Björck et al., 2001) suggests that the climate system may be more sensitive to solar-related forcing than was previously thought and that this may be an important mechanism for inducing sub-Milankovich scale variability. Such studies have not yet been related to glacier behaviour in Scandinavia but it will certainly be of great interest. The current interest in linking atmospheric and oceanic heat sources and sinks has obvious implications for both the behaviour of glaciers and the

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way in which glaciers can be used as a source of information for both past and future climatic effects. Some of these aspects have been mentioned previously with respect, for example, to continentality in Scandinavia. There are however still wider linkages, most importantly between the position of the Polar front over time, the location of the North Atlantic Deep Water – NAO effects as well as the possible effects of orbital forcing components of climate change. The locations of the glaciers in mainland Scandinavia clearly have a significant role to play in elucidating these linkages and providing long-term data. Not only because of the number of mass-balance measurements taken (Norway supplies about a quarter of the world’s total) but also because of its location vis-à-vis the North Atlantic and of Greenland and Svalbard with their longer-term ice core data. Ice cap modelling for Scandinavia, usually for the LGM or Younger Dryas, will not be reviewed here as this is generally related to mapping the extent of ice limits (e.g. Holmlund and Fastook, 1993; Fastook and Holmlund, 1994). It should be noted however that there are clearly close links to be made from the present day conditions through the atmospheric/oceanic forcing factors to these large scale models. In general terms, the location of Scandinavia has long been viewed in relation to the Polar Front and its variation over time. However, very little direct evidence has been obtained regarding the glaciers because of lack of ice-core data. There is however, good evidence from the Nordic seas using δ18O in dated sediment cores (Koç and Jansen, 1994) that there is a close link between the fluctuations of the polar front and sea-surface temperatures, the Nordic seas being very sensitive to insolation forcing. They suggest that the sea area is currently trending towards glacial conditions and has been since about 7,000 years ago. It is not known what effects this might have for the glaciers but it might be supposed that these would be a difference between those in the north and the south, especially of the maritime glaciers of Norway. Following the work of Rye et al. (1987), investigations by Nesje and Kvamme (1991) indeed suggest that the Jostedalsbreen ice cap disappeared during the warm Hypsithermal (8000-6000 BP). They calculated that the glaciation threshold was some 260 m (±100 m) above the present mountain plateau with area mean summer temperatures being some 2.5 - 3.0 ˚C warmer than present. Post the Hypsithermal, glacierisation started to be reactivated and the outlet glaciers reached their maximum Neoglacial extent in the LIA of the mid 18th Century. Work investigating the plateau glaciers in north Norway has yet to be undertaken although there are possibilities that, although outlet glaciers retreated rapidly at various times, the ice caps themselves need not have completely disappeared because there was little or no sediment production that reached the valley sediment sinks (Whalley et al., 1996; Rea et al., 1999).

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SUBGLACIAL OBSERVATIONS

Some mention should be made about the subglacial observations made in Norway over recent years as this relates to the overall significance of glaciers in our understanding of long-term climate change as well as glacial processes. The NVE has, in conjunction with the Statkraft, the State Power Company, developed a subglacial laboratory under approximately 200 m of ice at Engabreen (Svartisen, Nordland). This laboratory has been used in work on subglacial hydrology (Jackson, 2000; Lappegard G. et al., 1998) Kohler and others., 1998) and ice flow (Cohen et al., 1997) at the bed and is currently in use looking at subglacial deformation. Prior to the opening of this facility, some subglacial hydrology had also been undertaken (again with hydropower development) at Bondhusbreen (Hagen et al., 1983). The ice under Engabreen and Bondhusbreen is temperate and the till found has been produced by subglacial erosion on a bedrock base. However, evidence from the northern plateaus (e.g. in Lyngen on plateaus above about 1300m a.s.l.) shows the ice to be frozen to the bed. Both there and at lower altitudes where there are glaciers retreating off substantial plateaus (including Øksfjordsjøkelen at 900 m plateau, the ninth largest ice mass) substantial areas of blockfield exist and are also emerging as the glacier cover decreases. Studies of the material suggest chemical weathering of the bedrock to give a clay mineral suite that is indicative of warmer conditions than at present. The supposition is that these blockfields pre-date the Quaternary ice masses although somewhat different conclusions have been reached for similar work by Nesje in the south of Norway (Nesje et al., 1988). These observations related to past climatic and environmental conditions still have to be incorporated within our general understanding of the decay of the Scandinavian Ice Sheet and to the relatively small relicts of that ice mass visible today.

9.

CONCLUSIONS

The mass balance and associated work on Scandinavian glaciers is probably as extensive as anywhere in the world. The utility of these glacioclimatic records plus glacial chronological investigations allows not only comparisons from different climatic conditions in Scandinavia but testing of models of behaviour. Such models may relate to the behaviour of the glacier flow itself as perturbed by mass balance changes or they may be ways of modelling the mass balance itself in terms of meteorological/climatic parameters. All such modelling relies upon good field data and the papers reviewed briefly above show the importance of the longevity as well as

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quality of the Scandinavian data sets. With new techniques being added to the glaciological armoury and the ability to use them on an increasingly diverse range of glaciers, both spatially and with respect to size, the Scandinavian glaciers will remain a most important means of elucidating both past and future climate change. Furthermore, it is now becoming evident that the wide range of various ‘proxy’ data from both terrestrial and oceanic sources is also becoming better integrated with the mass balance data and the modelling of climate change. The glaciers of Scandinavia, by their nature, do not provide good ice-core records which go well back in time but the quality data which the glaciers do preserve in moraines, historical records as well as mass balance and snout retreat data for the last 150 years will continue to be important in elucidating climate change and climate variability.

10.

ACKNOWLEDGEMENTS

I wish to thank Erik Roland and the staff of Brekontoret, NVE for their assistance and hospitality during a stay at NVE while I held a Leverhulme Research Award and I thank the Leverhulme Trust for this. I also thank Per Holmlund, Atle Nesje and Wilfred Theakstone for their comments on a draft of this review.

11.

REFERENCES

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Holmlund, P. and Eriksson, M., 1989, The cold surface layer on Storglaciären. Geografiska Annaler, 71A, 241-244. Holmlund, P. and Jansson, P., 1999, The Tarfala mass balance programme. Geografiska Annaler, 81A, 621-631. Holmlund, P., Karlén, W. and Grudd, H., 1996a, Fifty years of mass balance and glacier front observations at the Tarfala Research Station. Geografiska Annaler, 78A, 105-114. Holmlund, P., Näslund, J.-O. and Richardson, C., 1996b, Radar surveys on Scandinavian glaciers, in search of useful climate archives. Geografiska Annaler, 78A, 147-154. Holmlund, P. and Schneider, T., 1997, The effect of continentality on glacier response and mass balance. Annals of Glaciology, 24, 272-276. Jackson, M., 2000, Svartisen subglacial laboratory. Norges vassdrags- og energidirektorat (NVE), Glasiohydrologiske undersøkelser, Oslo. Jansson, P., 1996, Dynamics and hydrology of a small polythermal valley glacier. Geografiska Annaler, 78A, 171-180. Karlén, W., 1973, Holocene glacier and climatic variations, Kebnekaise Mountains, Swedish Lapland. Geografiska Annaler, 55A, 29-63. Kawamura, T., Fujii, Y., Satow, K., Kamiyama, K., Izumi, K., Kameda, T., Watanabe, O., Kawaguchi, S., Wold, B. and Gjessing, Y., 1989, Glaciological characteristics of cores drilled on Jostedalsbreen, southern Norway. In: Proceedings National Institute for Polar Research Symposium on Polar Meteorology and Glaciology, Vol. 2, pp. 152-160. NIPR, Tokyo. Kjøllmoen, B., 2001, Glaciological investigations in Norway in 2000. Norwegian Water Resources and Energy Directorate (NVE), Oslo. Kjøllmoen, B., Olsen, H. C. and Svaerd, R, 2000, Langfjordjøkelen i Vest-Finnmark. Norges vassdrags- og energidirektorat (NVE), Glasiohydrologiske undersøkelser, Oslo. Knudsen, N. T. and Theakstone, W. H., 1997, Recent changes of the glaciers of Svartisen and Okstindan, Norway. Aahus Geoscience, 7, 113-128. Koç, K. N., Jansen, E. and Haflidason, H., 1993, Paleoceanographic reconstruction of surface ocean conditions in the Greenland, Icelandic and Norwegian Seas through the last 14ka based on diatoms. Quaternary Science Reviews, 12, 115-140. Koç, N. and Jansen, E., 1994, Response of the high latitude northern Hemisphere to orbital climate forcing: Evidence from the Nordic Seas. Geology, 22, 523-526. Koerner, R. M., 1997, Some comments on climatic reconstructions from ice cores drilled in areas of high melt. Journal of Glaciology, 43(143), 90-97. Kohler, J. and others., 1998, Effect of a Controlled Discharge Pulse on the Subglacial Drainage System and Ice flow at Engabreen, Northern Norway. Eos, 79(45), 274. Lappegard G., Kohler, J. and Hagen, J. O., 1998, Subglacial pressure variations beneath Engabreen, northern Norway. Eos, 79(45), 310. Laumann, T. & Reeh, N., 1993, Sensitivity to climate change of the mass balance of glaciers in southern Norway. Journal of Glaciology, 39(133), 656-665. Liestøl, O., 1967, Storbreen Glacier in Jotunheimen, Norway. Norsk Polarinstitutt, Skrifter, 141, 1-63. Matthews, J. A., 1984, Limitations of 14C dates from buried soils in reconstructing glacier variations and Holocene climate. In: Climate changes on a yearly to millennial basis (Ed. by N. A. Mörner and W. Karlén), pp. 281-290. Reidel, Dordrecht. Nesje, A. and Dahl, S. O., 1993, Late Glacial and Holocene glacier fluctuations and climate variations in western Norway: a review. Quaternary Science Reviews, 12, 255-261.

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Nesje, A. and Kvamme, M., 1991, Holocene glacier and climate variations in western Norway: evidence for early Holocene glacier demise and multiple Neoglacial events. Geology, 19, 610-612. Nesje, A., Lie, Ø. and Dahl, S. O., 2000, Is the North Atlantic Oscillation reflected in Scandinavian glacier mass balance records? Journal of Quaternary Science, 15, 587-601. Nesje, A., O., D. S., Anda, E. and Rye, N., 1988, Block fields in southern Norway: Significance for the Late Weichselian ice sheet. Norsk Geologisk Tidsskrift, 68, 149-169. Nye, J., 1997, The Cambridge Austerdalsbreen expeditions, 1954-1963. Annals of Glaciology, 24, 1-5. Nye, J. F., 1965, A numerical model of inferring the budget history of a glacier from its advance and retreat. Journal of Glaciology, 5(41), 589-607. Oerlemans, J., 1986, An attempt to simulate historic front variations of Nigardsbreen, Norway. Theoretical and Applied Climatology, 37, 126-135. Oerlemans, J., 1992, Climate sensitivity of glaciers in southern Norway: application of an energy-balance model to Nigardsbreen, Hellstugubreen and Alfotbreen. Journal of Glaciology, 38(129), 223-232. Oerlemans, J., 1997, A flowline model for Nigardsbreen: projection of the future glacier length based on dynamic calibration with the historic record. Annals of Glaciology, 24, 382-389. Olsen, L., Sveian, H. and Bergstrøm, B., 2001, Rapid adjustments of the western part of the Scandinavian Ice Sheet during the Mid and Late Weichselian - a new model. Norsk Geologisk Tidsskrift, 81, 93-118. Østrem, G., 1972, Height of the glaciation level in Northern British Columbia and Southeastern Alaska. Geografiska Annaler, 54A, 76-84. Østrem, G. and Brugman, M., 1991, Glacier mass balance measurements. Environment Canada, National Hydrology Research Institute, Saskatoon. Østrem, G. and Haakensen, N., 1993, Glaciers of Europe - Glaciers of Norway. In: Satellite image atlas of glaciers of the world, Vol. 1386-E (Ed. by R. S. Williams and J. G. Ferrigno), pp. 63-109. U.S. Geological Survey professional paper, Washington. Østrem, G., Haakensen, N. and Melander, O., 1973, Atlas over breer i Nord-Skandinavia, pp. 315. Norges vassdrags- og energiverk, Vassdragsdirektoratet, Meddelelse fra Hydrologisk avdeling. Østrem, G., Liestøl, O. and Wold, B., 1976, Glaciological investigations at Nigardsbreen, Norway. Norsk Geografisk Tidsskrift, 30, 187-209. Østrem, G., Selvik, K. D. and Tandberg, K., 1988, Atlas over breer i Sør-Norge. Norges vassdrags- og energiverk, Vassdragsdirektoratet, Meddelelse fra Hydrologisk avdeling, Oslo. Østrem, G. and Tvede, A., 1986, Comparison of glacier maps - a source of climatological information. Geografiska Annaler, 68A, 225-231. Øyen, P. A., 1908, Bidrag til braeegnes glacialgeologi. Nyt magasin for Naturvidenskap, 46, 301-379. Pohjola, V. A. and Rogers, J. C., 1997, Coupling between the atmospheric circulation and extremes of the mass balance of Storglaciären, northern Scandinavia. Annals of Glaciology, 24, 229-233. Rea, B. R., Whalley, W. B., Dixon, T. S. and Gordon, J. E., 1999, Plateau icefields as contributing areas to valley glaciers and the potential impact on reconstructing ELAs: a case study from the Lyngen Alps, North Norway. Annals of Glaciology, 28, 97-102.

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Rye, N., Nesje, A., Lien, R. and Anda, E., 1987, The late Weichselian ice sheet in the Nordfjord-Sunnmøre area and deglaciation chronology for Nordfjord, western Norway. Norsk Geografisk Tidsskrift, 41, 23-43. Schneeberger, C., Albrecht, O., Blatter, H., Wild, M. and Hock, R., 2001, Modelling the response of glaciers to a doubling in atmospheric CO2: a case of Storgläciaren, northen Sweden. Climate Dynamics, 17, 825-834. Schytt, V., 1959, The glaciers of the Kebnekajse-massif. Geografiska Annaler, 41, 213-227. Schytt, V., 1993, Glaciers of Europe - Glaciers of Sweden. In: Satellite image atlas of glaciers of the world, Vol. 1386-E, pp. 111-125. U.S. Geological Survey professional paper, Washington. Svenonius, F., 1910, Studien über den Kârso- und die Kebnesletscher nebst Notizen über andere Gletscher im Jukkasjärvigebirge. Sveriges geologiska undersökning, Ca5(1), 1-54. Theakstone, W., Raben, P. and Tørseth, K. (2000) Relations between winter climate and ionic variations in a seven-meter-deep snowpack at Okstindan, Norway 1995. Arctic, Antarctic and Alpine Research, 32, 189-196. Theakstone, W. H., 1990, Twentieth-century glacier change at Svartisen, Norway: the influence of climate, glacier geometry and glacier dynamics. Annals of Glaciology, 14, 283-287. Theakstone, W. H. and Jacobsen, F. M., 1997, Digital terrain modelling of the surface and bed topography of the glacier Austre Okstindbreen, Okstindan, Norway. Annals of Glaciology, 24, 148-151. Theakstone, W. H. and Knudsen, N. T., 1984, Recent changes of some glaciers of East Svartisen. Geografiska Annaler, 66A, 367-380. Tvede, A. M. and Laumann, T., 1997, Glacial variations on a meso-scale: examples from glaciers in the Aurland Mountains, southern Norway. Annals of Glaciology, 24, 130-135. Wallén, C. C., 1959, The Kårsa glacier and its relation to the climate of the Torne Träsk region. Geografiska Annaler, 41, 236-244. Whalley, W. B., 1973, A note on the fluctuations of the level and size of Strupvatnet, Lyngen, Troms, and the interpretation of ice loss from Strupbreen. Norsk Geografisk Tidsskrift, 27, 39-45. Whalley, W. B., Gordon, J. E., Gellatly, A. F. and Hansom, J. G., 1996, Plateau and valley glaciers in north Norway: responses to climate over the last 100 years. Zeitschrift für Gletscherkunde und Glazialgeologie, 31, 115-124. Winkler, S., 1996, Frührezente und rezente Gletscherstandsschwankungen in Ostalpen und West-/Zentralnorwegen. Trierer Geographischen Studien, 15, 580. Winkler, S., 2001, Neue Ergebnisse zur holozänen Gletscher- und Klimadynamik in Nordnorwegen. Norden, 14, 115-126. Winkler, S., 2002, A new Interpretation of the date of the 'Little Ice Age' maximum at Svartisen and Okstindan, northern Norway. The Holocene.

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PART III: THE CLIMATE AND ENVIRONMENTAL CHANGE RECORD OVER THE LAST 200 - 500 YEARS

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FOUR CENTURIES OF CLIMATIC VARIATION ACROSS THE TIBETAN PLATEAU FROM ICECORE ACCUMULATION AND δ18O RECORDS

Mary E. Davis and Lonnie G. Thompson

1.

INTRODUCTION

The Tibetan Plateau is an important component of the South Asian Monsoon system, but there are relatively few long, high-resolution records of precipitation and temperature variability from the Plateau itself. Only recently has it been recognized that the distribution and the chemistry of the precipitation across this large, elevated landmass is not uniform (AraguásAraguás et al., 1999; Tian et al., 2001; Liu and Lin., 2001). These studies have been based on data from precipitation and river water sample collection and analysis, and on patterns of meteorological and reanalysis data. Here, longer, high-resolution records of spatial and temporal variations of temperature and precipitation across the Tibetan Plateau since 1600 A.D. are presented from two ice cores. These ice-core sites are on the southern and northeastern edges of the Tibetan Plateau, thus providing wide spatial coverage (Figure 1). 145 L. D. Cecil et al. (eds.), Earth Paleoenvironments: Records Preserved in Mid- and Low-Latitude Glaciers, 145-161. © 2004 Kluwer Academic Publishers. Printed in the Netherlands.

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Figure -1. Map of western China and surrounding regions showing locations of the ice-core sites mentioned in the text and the mountain ranges in which they are located. The Tibetan Plateau is outlined by the dotted line.

In the boreal summer of 1987, three ice cores were drilled to bedrock on the Dunde ice cap (38o 06’N; 96o 24’E, 5325 masl) in the Qilian Mountains on the northeast edge of the Plateau. Ten years later, two cores were recovered to bedrock from the 7200 masl Dasuopu ice cap in the Himalayas (28o 23’N, 85o 43’E). The complete records of these cores and their climatic interpretations are presented in Thompson et al. (1989) and Thompson et al. (2000), respectively. They were analyzed for a variety of physical and chemical properties, including (but not limited to) oxygen isotopic ratios (δ18O) and reconstructed net balance (annual accumulation rate, or an). For

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the most recent parts of the records where seasonal stratigraphy was discernible, annual averages of δ18O and an were calculated, resulting in high-resolution records of climatic variation that span several hundred years. The records from these ice-core sites are now compared with other climate records and with each other to assess the effects of temperature and precipitation on δ18O values across the Tibetan Plateau, and to determine the spatial and high- and low-frequency temporal variations of temperature and precipitation since 1600 A.D., the extent of annual resolution for the Dunde record (for comparison Dasuopu is annually resolvable to 1440 A.D.).

2.

SEASONALITY IN THE TIBETAN PLATEAU ICE CORES

Figure -2. δ18O and concentration data from the (a) Dunde and (b) Dasuopu ice cores back to the depth corresponding to 1963 in each core, which is the time horizon marked by the ȕradioactivity peak in the Tibetan Plateau ice fields resulting from the Soviet Arctic thermonuclear tests.

Seasonal relationships between mineral dust concentration and δ18O in these two ice-core records are shown in Figure 2. In the Dunde ice, more negative (18O-depleted) δ18O values are generally contemporaneous with

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high dust concentrations, which occur in late winter and early spring (Davis, 2002). The opposite relationship occurs in the Dasuopu record from the central Himalayas; i.e. the high winter/spring dust concentrations are concurrent with less negative (18O-enriched) values in the core. Alternatively, during the monsoon season in the boreal summer, the precipitation that falls on the Dasuopu ice cap is more 18O-depleted and contains fewer dust particles. Nevertheless, for both these ice fields, over 70 percent of the annual precipitation falls in the summer. This dichotomy of seasonal δ18O values between the north and the south was discussed by Araguás-Araguás et al. (1998). Using the database from the IAEA/WMO Global Network “Isotopes in Precipitation” that spans from 41 to 14 years in Asia, depending on the location, the spatial distribution of the seasonal differences in δ18O shows a reversal between the two sides of the Plateau, with the dividing line at approximately 35oN. Using stable isotopes from a suite of river water and precipitation samples collected along a meridional transect on the Plateau, Tian et al. (2001) found a 15 permil decrease from north to south in summer values of δ18O. According to Tian and his colleagues, the spatial patterns of the oxygen isotopes, along with the deuterium excess measured on these water samples, suggest that the Plateau can be divided into two regions of varying degrees of monsoon influence, with the Tanggula Mountains marking the border. South of this range to the Himalayas, the summer moisture is monsoon-derived, although recycled during its transport over the Indian subcontinent, and to the north the influence of the summer monsoon is greatly diminished.

2.1

δ18O and accumulation records from Tibetan Plateau ice cores since 1600 A.D.

The δ18O profiles from the Dunde and Dasuopu ice cores are shown in Figure 3a and 3b as five-year averages. The Northern Hemisphere temperature anomaly data of Jones et al. (1998), also calculated in five-year averages, are shown in Figure 3c. All the records are broadly similar, with gradually increasing trends until the middle of the 19th century, then a steeper positive slope toward the present. The δ18O and the temperature anomalies peak in the middle 20th century, decrease slightly in the 1960s and 1970s, then increase toward the present.

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Figure -3. Five-year averages of δ18O values from the (a) Dunde and (b) Dasuopu ice cores back to 1600 A.D. The Northern Hemisphere temperature record of Jones et al. (1998), also calculated as five-year averages, is shown in (c).

The reconstructed accumulation records from these sites are displayed in Figure 4. The Dunde data show wet periods from 1600 to about 1760 and about 1880 to the present, and lower accumulation between 1760 and 1880. The Dasuopu record, however, has an inverse accumulation profile, with much higher rates during the period of lowest values in the northern site. In fact, the correlation coefficient between Dunde and Dasuopu (5-year averages) is –0.57, significant at the 99 percent level. The spatial patterns of accumulation between the north and south sides of the Tibetan Plateau that are presented by these ice cores suggests that the region affected by the monsoons, in which the more isotopically depleted snow falls in the summer, has an opposite precipitation history to that in the north, where the seasonal relationships of δ18O are also reversed.

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Figure -4. Five-year averages of accumulation from the (a) Dunde and (b) Dasuopu ice cores back to 1600 A.D. The Dasuopu d-excess, also calculated as five-meter averages, is shown in (c).

2.2

Temperature, precipitation and their influence on δ18O in Tibetan precipitation

Currently, there is a debate over the atmospheric and hydrological parameters that control the stable isotopic composition of rain and snow in low latitudes, including the Tibetan Plateau. In the case of δ18O in meteoric water in the northern part of Tibet, where the monsoon influence is weak or nonexistent, many believe that it is controlled mainly by temperature (Araguás-Araguás et al., 1998; Yao et al., 1996; Thompson et al., 1989).

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However, the controversy surfaces over the interpretation of stable isotopes in precipitation in the monsoon-dominated south. Because the summer precipitation in the Himalayas is more 18O-depleted than in the winter, Dahe et al. (2000), Shichang et al. (2000), and Tian et al. (2001) credit the “amount effect” as being the dominant influence on the isotopic fractionation. In the regions of Asia that lie within the domain of the Indian and Southeast Monsoons, Araguás-Araguás et al. (1998) noticed an inverse correlation between mean monthly δ18O in summer rainfall and the amount of precipitation. A simple explanation of the mechanism behind the amount effect in a monsoon regime is that the amount of precipitation is the controlling factor in the isotopic fractionation of oxygen according to the Rayleigh distillation model, since during precipitation the more 18O-enriched water vapor will initially condense and precipitate out, leaving more 18Odepleted water vapor behind to form the later condensate (Dansgaard, 1964). The longer a precipitation event occurs, the greater the amount of 18 O-depleted precipitation that falls. This applies to an air mass as it travels farther from its source such that progressively lower δ18O condensate forms and precipitates as the vapor moves inland. However, since the fractionation process is not always at equilibrium and the water vapor and the condensate are not isolated from the environment, the controls on δ18O in precipitation are more complex, especially in mountain regions. The ratio of 18O to 16O in alpine snowfall is also affected by temperature at the source and at the deposition site, vertical and horizontal transport pathways, and atmospheric circulation patterns (e.g. Grootes et al., 1989; Henderson et al., 2000; Araguás-Araguás et al., 2000; Thompson, 2001; Bradley et al., 2003). Taken together, the accumulation and oxygen isotope profiles from the Tibetan Plateau ice cores raise interesting questions about the influences of temperature and precipitation on stable isotopes. The δ18O of the Dunde site should be controlled primarily by temperature, since it is located in a more continental climatic regime (Araguás-Araguás et al., 1998; Tian et al., 2001). The comparison between the Dasuopu and Dunde δ18O records raises an issue: if the primary influence on the isotopic composition of snow in the north is temperature, while in the south it is monsoon intensity, why are the two records similar to each other? If the Dasuopu δ18O record is a proxy for monsoon precipitation intensity and is strongly influenced by the “amount effect”, then it should show an inverse relationship with its accumulation record. This would imply that the region of the Himalayas where Dasuopu is located has been experiencing progressively less snowfall since 1600 while the northeast Plateau region has been getting warmer. In fact, this has not been the case, as Figures 3b and 4b demonstrate. The most obvious disagreement occurred between 1810 and 1870, when the accumulation rate abruptly increased, but the δ18O did not decrease. In fact, the Dasuopu δ18O

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bears a much closer resemblance to the Northern Hemisphere temperature anomalies of Jones et al. (1998), which is discussed in Thompson et al. (2000) and is shown in Figure 3c. This presents a challenge to assertions about the role of the quantity of precipitation in the fractionation of oxygen in the Himalayan snow. The Dasuopu profile in Figure 3b is assumed to be indicative of the monsoon season record of δ18O variations on the basis that 70 percent or more of the annual precipitation falls in the summer season. To prove that this assumption is valid, the isotopic averages in the summer ice layers are compared with those in the annual layers. Because the accumulation is high and the rate of layer thinning is low in the upper 50 m of the Dasuopu ice cap, the annual cycles of δ18O can be separated into summer and winter averages over the most recent decades. As Figure 5a shows, the annual averages of δ18O in the Dasuopu record are reliable approximations of the monsoon averages, since the R2 between these two time series between 1949 and 1996 is 0.90. When these monsoon season δ18O values over this interval are compared with the annual accumulations (which are also closely related to the monsoon season amounts), the relationship between them, though present, is small (R2 = 0.09, Figure 5b). This is not to suggest that precipitation intensity in the Himalayas has no relationship to any stable isotope parameters. Deuterium was also measured in the Dasuopu ice samples, and from δH and δ18O, the deuterium excess (dexcess) was calculated. The d-time series shows an inverse relationship with the accumulation profile, particularly between 1800 and 1880 A.D., when the accumulation rate is high (Figure 4b,c). The d-excess (d=δ2H-8δ18O) is a measure of kinetic effects at both the source and the deposition of the moisture. It is lower when wind speeds are higher or when evaporation rates decrease at the source (Merlivat and Jouzel, 1979; Jouzel et al., 1982; Johnsen et al., 1989). In the Dasuopu ice core, the low d-excess that is contemporaneous with the high accumulation from 1810 to 1870 A.D is suggested by Thompson et al. (2000) to mean that the monsoon intensified greatly during this period.

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Figure -5. Regressions between (a) the average monsoon season δ18O and the annual average δ18O for the Dasuopu ice core from 1949 to 1996, and (b) the average monsoon season δ18O and the annual accumulation for the Dasuopu core from 1949 to 1996.

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PRECIPITATION SOURCES AND INFLUENCES

Figure -6. Dasuopu annual accumulation compared with rainfall amounts during June, July, August and September at Gauhati and Calcutta, India. All data are five-year averages. The Indian precipitation data is from the NOAA NCDC GCPS Monthly Station data set, available through the IRI/LDEO Climate Data Library (http://iridl.ldeo.columbia.edu).

The averages of the annual accumulation rates over the last four centuries on the Tibetan Plateau ice caps have been highly variable, from as little as 40 cm of ice per year on Dunde, up to 97 cm on Dasuopu. As mentioned above, most of the annual snow budget (70 to 80 percent) falls on these sites in the summer, so the records of δ18O and accumulation in Figures 3 and 4 can be considered as summertime climate histories. In the winter, the westerlies dominate at the 500 hecto Pascals (hPa) level throughout the Plateau, even along its borders, and the ice-core sites most likely receive their winter moisture from cyclonic activity that ultimately originates in the North Atlantic and travels through western and central Eurasia. Precipitation sources in the summer are more variable. The Himalayas, which are located well within the region of South Asian monsoon influence, receive their snow from the Indian Ocean through the Arabian Sea. Lin et al. (1990) and Shrestha et al. (2000) have noted two moisture trajectories

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over the Himalayas from the south: on the west side, water advects northward from the Arabian Sea over the Indian subcontinent, and the moisture for the eastern Himalayas travels from the Arabian Sea across India to the Bay of Bengal, then northward along the Brahmaputra River valley. The Dasuopu ice cap appears to be located close to the dividing point between these two regimes. Meteorological station records from Calcutta and Gauhati on the eastern edge of the Indian Peninsula (Figure 1) suggest increases in summer rainfall in the middle of the 19th century as seen in the Dasuopu accumulation record (Figure 6), although after this time the ice core and station data show no significant correlations. Tracking the summer sources of snow for the north Plateau is more complicated. The degree to which this region of Tibet is currently affected by the summer monsoon is uncertain. The Dunde ice cap is located in an area that Winkler and Wang (1993) assert is affected primarily by air masses that travel through Central Asia in the summer. Tian et al. (2001) have discovered that north of the Tanggula Mountains the river water chemistry (high δ18O and d-excess) indicates that precipitation is the result of local convective activity resulting from continental moisture recycling. There is evidence that the accumulation on the Dunde ice cap is linked to atmospheric processes in the North Atlantic. Similarities between the accumulation record, the July and August Icelandic Low sea level pressure variations since 1824, and the July and August North Atlantic Oscillation index since 1865 (Hurrell, 1995; Jones et al., 1997) are illustrated in Figure 7. Although these data appear to agree broadly, the relationship of high accumulation/high Icelandic Low sea-level pressure (SLP)/low NAO index is not what one might expect if the North Atlantic is the summertime source of moisture for the northern part of the Plateau. Liu and Yin (2001) analyzed meteorological data from an array of stations to determine that the summer indices of the NAO are related to the precipitation patterns over the eastern Tibetan Plateau. During negative phases of the NAO, southerly winds in the southern Tibetan Plateau and northerly winds in the northern Tibetan Plateau are intensified, resulting in lower precipitation in the north and higher precipitation in the south. However, comparisons of Figures 4 and 7 show the opposite configuration in the case of Dunde and Dasuopu, that is, during periods of low July and August NAO indices (and high Icelandic Low SLP’s), the accumulation rates over Dunde in the north increased, while on Dasuopu in the south they decreased.

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Figure -7. Relationship between the average of July and August Icelandic Low sea level pressure (heavy line), average July and August North Atlantic Oscillation Index (light line), and the annual accumulation from the Dunde ice core. All data are calculated as five-year averages.

An explanation for this may be provided by Fu et al. (1999), who noted that during summers of weakened westerlies from the North Atlantic to western Asia (which is a situation described by high NAO indices) the Asian monsoon troughs strengthen. This is more consistent with the Tibetan Plateau ice-core records; for example, the summertime Icelandic Low SLP was low (and the NAO was high) from the early to the middle 19th century, contemporaneously with the accumulation increase in Dasuopu which may have been caused by a strengthened monsoon. At the same time, the accumulation decrease in the Dunde record may have resulted from weakened cyclonic activity from the west. If these relationships between the ice cores and the North Atlantic summertime atmospheric conditions persist further back in the record, then these accumulation data from the north and south sides of the Tibetan Plateau may provide information on the variability of the summer NAO that spans several centuries.

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Comparison of Tibetan Plateau precipitation histories with other proxy records

There are few high-resolution climate records that have been recovered from the Tibetan Plateau and surrounding regions. One source of information comes from tree rings, such as the composite of seven records from the Tibetan Plateau (Wu, 1995) that is shown in Figure 8, along with the Dunde accumulation record. The broad, low-frequency variations are similar between the two records, especially the middle 19th century low, the increases in the 20th century, and the most recent abrupt drop in precipitation. Unfortunately, Wu does not indicate from where on the Tibetan Plateau these records come, but a description by Li (1985) of a precipitation record that was reconstructed from a set of tree-ring chronologies from the east Tien Shan in northwest China indicates that the dry periods here were from 1685-1725, 1813-1890, and from 1927 to the end of the record. This generally agrees with the time series shown in Figure 8.

Figure -8. Comparison between the precipitation history from the Tibetan Plateau that was reconstructed from tree rings (Wu, 1995) and the five-year averages of the Dunde accumulation back to 1606 A.D. The tree-ring data are presented as departures from the mean of the latest 30 years.

A high-resolution marine record of monsoon intensity provides a much different scenario than that presented by the Dasuopu ice core. Increasing

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abundances of G. bulloides in shallow marine cores studied by Anderson et al. (2002) indicate that upwelling (and thus wind speed which is indicative of monsoon strength) has been intensifying since 1600 A.D. Since the middle of the 19th century this trend has been opposite to the Dasuopu accumulation profile. The marine record has broadly tracked the Northern Hemisphere temperature curves of Mann et al. (1998) and Jones et al. (1998), and the δ18O record from Dasuopu. The link between the monsoon strength and these temperatures is made through the North Atlantic climate and its influence on the extent of Asian snow cover (Overpeck et al., 1996), which in turn may affect the intensity of the South Asian Monsoon (Barnett et al., 1988; Vernaker et al., 1995). However, another scenario involves the strengthening of the tropical hydrology resulting from the increasing Northern Hemisphere temperatures. The increase in sea surface temperatures would lead to greater evaporation rates, which would introduce more moisture in the lower troposphere. Higher surface temperatures would also increase the thermal gradient between land and ocean in the summer, which drives the Asian monsoon system (Hu et al., 2000).

4.

CONCLUSIONS

The Tanggula Mountains, which span east-west across the Plateau at 32o to 33oN, mark the location of a precipitation transition between the north and south sides of the Tibetan Plateau. North of this latitude, the accumulation rate is low, 18O-enriched precipitation falls in the summer (Figure 2, this paper; Araguás and Araguás et al.,1998), and the d-excess is high (Tian et al., 2001). To the south, the annual accumulation is higher, the 18O-enriched precipitation is deposited in the winter, and the d-excess is lower. In addition, the loading pattern of the first unrotated principle component of the summertime precipitation anomalies (Liu and Yin, 2001) are out of phase between north and south of approximately 33oN across the eastern part of the Plateau. Many investigators, such as Tian et al. (2001) and Araguás and Araguás et al. (1998) have concluded that the Tanggula Mountains are a barrier to the northward expansion of the influence of the South Asian Monsoon, while to the north recycled rainfall from continental precipitation processes dominate. Araguás-Araguás et al. (1998) noted that the line that separates the different characteristics of precipitation in the north and the south sides of the Plateau occurs at the summertime northernmost extent of the Intertropical Convergence Zone over this region. The spatial and temporal comparisons of the stable isotope and accumulation records from the ice cores recovered from the north and the south of the “Tanggula transition” tend to agree with the stable isotope

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analyses of water and precipitation samples from across the region, and with the patterns of precipitation variations derived from meteorological station data. Several studies of the stable-isotope chemistry of the precipitation in the South Asian Monsoon region link the “amount effect” with the low summertime values of δ18O, at least on monthly time frames, but over semidecadal and longer time scales the strongest links appear to be with atmospheric temperature. Currently, an ice core that was drilled in the Tangulla Mountains in 2000 is being analyzed in the same manner as the Dunde and Dasuopu cores, and it is hoped that the stable isotope and accumulation history from this region will help refine the northern extent of the monsoon influence over time. In the final analysis, no single climate record will provide the complete picture of the history of so large, complex, and spatially variable a system as the South Asian Monsoon. Since the summer precipitation from this oceanatmosphere interaction is so vital to such a large, densely populated region, it is important that more information be gathered and analyzed on its spatial and temporal patterns in order to help understand the processes that are responsible for the variability in these patterns. It is also important to be able to understand the effects of the recent warming on the intensity and distribution of the monsoon precipitation.

5.

ACKNOWLEDGEMENTS

We would like to thank Dr. Yao Tandong and the staff of the Laboratory of Ice Core and Cold Regions Environment, Cold and Arid Regions Environmental and Engineering Research Institute, Lanzhou, China for their collaboration in the Dunde and Dasuopu ice core programs. We would also like to thank Dr. N. Gundestrup of the University of Copenhagen for the analyses of the δ18O in the Dunde ice core shown in Figure 2, and Pin-Nan Lin for the analyses of the Dasuopu ice cores for stable isotopes. John Bolzan reconstructed the accumulation profiles for the Dunde ice core, and Keith Henderson produced the Dasuopu accumulation data. The Tibetan Plateau ice-core programs were funded by the National Science Foundation’s Office of Climate Dynamics and Division of Polar Programs (Dunde), the NSF-ESH program (Dasuopu), The National Geographic Society, and The Ohio State University. This is Byrd Polar Research Center contribution number 1288.

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REFERENCES

Anderson, D.M., Overpeck, J.T., and Gupta, A.K., 2002, Increase in the Asian Southwest Monsoon during the past four centuries. Science 297, 596-599. Araguás-Araguás, L., Froehlich, K., and Rozanski. K., 1998, Stable isotope composition of precipitation over southeast Asia. Journal of Geophysical Research D 103, 28,72128,742. Araguás-Araguás, L., Froehlich, K., and Rozanski, K., 2000, Deuterium and oxygen-18 composition of precipitation and atmospheric moisture. Hydrological Processes 14, 13411355. Barnett, T. P., Dumenil, L., Schlese, U., and Roeckner, E., 1988, The effect of Eurasian snow cover on global climate. Science 239, 504-507.Bradley, R.S., M. Vuille, D. Hardy, and L.G. Thompson (2003) Low latitude ice cores record Pacific sea surface temperatures. Geophysical Research Letters, 30, 10.1029/2002GL016546 Bradley, R.S., Vuille, M., Hardy, D.R., and Thompson, L.G., 2003, Low latitude ice cores record Pacific sea surface termperatures. Geophysical Research Letters, 30, 1174-1177; doi:10.1029/2002GL016546 Dahe, Q. and 9 others, 2000, Evidence for recent climate change from ice cores in the central Himalaya. Annals of Glaciology 31, 153-158. Dansgaard, W., 1964, Stable isotopes in precipitation. Tellus 16, 436-468. Davis, M.E., 2002, Climatic interpretations of eolian dust records from low-latitude, highaltitude ice cores. Ph.D. Dissertation, The Ohio State University. Fu, C., Diaz, H.F., Dong, D., and Fletcher, J.O., 1999, Changes in atmospheric circulation over Northern Hemisphere Oceans associated with the rapid warming of the 1920s. International Journal of Climatology 19, 581-606. Grootes, P.M., Stuvier, M., Thompson, L.G., and Mosley-Thompson, E., 1989, Oxygen isotope changes in tropical ice, Quelccaya, Peru. Journal of Geophysical Research D 94, 1187-1194. Henderson, K.A., Thompson, L.G., and Lin, P.N., 2000. Recording of El Niño in ice core δ18O records from Nevado Huascarán, Peru. Journal of Geophysical Research D, 104, 31,053- 31,065. Hu, Z.-Z., Latif, M., Roeckner, E., and Bengtsson, L., 2000, Intensified Asian summer monsoon and its variability in a coupled model forced by increasing greenhouse gas concentrations. Geophysical Research Letters, 27, 2681-2684. Hurrell, J.W., 1995, Decadal trends in the North Atlantic Oscillation and relationships to regional temperature and precipitation. Science, 269, 676-679. Johnsen, S.J., Dansgaard, W., and White, J.W.C., 1989, The origin of Arctic precipitation under present and glacial conditions. Tellus, Series B, 41B, 452-468. Jones, P.D., Jonsson, T., and Wheeler, D., 1997, Extension to the North Atlantic Oscillation using early instrumental observations from Gibralter and South-west Iceland. International Journal of Climatology, 17, 1433-1450. Jones, P.D., Briffa, K.R., Barnett, T.P., and Tett, S.F.B., 1998, High-resolution palaeoclimatic records for the last millennium: interpretation, integration and comparison with general circulation model control-run temperatures. Holocene, 8, 455-471. Jouzel, J. and Merlivat, L., 1982, Deuterium excess in an East Antarctic ice core suggest higher relative humidity at the oceanic surface during the last glacial maximum. Nature, 299, 688-691.

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Jouzel, J. and Merlivat, L., 1984, Deuterium and oxygen-18 in precipitation: modeling of the isotope effects during snow formation. Journal of Geophysical Research, 89, 11,74911,757. Li, J., 1985, Climatic variations of Xinjiang in recent 3000 years. Papers on Quaternary Research in Arid Area of Xinjiang, 1-7 (in Chinese). Lin, Z. and Wu, X., 1990, A preliminary analysis about the tracks of moisture transportation on the Qinghai-Xizang Plateau (in Chinese with English abstract). Geographical Research, 9, 30-49. Liu, X. and Lin. Z.Y., 2001, Spatial and temporal variation of summer precipitation over the eastern Tibetan Plateau and the North Atlantic Oscillation. Journal of Climate, 14(13), 2896-2909. Mann, M.E., Bradley, R.S., and Hughes, M.K., 1998, Global-scale temperature patterns and climate forcing over the past six centuries. Nature, 392, 779-787. Merlivat, M. and Jouzel, J., 1979, Global climatic interpretation of the deuterium-oxygen-18 relationship for precipitation. Journal of Geophysical Research, 84, 5029-5033. Overpeck, J.T. and 17 others, 1997, Arctic environmental change of the last four centuries. Science, 278, 1251-1255. Shichang, K, Wake, C.P., Dahe, Q., Mayewski, P.A., and Yao, T., 2000, Monsoon and dust signals recorded in the Dasoupu glacier, Tibetan Plateau. Journal of Glaciology, 46, 222226. Shrestha, A.B., Wake, C.P., Dibb, J.E., and Mayewski, P.A., 2000, Precipitation fluctuations in the Nepal Himalaya and its vicinity and relationship with some large scale climatological parameters. International Journal of Climatology, 20, 317-327. Tian, L.,. Masson-Delmotte, V., Stievenard, M., Yao, T., and Jouzel, J., 2001, Tibetan Plateau summer monsoon northward extent revealed by measurements of water stable isotopes. Journal of Geophysical Research D, 106, 28,081-28,088. Thompson, L.G., 2001, Stable isotopes and their relationship to temperature. In: Geological Perspectives of Global Climate Change (L.C. Gerhard, W.E. Harrison, and B.M. Hanson, eds.) AAPG Studies in Geology No. 47, Tulsa, pp. 99-120. Thompson, L.G. and 9 others, 1989, Holocene-Late Pleistocene climatic ice core records from Qinghai-Tibetan Plateau. Science 246, 474-477. Thompson, L.G., Yao, T., Mosley-Thompson, E., Davis, M.E., Henderson, K.A., and Lin, P.N., 2000, A high-resolution millennial record of the South Asian Monsoon from Himalayan ice cores. Science 289, 1916-1919. Vernekar, A.D., Zhou, J., and Shukla, J., 1995, The effect of Eurasian snow cover on the Indian monsoon. Journal of Climate 8, 248-266. Winkler, M.G. and Wang, P.K., 1993, The Late-Quaternary vegetation and climate of China. In: Global Climates since The Last Glacial Maximum (H.E.Wright, Jr., J.E. Kutzbach, T. Webb III, W.F. Ruddimann, F.A. Street-Perrott, and P.J. Bartlein, eds.) University of Minnesota Press, Minneapolis, pp. 221-261. Wu, X.D., 1995, Dendroclimatic studies in China. In: Climate since A.D. 1500 (R.S. Bradley and P.D. Jones, eds.), Routledge, New York, pp. 432-445. Yao, T., Thompson, L.G., Mosley-Thompson, E., Zhihong, Y., Xingping, Z., and Lin, P.N., 1996, Climatological significance of δ18O in north Tibetan ice cores. Journal of Geophysical Research D 101, 29,531-29,537.

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CLIMATIC CHANGES OVER THE LAST 400 YEARS RECORDED IN ICE COLLECTED FROM THE GULIYA ICE CAP, TIBETAN PLATEAU

Yao Tandong and Yang Meixue

1.

INTRODUCTION

Precipitation is a major factor in climatological studies. Ice archived in glaciers and ice caps is an important source for records of past precipitation and climate although it is often limited by the continuity of the available data at any given site. The older the age of an ice core, the more obvious the limitation is. Accumulation recorded in ice cores is one of the most reliable indexes for the recovery of past precipitation. The accumulation of snow and ice is an indirect record of the precipitation on a glacier. Therefore, a highresolution record of ice-core accumulation may provide a continuous record of long-term regional precipitation. However, the accumulation can vary with elevation, as precipitation above snow line is usually much larger than precipitation in the ablation area. It follows that changes in relative humidity occur with changes in elevation. In our study area, the Qinghai-Tibet Plateau, relative humidity is quite different from that found in eastern China. The relative humidity in the Tibetan Plateau does not decrease with an increase of altitude as it does in eastern China, but increases within a certain altitude range and reaches the highest humidity at about 400 hPa (hectapascals, a unit of air pressure that is equal to 1 millibar). After the point in altitude that this largest humidity is reached, the relative humidity decreases with increasing altitude. (Figure 1). Summer is a rainy season in the Qinghai-Tibet Plateau and the mean height of the convective winds over the Plateau during this season is about 400 hPa. The precipitation captured at and near this height is larger than that at lower altitude. The top of Guliya ice cap is nearly 7000 masl and lies just within this high humidity layer. Thus, the amount of the precipitation in the area of the summit of the Guliya ice cap should be larger than the precipitation at meteorological stations located at lower altitudes 163 L. D. Cecil et al. (eds.), Earth Paleoenvironments: Records Preserved in Mid- and Low-Latitude Glaciers, 163-180. © 2004 Kluwer Academic Publishers. Printed in the Netherlands.

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Climatic Changes over the Last 400 Years

Figure -1. Changes in relative humidity with changes in altitude for the summer season in the Qinghai-Tibet Platteau.

The Guliya ice cap is the largest polar-type ice cap in the Tibetan Plateau, with an area of 376 square kilometers (km2) and is located at 35°17’ north latitude and 81°29’ east longitude (Thompson et al., 1995; Yao et al.,1995b; and Yao et al, 1997). In 1992, three ice cores with respective lengths of 308.7 m, 93.2 m and 34.5 m were collected successfully between 6200 and 6700 masl The deepest core collected (308.7 m) reached the bedrock. The 308.7-m ice core at 6200 masl was collected (Figure 2) using an electromechanical drill from a dry borehole to a depth of 200 m. At that depth (200 m), a thermal drill with an alcohol-water mixture was used to complete the core to bedrock at a depth of 308.7 m. Measured temperatures in the borehole were –15.6 degrees Celsius (°C) at a depth of 10 m below the surface, -5.9°C at a depth of 200 m, and –2.1°C at the ice-bedrock interface. Mass balance calculations made in 1990 and 1991 indicate that the ice cap receives accumulation of about 200 mm weq per year (Thompson et al, 1997; Yao et al , 1997). The field and laboratory methods used in this study

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have been described elsewhere (Thompson et al 1995, 1997; Yao et al, 1996).

Figure -2. The locations of the Guliya (G) and Dunde (D) ice caps and the site where the 308.6 m Guliya ice core was drilled. Drill site elevation is at 6200 masl. (Modified from Thompson et al, 1997).

Stable isotopes of oxygen archived in ice cores have been used as proxies for historical temperature recontruction in the Guliya ice core. The results of these studies indicate that there is a positive relationship between delta oxygen-18 (δ18O) and temperature on the Tibetan Plateau (Yao et al, 1995a, 1995b, 1996). Additionally, glacial accumulation on the Guliya ice cap was measured by using three methods: (1) accumulation stakes at the glacial surface, (2) documentation of visible stratigraphy in snow pits, and (3) insoluble particulate gross-beta and tritium radioactivity (Thompson et al, 1995). Yao et al. (1996) concluded from these investigations that net accumulation on the Guliya ice cap closely matches the measured precipitation.

2.

CALCULATION OF GLACIAL ACCUMULATION FROM THE GULIYA ICE CORE

Although annual layer thickness may be identified in an ice core, this may not represent the original snowfall accumulated at the glacial surface. This is, in part, a result of glacial deformation. Model and field validation may be necessary to reconstruct the original annual accumulation. In recent years, several snow and ice accumulation validation models have been

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discussed in the literature including, Nye (1963), Reeh (1978), Willans (1979), and Raymond (1983). The Nye (1963) model was the first to attempt to validate glacial accumulation compared to field measurements. Nye proposed that the bottom of the ice at a glacier was frozen to the bedrock underneath and the strain rate along any vertical line within the glacier was the same at any point in time. Further, he proposed that the amount of strain on any annual layer within the glacier is equal to the total strain on the glacier under the annual layer in question according to the following equation:

λi λ0 = Y H

(1)

Where Ȝi is the thickness of the glacier layer i in the ice core, Ȝ0 isthe thickness being validated, Y is the distance from layer i to the bottom of the glacier, and H is the total thickness of the glacier. The units are weq. Reeh (1978) validated the annual records of ice thickness and adjusted the layer thickness in such a way as to take into account the position of the drill site with respect to the direction of ice flow and the magnitude of ice strain due to flow. Additionally, Reeh (1978) transformed the annual layer thickness of three ice cores collected in Greenland (Dye-3, Milcent, and Crete) into accumulation. Using glacier-strain network data collected at Byrd Station, Antarctic, Willans (1979) constructed a glacier flow model and also derived an accumulation validation equation. That equation is: ti

λi = λ0 exp ³ e(t ) dt

(2)

t0

Where e(t) is the vertical strain rate of the glacier with time, ti and t0 are the time associated with Ȝi (the thickness of the glacier layer the i) and Ȝ0 (the thickness being validated) respectively. This equation was used to transform the annual layer thickness of an ice core into accumulation. Raymond (1983) also developed a formula that relates the vertical velocities of the glacier surface with corresponding layers (i) in an ice core. This equation was used to calculate annual layer thickness: n +1 § V0 · ° § Y · ª 1 1 § Y · º ½° ¸¸ • ®1 − ¨1 − ¸ • «1 + 1 − − ¨ ¸ »¾ © Vs ¹ °¯ © H ¹ ¬« n + 1 n + 1 © H ¹ ¼» °¿

λi = λ0 ¨¨

(3)

Where n is the index of the glacial-flow law and usually is taken as 3 for the simplicity of the calculation (Glan, 1958 and Raymond, 1983) and Ȝi is the thickness of the glacier layer i in the ice core, Ȝ0 isthe thickness being

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validated, Y is the distance from layer i to the bottom of the glacier, and H is the total thickness of the glacier. In order to calculate the annual ice-layer thicknesses in the Guliya ice core, we have developed a more rigorous validation model. In order to calculate the net accumulation and the variation in annual layer thickness, we used a new equation to estimate the optimal correlation simulation resulting in a more realistic fit to the field observations for our study area. That equation is:

§ Y · λi = λ0 ¨1 − ¸ © H¹

p +1

(4)

Where, again, n is the index of the glacial-flow law usually taken as 3 for the simplicity of the calculation (Glan, 1958, Raymond, 1983) and Ȝi is the thickness of the glacier layer i in the ice core, Ȝ0 is the thickness being validated, Y is the distance from layer i to the bottom of the glacier, and H is the total thickness of the glacier. The results of comparing calculations of ice-layer thickness in correlation with actual field measurements, we found that the glacial-flow law term (n) is about 2.369 for the Guliya ice core. When calculating accumulation as reflected in the ice core, we assumed that the glacier was stable in the past and the change of the horizontal strain in the glacier within this thickness was linear and without abrupt change. Thus, the glacial velocity vertical to the glacial surface is equal to the annual layer thickness. The area of the Guliya Ice Cap is more than 300 km2 and the area at the top is more than 100 km2. According to the debris at the terminal end of the glacier, the ice cap is stable. There is no abrupt change in ice layers from the top to the bottom. Thus, the features of the Guliya Ice Cap satisfy the assumptions made for our calculations. We can use eq. (4) and n=2.369 to calculate the precipitation over the last 400 years (Figure 3). It can be seen that there were great changes of the annual precipitation recorded in accumulation of the Guliya ice core. The maximum precipitation recorded in the ice core is an order of magnitude greater than the minimum (Figure 3a). The variation of a 10-year averaged precipitation value could reflect the changes in the 10-year time scale and reveal a trend (Figure 3b). However, in Figure 3c, an eleven-point moving average of the 10-year averaged values (Figure 3b), the trend of the 100-year time scale is more distinguishable. Since approximately the year 1571, two periods of small precipitation (the years from about 1571 to 1681, and from 1811 to 1901) and two periods of large precipitation (the years from about 1691 to 1800, and from 1911 to 1990) are suggested by the plot of the eleven-point running average (Figure 3c). The average precipitation in the two elevated precipitation periods was 222 mm, while the average precipitation in the two small precipitation periods was 181 mm or 21 percent.

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Climatic Changes over the Last 400 Years

Figure -3. The annual fluctuation of precipitation recorded in the Guliya ice core in the past 400 years. (a) Annual fluctuation; (b) 10-year average; (c) 11-point moving average of 10year average.

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169

A COMPARISON BETWEEN THE PRECIPITATION VARIATIONS OF THE GULIYA ICE CAP AND ITS VICINITY

Recently, our team drilled an ice core (Tanggula ice core) on the Dongkemadi Glacier in the Tanggula Mountains in eastern Qinghai-Tibet Plateau (Yao et al, 1993). This ice core is relatively short in length and the precipitation (from accumulation) since the year 1940 was retrieved. Figure 4 shows a comparison between the precipitation recorded in the accumulation of the Tanggula ice core and that in the Guliya ice core since 1940. There is evidence of a phase difference between the accumulations as a result of precipitation recorded in the two ice cores. However, their variation trends are consistent with what is known about precipitation in the northern hemisphere.

Figure -4. A comparison between the precipitation recorded in the Guliya ice core and the Tanggula ice core (5-point moving average of annual precipitation). Curve 1, Guliya ice core; curve 2, Tanggula ice core.

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Climatic Changes over the Last 400 Years

Figure -5. The secular trend of precipitation recorded in Gulya ice core (curve 1) and Dunde ice core (curve 2) (11-point moving average of 10-year average). Curve 1, Guliya ice core; curve 2, Dunde ice core.

The Dunde ice core was obtained from the Dunde Ice Cap, in the northern Tibetan Plateau. Using this ice core, the precipitation trend since the early 17th century was analyzed in detail by Yao et al (1990, 1992). The large difference between the Guliya ice core and the Dunde ice core on short time scales was observed. However, on a century time scale, the correlation is quite good. Figure 5 gives the 11-point moving average of the 10-year averaged accumulation in the Guliya ice core and in the Dunde ice core. The variation trends of precipitation recorded in the two ice cores are the same on a 100-year time scale. The only difference is that the precipitation in the Dunde ice core is larger than that in the Guliya ice core. Using the available data, Bradley et al (1987) systematically studied the precipitation variation in the northern hemisphere since the year 1850. Figure 6 gives a comparison between the precipitation recorded in the Guliya ice core and that in the northern hemisphere. The major trends of these two precipitation series are similar but their differences are also significant. The smaller precipitation periods in the northern hemisphere appeared approximately in the years 1860, 1910, and 1940, while those in the Guliya ice core appeared approximately in the years 1870, 1920, and 1950. There is a time lag of about 10 years in the northern hemisphere record as compared to the Guliya ice-core record.

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Figure -6. The precipitation in the northern hemisphere and in the Guliya ice core (5-point moving average of 10-year average). Curve 1, Guliya ice core; Curve 2, northern hemisphere (average).

Although the precipitation variation in the Guliya ice core lagged behind that in the northern hemisphere for about 10 years in secular trend, the extreme precipitation events in both the northern hemisphere and the Guliya ice core occurred simultaneously. After examining the extreme precipitation events over the past 100 years, Bradley found that the extremely dry years were 1912 and 1913 and the extremely wet years were 1953, 1954, 1956, and 1957 (Yao et al, 2000). Table 1 lists the precipitation in the extreme years for the northern hemisphere and for the Guliya ice cap. Table -1. Comparison between the extreme precipitation events in the northern hemisphere and the precipitation in the Guliya ice core. Note: the average precipitation on Guliya Ice Cap during 1400-1991 A.D. is 200 mm. Extreme Extremely wet year (A.D.) Extremely dry year (A.D.) precipitation events in northern 1953 1954 1956 1957 1912 1913 hemisphere Precipitation 196 402 224 408 42 42 recorded in Guliya ice core (mm) Anomaly

-4

+202

+24

+208

-158

-158

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4.

Climatic Changes over the Last 400 Years

ENSO EVENTS RECORDED IN THE GULIYA ICE CORE

In the last few decades, many reports have suggested that, in the lower latitudes, sometimes even in the mid and high latitudes, the annual variations of precipitation are related to the events of El Niño and the Southern Oscillation (ENSO), which is caused by large-scale interactions between the ocean and the atmosphere (Rasmusson et al, 1983; Ropelewski et al, 1987; Bradley et al, 1987). In most regions at mid- and low-latitudes of the earth, the ENSO events have significantly influenced annual climatic changes (Whetton and Rutherford, 1994). ENSO events are aperiodic, which occur with a frequency of 2 to 10 years (Rasmusson et al, 1982). A generally-used index of the ENSO event (Southern Oscillation Index [SOI]) is calculated from the atmospheric pressure difference between Tahiti and Darwin (Walker, 1923; Walker and Bliss, 1932). Over the last decade, ENSO events during the pre-instrumental period have been examined by using proxy data such as precipitation and other variables (Diaz et al 1992). However, the most detailed chronological analysis of such ENSO events came from the Peruvian coastal flooding records that are normally associated with the occurrence of El Niño events (Hamilton et al, 1986; Quinn et al 1987; Quinn et al, 1992). The QN chronology (Quinn and Neal, 1992) extended from the years 1525 to 1987 suggested that the El Niño phenomena had occurred throughout the Spanish colonization in Latin American with low variation frequency. The teleconnection chronology of El Niño and La Niña was produced from the pattern feature of ENSO on a global scale (Whetton and Rutherford,1994). However, different chronologies have been reported (Whetton and Rutherford, 1994;Quinn et al, 1987; Quinn et al, 1992). The studies in relation to the impact of ENSO events on climatic changes in China (Wang et al, 1981; Guo et al, 1987) suggested that a long-term connection existed between the SOI and the position and the strength of the west pacific subtropical high-pressure system. This has a great effect on the amount and timing of precipitation in eastern China. For example, the relationship between precipitation variation in semi-arid regions of northern China and ENSO events suggests a precipitation deficit in northern China associated with El Niño years (Wang et al, 1990).

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Figure -7. The annual variations of net accumulation in Guliya ice core since 300 BP. Smoothed curves constructed using a ten-year Gaussian filter are superimposed on each curve. Deviations of the raw data from smoothed curves (i.e. high passed filtered data) were used in subsequent analysis.

To better understand and predict the ENSO events, the past ENSO records are needed. Fortunately, the ice cores that preserved the paleoclimate information can provide long-term detailed records in relation to these extreme events (Thompson et al, 1984). In this paper, we focus our attention to discussing the relationships between ENSO events and the climate records (notably, precipitation and δ18O ) in the Guliya ice core, Tibetan Plateau. Figure 7 shows the annual variations of net accumulation in the Guliya ice core over the past 300 years. In order to remove the long-period variations we re-expressed the series as anomalies from a smoothed version of the series. The smoothing used a 10-year Gaussian manipulation, and the ends of the series were padded so that all data could be used (using the method of Jones et al, 1986 and Whetton and Rutherford, 1994). The smoothed series are shown in Figure 7. The deviations (as anomaly records) of the raw data from the smoothed curved (i.e. high-pass filtered data) (Figure 8) were used in subsequent analyses to ensure that any marked trends are not simply due to long-period fluctuations in the data sets. In the climate time series analysis, the accumulative anomaly chart is often used to determine the trend of a period and to determine whether a negative anomaly or a positive anomaly is dominant (Whetton and Rutherford, 1994). In this study, we used this technique to compare net accumulation with ENSO event chronology. If we use the QN El Niño chronology, from 1690 to 1987 (nearly 300 years), there are a total of 87 El Ni ño years. Of the 87 years, 60 correspond to negative precipitation anomalies (Figure 8). In other words, the precipitation is reduced in these 60 El- Niño years. In the years from 1844 to

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1915, there were 27 El-Niño years. Of these 27 El-Niño years, 20 were associated with negative precipitation anomalies. This was especially true during the period from 1844 to 1862, when there were a total of 9 El-Niño events that occurred consistently with reduced precipitation on the Guliya ice cap. Not all of the years with reduced precipitation must necessarily be El Niño events as precipitation amounts can be also affected by other factors. In addition, some events do not manifest typical ENSO-type behavior in the mid-latitudes. For example, Guliya precipitation anomalies are not consistently negative during El Niño. In fact, 25 El Niño years between 1709 and 1979 have precipitation anomalies that are positive. Among these 25 El NinѺo years, 12 are not listed in Whetton and Rutherford’s chronology (1994).

Figure -8. The anomalies of the net accumulation in Guliya ice core. In this figure, the anomalies are the deviations of the raw data (net accumulation and δ18O in Guliya ice core) from smoothed curves (i.e. high-pass filtered data). Bolded bars are QN El Niño years.

Figure 9 shows the cumulative anomalies of net accumulation and δ18O in QN El Niño years. The anomalies were used only for the years contained in the El Niño chronology; anomalies for other years were not used. Where a series in this plot shows a downward trend, there is a tendency for negative anomalies in the series to be associated with El Niño occurrence. This may indicate changes over time in the strength of any relationship between El Niño and the regional precipitation series. Note also that high-pass filtering has been applied to the continuous data sets before this analysis to ensure that any marked trends are not simply due to long-period fluctuations in the data sets. It is the slope of the accumulative anomaly that is important rather than its position relative to the zero intercept.

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Figure -9. Cumulative anomalies from the net accumulation in Guliya for the years of QN El NinѺo. Starting from the first year in each series, anomalies are used only for the years contained in the QN El Niño chronology; anomalies for other years are not used. A downward trend in a curve represents a tendency for negative anomalies in the series to be associated with QN El Niño. Note that the slope of the cumulative anomaly curve is important rather than its position relative to zero.

For precipitation recorded in the Guliya ice core, there was good evidence of association between reduced precipitation and El Niño years from 1770 to the present, except for a marked reversal in the relationship around 1815-1835 (Figure 9). Prior to 1770, there was no marked tendency for the accumulative precipitation anomaly. Whetton et al (1996) pointed out that there was a tendency for negative anomalies except for the years from 1710 to 1760 when the relationship was reversed (Yang et al, 2000). For the temperature recorded in the Guliya ice core, there was also evidence of an association between low temperature and El Niño years from about 1690 to 1940 except for a marked reversal in the relationship between the years 1815-1835 and 1865-1880. From 1940 to the present, a marked reversal in the relationship exists. The statistical significance of any relationships can be tested with a Mann-Whitney U-test by comparing values during QN El Niño years with the values in the full accumulation series. A one-tailed test is appropriate in

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testing the expected negative trend (Whetton and Rutherford, 1994). A onetailed statistical test was applied to three subdivisions of our data (for the years 1690 to 1991, for 1690 to 1840, and for 1841 to 1991). The results are given in Table 2. Table -2. The average anomaly of the Guliya ice core series in QN statistics are based on the Mann-Whitney U-test. The significance is p<0.1(*), p<,0.05(**), p<0.01(***), p<0.005(****). Series Whole period 1690-1840 (1690-1991) Precipitation -2.23 ** -0.32 18

δ O

-0.10

-0.43

El Niño years. The shown by asterisks: 1841-1991 -2.93 **** -0.27

For the entire period of record (1690-1991), the precipitation recorded in the Guliya ice core in El Niño years shows a significant decrease in probability (p<0.05), it is especially true for the period from 1841 to 1991 (p<0.005). However, the negative anomaly trend of the δ18O record is poorly related to the QN El Niño years, especially from 1940 to 1991 where it is markedly positive. In order to confirm the relationship between ENSO events and the climate anomaly, the SOI from 1882 to 1991 were used to compare with the precipitation in Guliya (Figure 10). The SOI are monthly values. In most cases, the reduced precipitation recorded in the Guliya ice core was associated with the lower SOI from autumn to spring and vice versa. The correlation coefficient between the SOI (an average from October through February) and the precipitation anomalies is 0.28 for the entire period of record (from 1690 through 1991) and is 0.37 for the period from 1935 through 1991.

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Figure -10. Comparison of the S0I with the anomalies of the precipitation recorded in Guliya ice core during 1882-1991. The S0I is monthly variation and the smoothed curve constructed using 12-month moving average also shown on Figure 10.

5.

CONCLUSIONS

Generally, the precipitation recorded at the meteorological station in our study area is influenced by local conditions, especially by the local climate or topography. Thus, scientists often pay more attention to temperature than to precipitation in climate studies. However, Bradley’s study demonstrated that the precipitation variation also exhibits significant similarity in a hemispheric context. Then, the question arises as to whether or not basic precipitation features (amount, timing, etc.) can be established for a given region. Based on the precipitation records from ice caps on the Qinghai-Tibet Plateau presented here, the answer to that question is yes. Comparing our ice-core record with the

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records from the meteorological station located at lower altitude, these ice caps are not influenced by local climate and micro-topography. Those ice caps located in the middle-upper part of the convection layer receive and collect all kinds of atmospheric information and are the real natural archive of the atmospheric precipitation record. The precipitation record in the Guliya ice core is important not only in global temperature studies, but also in the study of large-scope precipitation features. The negative anomalies of the precipitation recorded in Guliya ice core are significantly correlated with ENSO events. Thus, the ice-core records can be used as a proxy to study global climate anomalies associated with ENSO events. The variation of the precipitation recorded in the Guliya ice core show a teleconnection with ENSO events. Such teleconnections, although statistically significant, are relatively weak for some periods of record.

6.

REFERENCES

Bradley, R.S., Diaz, H.F., Eischeid P.T. et al., 1987, Precipitation fluctuation over the northern hemisphere and areas since the mid-19th century, Science, 237:171. Bradley, R.S., Diaz, H.F., Kiladis, G.N., et al, 1987, ENSO signal in continental temperature and precipitation records. Nature,327, pp. 497-501. Diaz, H.F. and Markgraf, 1992, El Niño: historical and palaeoclimatic aspects of the Soutjhern Oscillation. CUP. Cambridge, pp.175-192. Glan, J.W., 1958, The flow of ice law. A discussion of the assumptions made in glacier theory, their experimental foundations and consequences. IASH, 47, pp. 171-183. Guo, Q., 1987, The eastern Asian monsoon and the southern oscillation :1871-1980. The climate of China and Global climate, Ocean Press, 249-255. Hamilton, K.and Garcia, R.R., 1986, El Niño / Southern Oscillation events and their association mid latitude teleconnections. Bull. Am. Metero. Soc., 67,1354-1361. Hansen, J. and S.,Lebedeff, 1987, Global trends of measured surface air temperature. J. Geophys. Res, 92 D11,13,345-13,372. Jones,P.D., Raper,S.C.B., Bradley,R.S., Diaz, 1986, Northern hemisphere surface air temperature variations:1851-1984. J.Clim.Appl.Meteorol. 25,161-179. Li, Y.F., Yao, T.D., Huang, C.L., 1993, Spatial variations of chemical species in Guliya Ice Cap, J. Glaciology and Geocryology, 15(3):467. Nye, J.F., 1963, Correction factor for accumulation measured by the thickness of the annual layers in ice sheet, J. Glaciol. 4:785. Quinn, W. H. , Neal, V.T., and Antuniez et al, 1987, El Niño occurrences over the past four and half centuries. J. Geiphys. Res., 92,14,449-14,461. Quinn, W.H. and Neal ,V.T. . 1992 . The historical record of El Niño events, in Bradley, R.S. and Jones, P.D., (eds.) Climate since 1500A.D., Routledge, London, PP.623-648. Rasmusson, E.M., and Wallace, J.M., 1983, Meteorological aspects of the El Niño / Southern Oscillation. Science, 222,1195-1202. Rasmusson, E.M., and Carpenter, T.H., 1982, Variation in tropical sea surface temperature and surface wind fields associated with the Southern oscillation/ El Niñ o. Mon. Wea. Rev., 110,354-384.

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Raymond, G.F., 1983, Deformation in the vicinity of ice divides, J. Glaciolo, 29:357. Reeh, N., Fisher, D.A., Hammer, C.U., and Thomsen, H.H., 1987, Use of trace constituents of test flow models for ice sheets and ice caps, IAHS, pp 170-299. Ropelewski, C.F., and Halpert, M.S., 1987, Global and regional scale precipitation patterns association with the El Niño/ Southern oscillation. Mon. Wea. Rev., 115,1606-1626. Thompson, L.G., 1992, Ice core evidence from Peru and China, in Bradley, R.S., and Jones, P.D. (eds.), Climate since 1500 A.D., Routledge, London, pp. 517-548. Thompson, L.G., Thompson, E.M., Davis M.E., et al. 1995, A 1000 year climate ice core record from the Guliya ice cap, China: its relationship to global climate variability. Ann. Glacio. 21,175-181. Thompson, L.G., Mosley-Thompson, E.and Amao, B.M., 1984, El Niño -southjern oscillation events recorded in the stratigraphy of the Quelccaya ice cape, Peru. Science,226,50-53. Thompson, L.G., Yao, T., Davis M.E., et al., 1997, Tropical climate instability: the Last Glacial Cycle from a Qinghai-Tibetan ice core. Science, 276,1821-1825. Walker, G.T., 1923, Correlation in seasonal variations of weather. Part VIII: A preliminary study of world weather. Memoirs of the Indian Meteorological Department, 24, pp. 75131. Walker, G.T., and Bliss, E.W., 1932, World weather V. Memoirs of the Royal Meterological Society, 4, pp. 53-84. Wang, S.W. and Zhao, Z.C., 1981, Droughts and floods in China, 1470-1979. Climate and history, Wigley et al , eds., Cambridge University Press,271-288. Wang,W.C., Li, K.R., 1990, Precipitation fluctuation over semiarid region in northern China and the relationship with El Niño /Southern Oscillation, Am. Meteo. Soc., 30, 769 - 783. Whetton, P., Allan, R., and Rutherfurd, I., 1996, Historical ENSO teleconnections in the eastern hemisphere: comparison with latest El Niño series of Quinn. Climatic change 32:103-109. Whetton, P. and Rutherfurd, I., 1994, Historical ENSO telecoonections in the eastern hemisphere. Climatic change, 28,221-253. Willans, I.M., 1979, Ice flow along the Byrd station strain network, Antarctic, J. Glaciol, 24:15. Yang, M, Yao T, He Y and Thompson, L.G., 2000, ENSO events recorded in the Guliya ice core. Climatic Change, 47: 401-409. Yao T.D. and Thompson, L.G., 1992, Trends and features of climatic changes in the past 5000 years by the Dunde ice core, Ann. Glaciol., 16:21. Yao Tandong, Jiao Keqing, Yang Zhihong, 1995b, The climate variation since Little Ice Age recorded in Guliya ice core. Science in China ( Series B), 25(10):1108-1114. Yao, T., Thompson, L.G., and Jiao, K., 1995a. Recent warming as recorded in the QinghaiTibetan cryosphere. Annals of Glaciology, 21,196-200. Yao, T., Thompson, L.G., Shi, Y, et al., 1997, Climate variation since the Last Interglaciation recorded in the Guliya ice core. Science in China, 40(6):662-668. Yao, T.D., Jiao, K.Q., Li Z.Q. et al., 1992, Glaciological study on Guliya ice Cap, J. glaciology and Geocryology, 14(3):233. Yao, T, Jiao K, Yang M., 2000, Precipitation variations recorded in Guliya ice core int the past 400 years. Progress in Natural Science, 40(4): 288-293 Yao, T.D., Jiao, K.Q., Li, Z.Q. et al., 1994, Climatic and environmental records in Guliya Ice Cap, Science in China (B), 24(7):766. Yao, T.D., Pu, J.C., Liu, J.S. et al., 1993, Climatology study on Tanggula ice core. In: The glacial Climate and Environment in Qinghai-Tibet Plateau. (eds. Yao, T.D. and Ageta, Y.), Beijing: Science Publication, 16-22.

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Yao, T.D., Xie Z.C., Wu, X.L., 1990, The climatic records since Little Ice Age in Dunde ice core, Science in China (B), 11:1196. Yao,T., Thompson, L.G., Qin, D., et al., 1996, Variation in temperature and precipitation in the past 2000a on the Xizang (Tibet) Plateau----Guliya ice core record. Science in China (Series D), 39(4):425-433.

EVIDENCE OF ABRUPT CLIMATE CHANGE AND THE DEVELOPMENT OF AN HISTORIC MERCURY DEPOSITION RECORD USING CHRONOLOGICAL REFINEMENT OF ICE CORES AT UPPER FREMONT GLACIER

Paul F. Schuster, David L. Naftz, L. DeWayne Cecil , and Jaromy R. Green

1.

INTRODUCTION

Paleoclimatic and paleoenvironmental ice-core records are not common from mid-latitude locations in the contiguous U.S.A. Although excellent paleo-records exist for the high latitudes (Hammer, 1980; Lyons et al., 1990, Dansgaard et al., 1993, Taylor et al., 1993a; Clausen et al., 1995, Alley et al., 1997, Johnsen et al., 1997, Jouzel et al., 1997, Mayewski et al., 1997, Taylor et al., 1997, White, J.W.C, et al., 1997, Zielinski et al., 1997), icecore records from polar regions may be considered proxy indicators of climatic and environmental change in the mid latitudes. Unlike polar ice cores which are more likely to preserve visual, chemical and isotopic stratigraphy with sub-annual resolution, visual stratigraphy and sub annual isotopic resolution are generally not apparent in mid-latitude ice cores. In addition, meltwater percolation can influence chemical and isotopic stratigraphy of alpine glaciers from mid-latitude ice cores by “damping” the environmental signal (Wagenbach, 1989). Despite these problems, Naftz (1993), Naftz et al., (1994), Naftz et al. (1996) and Cecil and Vogt (1997), through chemical lines of evidence, indicated that the Upper Fremont Glacier (UFG) in the Wind River Range, Wyoming, U.S.A., (43° 07’ 37” N, 181 L. D. Cecil et al. (eds.), Earth Paleoenvironments: Records Preserved in Mid- and Low-Latitude Glaciers, 181-216. © 2004 Kluwer Academic Publishers. Printed in the Netherlands.

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109° 36’ 54” W) (Figure 1) is suitable for reconstruction of isotopic and chemical paleoclimate and environmental change records. For a temperate glacier, the UFG has a combination of qualifications conducive to preserving paleoenvironmental signals. The drill-site elevation is 4100 m. Minimum, maximum and average annual air temperatures during 5 years of record were –36°C, 13°C, and –7°C respectively. Temperature profiles from snow pits, conducted on an intermittent basis on the UFG, indicated that the snow pack was typically isothermal at 0°C during the summer months. During the winter months, the snow pack was below 0°C, ranging from -7°C to -2°C. The net accumulation rate is 96 cm snow weq/yr, based on the vertical position of the 1963 tritium (3H) fallout record. The glacial surface gradient is nearly level, reducing crevassing and fracturing of the ice strata (Naftz, 1993). These characteristics reduce the potential for meltwater to alter any paleoenvironmental signal. Because the remoteness of the site limits the influence of local sources of atmospheric mercury (Hg) deposition to the UFG, the location is favorable for measuring historical regional and global deposition of Hg and other chemical constituents from the atmosphere. Naftz (1993) recovered a grasshopper leg near the base (152-m depth) of the UFG which yielded a carbon-14 (14C) age of 221± 95 years before present. Based on present-day accumulation and ablation rates, ice deposited near the bottom of the glacier formed from snow which fell between about A.D. 1716 and 1820. This chronology, albeit informative, did not provide the accuracy needed to date abrupt climate change or interpret paleoenvironmental signals on, at least, a decadal time scale during the last 270 years. Further study, direct current electrical conductivity measurements (ECM), scanning electron microscopy (SEM), and energy dispersion analysis (EDA) were performed, Hg concentrations were measured, and isotopic and chemical data were reexamined to: (1) identify volcanic fallout in the ice cores from historical volcanic eruptions, (2) support and refine previous low-resolution chronological estimates of the UFG ice cores, (3) constrain the time of the Little Ice Age (LIA) transition in alpine regions of Central North America, and (4) reconstruct an historical record of atmospheric Hg deposition. The results discussed here will, hopefully, increase awareness and urgency of recovering ice cores from mid-latitude alpine glaciers for paleoclimate studies before these records disappear or are severely compromised by meltwater effects in response to a warming climate With the addition of the ECM and the Hg data, the UFG ice core data set is the most extensive of its kind from a glacier in the contiguous U.S.

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Figure -1. Map showing the location of the 1991 and 1998 ice-core drilling sites. Each site was located at about the same altitude separated by about 220 meters

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Figure -2. Location of the Upper Fremont Glacier showing very little impact from upwind local sources of atmospheric Hg

To underscore the urgency to recover these unique records, increasing global temperatures are threatening the existence and integrity of mid- and low-latitude glaciers, which are receding rapidly. Thus, any future research involving mid-latitude glaciers must be prompt. Key to the advancement of understanding environmental change on a global scale through the interpretation of mid-latitude ice-core data is the establishment of global linkage with ice cores from other regions of the world. If recession continues at these rates, the Dinwoody Glacier, about 3 kilometers north of the UFG, will be gone in about 20 years (Marston et al, 1991). Other estimates indicate that the remaining glaciers in Glacier National Park, Montana will no longer exist in 50 to 70 years (Meier, 1998) and high alpine glaciers in the Andes of South America (i.e. Quelccaya) will be severely compromised by meltwater processes (Thompson, 1985). These irreplaceable paleoenvironmental resources may literally melt away in the near future releasing an additional and potentially large reservoir of Hg and other contaminents trapped in snow and ice to the environment.

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185

METHODS

Continuous ice cores were collected from the UFG in 1991 and again in 1998. To prevent melting, the core tubes were stored at 0°C in snow vaults on the glacier. After the ice cores were removed from the aluminum core barrel, one-meter sections were quickly sealed in polyethylene bags and placed in core tubes by personnel wearing Tyvek1 suits and powder-free Latex1 gloves. During the entire process, “clean hands” protocol was used (USEPA, 1996). Immediately after drilling was complete, the cores were transported to a freezer truck via a 10-minute helicopter flight to freezer storage. All ice-core sectioning was performed in the National Ice Core Lab (NICL) exam room in Denver, Colorado. The cores were equilibrated to exam room temperature (-24° C) for 24 hours prior to processing. Powderfree polyethylene gloves were used at all times when handling the core.

2.1

Direct current electrical conductivity measurements (ECM)

Ice core sections, cut lengthwise, were placed in a core tray flat side up and secured. A clean surface (2.5 cm wide by 1 cm deep) was routed into the ice and ECM were taken for the entire length of the core (Hammer,1980; Taylor et al., 1992). A pair of ECM electrodes (1 cm apart) were drawn down the ice at a constant velocity (5 cm/sec). The electrodes conducted a current of 2000 volts into the ice that was measured every millimeter. Breaks in the core were recorded with an on/off toggle switch as “unusable” data. Once an ECM profile was recorded from a core section and the greatest signal to noise ratio was achieved, the data were reduced and processed to generate a continuous profile.

2.2

Mercury

All ice-core sectioning took place in the clean –24°C environment of NICL. The cores, totaling 160 meters in length, were cut into 7-cm sections with a stainless steel band saw cleaned with methanol. A total of 57 samples and 40 samples were removed from the 1991 and 1998 cores, respectively. The sections were placed in Hg-free clean polyethylene bags and shipped frozen on dry ice overnight to the USGS’s Wisconsin District Mercury Research Lab (WDMRL) in Middleton, Wisconsin. Temperature recorders were placed in the shipping containers and the temperature inside the containers during shipment did not exceed 0° C.

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At the WDMRL, ice samples were removed from the bags with gloved hands and rinsed with about 50 ml of WDML de-ionized water (total Hg concentration of 0.1 ng/L) to remove any potential contamination from field procedures. After rinsing, the samples were placed in Hg-clean (USEPA, 1999) Teflon1 jars. Four milliliters of bromine monochloride (BrCl) were added to oxidize all species of Hg to Hg(II), the jars were capped, and the samples allowed to thaw at room temperature. Samples ranged in volume from 25 to 75 ml because of demands for ice to address other research studies, paleoclimatic and interests (i.e. chlorine-36 (36Cl) paleoenvironmental studies). Fifty ml of Hg-free rinse water was used to remove the potentially contaminated outer core layer. Once thawed, the liquid was transferred to Hg-clean Teflon bottles that were placed in an oven at 50°C overnight to ensure complete oxidation of all mercury species. Analysis for total Hg was performed with Dual Amalgamation Cold Vapor Atomic Fluorescence Spectrometry (USEPA, 1999) with a method detection limit of 0.04 ng/L (USEPA, 1990).

2.3

SEM and EDA

Thirty subsamples of clearly visible dust layers were cut from the ice core using a conventional band saw that had been cleaned with methanol. These samples were placed immediately in 6-mil poly bags and sealed. They were melted at room temperature inside the poly bags. Each bag was moved to a Class 100 laminar flow hood, opened, and poured into a precleaned Teflon1 filter holder. Each sample was filtered through cellulose acetate filter paper with an effective pore size of 8 µm. The particulates were collected on the filter paper surface, placed inside a plastic container and dried at room temperature. A small section (approximately 1 cm by 1 cm) was cut and mounted with carbon tape and sprayed with carbon dust for analysis on a JSM-5800LV SEM. Spectrums were obtained from particles of interest using EDA.

2.4

Sulfate and chloride

Concentrations of sulfate (SO42-) and chloride (Cl-) were determined by ion exchange chromatography (Fishman and Friedman, 1989). Because of the very low ionic strength of the ice samples, the method was modified by increasing the sample-loop size and using ultrapure materials for standard preparation. A complete data set of major ion chemistry and quality assurance for the 1997 UFG ice core is given in Naftz (1993).

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Quality assurance/quality control 2.5.1 ECM

A 1-m meter core section was randomly chosen and the ECM was performed in replicate. A bivariate scatter plot (Figure 3) showed an approximate 1:1 correlation with an R2 of 0.988 indicating high instrument precision and reproducibility. Additionally, deionized-distilled water (18 MΩ) was frozen to -24° C in the form of a 0.25-m ice core and prepared for ECM using the same methods employed for the UFG ice core. The mean value of the deionized-distilled water ECM profile was an order of magnitude lower than the mean ECM value and two to three orders of magnitude lower than any significant ECM peaks measured on the entire length of the UFG ice core profile. 2.5.2 Hg Typically, ice cores are recovered with electro-mechanical drills and deep polar cores require the use of liquid lubricants such as fuel oil or antifreeze to keep the drill hole open. The UFG cores, located about 220 meters apart roughly along the same contour elevation, were recovered with a 7.6-cm diameter thermally heated aluminum coring device (Naftz, 1993). The thermal drilling process excludes the use of lubricants, reducing the potential for Hg contamination. The drill winch cable was a Kevlar1 braid protected by an outer Nylon1 braid. All bolts on the core barrel were stainless steel and silver was a major component of the thermal blade. The metals exposed to the ice core surface during recovery and processing were composed of aluminum and stainless steel. To assess possible Hg contamination from these metals, "veneer experiments" similar to those reported for polar ice (Boutron et al., 1994) were performed on a 20-cm piece of archived UFG ice (circa 1855) from a depth of 103 meters. Four successive ~1 cm layers (i.e. concentric rings) were scraped from the sample using a titanium blade onto a Teflon working platform inside a laminar-flow hood. Powder-free Latex1 gloves were worn throughout the procedure (USEPA, 1996). The scrapings were collected into acid-boiled 250-ml Teflon1 bottles and allowed to melt at room temperature. The melted samples, ranging in volume from 51 to 71 ml, were analysed for Hg using internationally adopted and proven analytical methods (USEPA, 1999). Rinse water from the titanium blade was measured at 0.35 ng/L. Selected metal concentrations (including rare earths) were measured by ICP-MS. Mercury concentrations from the veneer experiment ranged from 9.3 to 11.2 ng/L with no obvious trend in those data, and the average concentration was

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Evidence of Abrupt Climate Change

within 6 percent of the composite sample taken from the same horizon for the development of the Hg profile (Figure 4). The uniform concentrations through the thickness of this ice core sample suggest two possibilities; 1) the source of Hg is from atmospheric deposition and represents an uncontaminated signal, or 2) a Hg source from the core barrel has penetrated the entire thickness of the core. The latter is unlikely for three reasons; 1) Hg(0) is sparingly soluble and, with respect to other solutes, the solubility of Hg(II) and Hg(0) is low, reducing movement in ice, a solid-phase exchange process would be required, 2) the significant variations in Hg concentrations observed throughout the length of the core would be masked or dampened by a constant source of contamination, and most significantly, 3) if a contamination source existed, concentrations of these constituents would decrease from outer to inner core (Boutron, et al., 1994); this trend was not observed in the UFG ice core.

Figure -3. Bivariate scatter plot of a randomly chosen 1-m core section. Electrical Conductivity Measurements (ECM) were run in replicate.

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Figure -4. Concentrations of mercury from “veneer experiments” showing no trends into the ice core. Each consecutive radius is an approximate 1-cm layer of ice scraped from the core using a titanium blade.

Aluminum and zinc, two major components of the core barrel and saw blade, also showed no decreasing trends from the outer layers to the ice core center (Figure 5). Moreover, silver, a major component of the thermal blade, and chromium, a component in stainless steel, were not detected. Aluminum did not correlate with the rare earth elements La, Ce, and Nd in the outer most layer as a function of radius from the center suggesting that a fraction of this aluminum is from the core barrel (Figure 6). The next five layers to the center of the core, however, show aluminum is correlated to these rare earth elements, indicating the source is natural or crustal earth. Although the potential for Hg contamination exists, based on these results, removal of the outer layer of the ice core samples greatly reduces the potential for Hg contamination from recovery and processing techniques used for the UGF ice cores.

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Figure -5. Concentrations of aluminum and zinc from “veneer experiments”. Each consecutive radius is an approximate 1-cm layer of ice scraped from the core using a titanium blade.

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Figure -6. Concentrations of aluminum and selected rare earth elements showing a correlation of rare earth elements to aluminum, suggesting the source of aluminum is natural or crustal.

Quality control (QC) check samples were analyzed at the beginning of the run, at least every tenth sample, and at the end of the run to establish daily statistical control. QC checks were prepared with WDMRL de-ionized water and a known amount of Hg standard from a source other than that used for standardization. The QC standards measure any possible instrument drift and provide an external check on the accuracy of the calibration standards. Four jar blanks (process blanks) were run during the period of analysis. A jar blank consisted of brominated de-ionized water that was allowed to sit in a clean jar for the time it took for the ice to melt, then transferred into a 1 Teflon bottle and treated like a sample. Results from the jar blanks were used to determine the contribution of Hg from the oxidant BrCl and any Hg sources from the jars. The blanks ranged in concentration from 0.30 to 0.86 ng/L (mean = 0.66, standard deviation = 0.25, n=4). After blank subtraction of the mean blank value, the lowest total Hg concentration from 97 samples was 1.21 ng/L, which is still significantly above the highest blank value. All samples were analyzed in duplicate. If the percent difference between the two duplicates was greater than 10 percent, the sample was analyzed a third time. In all cases, the relative standard deviation between the 3 replicates was less than 10 percent. On each workday, at least one sample was spiked and the percent recoveries ranged from 90 to 111 (mean = 99, standard deviation = 6, n= 15).

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3.

RESULTS AND DISCUSSION

3.1

Chronology of the UFG

Annual dust layers from the UFG were not always visible, thereby making visual age-dating methods unreliable. In the absence of visually identifiable annual dust layers in the UFG ice core, other estimates of age must be employed. This research employs chemical age-dating techniques as a means of determining the chronology of the UFG ice core. Naftz (1993) proposed an age-depth profile for the UFG ice core. However, a large error is associated with this profile. Initial analysis of the ice core by Naftz (1993) indicated the bottom of the core was approximately 250 years old based on a 14C date for a grasshopper leg found deep in the core. However, the possibility there could have been post-depositional movement of the material within the glacier at a depth where the leg was found and the ± 95 year error associated with the 14C date pointed to a need to obtain other information about the age of the ice within the core at different depths. ECM and a reexamination of sulfate and chloride concentrations provided additional means of determining the age of the core. The ECM is a direct measurement of the acidity of the ice (Hammer, 1980; Taylor et al., 1992). Volcanic eruptions and other sources of acid, including anthropogenic, increase the acidity of precipitation, resulting in increased ECM of the ice. In addition, increased sulfate and chloride concentrations in the ice are common indicators of fallout from a volcanic eruption (Hammer, 1980). Strong ECM signals coupled with sulfate and chloride concentration peaks in the ice core profile suggest that the snow falling on the glacier surface shortly after the volcanic eruption contained volcanic fallout.

3.2

ECM, sulfate, chloride, and oxygen-18 profiles

The ECM, SO42-, and Cl- from the UFG ice core are shown in Figure 7. Timeline events are noted to the right of the figure and will be discussed later. Also included are the ice-core depths and profile ages (A.D.). Approximately, the top 15 m of core are firn, (a permeable aggregate of small grains of ice with a density less than 0.55 g/cm3, Naftz, [1993]) which produces a relatively low ECM signal due to lack of conductivity through interstitial pore spaces. The first 10 m of firn chemistry was not included in the profiles because it is not the focus of this research. Below the firn/ice transition (about 15 m), 95 percent of all significant Cl- peaks (>0.05 mg/l) and SO42- peaks (>0.01 mg/l) were coincident with significant ECM peaks (>4.0×10-9 S/cm). From approximately 15 to 108 m, ECM amplitude and

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variance are relatively high, effectively obscuring all but the strongest peaks. From about 108 to about 150 m, the amplitude and variance decreases significantly and ECM peaks are easily visible. The ECM profile was compared to the delta oxygen-18 (δ18O) profile measured by Naftz (1993) (Figure 8). Naftz’s work showed an abrupt shift in the amplitude and an increase in the variance of the δ18O signal, exceeding values expected from lateral variability (Naftz, 1994) at a depth coincident with the decrease in ECM signal amplitude (Figure 8c). The bottom 10 m of the ice core differ significantly in chemical and isotopic composition, suggesting post-depositional alteration (Figure 7a, 8a). Based on visible dust layers, ECM, SO42- and Cl- profiles, thirty suspected volcanic fallout layers were isolated and analyzed by SEM and EDA. With the exception of one layer at about 71 m, material resembling volcanic fallout has not been currently identified. The layer at 71 m, however, contains particles that resemble volcanic glass shards or pumice. The EDA’s suggest these particles are pumiceous but these data are only qualitative and inconclusive. According to the proposed chronology, global volcanic activity was highly active during this period (White et al., 1997). Microprobe analyses and a newly developed particle laser counting technique to identify volcanic fallout in the UFG ice core are the focus of continuing research. 3.2.1 Factors influencing volcanic fallout to a glacier surface The atmospheric transport of volcanic materials is a complex process complicated by a number of factors. The explosivity of the event, meteorological conditions and seasonal timing, and the location of the eruption relative to the glacier all lead to varying transport histories. For example, material from strong events in the Southern Hemisphere may not be deposited on a Northern Hemisphere glacier surface for up to two years, thus complicating the investigated chronological history of the ice (White et al., 1997).

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Figure -7. Measurement profiles of chloride (a), electrical conductivity measurement (ECM) (b), and sulfate (c) of the Upper Fremont Glacier ice core showing isotopic timeline events and volcanic events speculated by significant ECM peaks and coincident sulfate and chloride peaks.

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Figure -8. Profiles of δ18O (a), ECM (b) and a 999-point running average of ECM data (c) showing a significant decrease in variance at a depth of about 108 meters.

The Volcanic Explosivity Index (VEI) proposed by Newhall and Self (1982) ranks volcanic eruptions from 1 to 8 based on the magnitude, intensity, dispersive power, and destructiveness of the eruption. It does not classify explosivity based on the composition of the eruptive plume. Events such as Krakatau and Tambora which are believed to have injected material into the stratosphere are assigned values of 4 or greater. It is these eruptions plus 11 select VEI 3 events and 4 select VEI 2 events that have been used to refine the chronology of the UFG ice cores (Figure 9, Table 1). Complicating the identification of volcanic events in an ice core by ECM, SO42- and/or Cl- peaks is that the strongest events (VEI 6-8) tend to be silicic and generally produce large amounts of fine ash but are low in sulfur and chlorine (Rampino and Self, 1982). Eruptions which are basaltic in

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composition are generally less explosive, but are more sulfur and chlorine rich. Mildly explosive, but sulfur and chlorine rich events may express themselves in the ECM and chemical ice core data disproportionate to their VEI ranking compared to the more silicic but explosive VEI 6-8 events. Complicating the transport histories of volcanic eruptions even further are the meteorological conditions and seasonal timing of a volcanic event. Chemical and aerosol species must reach the stratosphere for circumpolar, and especially inter-hemispheric, transport. Once these species reach the stratosphere they are eventually brought back into the troposphere due to gravity and the effects of the springtime tropopause fold, thus subjecting them to scavenging and both wet and dry deposition (White et al., 1997). The location of the eruption relative to the glacier also plays an important part in the amount of volcanic fallout material accumulating on the glacial surface. Distant eruptions, which may be sulfur and chlorine rich, but do not inject those aerosols into the stratosphere, are unlikely candidates for identification in the ice. Distant eruptions which do inject material into the stratosphere may be weakly represented in the ice due to long range dispersion of the chemical species. Eruptions closer to the sample site which are low in sulfur and chlorine and only inject material into the troposphere may be over represented. Therefore, identification of geographically near, yet weak (
3.3

Post depositional chemical alterations

Mid-latitude glaciers are subject to meltwater processes during the summer months (Wagenbach, 1989). Periodic exposure of the snow surface and firn to meltwater percolation during exceptionally warm summers can deplete particulates and ions resulting in a damped chemical signal in the ice. A chemical elution sequence (SO42->NO3->NH4+>K+>Ca+2> Mg+2>H3O+>Na+>Cl- ) is observed in glacial ice (Brimblecombe et al., 1985). The elution sequence describes the order in which chemical constituents are removed from the ice by meltwater processes. This same elution sequence has been observed in the ice cores from the Wind River Range in Wyoming (Naftz, 1993).

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Table -1. Major volcanic events with potential for global volcanic fallout to the Upper Fremont Glacier surface (modified from White, 1997). Ref # for Fig. 7 1 2 3 4 5 6 6 7 8 9 9 9 10 11 12 12 13 14 15 16 16 16 17 18 18 19 20 21 22 22 23 24 25 26 27 28 29 30 30 30 31

Name of Volcanic Event

Location

Mt. St. Helens Washington Westdhal Alaska Hekla Iceland Fernandina Galapagos Is. Scheveluch Kamchatka Surtsey Iceland Mt. Agung Bali Bezymianny Kamchatka Hekla Iceland Mt. Vesuvius Italy Toya Japan Paricutin Mexico Marapi Java Mayon Philippines Raikoke Ryukyu Is. Iriomote-Jima Japan Katla Iceland Katmai Alaska Ksudach Kamchatka Santa Maria Guatemala Mt. Pelee Martinique Soufriere St. Vincent Mayon Philippines Mayon Philippines Bogoslof Alaska Krakatau Indonesia Cotopaxi Ecuador Scheveluch Kamchatka Chikurachki Kurile Is. Mayon Philippines Babuyan Claro Philippines Kliuchevskoi Kamchatka Galung Gung Java Okmok Alaska Tambora Indonesia Unknown -Miyi-Yama Java Skaptarjokull Iceland Asama Japan Lakigigar Iceland Medvezhii Kurile Is.

Lat/Long 46.02/122.18W 54.52N/164.65W 63.98N/19.70W 0.37S/91.55W 56.78N/161.58E 63.30N/20.62W 8.342S/115.51E 56.07N/160.72E 63.98N/19.70W 40.82N/14.43E NA*/ NA 19.48N/102.25W 7.54S/110.44E 13.26N/123.68E 48.25N/153.25E 27.88N/128.25E 63.63N/19.03W 58.27N/155.16W 51.83N/157.52E 14.76N/91.55W 14.82N/61.17W 13.33N/61.18W 13.26N/123.68E 13.26N/123.68E 53.93N/168.03W 6.10S/105.42E 0.65S/78.43W 56.78N/161.58E 50.33N/155.46E 13.26N/123.68E 19.50N/121.95E 56.18N/160.78E 7.25S/108.05E 53.42N/168.13W 118.00E NA NA/ NA NA/ NA 36.40N/138.53E 64.08N/18.25W 45.38N/148.80E

VEI† Year AgeɁ 5 3 4 4 4 3 4 5 4 3 NA 3 2 3 4 2 4 6 5 6 4 4 3 3 3 6 4 5 4 3 4 4 5 3 7 NA NA NA 4 4 2

1980 1978 1970 1968 1964 1963 1963 1956 1947 1944 1944 1944 1931 1928 1924 1924 1918 1912 1907 1902 1902 1902 1897 1888 1888 1883 1877 1854 1853 1853 1831 1829 1822 1817 1815 1809 1793 1783 1783 1783 1778

11 13 21 23 27 28 28 35 44 46 46 46 60 63 67 67 73 75 84 89 89 89 94 103 103 108 114 136 137 137 160 162 169 174 176 182 198 208 208 208 213

Peaks (ECM SO4, Cl)

Sulfate Peak in GISP2†

Y,Y,Y Y,Y,Y Y,Y,N Y,Y,N Y,Y,N Y,Y,N Y,Y,N Y,Y,Y Y,Y,N Y,Y,Y Y,Y,N Y,Y,N Y,N,N Y,N,N Y,Y,N Y,Y,N Y,Y,Y Y,Y,N Y,Y,N Y,Y,Y Y,Y,Y Y,Y,Y Y,N,N Y,Y,N Y,Y,N Y,Y,Y N,Y,N Y,N,N Y,N,N Y,N,N Y,Y,N N,Y,N Y,Y,N Y,Y,N Y,Y,Y Y,N,N Y,N,N Y,N,N Y,N,N Y,N,N Y,N,N

? Y Y N Y N N Y Y N N N N N Y N Y Y Y Y Y Y N N N Y N Y Y N Y Y N N Y Y N N N Y N

Depth in Fremont ice core (m) 16 17 22 22.5 26 30 30 41 52 54 54 54 65 67 71 71 74 75 76 78 78 78 83 84 84 88 90 100 101 101 113 114 118 122 123 126 127 130 130 130 131

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Table -2 continued. Major volcanic events with potential for global volcanic fallout to the Upper Fremont Glacier surface (modified from White, 1997). Ref # Name of for Location Volcanic Fig. Event 7 32 Opala Kamchatka 33 Hekla Iceland 34 Tarumai Japan

Lat/Long 52.43N/157.45E 63.98N/19.70W 42.68N/141.38E

†

Ɂ

VEI

Year

Age

2 4 5

1776 1765 1739

215 226 252

Sulfate Peaks (ECM Peak in SO4, Cl) GISP2† Y,N,N Y,N,N Y,N,N

N Y N

Depth in Fremont ice core (m) 132 138 144

*NA, data not available †VEI, Volcanic Explosivity Index Ɂ age is given in years BP (before the present date of 1991) TTGISP2, Greenland Ice Sheet Project 2

Therefore, any chemical signal preserved in the ice, albeit possibly shifted, may have potentially been much larger during deposition and subsequently eluted (damped). Past research, (Naftz, 1993; Naftz et al., 1996), and (Cecil and Vogt, 1997) have shown that the stratigraphy of the ice is generally preserved. An example of post-depositional alteration in the UFG ice core may be the deposition of the Mt. St. Helens volcanic ash layer. A series of very large ECM peaks were recorded at a depth of about 15 m. Based on present-day accumulation rates and isotopic age dates in the upper 30 m of ice, the Mt. St. Helens ash layer should be near a depth of 15 m. Although the Mt. St. Helens ash layer was identified by SEM and EDA in Knife Point Glacier about 4 miles to the southeast of Fremont Glacier (Naftz, 1993), no ash layer was visible near that depth in the UFG ice core, suggesting meltwater processes were conducive to removing particulates from the snow surface but leaving behind an acid signal detected by ECM. Another example of post-depositional alteration may be the abrupt change in chemical and isotopic composition in the bottom 10 m of the ice core. Chloride concentrations at a depth of 150-155 m exceeded the highest Clconcentrations throughout the rest of the core by greater than 2 times (Naftz, 1993). The elevated Cl- concentration and lack of ECM signal in this section very close to bedrock suggest the source of Cl- is not atmospheric. It is likely that the processes responsible for these chemical and isotopic changes are post-depositional and are due to chemical interactions at the ice-bedrock interface.

3.4

Chronological refinement

Identifying fallout from a specific volcanic event in an ice core must be based on some chronological reference point already verified. SEM and EDA results were inconclusive. Particles suspected to be volcanic ash

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fallout were not large enough or in sufficient quantity for analysis and identification by microprobe. Therefore, chemical age dates determined from 3H, 36Cl and 14C (Naftz, 1993, Cecil and Vogt, 1997) were first used to generate a “preliminary” age-depth profile of the UFG ice core. The profile predicted an age of 1881 A.D. at a depth of 88 m where the largest ECM peak in the core occurred. The two largest eruptions in recorded history, Krakatau (1883 A.D.) and Tambora (1815 A.D.), albeit in the southern hemisphere 20,000 km from the UFG, reached well into the stratosphere with global effects. The ship, Medea, measured the Krakatau eruption column height to be up to 26 km (Self, 1992). The Tambora event, perhaps the largest eruption in the last 10,000 years, injected volcanic material to a height estimated to be as high as 44 km into the stratosphere (Sigurdsson and Carey, 1987). Historical observations of remarkable sunsets in Europe, North America and Hawaii and optical effects for up to two years after each eruption were another indication that the dust columns reached the stratosphere (Lamb, 1970). Fallout from the Tambora and Krakatau events has been identified in Antarctic and Greenland ice cores (Delmas et al., 1992; Kohno et al., 1999). The similarity between the preliminary age-depths of the two largest ECM peaks (based on chemical age dates) and the ages the Krakatau and Tambora events indicate these ECM peaks and associated SO42- and Clpeaks resulted from atmospheric fallout of these eruptions. No other volcanic events within ± 20 years of these eruptions were comparable in magnitude (Table 1). Table 1 lists the significant volcanic events in the past 250 years most likely to have deposited volcanic fallout to the glacier surface, thus increasing the ECM, SO42- and Cl- signals. The criteria used to select these volcanic events include VEI, proximity to UFG, and correlation of SO42peaks in the UFG core to core from the Greenland Ice-Sheet Project 2 (GISP2) (White, 1997; Zeilinski et al., 1997). Using the known age dates based on data from 3H, 36Cl, 14C and the Krakatau and Tambora events, the remaining volcanic events listed in Table 1 were assigned to ECM peaks (which generally were coincident with SO42- and Cl- peaks) in the ice core and listed as additional timeline events in the right margin of Figure 7. A refined age-depth profile was fitted to the data and plotted in Figure 9. Because of the uncertainty associated with the 14C date (221 ± 95 years), the last 7 volcanic events (prior to Tambora, 1815) and the 14C age date were not used to generate the refined age-depth profile. The polynomial fit for the refined age-depth profile is as follows: Age in years = 0.00739 (D)2 + 0.5558(D)

(1)

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The refined age-depth profile, where D is depth in meters, yielded an age of 1885 A.D. for the 1883 A.D. Krakatau event and 1811 A.D for the 1815 A.D Tambora event. At a depth of 152 m, the refined profile shows good agreement (1736 A.D.) with the 14C age date (1729 A.D. ± 95 years). The addition of volcanic time events to isotopic and chemical time events, supports and refines the age-depth profile (Equation 1) of the UFG ice core developed in previous work (Naftz, 1993; Naftz et al., 1996).

Figure -9. Plot of reported volcanic events and isotopic age dates used to generate a polynomial fit for an age-depth profile of the Upper Fremont ice core.

3.5

Abrupt Climate Change

Visual inspection of the δ18O and ECM profiles indicated changes in variance at a depth between 95 and 110 m (Figure 9). Although it is still unclear why this relationship exists, it has been documented in other cores (Taylor, 1993b). One possible explanation is that a colder climatic period such as the LIA will typically be windier and dryer (Taylor, 1993a). In the Western U.S.A., this would increase the amount of calcium-magnesium rich wind-blown dust in the atmosphere, thus neutralizing precipitation acidity and decreasing the ECM signal as seen in the ice-core profile below 108 m (Taylor et al, 1996). Colder, drier climates aid to preserve the seasonal δ18O signal, thus increasing the δ18O variability. Warmer climatic periods are typically wetter and less windy, minimizing calcium-rich neutralizing wind-

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blown material, thus causing the ECM signal to increase in response to increased precipitation acidity. Conversely, the δ18O signal will be damped due to increased melting during the summer months (Wagenbach, 1989). The shift in δ18O signal and increase in variance in the UFG ice core is an indication of a change in climate (LIA) (Naftz et al., 1996). However, based on δ18O data alone, the time of the LIA termination is unclear. The refined age-depth profile indicates that, at this depth, 15 m of ice represents about 30 years. The sampling interval for δ18O is sufficiently large (20 cm) that it is difficult to pinpoint the LIA termination. The ECM data set contains over 125,000 data points at a resolution of 1 data point per mm of ice core. A 999-point running average was calculated and plotted on Figure 8. The plot indicates the variance of the ECM data decreases significantly at about 108 m. To determine if these increases in δ18O variance and decreases in ECM variance were coincident, a series of f-tests (Walpole and Myers, 1985) were performed at 5 depths between 80 and 110 m (Table 2). Results from running average and the f-tests indicate that the variance from high to low in the ECM data occurs at about 108 m. At this depth, the age-depth profile predicts an age of 1845 A.D. This predicted age-depth of the LIA termination also coincides with a trend of increasing radial growth of tree cores collected 2 km from the ice core site starting at about 1840 A.D. Also beginning at this time was industrialization in the Western U.S.A. It is possible, however unlikely, that the high ECM signal amplitude and variance could be due to anthropogenically derived acids produced by pollution. Below 108 m (1845 A.D.), prior to industrialization, background ECM levels decrease and ECM peaks are easily visible. These peaks are most likely due to volcanically derived sources of acid from major volcanic events. It has also been shown that sources such as fire-related biomass burning can neutralize precipitation acidity, causing a decrease in the background ECM signal (Taylor, 1996). The decrease in ECM signal below 100 m, however, is sustained over more than 100 years, thus making biomass burning an unlikely cause to the observed decrease in ECM signal. Glaciological processes could also account for changes in ECM (and δ18O) variance observed at 108 m. Specifically, ice at the bottom of the core could have been from snow that fell at a location much closer to the headwall of the cirque. It is possible, but not probable, that this snow contained a different chemical signature due to its proximity to the headwall. However, there were no significant changes in δ18O values during a snow surface grid sampling program (Naftz et al., 1994). Moreover, the headwall rocks are of igneous/metamorphic origin and would not generate significant dust input. Finally, based on SEM findings of diatoms throughout the ice core and visual observations of dragonflies on the glacier surface, the origin of particle dust is most likely from the Green River basin. This is not

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surprising, considering Fremont peak (4000 m) intersects the bottom of the high-altitude jet stream which could transport particles long distances. These processes would result in a homogeneous distribution of dust over the glacier surface. For these reasons, a change in climate, not a change in snowfall location, is suggested as the trigger for the changes seen in the ECM and δ18O variance at 108 m. Ice cores from polar regions have shown that climate can change rapidly (2-3 years) (Taylor, 1993a). Results from the UFG indicate that a rapid climate shift can also occur at mid latitudes. The termination of the LIA was abrupt, with a major climatic shift to warmer temperatures around 1845 A.D. and continuing to present day. Prediction limits (error bars) calculated for the profile ages predicted along the length of the core (Walpole and Myers, 1985) were ± 10 years (90 percent confidence level). Confidence limits were much lower, about two to three years throughout the length of the core. Thus, a conservative estimate for the time taken to complete the LIA climatic shift to present-day climate is about 10 years, suggesting the LIA termination in alpine regions of central North America may have occurred on a relatively short (decadal or less) time scale.

3.6

Mercury concentrations in the UFG

The remote location and high elevation of the UFG (Figure 1) should reduce the contributions from local anthropogenic influences of atmospheric Hg (Figure 2) (USEPA, 1997). Thus, Hg concentrations in ice cores from the UFG reflect regional and global atmospheric inputs. Pre-industrial or background Hg concentrations not influenced by volcanic activity, in the UFG ice cores (Figure 10) were similar to those found in Antarctic ice and Greenland ice (Vandal et al., 1993; Boutron et al., 1998), indicating that the ranges of Hg concentrations found in the UFG are not an artifact of contamination but rather reflect natural and anthropogenic deposition of atmospheric Hg at this latitude. Total Hg measured in 97 ice core samples spanning 160 meters provided an average Hg profile resolution of three years. The detailed chronology of the UFG cores, coupled with analytical advances in measuring trace levels of Hg, together with a 3-year profile resolution, provide a clear and direct measure of historical natural and anthropogenic contributions to atmospheric Hg deposition. Furthermore, the continuity of the Hg profile from the 1991 core to the 1998 core indicates Hg is preserved in the ice (Figure 10). By integrating the peak areas identified as separate atmospheric sources of Hg, the relative contributions of these sources were quantified. Eighteen pre-industrial (before 1840) measurements of Hg were used to extrapolate a background value (4 ng/L) through the ice-core record. Background

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concentrations contributed 42 percent of the total Hg in the ice core during its 270-year record.

Figure-10. A. Profile of historic concentrations of Hg in the Upper Fremont Glacier. A conservative concentration of 4 ng/L was estimated as pre-industrial input and extrapolated to 1993 as a background concentration. Age-depth prediction limits are ±10 years (90 percent confidence level); confidence limits are 2-3 years (Schuster et al., 2000). Inset B; Hg production during the California Gold Rush (adapted from Figure 5 in Alpers and Humerlack, 2000). Inset C; World production of Hg in tons per year during the last century. Adapted from Figure 4b in Engstrom and Swartz, 1997).

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Evidence of Abrupt Climate Change 3.6.1 Volcanic sources of mercury

Volcanic eruptions are a known atmospheric Hg source (Varecamp and Buseck., 1981; Varecamp and Buseck., 1986); however, their importance on a global scale has remained unresolved. Three distinct peaks in the ice-core Hg profile are coincident (within the chronology prediction limits of ±10 years) with increased Cl and SO42- concentrations and Electrical Conductivity Measurements (ECM) (Figure 11) (Schuster et al., 2000). These natural geologic events were point sources in terms of Hg origin but were followed by global scale deposition. The Mount St. Helens eruption (1980), although orders of magnitude smaller in scale, was only 600 km distant and directly upwind of the UFG, blanketing the region with volcanic ash (Naftz, 1993). The proximity of Mount St. Helens to the UFG qualifies the corresponding Hg peak as a regional Hg source. The peak’s superposition on elevated concentrations due to near-peak anthropogenic Hg emissions resulted in the profile’s highest measured Hg concentrations (Figure 10). Differences in Hg loads among the three volcanic peaks may have been due to differences in volcanic dust compositions as indicated by differences in Cl , SO42-, and ECM peaks. Whether the volcanic source of Hg was regional, global or altered by post-depositional processes, it is clear that these globally impacting natural events have “punctuated” the historical Hg record in the UFG and likely elsewhere. Integrating the peak areas attributed to atmospheric Hg input due to volcanic activity with regional and global impact (Figure 10), these natural atmospheric Hg sources were quantified. During the past 270 years, three major volcanic events (Tambora, Krakatau, and Mount St Helens) contributed six percent of the total Hg measured in the ice cores. It is likely, however, that six percent is an underestimate. There are three main possibilities for this underestimate: 1) there have been numerous smaller volcanic events (White et al., 1997) during the past 270 years. Some of these events undoubtedly had some global impact but the volcanic signal was likely masked by the background or anthropogenic signal, 2) only 6.7 m of a total length of 160 m of ice was sampled for Hg throughout the length of the core. The Hg signal from a volcanic source is of short duration (1-2 years). Thus, it is likely that some volcanic events were not sampled, and 3) it is also possible that elution processes (described earlier) dampened the volcanic Hg signal of the three major volcanic eruptions identified in the UFG ice core.

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Figure-11. Profiles for Hg compared to chloride, sulfate, and ECM (Electrical Conductivity Measurements). The y-axis is scaled with the age-depth relationship, thus giving the Hg profile a slightly different appearance from Figure 10. ECM is a measure of the acidity of the ice. A correlation among chloride, sulfate, ECM, and Hg is a strong indication of a volcanic source. Age-depth prediction limits are ∀10 years (90 percent confidence level); confidence limits are 2-3 years (Schuster et al., 2000). Adapted from Figure 3 in Schuster, et al., 2000).

3.6.2 Anthropogenic sources of mercury Mercury was used on a large scale to recover gold from mining operations throughout the western U.S.A. beginning around 1850. These activities peaked around 1860 and then again around 1877 (Figure 10, Inset B) (Alpers and Hunerlach, 2000). The bi-modal nature of these activities was reflected in the ice-core Hg profile, showing significant increases coincident with peak Hg production in California during this period. The age-depth prediction limit for the UFG ice cores is ±10 years, thus accounting for the slight offsets among Figure 10A and Insets B and C.

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Mercury production decreased significantly in 1884 with the introduction of legislation (The Sawyer Decision) (Alpers and Hunerlach., 2000) that greatly reduced the use of Hg for gold extraction in California. A precipitous drop in UFG ice-core Hg concentrations coincided with this period. Most sediment core studies do not indicate an increase in Hg concentrations coincident with the start of the California Gold Rush. There are some studies, however, that do record a “jump” in the sediment Hg profile circa 1850 (Heyvaert et al., 2000; Bindler et al., 2001). Nriagu (1994) explains most of the Hg would have blown west, describing this transport as a “grasshopper-like dispersal pattern”. The mercury-gold amalgamation practices during the California Gold Rush during the mid to late 1800’s were unregulated and unrivaled by any other mining activity up to that time (Alpers, 2000). During this time, unknown amounts of Hg were volatilized to the atmosphere. The depositional pattern of atmospheric Hg from this source would be, in large part, dependent on storm trajectories and jet stream patterns for which there is obviously no data for that period. Based on; 1) today’s general knowledge that it is not uncommon for storm trajectories and the jet stream to migrate north and south, 2) the UFG’s proximity to the California mining belts, and 3) the magnitude of the estimates of Hg volatilized into the atmosphere for 30 years (circa 18491884), it is suggested that the source of elevated Hg concentrations measured in the UFG ice core coincident with the same time period is Hg from California mining activities (Figure 10). If the source of these elevated Hg concentrations is from California gold-mining activities, as suggested by Figure 10, then the integration of the profile indicates that the Hg-gold amalgamation activities during the California Gold Rush contributed 13 percent of the total Hg in the 270-year ice core record. These data suggest that the California Gold Rush had a significant regional impact in terms of atmospheric Hg deposition in the western U.S.A. th At the turn of the 20 century, atmospheric Hg levels remained elevated compared to pre-industrial (before circa 1840 A.D.) or background values. Increases in anthropogenic Hg emissions during the past century have been attributed mainly to coal-burning power plants, waste incineration, and chlor-alkali plants (Hanisch,1998; Enstrom et al., 1997; Nriagu, 1988). The next significant increase in ice-core Hg concentrations coincided (within the ~10 year prediction limits) with increased global Hg production (Figure 10, Inset C) most likely in response to industrial mobilization for World War II. There was a post-WWII decline in global Hg production once again coincident with decreases in ice-core Hg concentrations. The last half of the th 20 century up until 1990 shows a consistent increase in both global Hg production and Hg concentrations in the UFG ice cores.

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Volcanic eruptions contributed to the global Hg pool for brief periods (<2 years), and thus, cannot account for the substantial increase in ice-core Hg concentrations during the last century. The volcanic inputs, albeit competitive with industrial inputs, were short in comparison to the chronic levels of elevated Hg concentrations during the last 100 years, indicating that anthropogenic inputs have had the greatest influence on the atmospheric Hg deposition record in the UFG. During the past 270 years, anthropogenic inputs contributed 52 percent of the Hg accumulation in the core. More significantly, during the last 100 years, anthropogenic sources contributed 70 percent of the total Hg input. A post-1990 decline in the ice-core Hg concentrations is discussed below. 3.6.3

Mercury deposition rates to the UFG

Historical Hg deposition rates were calculated from Hg concentrations measured in the ice cores (Table 2). The calculated rates of deposition assume an average accumulation rate of 1 m of ice to the UFG per year. Obviously, this rate varies from year to year. However, based on average measurements of accumulation and ablation rates (Naftz et al., 1996; Naftz, 1993) this estimate is not unreasonable. Moreover, up to 50 percent of seasonal snowfall accumulation is lost through ablation (Naftz, 1993). This process, although difficult to quantify, would, most likely, lead to an underestimate of Hg deposition calculated from concentrations in the ice core. There is a down-core change in the age-depth relationship due mostly to glacial flow processes leading to layer thinning with depth. Basically, the same 7-cm section of ice core sample represents more time with depth. Considering the calculation of Hg deposition rates and utilizing the agedepth relationship (Schuster et al., 2000 ), a ratio (change in age/change in depth) was calculated and applied to Hg deposition results to develop corrected Hg deposition rates using Equation 2: (Ai – Ai-1)/(Di – Di-1)

(2)

where A is the calculated age in years (Cecil et al., 1997), D is the ice-core depth (m), and i denotes the sequential Hg sample (1 to 97). At the base of the core (the 97th Hg sample), the ratio is 2.88. Thus, at this depth, one meter of ice represents approximately 2.88 years. Equation 1 was applied to Hg deposition rates as a correction factor to compensate for down-core changes of the age-depth relationship (due to thinning) on each 7-cm sample.

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Table -3. Mercury (Hg) deposition measured among 3 sample media. Site Sample Episode Year (s) Average Hg Deposition* Change From Media (A.D.) (ng/L) (µg/m2/yr) Pre-industrial (Fold) UFG ice Clean Air Act 1986-93 9 11.4 11 UFG ice Industrial Max. 1984 20 20.3 20 UFG ice Mt. St. Helens 1980 11† 12.7§ 12 UFG ice Industrial 1900-93 10 11.0 11 UFG ice WWII 1938-46 7 4.73 5 UFG ice Krakatau 1883 21‡ 18.2§ 18 UFG ice Gold Rush 1850-78 8 4.84 5 UFG ice Tambora 1815 10‡ 8.60§ 8 UFG ice Pre-industrial 1719-1847 3 0.78 NAŒ Reference Minnesota wet ppt¶ 26 1997-99 14 6.99 7 Colorado wet ppt 26 1999 10 9.20 9 <1880 NA 80.0 NA Minnesota lake sed 28 >1880 NA 170 2 <1850 NA 3.70 NA Minnesota lake sed 25 modern NA 12.5 3 <1750 NA 2.00 NA Arctic lake sed 37 1980 NA 12.5 6 <1850 NA 5.00 NA New York lake sed 35 modern NA 8.90 2 <1850 NA 7.60 NA California lake sed 36 >1980 NA 38.0 5 *Deposition calculated using age-depth correction factor †Pre-industrial and industrial inputs subtracted to isolate volcanic signal; maximum input reported ‡Pre-industrial input subtracted to isolate volcanic signal; maximum input reported § Age-depth correction factor not used to calculate deposition rate Œ Not Applicable or Not Available ¶ wet precipitation **Sediment # Change measured from "pre-industrial" dated cores from cited study

The ratio attained 1 at about 35 meters of depth. It is assumed there is no change in the age-depth relationship (due to thinning) from 35 meters to the top of the core (one meter of ice ~ one year). The residence time of Hg(0) in the atmosphere is on the order of a year (Morel et al., 1998). Thus, deposition from volcanic sources represents, at most, a one-year period. Therefore, volcanic deposition values were calculated and reported without age-depth correction factors. Also, pre-industrial (background) and industrial inputs were subtracted from the calculated volcanic deposition to isolate the volcanic signal. Using maximum input from each volcanic event and the conditions described above and deposition rates from the volcanic events

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identified here, an 8 to 18-fold increase in Hg deposition over background due to globally impacting volcanic activity is indicated. Again, assuming an accumulation rate of one meter of ice per year (Naftz, 1993) at the top of the ice core and accounting for changes in the agedepth relation down-core (Schuster et al., 2000), there was a 20-fold increase from pre-industrial times to an “industrial maximum” circa 1984. During the last century, the average increase due to industrialization was 11 fold. Analysis of sediment cores from lakes (Heyvaert et al., 2000; Swain et al., 1992, Lorey et al., 1999; Meger, 1986) and precipitation (NADP/MDN, 2001) also indicate increases in atmospheric Hg deposition (2 to 9 fold) since the 1700’s. The increase in Hg deposition rates from pre-industrial times to the mid-1980s, as indicated by the ice cores, are up to ten times higher than increases determined from sediment cores and precipitation. Recent work indicates that ice-core response to changes in global atmospheric cycling masses and deposition may be affected by postdepositional processes. A recent study indicates Hg in snow packs is susceptible to reemission due to photochemical redox reactions, resulting in reductions of Hg levels by 54 percent within 24 hours after deposition (LaLonde et al., 2002). If this process does occur at the UFG, the estimated Hg deposition rates calculated from the UGF ice cores could be underestimated by as much as one half. On the other hand, recent work has also shown that mercury deposition may be affected by altitude, resulting in increases in atmospheric Hg deposition. Although the mechanisms are unclear, the work concluded that there is a positive relationship of altitude to Hg loading in snow. Work in the Wasatch and Teton ranges near the UFG indicate that annual Hg accumulation rates increase from 100 to 175 percent with an elevation gain of 1000 meters (Susong et al., 1999). In addition, recent work on Denali (Mt. McKinley) in Alaska (Krabbenhoft, to be submitted for publication) showed a 30 to 75-fold increase in Hg concentrations in the surface snow with an elevation gain of about 5500 meters; the ice-core site on the UFG is at an elevation of 4100 meters (Figure 1). It appears the altitude effect is much larger than the reemission processes indicated by LaLonde et al (2002). This may by why there are measurable and distinct volcanic and anthropogenic Hg signals in the UFG ice cores and why this profile differs greatly from those found in sediment cores. The nearly 50 percent decline in mercury accumulation at the top of the ice core compares very favorably in magnitude with independent estimates of recent global declines of mercury production and use (Engstrom et al., 1997, Pacyna et al., 2001). Lake sediments, on the other hand, retain only a small fraction of the total Hg deposition, and the remainder is generally recycled back to the lake (Hurley et al., 1994). Moreover, uncertainties such as sediment focusing associated with using sediment cores

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to estimate accumulation rates prevent simple comparisons of the two methods. 3.6.4 Estimation of ‘global impact’ volcanic mercury deposition An estimated 21 km of volcanic material was ejected during the 1883 Krakatau eruption (Self et al., 1981). The 1815 Tambora event produced a bulk volume of approximately 150 km3 of pumice and ash (Stothers, 1984). Assuming the ejecta and gases reached the stratosphere and were distributed evenly over the earth’s hemisphere (Rampino, 1982), an estimation of the atmospheric deposition attributed to these globally impacting volcanic events can be calculated with the following equation: 3

Hgvol (:g/m2) = (Veje (cm3) x ∆plu (g/cm3) x 1/∆str (g/m3) x Hgplu (:g/m3) )/Ahem (m2)

(3)

where Hgvol is the atmospheric deposition from a globally impacting volcanic eruption, Veje is the volume of volcanic ejecta, ∆plu is the density of the volcanic plume, ∆str is the density of air at 5000 m elevation (pressure ~ 0.4 atmospheres, and average air temperature in the volcanic plume is ~ -20ΟC), Hgplu is the concentration of Hg in the plume, and Ahem is the area of the earth’s hemisphere. Based on previous work (Lepel et al., 1978; Siegel, et al., 1978; Phelen et al., 1982; Unni et al., 1978; Fruchter et al., 1980), the concentration of Hg in an atmospheric volcanic plume or volcanic fumarolic gases can range 3 from 1 to > 7000 g/m . For the sake of argument, a conservative value of 48 3 3 g/m (Unni et al., 1978) and a fine ash density of 1 g/cm (Rampino et al., 1982) were used in Equation 3. Assuming conditions at 5000 meter elevation (the approximate lower limit of the stratosphere), the estimated Hg 2 deposition for the Tambora eruption is 25.6 g/m ; approximately three times the estimated Hg deposition calculated from Hg concentrations in the ice 2 core (8.6 g/m ). In Equation 3, if the atmospheric deposition (Hgvol) is set equal to the estimated value from the ice core and the equation solved for the 3 concentration of Hg in the volcanic plume (Hgplu), a value of 16 g/m is calculated. Applying the same assumptions to the Krakatau eruption, 2 atmospheric Hg deposition is estimated to be 3.6 g/m ; almost five times less 2 than the deposition calculated from ice-core Hg concentrations (18.2 g/m ). Again, setting atmospheric deposition equal to the Hg deposition estimated from the ice core in the equation and solving for the concentration of Hg in 3 the volcanic plume (Hgplu), a value of 244 g/m is calculated. Based on a limited number of studies measuring Hg concentrations in volcanic plumes, the estimate of Hg deposited from the Krakatau volcanic plume calculated here is comparatively high. The measurements made in

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previous studies (Lepel et al., 1978; Siegel, et al., 1978; Phelen et al., 1982; Unni et al., 1978; Fruchter et al., 1980), however, suggest large ranges of Hg concentrations in volcanic ash plumes are possible. This estimation, though over-simplified, demonstrates that the Hg deposition calculated from concentrations in the ice core attributed to the globally impacting volcanic eruptions of Tambora and Krakatau are not unreasonable. While individual volcanic events lead to short-term deposition rates similar to the industrial maximum (Table 2), the brief duration of the events limits their importance in overall deposition. 3.6.5

Recent declines in atmospheric mercury deposition

Since the industrial maximum (circa 1984), Hg concentrations in the UFG ice core have declined from the 20-fold increase since pre-industrial times to an 11-fold increase during the 1990s. This decline is corroborated by recent declining trends observed in dated sediment cores (Bindler et al., 2001; Engstrom et al., 1997; Norton et al., 1997) and precipitation (Susong et al., 1999). The declining trends recorded during the last ten years of icecore record are consistent with the last 7 years of precipitation data (USEPA, 1997). The top 10 m of the ice core have a calculated average deposition rate of about 1 µg/m2. The data shown in Figure 2 shows the UFG region receiving 1-3 µg/m2. The recent declines may be in response to emission controls implemented through the United States Clean Air Act of 1970 and the Clean Air Amendment of 1990 requiring pollutant scrubbers that also likely remove a fraction of the Hg in flue gasses. If so, the results presented here suggest that further reductions are achievable.

4.

CONCLUSIONS

ECM results, along with isotopic and chemical data from the UFG, support and refine chronological estimates of the ice core based solely on isotopic age dates. Using combined chemical and isotopic age dates (3H, 36 Cl, 14C, Krakatau and Tambora volcanic events), additional historical volcanic events were assigned to ECM peaks in the ice core and used as additional timeline events to calculate a refined age-depth profile. The refined profile yielded an age of 1885 A.D. for the 1883 A.D. Krakatau event and 1811 A.D for the 1815 A.D Tambora event. At a depth of 152 m, the refined profile also shows good agreement (1736 A.D.) with the 14C age date (1729 A.D. ± 95 years). An increase in variance and shift in the δ18O signal of the UFG ice core is an indication of a change in climate known as the LIA. However, the

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sampling interval for δ18O is sufficiently large (20 cm) that it is difficult to pinpoint the LIA termination based on δ18O data alone. Based on past research, a change in the δ18O variance is generally coincident with a change in ECM variance. A 999-point running average of the ECM data set and results from f-tests indicate the variance of the ECM data decreases significantly at about 108 m. At this depth, the age-depth profile predicts an age of 1845 A.D. These data indicate that the termination of the LIA was abrupt with a major climatic shift to warmer temperatures around 1845 A.D. and continuing to present day. Prediction limits (error bars) calculated for the profile ages were ± 10 years (90 percent confidence level). Confidence limits were much lower at about two to three years throughout the length of the core. Thus, a conservative estimate for the time taken to complete the LIA climatic shift to present-day climate is about 10 years, indicating that a major climatic shift in alpine regions of central North America’s recent history (250 years before present) was on a decadal scale. The ice cores also contain a record of total atmospheric Hg deposition. The record indicates that major atmospheric releases of both natural and anthropogenic Hg from regional and global sources. Integrated over the past 270-year ice-core history, anthropogenic inputs contributed 52 percent, volcanic events 6 percent, and background sources 42 percent. More significantly, during the last 100 years, anthropogenic sources contributed 70 percent of the total Hg input. Unlike the 2 to 7-fold increase observed from pre-industrial times (before 1840) to the mid 1980s in sediment-core records, the UFG record indicates a 20-fold increase for the same period. The sediment-core records, however, are in agreement with the last 10 years of this ice-core record, indicating declines in atmospheric Hg deposition.

5.

REFERENCES

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VARIATIONS BETWEEN δ18O IN RECENTLY DEPOSITED SNOW AND ON-SITE AIR TEMPERATURE, UPPER FREMONT GLACIER, WYOMING

David L. Naftz, David D. Susong, L. DeWayne Cecil, and Paul F. Schuster

1.

INTRODUCTION

Oxygen isotopic ratios (δ18O) in ice cores have been used extensively to reconstruct past climate trends (Covey and Haagenson, 1984; Charles et al.; 1994; Lorius et al., 1990; Cuffey et al., 1994; and Jouzel et al., 1997). Recent investigations have documented positive correlations between sitespecific δ18O values in snow and ice to on-site average air temperature (TA) during snow accumulation events in polar regions (Thompson et al., 1994; Shuman et al., 1995). To date (2003), only limited observations link the δ18O values in ice and snow samples to on-site air temperature at high-altitude, mid- and lowlatitude ice-coring sites (Davis et al., 1995; Yao et al., 1996; Yao et al., 1999; and Naftz et al., 2002). Work by Naftz et al., (2002) developed a series of site-specific transfer functions between on-site air temperature and δ18O values in snow deposited at a mid-latitude ice coring site established at over 4,000 m above sea level (masl) on Upper Fremont Glacier (UFG), Wyoming. The site-specific transfer functions developed from this site were used in conjunction with δ18O values from ice cores obtained from UFG to reconstruct changes in air temperature since the early 1700s. Because UFG is a remote, high-altitude site, it is not possible to physically collect discrete snow samples from individual storm events for δ18O analysis. Instead, the site was visited one to three times per year and snow pits were excavated and used to sample the accumulated snowpack. The timing and amount of each snow accumulation event in the excavated snow pits on UFG was determined from snow pillow data from Cold Springs SNOTEL site, approximately 22 km northeast from UFG, and used in the development of site-specific transfer functions (Naftz et al., 2002). This method of determining the timing and relative amounts of accumulation events on UFG relied on the following assumptions: (1) snow redeposition 217 L. D. Cecil et al. (eds.), Earth Paleoenvironments: Records Preserved in Mid- and Low-Latitude Glaciers, 217-234. © 2004 Kluwer Academic Publishers. Printed in the Netherlands.

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from wind events and melting was minimal after each accumulation event; (2) the same storms impact both sites with the same relative intensities; and (3) all precipitation on the UFG is in the form of snow and, therefore, accumulates and can be used as a proxy for cumulative precipitation. None of these assumptions were valid 100 percent of the time. To better record the on-site snow accumulation and redeposition on UFG during 1999-2000, an hourly record of snow depth was obtained using an ultrasonic sensor (Figure 1). Instead of relying on the SNOTEL data as a proxy record for snow accumulation on UFG, on-site data were possible. The objectives of this chapter are to: (1) investigate and model the transfer function between δ18O values in snow and the corresponding air temperature using the continuous, on-site snow-depth and air temperature monitoring equipment installed on UFG (Figure 2); (2) compare the transfer function developed from the on-site snow-depth sensor to transfer functions developed using off-site snow accumulation data; and (3) reconstruct and compare air temperatures from δ18O values in UFG ice cores using both transfer functions.

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Figure -1. Ultrasonic depth sensor during May 2000, Upper Fremont Glacier, Wyoming.

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Figure -2. Location of meteorological station, snow-pit sampling site, snow-depth sensor, and ice-coring sites, Upper Fremont Glacier, Wind River Range, Wyoming.

2.

BACKGROUND INFORMATION AND METHODOLOGY

2.1

Background Glaciological Data

Background data on UFG were collected from 1990-91 (Naftz and Smith, 1993). Radio-echo sounding (Watts and England, 1977; Trombley, 1986) was used to determine ice thickness and bedrock topography. Ice thickness ranged from 60 to 172 m in the upper half of the glacier during 1990. The 10-m borehole temperatures (4-day equilibration period) indicated that the ice is at the pressure melting point (0 + 0.4 oC). Annual ablation (including snow, firn, and ice) measured during 1990-91 averaged 0.93 meters per year (m/y). Densification processes proceed rapidly at the site, with densities

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2 exceeding 8.5 x 10 kilograms per cubic meter (kg/m3) at depths 14 m below the surface. Ice velocity decreases in a downslope direction, ranging from 0.8 to 3.1 m/yr. Mean air temperature monitored on UFG from July 11, 1990 o to July 10, 1991 was -6.9 C. snow-pit sampling at the site indicated preservation of the annual δ18O signature during the initial summer melt season followed by dampening of the signal in subsequent melt seasons (Naftz, 1993; Naftz et al., 1993). A detailed age-depth profile for UFG was determined using continuous electrical conductivity measurements (125,000 data points) on the ice core collected in 1991 (Schuster et al., 2000) in combination with carbon-14 age dating of insect parts found in the lower part of this core (Naftz et al., 1996). Based on these data, ice from the bottom of the 1991 ice core (162 m below the surface) was probably deposited as snow prior to 1710 A.D. The existence and preservation of paleoenvironmental records is well documented in the ice cores collected from UFG. The δ18O record in the ice core was used to document the rapid termination of the Little Ice Age (LIA) in northwestern Wyoming and provide a global linkage to this rapid shift with the δ18O record from the Quelccaya Ice Cap in Peru (Naftz et al., 1996). Electrical conductivity logging of the UFG ice core identified a detailed record of outfall from historic volcanic eruptions and further refined the timing of the climatic shift observed during the termination of the LIA (Schuster et al., 2000; Schuster et al). A detailed record of radioactive fallout from above-ground nuclear testing was preserved and identified in the UFG ice core by Naftz (1993); Naftz et al., (1996); and Cecil and Vogt (1997). The UFG ice-core also contains the first and most comprehensive atmospheric mercury deposition record of its kind, currently available in North America (Schuster et al., 2002; Krabbenhoft and Schuster, 2002; Schuster et al)

2.2

Data Collection

Site specific air temperature adjacent to UFG was monitored with an automated weather station (WS) installed in a boulder field approximately 100 m to the north of UFG at an altitude of 3,960 m (Figure 3). Sensor height was approximately 2 m above land surface and it was shielded with a RM Young 12 plate grill radiation shield. The temperature sensor was factory calibrated and was periodically checked for accuracy at the field site with a hand held digital thermocouple probe. The WS was operated from July 1990 through August 1991 and from September 1997 through March 2001. Relative humidity, wind speed and direction, and solar radiation (band width) also were measured each minute and compiled into hourly averages. Changes in snow depth and air temperature on UFG were

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continuously monitored with a Judd snow-depth sensor installed on UFG at an altitude of 3,990 m in September 1997 (Figure 1). This ultrasonic depth sensor was suspended on a horizontal bar above the snow surface and emitted a sonic signal. The travel time for the sonic pulse to the snow surface and return was measured, and the distance to the surface was calculated after correcting for air-temperature effects on the speed of sound. The sensor recorded snow depth and air temperature every 60 minutes.

Figure -3. Automated weather station installed near Upper Fremont Glacier, Wyoming.

Snow samples were collected on UFG from a vertical trench face in a snow pit excavated to the firn layer from the previous accumulation cycle (1998-99). Seven to 10 centimeters (cm) of snow were composited in precleaned, 500 milliliter (ml) Nalgene wide-mouthed sample bottles. Snow density was determined by collecting snow at 10-cm intervals from trench faces in 1,000-cm3 samplers and weighing the samples on a portable

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electronic balance. Ice-core samples from UFG were collected according to methods outlined by Naftz et al. (1996).

2.3

Sample Processing and Analysis

After melting, the snow samples were filtered (0.45 micrometer (µm)), placed in a glass vial, sealed with a polyseal cap and coated with ParaFilm. 18 The δ O value of each sample was determined by using the method developed by Epstein and Mayeda (1953) at the U.S. Geological Survey Stable Isotope Laboratory in Menlo Park, California, and reported relative to Standard Mean Ocean Water (SMOW) in permil notation.

3.

RESULTS AND DISCUSSION

3.1

Depth-Sensor, Snow Pit, and Weather Station Data

As noted in Naftz et al. (2002), the transfer functions developed by using storm occurrence and amount data from off-site measuring stations do not account for possible localized effects occurring on UFG. These effects include wind erosion, redeposition of snow from previous storm events, or significant differences in the relative timing and magnitude of storms impacting UFG relative to the off-site SNOTEL site. Because of these localized differences, transfer functions developed from on-site versus offsite snow accumulation data could be different. A continuous record of snow depth was recorded from late-September 1999 through early-May 2000 (Figure 4). On May 5, 2000, a snow pit adjacent to the depth sensor was excavated and sampled for δ18O values in 10-cm increments (Figure 4). The measured depth in the snow pit was 137cm, and sampled the accumulated snowpack from late-September 1999 through early-May 2000. The accumulated snow depth measured by the depth sensor was 138 cm (Figure 4), indicating the same time period of accumulation was sampled by the snowpit.

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Figure -4. Cumulative snow depth and average air temperature during the 1999-2000 accumulation period compared to changes in isotopic composition in snow pit FRE-00-1, Upper Fremont Glacier, Wyoming. Preserved snow layers designated as L1 through L7

Based on on-site snow depth measurements, seven discrete storm events were preserved from postdepositional processes in the snow pit (Figure 4). Storm events that were preserved in the snow pack were identified as periods with continuous increases in snow depth. Because of wind removal, only seven storm events remained intact during the study period. The mean air temperature during each of the seven accumulation periods was determined by averaging the hourly air temperature recorded at the depth sensor during each accumulation period. Changes in the δ18O values in the snow-pit samples agree with changes in the air temperature during each storm event (Figure 4). For example, the most negative δ18O value measured in the snow pit (30 to 40 cm depth interval) was deposited during January and February, coinciding with some of the coldest air temperatures during the winter, ranging from –13.1 to –14.7 oC. In contrast, the least negative δ18O value measured in the snow pit (130-137 cm depth interval) was deposited in midApril, coinciding with the warmest air temperature (–7.8 oC) recorded during a preserved snow accumulation event.

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Storm Occurrence and Post Depositional Processes

The continuous monitoring of snow depth on UFG indicates that selected accumulation events may have been completely removed after deposition, probably by wind (Figure 4). On-site air temperatures recorded on UFG from late-September 1999 through early-May 2000 were not high enough for melting to cause the observed decreases in snow depth recorded by the depth sensor. The depth-sensor data do not allow for a clear separation between decreases in snow depth caused by wind removal or settling; however, instances when snow depth returns to pre-storm measurements is probably dominated by wind removal processes. For example, accumulation events that occurred between early October 1999 and early February 2000, may have been subjected to a higher proportion of wind removal than other storm events (Figure 5). Lower air temperatures during this time period likely resulted in a less dense snow that was more vulnerable to wind removal. High wind speeds that occur between accumulation events (Figure 5) do not appear to cause significant snow removal, probably because of the development of a wind or sun crust on the snow surface. On the basis of these observations, accumulation events that occur during lower air temperatures may be more susceptible to wind removal, potentially biasing the δ18O values in the snow pit (and ultimately the ice core) toward periods of higher air temperatures (late season snowfall).

Figure -5. Average daily wind speed and hourly snow depth measured on Upper Fremont Glacier, Wyoming from October 1999 through early May 2000.

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Seasonally influenced wind removal of snow deposited on UFG was further investigated by comparing the cumulative precipitation measured daily at the Cold Springs SNOTEL site to the cumulative precipitation of snow accumulation events preserved on UFG. The Cold Springs SNOTEL site is located below tree line and is not as susceptible to wind removal events compared to UFG. The annual precipitation at the Cold Springs SNOTEL site is typically less than the annual precipitation measured on UFG (Naftz et al., 2002). Relative changes in the cumulative precipitation at the two sites is supportive of early season snowfall (before February 2000) being more susceptible to wind removal compared to snowfall between February and early May 2000 (Figure 6). From early October 1999 through at least mid-January 2000, the cumulative precipitation on UFG was approximately equal to the cumulative precipitation measured at the Cold Springs SNOTEL site. After early February 2000, there was a significant increase in the cumulative precipitation measured on UFG relative to the SNOTEL site (Figure 6) and by mid-April 2000, UFG cumulative precipitation exceeded SNOTEL precipitation by about 13 cm. The mechanism for this increased cumulative precipitation on UFG is not definitive, but is probably the result of warmer air temperatures allowing the snow to be less susceptible to wind erosion and perhaps the redeposition of snow from the higher elevation ridge tops surrounding UFG. Other processes could also account for the observed differences, for example, seasonal changes in moisture sources; however , this is beyond the scope of this research.

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Figure -6. Cumulative precipitation measured daily at the Cold Springs SNOTEL site compared to cumulative precipitation of preserved snow accumulation events on the Upper Fremont Glacier, Wyoming.

The seasonally biased snowpack on UFG could also bias the isotopic record and resulting air temperature reconstructions. For example, an annual snowpack on UFG that had a lighter than average δ18O signal could indicate a decrease in the annual air temperature; however, it could also indicate a year with less wind, resulting in the preservation of a higher proportion of mid-winter snowfall, with a lighter isotopic signature. This and other meteorological influences could contribute to changes in the δ18O values in an annual snowpack at UFG as well as other high elevation ice-coring sites.

3.3

Modeling Air Temperature with δ18O Variation in Snow

The δ18O value of each preserved snow layer was determined using the detailed depth data recorded by the on-site snow-depth sensor in combination with the density of each snow-pit sample (Figure 7). Each preserved snow layer, designated L1 through L7, was assigned a mean air

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temperature on the basis of measured air temperature on UFG during the accumulation period. Because the snow pit was sampled in 10-cm composites, the δ18O value for each of the seven snow layers was determined using density-weighted averages according to the sample interval (Figure 7).

Figure -7. Weighted mean δ18O values from snow-pit samples in relation to air temperatures measured during snowfall events using on-site depth sensor data (a) and Cold Springs SNOTEL data (b), Upper Fremont Glacier, Wyoming. Linear trends between mean air temperature and weighted mean δ18O values are from snow-pit samples using on-site depth sensor data (c) and Cold Springs SNOTEL data (d)

A transfer function between δ18O and mean air temperature was developed using the depth-sensor and snow-pit data (Figure 7). The R2 value for the transfer function was 0.71 and was statistically significant (p = 0.0179). The slope (δ18O /air temperature) of the transfer function is 1.350. The transfer function developed from the depth-sensor data was compared to the transfer function developed using the Cold Springs SNOTEL data as a proxy indicator for the timing and amount of snow accumulation events on UFG during the late-September 1999 through early-May 2000 period (Naftz et al., 2002). The R2 value for this transfer function was 0.65 and was statistically significant (p <= 0.0001). The slope of the transfer function developed using the SNOTEL data (0.688) was about half of the slope of the transfer function developed using the on-site depth-sensor data (1.350).

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One reason for the significant difference in slope between the two transfer functions is the wind removal of early season snowfall measured by the on-site depth sensor. Without an on-site record of snow accumulation and removal, the SNOTEL data indicated 18 percent of the seasonal precipitation (late September 1999 to early May 2000) on UFG occurred during early season accumulation events (9/25/99 to 11/15/99). In contrast, the depth sensor record indicated that only 3 percent of the seasonal precipitation was actually preserved on UFG during the same early season accumulation cycle. The transfer function developed with the SNOTEL accumulation data assumed that the snow from the lower parts of the snow pit were associated with the warmer air temperatures that occurred during the early season snowfall events. This resulted in a larger range in air temperature for the transfer function developed with the SNOTEL data relative to the transfer function developed with the on-site depth sensor (Figure 7). Based on these observations and processes, it is clear that the transfer function developed with the on-site depth sensor best represents the changes in δ18O values with air temperature during the 1999-2000 accumulation season.

3.4

Application of Transfer Functions to Ice-Core Data

Before obtaining the on-site depth sensor record, a transfer function developed with the off-site SNOTEL data was used in combination with δ18O values from the UFG ice core to reconstruct trends in average air temperature over the past 300 years (Naftz et al., 2002). On the basis of this temperature reconstruction, an average temperature increase of approximately 3.5 oC occurred on UFG since the mid- to late-1960s. Results presented in the previous section indicate that continuous, on-site monitoring of snow depth can improve the reliability of transfer functions relative to transfer functions constructed with off-site SNOTEL data reported in Naftz et al. (2002). The TAs reconstructed from the UFG ice-core δ18O values using the transfer function derived from the SNOTEL data were compared to the TAs derived from the on-site depth sensor data (Figure 8). Although still substantial, the increase in TA using the transfer function developed from the on-site depth sensor is less than the increase calculated using the transfer function developed with the SNOTEL data. Instead of a TA increase of approximately 3.5o C since the mid- to late-1960s (Naftz et al., 2002), a TA increase of 1.5o C was determined (Figure 8) with data from the on-site depth sensor. This increase is much higher than the observed global mean temperature increase of approximately 0.7 oC during the entire 20th Century. The reconstructed temperature increase from UFG of 1.5 oC is comparable to

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the temperature increase of + 2 oC for sites in the European Alps since the early 1980s (Haeberli and Beniston, 1998); + 2.2 oC at a high-altitude site in central Asia from 1962 to 1990 (Mikhalenko, 1997); and + 4 oC in the western Arctic during the 20th Century (Sin’kevich, 1991). The ice-core data indicate that UFG and adjacent alpine areas in the Wind River Range, Wyoming, may be warming at much faster rates than the global average.

Figure -8. Comparison of reconstructed air-temperature trends (20-sample running mean) using the composite transfer function developed from SNOTEL data (Naftz et al., 2002) compared to reconstructed air temperatures using the on-site depth sensor. The δ18O values used in the reconstruction were obtained from the 1998 ice core collected from Upper Fremont Glacier, Wyoming and represent the time period from 1950-1995.

Further increases in temperature were detected by reconstructing TA values from deeper sections of the core. Previous studies (Naftz et al., 1996; Schuster et al., 2000) indicate that the ice-core section from about 102 to 150 m below the surface corresponds to snow deposited during the LIA from approximately 1740 to 1860 A.D. The mean δ18O value in ice from this section of the core is -19.85 permil (n = 248). Using the transfer function developed from the SNOTEL data, the mean δ18O value during the LIA corresponds to a mean air temperature of -11.5 oC (Naftz et al., 2002). In

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contrast, the transfer function developed from the on-site depth sensor data for the same section of ice core corresponds to a mean air temperature of – 12.1 oC (Figure 8). Using this mean air temperature of –12.1 oC, an increase in TA of approximately +2 oC has occurred on UFG from the present to the LIA (Figure 8). The difference in reconstructed TAs serves as a good example for showing the importance of snow removal and accumulation cycles in the development of a transfer function between TA and δ18O of the snow samples. In less remote areas where personnel can collect snow samples from discrete storm events, the development of transfer functions can be relatively straightforward. In contrast, logistical constraints at remote icecoring sites do not permit the direct collection of snow samples from discrete storm events. Instead, samples from intermittent snow pits must be combined with weather station and snow-depth data for the development of transfer functions relating δ18O to TA. Results from the UFG 1999-2000 accumulation cycle indicate that transfer functions developed using on-site depth-sensor records best represent changes in δ18O values with air temperature.

4.

CONCLUSIONS

Localized transfer functions relating TA and δ18O values in recent precipitation at ice-coring sites are needed to determine long-term changes in TA and δ18O values preserved in the ice. Transfer functions relating TA and δ18O values in precipitation are difficult to determine at remote ice-coring sites, such as UFG, because it is not possible to physically collect discrete snow samples from individual storm events. Instead, snowpits are sampled one to three times per year for δ18O analysis and the on-site TA is measured using an automated WS. The timing and amount of each precipitation event must be determined from automated, depth sensor equipment. A transfer function, developed from a snow-depth sensor installed on UFG, was compared to a transfer function developed with snow-depth data from the Cold Springs SNOTEL site, approximately 22 km from UFG. A continuous record of snow depth was recorded from late-September 1999 through early-May 2000. In early May 2000, a snow pit adjacent to the depth sensor was excavated and sampled. The most negative δ18O values in the snowpit samples corresponded to the snowfall events that occurred during the coldest air temperatures recorded by the on-site WS. Data from the onsite depth sensor indicate snowfall events occurring between early October 1999 and early February 2000, were subjected to a higher proportion of wind removal than other events, probably due to lower air temperatures. Relative

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changes in the cumulative precipitation between UFG (above tree line) and the Cold Springs SNOTEL (below tree line) site is supportive of early season snowfall (before February 2000) being more susceptible to wind removal compared to snowfall between February and early May 2000. A transfer function was developed using the depth-sensor and snow-pit data (R2 = 0.71, p = 0.0179) and compared to the transfer function developed using the Cold Springs SNOTEL data (R2 = 0.65, p <= 0.0001). The slope of the transfer function developed using the SNOTEL data (0.688) was about half of the slope of the transfer function developed using the on-site depthsensor data (1.350). A reason for the significant difference in slope is the wind removal of early season snowfall measured by the on-site depth sensor. The transfer function developed with the on-site depth sensor best represents the changes in δ18O values with air temperature during the 1999-2000 accumulation season. The transfer function developed with on-site depth sensor data was used in combination with δ18O values from the UFG ice core to reconstruct trends in average air temperature over the past 300 years. Based on this reconstruction, an increase in TA of approximately +2 oC has occurred on UFG from the LIA to the present. This is less than the increase in TA determined for the same time period using the transfer function developed with SNOTEL data. The difference in reconstructed TAs serves as a good example for showing the importance of snow removal and accumulation cycles in the development of a remote site transfer function at ice-coring sites. Results from UFG (1999-2000 accumulation cycle) indicate that onsite depth sensor records are best for developing transfer functions between δ18O and TA at remote ice-coring locations.

5.

REFERENCES

Cecil, L.D. and Vogt, S, 1997, Identification of bomb-produced chlorine-36 in mid-latitude glacial ice of North America, Nucl. Instrum. Methods Phys. Res., Sect. B, 123, 287-289. Charles, C.D., Rind, D., Jouzel, J., Koster, R.D., and Fairbanks, R.G., 1994, Glacialinterglacial changes in moisture sources for Greenland: Influences on the ice core record of climate, Science, 263, 508-511. Covey, C., and Haagenson, P.L., 1984, A model of oxygen isotope composition of precipitation; implications for paleoclimate data, Journal of Geophysical Research, 89, D3, 4,647-4,655. Cuffey, K.M., Alley, R.B., Grootes, P.M., Bolzan, J.M., and Anandakrishnan, S., 1994, Calibration of the δ18O isotopic paleothermometer for central Greenland, using borehole temperatures, Journal of Glaciology, 40, 341-349. Davis, M.E., Thompson, L.G., Mosley-Thompson, E., Lin, P.N., Mikhalenko, V.N., and Dai, J., 1995, Recent ice-core climate records from the Cordillera Blanca, Peru, Annals of Glaciology, 25, 225-230.

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Epstein, S., and Mayeda, T., 1953, Variation of the O-18 content of waters from natural sources, Geochimica et Cosmochimica Acta, 4, 213-224. Haeberli, W. and Beniston, M., 1998, Climate change and its impacts on glaciers and permafrost in the Alps, Ambio, 27, 258-265. Jouzel, J., Alley, R.B., Cuffey, K.M., Dansgaard, W., Grootes, P., Hoffmann, G., Johnsen, S.J., Koster, R.D., Peel, D., Shuman, C.A., Stievenard, M., Stuiver, M, and White, J., 1997, Validity of the temperature reconstruction from water isotopes in ice cores, Journal of Geophysical Research, 102, 26,471-26,487. Krabbenhoft, D.P. and Schuster, P.F., 2002, Glacial ice cores reveal a record of natural and anthropogenic atmospheric mercury deposition for the last 270 years, U.S. Geological Survey Fact Sheet FS-051-02, 2 p. Lorius, C., Jouzel, J., Raynaud, D., Hansen, J., and Treut, H., 1990, The ice-core record: Climate sensitivity and future greenhouse warming, Nature, 347, 139-145. Mikhalenko, V.N., 1997, Changes in Eurasian glaciation during the past century: Glacier mass balance and ice-core evidence, Annals of Glaciology, 24, 283-287. Naftz, D.L., 1993, Ice-core records of the chemical quality of atmospheric deposition and climate from mid-latitude glaciers, Wind River Range, Wyoming, Ph.D. thesis, Colorado School of Mines, Golden, Colorado. Naftz, D. L., Michel, R.L., and Miller, K.A., 1993, Isotopic indicators of climate in ice cores, Wind River Range, Wyoming, in Swart, P. K., K. C. Lohmann, J. McKenzie, and S. Savin, (eds.). Climate Change in Continental Isotopic Records, Geophysical Monograph 78, Washington, D.C., American Geophysical Union, 55-66. Naftz, D.L., and Smith, M.E., 1993, Ice thickness, ablation, and other glaciological measurements on Upper Fremont Glacier, Wyoming, Physical Geography, 14, 404-414. Naftz, D.L., Klusman, R.W., Michel, R.L., Schuster, P.F., Reddy, M.M., Taylor, H.E., Yanosky, T.M., and McConnaughey, E.A., 1996, Little Ice Age evidence from a southcentral North American ice core, U.S.A., Arctic and Alpine Research, 28, 35-41. Naftz, D.L., Susong, D.D., Schuster, P.F., Cecil, L.D., Dettinger, M.D., Michel, R.L., and Kendall, C., 2002, Ice-core evidence of rapid air temperature increases since 1960 in alpine areas of the Wind River Range, Wyoming, United States: Journal of Geophysical Research, vol. 107, no D13, doi: 10.1029/2001JD000621. Schuster, P.F., White, D.E., Naftz, D.L., and Cecil, L.D., 2000, Chronological refinement of an ice core record at upper Fremont Glacier in south central North America, Journal of Geophysical Research, 105, 4,657-4,666.

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Schuster, P.F., Krabbenhoft, D.P., Naftz, D.L., Cecil, L.D., Olson, M.D., Dewild, J.F., Susong, D.D., and Green, J.R., 2000, Atmospheric mercury deposition during the last 270 years: A glacial ice core of natural and anthropogenic sources, Environmental Science and Technology, 36, 2,303-2,310. Shuman, C.A., Alley, R.B., Anandakrishnan, S., White, J.W.C., Grootes, P.M., and Stearns, C.R., 1995, Temperature and accumulation at the Greenland Summit: Comparison of highresolution isotope profiles and satellite passive microwave brightness temperature trends, Journal of Geophysical Research, 100, 9,165-9,177. Sin’kevich, S.A., 1991, Climate warming in the Twentieth century as reflected in Svalbard ice cores: Glaciers-Ocean-Atmosphere: Interactions, International Association of Hydrological Sciences Publication No. 208, 257-267. Thompson, L.G., Peel, D.A., Mosley-Thompson, E., Mulvaney, R., Dai, J., Lin, P.N., Davis, M.E., and Raymond, C.F., 1994, Climate since A.D. 1510 on Dyer Plateau, Antarctic Peninsula; evidence for Recent climate change, Annals of Glaciology, 20, 420-426. Trombley, T.J., 1986, A radio echo-sounding survey of Athabasca Glacier, Alberta, Canada, M.S. thesis, University of New Hampshire, Durham, New Hampshire. Watts, R.D. and England, A.W., 1977, Radio-echo sounding of temperate glaciers: ice properties and sounder design criteria, Journal of Glaciology, 17, 39-48. Yao, T., Thompson, L.G., Mosley-Thompson, E., Zhihong, Y., Xingping, Z., and Ping-Nan Lin, 1996, Climatological significance of the δ18O in north Tibetan ice cores, Journal of Geophysical Research, 101, 29,531-29,537. Yao, T., Masson, V., Jouzel, J., Stievenard, M., Weizhen, S., and Keqin, J., 1999, Relationships between δ18O in precipitation and surface air temperature in the Urumqi River Basin, east Tianshan Mountains, China, Geophysical Research Letters, 26, 3,4733,476.

Summary

L. DeWayne Cecil

The majority of the world’s population lives at mid-latitudes. It is vital, therefore, to understand how the climate and the environment are changing on local, regional, and global scales. It is also imperative that we understand the impacts we as people are having on our water and other natural resources as the population continues to grow at an alarming rate. In this first of its kind volume, a wide ranging discussion of the scientific value of various records archived in mid-latitude glaciers provides a unique opportunity to see how our environment has changed in the past, how it is changing today, and to project possible changes in the future. The papers in this book focus on the archived record in glacial ice worldwide that represents approximately the past 500 years. The study of mid-latitude glacial ice is a difficult and arduous task. Study sites presented in this volume range from the Rocky Mountains of western North America to the Tien Shan and Himalayan Mountains of Central Asia, from the Andes of South America to the glaciers and mountains in Scandinavia. New analytical and interpretive techniques that lower analytical detection levels and expand our ability to study glacial records were presented (Schwikowski, Naftz et al., and Schuster et al.). This information was applied to societal problems ranging from understanding climate change signals and making future climate predictions (Yao and 235 L. D. Cecil et al. (eds.), Earth Paleoenvironments: Records Preserved in Mid- and Low-Latitude Glaciers, 235-241. © 2004 Kluwer Academic Publishers. Printed in the Netherlands.

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Summary

Yang, Naftz et al., Whalley, and Davis and Thompson) to documenting evidence of industrial chemical and natural fallout (Schwikowski, Kreutz et al., Wake et al., and Schuster et al.). For example, in Naftz et al., a transfer function was developed for the Upper Fremont Glacier (UFG) in Wyoming, U.S.A., using snow depthsensor and snow-pit data that was then compared to a transfer function developed using a nearby SNOTEL dataset. The slope of the transfer function developed using the SNOTEL data (0.688) was about half of the slope of the transfer function developed using the on-site glacier depthsensor data (1.350). A reason for the significant difference in slope was attributed to wind removal of early season snowfall measured by the on-site depth sensor. The transfer function developed with the on-site depth sensor best represents the changes in stable isotopic (į18O) values with air temperature during the 1999-2000-accumulation season at this glacial site. The transfer function developed with on-site depth sensor data was used in combination with į18O values from the UFG ice core to reconstruct trends in average air temperature (TA) over the past 300 years. Based on this reconstruction, an increase in TA of approximately +2 oC has occurred on UFG from the Little Ice Age (LIA) to the present. This is less than the increase in TA determined for the same time period using the transfer function developed with SNOTEL data. The difference in reconstructed TAs serves as a good example for showing the importance of snow removal and accumulation cycles in the development of a remote-site transfer function at ice-coring sites. Results from this study on the UFG indicated that on-site depth sensor records are best for developing transfer functions between į18O and TA at remote ice-coring locations. Another study presented in this volume demonstrated the effects of snow removal, snow accumulation, and changes to an archived ice record due to meltwater percolation that can be a major factor at these mid-latitude sites. In Schotterer et al., post-depositional processes that remove, redistribute, and change the isotopic and chemical information about climatic and environmental change that arrives with the snow flakes falling on a glacier's surface were documented. It was shown that sublimation hardens a fresh snow cover and may counteract wind scour. Refreezing of melt has the same effect, but sublimation and melting can also change the accumulated information. From the examples considered, it was concluded that wind scour is probably the most important process at high-altitude drilling sites because it may disturb or even effectively remove seasonal cycles. Schotterer et al., stated that careful site selection and exploratory process studies prior to deep drilling could prevent unfortunate surprises. Saddle sites with channeling of winds and steep crests are obviously less well suited than more open summit sites. If a cold glacier site is excessively wind-

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exposed, a more sheltered temperate glacier nearby might offer a superior record, at least for acquisition of a stable-isotope record. It was further shown that chemical species are affected by postdepositional processes in a much more complex manner than stable isotopes. The latter are fundamental constituents of water molecules and interpretation of isotope records benefits from quantitative, physically based understanding of their behavior during phase transitions. In spite of this knowledge, no foolproof strategy exists to anticipate or avoid such complications. Glaciers are dynamic and open systems for precipitation. Wind, sun, cloudiness, precipitation rate, temperature distribution and a myriad of other factors change as climate changes, and so does the information on climatic and environmental change. This study illustrated the difficulty of interpreting signals in ice cores. Schikowski, and Schuster et al. gave examples of the effects of environmental laws and voluntary controls on industrial emissions as recorded by fallout onto glaciers. Paleo-records from Alpine glaciers allow the reconstruction of European atmospheric pollution history (Schikowski). These records show a generally consistent picture of a vastly altered atmospheric composition due to anthropogenic emissions. Typically following the onset of industrialization, increased concentrations of major aerosol components (e.g. sulfate, nitrate, ammonia, black and elemental carbon), trace constituents (e.g. lead and other heavy metals), gaseous species (including the greenhouse gas methane), and radionuclides (tritium, cesium-137, chlorine-36, and iodine-129) were observed in the ice cores. However, on a positive note, Alpine glaciers are also a reliable indicator of the progress achieved in environmental protection. Since approximately 1970, many of the pollutants mentioned here show a clear tendency to lower concentrations, which is a direct result of various air quality measures such as the use of filtering units in power plants, in incineration plants and in industry, the increased use of oils with low sulfur content, as well as the introduction of catalytic converters and lead-free gasoline. In Schuster et al., the ice cores collected from the UFG contained a record of total atmospheric mercury (Hg) deposition at this North American site. The record indicates that major atmospheric releases of both natural and anthropogenic Hg from regional and global sources. Integrated over the past 270-year ice-core history, anthropogenic inputs contributed 52 percent, volcanic events 6 percent, and background sources 42 percent. More significantly, during the last 100 years, anthropogenic sources contributed 70 percent of the total Hg input. Unlike the 2 to 7-fold increase observed from pre-industrial times (before 1840) to the mid 1980s in sediment-core records, the UFG record indicates a 20-fold increase for the same period. The sediment-core records, however, are in agreement with the last 10 years of

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this ice-core record, indicating declines in atmospheric Hg deposition. The general decline in Hg concentrations in the glacial ice since about 1980 were attributable to voluntary emission controls and the enactment of environmental laws. Davis and Thompson presented an important piece of information concerning the management of water resources in densely populated Central Asia. From the Tanggula Mountains in China, which span east-west across the Tibetan Plateau at 32o to 33oN, the location of a precipitation transition between the north and south sides of the Plateau was documented. North of this latitude, the accumulation rate is low, 18O-enriched precipitation falls in the summer and the deuterium excess is high. To the south, the annual accumulation is higher, the 18O-enriched precipitation is deposited in the winter, and the deuterium excess is lower. In addition, the loading pattern of the first unrotated principle component of the summertime precipitation anomalies are out of phase between north and south of approximately 33oN across the eastern part of the Plateau. Many investigators have concluded that the Tanggula Mountains are a barrier to the northward expansion of the influence of the South Asian Monsoon, while to the north recycled rainfall from continental precipitation processes dominate. It was noted that the line that separates the different characteristics of precipitation from the north to the south sides of the Plateau occurs at the summertime northernmost extent of the Intertropical Convergence Zone over this region. It was further concluded that the spatial and temporal comparisons of the stable isotope and accumulation records from the ice cores recovered from the north and the south of the “Tanggula transition” tend to agree with the stable isotope analyses of water and precipitation samples from across the region. Several studies of the stable isotope chemistry of the precipitation in the South Asian Monsoon region link the “amount effect” with the low summertime values of δ18O on monthly time frames. However, over semidecadal and longer time scales the strongest links appear to be with atmospheric air temperature. It is hoped that the stable isotope and accumulation history from this region will help refine the northern extent of the monsoon influence over time and facilitate a better understanding of water resources and their fate in heavily populated Central Asia. In Wake et al. samples from firn, ice cores, and snow pits were collected for analysis of major ion concentrations from high-elevation accumulation zones from six glaciers in the Himalaya; three on the north side and three on the south side of the main crest. Depth-age relationships, established via annual layer counting and identification of marker horizons, allowed for calculation of annual accumulation and fluxes of major ions. Confirming previous studies, concentrations and fluxes of dust related major ions were 2 to 3 orders of magnitude greater for northern slope sites compared to the

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southern slope sites. This study defined the relatively short distance over which this change occurs with the boundary defined by the main crest of the Himalaya. This crest not only defines different climatic regimes, but also different air masses. Glaciers to the north of the main crest are strongly influenced by dust derived from the arid regions in central Asia. As a result, glaciochemical records that document changes in the dust source regions and/or changes in atmospheric circulation patterns that transport dust across central Asia can and should be developed from north-slope glaciers. While regional sources of pollution are derived primarily from the south, the spatial variation of pollution related aerosols does not display a strong gradient across the Himalaya. In fact, annual fluxes of nitrate and sulfate in Himalayan snow are comparatively low. This is in part due to the scavenging of aerosols in summer monsoon air masses by precipitation during transport to the Himalaya. Thus the pollution signal stored in Himalaya glaciers, especially those on the southern slope, is primarily a function of atmospheric transport and the amount and intensity of precipitation. Extracting pollution related signals from these glaciochemical records requires that changes in atmospheric circulation and precipitation be taken into account. Nitrate and sulfate deposition on the north-slope glaciers are closely linked to dust deposition. Extracting anthropogenic signals from the dust dominated glaciochemical record stored in north slope glaciers will require innovative approaches, such as developing detailed records of trace elements, rare earth elements, and sulfur isotopes. Multi-parameter glaciochemical records provide the basis for identifying potential sources and transport pathways for major ions and thereby identifying the causes for spatial variability in major ion records preserved in mid-latitude glaciers. Improving our understanding of these processes is critical for increasing our confidence in paleoclimatic and environmental reconstructions from ice core records in these regions. The mass balance and associated work on Scandinavian glaciers is probably as extensive as anywhere in the world. The utility of these glacioclimatic records plus glacial chronological investigations allows not only comparisons from different climatic conditions in Scandinavia but testing of models of behavior. Such models may relate to the behavior of the glacier flow itself as perturbed by mass balance changes or they may be ways of modeling the mass balance itself in terms of meteorological and climatic parameters. All such modeling relies upon good field data and the papers reviewed by Whalley briefly show the importance of the longevity as well as quality of the Scandinavian data sets. With new techniques being added to the glaciological armory and the ability to use them on an increasingly diverse range of glaciers, both spatially and with respect to size, the

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Scandinavian glaciers will remain a most important means of elucidating both past and future climatic change. Furthermore, it is now becoming evident that the wide range of various ‘proxy’ data from both terrestrial and oceanic sources is also becoming better integrated with the mass balance data and the modelling of climate change. The glaciers of Scandinavia, by their nature, do not provide good ice-core records which go well back in time but the quality data which the glaciers do preserve in moraines, historical records as well as mass balance and snout retreat data for the last 150 years will continue to be important in elucidating climate change and climate variability. Generally, precipitation recorded at a meteorological station in a given area is influenced by local conditions, especially by the local climate or topography as demonstrated in the paper by Yao and Yang. These authors pointed out that scientists often pay more attention to temperature than to precipitation in climate studies. However, it has been demonstrated that the precipitation variation also exhibits significant similarity in a hemispheric context. Then, the question arises as to whether or not basic precipitation features (amount, timing, etc.) can be established for a given region. Based on the precipitation records from ice caps on the Qinghai-Tibet Plateau presented by Yao and Yang, the answer to that question is yes. Comparing our ice-core record with the records from the meteorological station located at lower altitude, these ice caps are not influenced by local climate and micro-topography. Those ice caps located in the middle-upper part of the convection layer receive and collect all kinds of atmospheric information and are the real natural archive of the atmospheric precipitation record. The precipitation record in the Guliya ice core is important not only in global temperature studies but also in the study of large-scope precipitation features. The negative anomalies of the precipitation recorded in the Guliya ice core are significantly correlated with El NinҔo Southern Oscillation (ENSO) events. Thus, the ice-core records can be used as a proxy to study global climate anomalies associated with ENSO events. The variation of the precipitation recorded in the Guliya ice core shows a teleconnection with ENSO events. Such teleconnections, although statistically significant, are relatively weak for some periods of record. Based on data collected from fresh snow, snow pit, crevasse wall, and ice-core samples during the 2000 Tien Shan expedition in Kyrghyzstan, Kreutz et al., concluded that limited spatial and elevational variability exists in fresh snow isotope and major ion data from the Inilchek Glacier, central Tien Shan Mountains and vicinity. There was no apparent trend with increasing elevation for any of the parameters measured, and the intra-event variability in delta deuterium (įD) ratios and major-ion concentrations was much smaller than the inter-event variability. Based on the lack of isotope

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variability with elevation, the authors concluded that snow formation occurs from a horizontal cloud base across the Inilchek Basin. This model is supported by multi-year snowpit, crevasse wall, and ice core data, which also do not display significant elevational variability in mean isotope or major ion data. Isotope profiles from two deep ice cores have similar mean values and variability (±1 standard deviation), and there is no long-term trend in either įD profile. Based on estimated accumulation rates, the interannual-scale variability in the core profiles is significantly greater than the spatial variability during individual fresh snow events. Given the lack of isotope/elevation trend in fresh snow and snowpit data, we conclude that a majority of the down core isotope variability reflects climate changes rather than ice flow effects or deposition noise. These conclusions have implications for the interpretation of the highresolution isotope and major-ion records eventually developed from the two deep Inilchek ice cores. First, high-resolution records from the two cores can be used for detailed evaluations of time-series signal-noise ratios related to small-scale deposition processes. Second, given the high accumulation rate at the site it may be possible to reconstruct climate and environmental variability on an event basis. The large changes in fresh snow glaciochemistry, among other events, imply that such variability might be apparent in the deep-core record. Lastly, given the different isotope/ion relationships observed in the fresh snow events, it might be possible to use the down-core records to investigate specific moisture source and transport pathways through time. With the relatively recent increase in glacial studies at mid-latitudes as presented in this volume, it more evident than ever before that humans are impacting natural resources and global and regional climate. The mere fact that temperate glacial sites are essentially all receding and disappearing should alarm us all to the need to collect and archive ice and data from the most optimum sites in the most optimum ways (as pointed out by Green and Cecil, and Thompson) as rapidly as we can. The studies presented in this volume demonstrate that we have the tools and teams in place to accomplish this task and to preserve these fragile environmental archives before it is too late.

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Index

accelerated sea-level rise, xxv accelerator mass spectrometry, xvi, xxv, 20 accumulation, ix, xvi, 9, 11, 14, 30, 31, 39, 42, 43, 44, 45, 46, 47, 49, 51, 53, 55, 57, 58, 62, 71, 72, 76, 78, 83, 85, 87, 88, 90, 92, 94, 100, 119, 129, 132, 145, 146, 148, 149, 150, 151, 152, 153, 154, 155, 156, 157, 158, 159, 163, 164, 165, 166, 167, 169, 170, 173, 174, 175, 178, 182, 198, 207, 209, 212, 217, 218, 222, 223, 224, 225, 226, 227, 228, 229, 231, 232, 234, 236, 238, 241 Accumulation, ix, 41, 49, 56, 133, 163 acid rain, xx, 22 acoustic televiewer, 29 aerosols, xx, 8, 25, 90, 91, 92, 96, 196, 214, 239 air temperature, 66, 75, 79, 136, 178, 210, 217, 218, 221, 224, 227, 228, 229, 230, 231, 232, 233, 234, 236, 238 Ak-kem, 53, 54 Ak-tru, 53, 55

Alaska, 5, 11, 142, 197, 209 Aletschhorn, 96 alkalinity, 87 Alpine, vii, viii, 35, 57, 58, 59, 94, 95, 96, 97, 100, 103, 108, 110, 114, 115, 116, 117, 118, 119, 140, 143, 214, 233, 237 alpine glaciers, 10, 39, 42, 57, 181, 182, 184, 216 Alps, viii, xx, 17, 34, 41, 42, 44, 57, 59, 78, 95, 96, 97, 98, 100, 104, 105, 111, 116, 117, 118, 119, 142, 230, 233 altitude effect, 40, 44, 56, 77, 78, 209 Ammonium, 52, 87 amount effect, 94, 151, 159, 238 AMS, xxv, 20, 27, 35 Andes, viii, xx, 6, 8, 9, 10, 11, 17, 40, 41, 42, 48, 53, 57, 58, 59, 184, 235 Anions, 87 annual, 7, 9, 11, 15, 24, 30, 33, 39, 40, 43, 44, 46, 47, 49, 51, 53, 54, 55, 70, 83, 84, 85, 86, 87, 89, 91, 92, 97, 98, 99, 100, 112, 125, 126, 146, 148, 152, 153, 154, 156, 158, 165, 166, 167, 168, 169, 172, 173,

244 178, 181, 182, 192, 209, 216, 221, 226, 227, 238, 239 annual accumulation rate, 9, 146 annual layers, 7, 39, 40, 43, 97, 152, 178 Arabian Sea, 11, 154 Arctic, 35, 94, 126, 128, 129, 134, 140, 143, 147, 160, 161, 208, 214, 230, 233 Asian Monsoon, 9, 11, 15, 145, 158, 159, 161, 238 ASRL, xxv Atlantic Ocean, 11 augers, 18 Barents Sea, 133 Bay of Bengal, 155 BC, xxv, 14, 106, 107 becquerel, xxvii Belukha, 53, 55, 58 Bernese Alps, 96, 97 biomass burning, 22, 106, 201 black carbon, xxv, 106, 107 Bolivia, 8, 49, 58 BPRC, xxv, 4, 6 Bq, xxvii, 114 Byrd Polar Research Center, xi, xxv, 4, 159 carbon dioxide, xxvii Carbon-14, 27 CDC, xxv Celcius, xxvii centimeter, xxvii, 21, 39 Cerro Tapado, 41, 49, 50, 51, 57 Chago Glacier, 83 Chile, 41, 49 Chimborazo, 49, 51, 52 Chlorine-36, 26, 34 chronology, 18, 21, 22, 24, 27, 33, 128, 139, 143, 172, 173, 174, 175, 182, 192, 193, 195, 202, 204

Index CI, xxv, 134 Climate Diagnostics Center, xxv cm, xxvii, 20, 21, 41, 43, 44, 50, 51, 53, 63, 84, 86, 154, 182, 185, 186, 187, 189, 190, 192, 201, 207, 212, 222, 223, 224, 226, 228 CO2, xxvii, 134, 143 Coal, 110, 113 Col de la Brenva, 96 Col de Valsorey, 96 Col du Dôme, 96, 98, 104, 105, 109, 113 Col Major, 96 Colle Gnifetti, 17, 41, 43, 44, 57, 59, 96, 98, 99, 100, 102, 103, 106, 107, 108, 109, 117, 119 Columbia, 11, 142 continentality index, xxv, 134 corals, 3, 33 cubic kilometers, xxviii Dasuopu, 6, 8, 9, 11, 78, 146, 147, 148, 149, 150, 151, 152, 153, 154, 155, 156, 157, 159 Dasuopu ice cap, 148, 152, 155 Dating, ix, 32, 97, 117, 119, 213 delta, xxviii, 23, 65, 165, 193, 240 delta deuterium, xxviii, 240 Delta Notation, 23 delta oxygen, 165 Denmark, 126 Densification, 220 deuterium excess, 46, 50, 51, 53, 54, 55, 148, 152, 238 Dhaulagiri Himal, 83 diffusion, 39, 58 Dinwoody Glacier, 184 DOE, xxv Dongkemadi Glacier, 169 drilling, xix, xxi, 4, 12, 18, 19, 35, 40, 41, 42, 43, 44, 49, 51, 53, 54,

Index

245

55, 56, 57, 58, 96, 97, 115, 116, 117, 183, 185, 187, 236 drills, 18, 19, 187 Dunde, 6, 9, 11, 17, 35, 146, 147, 148, 149, 150, 151, 154, 155, 156, 157, 159, 165, 170, 179, 180 dust, 4, 6, 7, 8, 9, 14, 21, 24, 33, 49, 61, 67, 70, 75, 76, 78, 79, 81, 87, 89, 90, 92, 93, 94, 98, 101, 102, 103, 113, 116, 117, 119, 147, 160, 161, 186, 192, 193, 199, 200, 201, 202, 204, 213, 238, 239

Far East Rongbuk Glacier, 83, 88 Ferghana Valley, 75 Fiescherhorn Plateau, 43, 45, 46 Fieschersattel, 96 firn cores, 41, 57 firn/ice transition, 44, 192 fjord, 136 Folgefonni, 124, 125, 126, 132, 135 forest fires, 22, 24, 33 fractionation, 23, 40, 46, 55, 57, 151, 152 France, 57, 102, 104, 107, 213

East Africa Rift Valley, 10 EC, xxv, 106, 107 ECLIPSE, 63 ECM, xxv, 21, 182, 185, 187, 188, 192, 193, 194, 195, 196, 197, 198, 199, 200, 201, 202, 204, 205, 211, 212 Ecuador, 49, 197 EDA, xxv, 182, 186, 193, 198 El Niño, xx, xxv, 3, 11, 12, 14, 15, 41, 49, 160, 172, 174, 175, 176, 178, 179 ELA, xxv, 124, 136 electrical-conductivity measurements, xxv, 21 elemental carbon, xxv, 106, 107, 116, 237 energy balance, 30, 34, 131, 135 energy dispersion analysis, xxv, 182 ENSO, ix, xxv, 11, 13, 172, 173, 174, 176, 178, 179, 240 equilibrium line altitude, xxv, 124, 132 Europe, xx, 13, 32, 42, 90, 95, 97, 102, 103, 106, 107, 108, 116, 117, 118, 122, 142, 143, 199 Evaporation, 40 Everest, 66, 78, 83, 94

g/cm3, xxvii, 39, 192, 210 GCM, xxvi General Circulation Model, xxvi Germany, xii, 102, 104, 105, 107 Glacier National Park, 184 Global Network for Isotopes in Precipitation, xxvi, 39 Global Ocean Atmosphere Land System, xxvi Global Ocean Observing System, xxvi GNIP, xxvi, 39, 54, 57 GOALS, xxvi Gongga Massif, 66 GOOS, xxvi Grand Combin, 96, 97 Greenland, xxi, 31, 33, 35, 91, 94, 98, 133, 137, 141, 166, 198, 199, 202, 212, 213, 214, 215, 216, 232, 234 Grenzgletscher, 43, 44, 47, 48, 52, 57, 96, 98, 110, 114, 115, 116 Gulf of Bothnia, 133

excess cations, 83

Hailougou Glacier, 66

Guliya, ix, 6, 8, 9, 79, 163, 164, 165, 167, 168, 169, 170, 171, 172, 173, 174, 175, 176, 177, 178, 179, 180, 240

246 half-life, 27 hecto Pascals, xxviii, 154 Hg, 24, 182, 184, 185, 186, 187, 189, 191, 202, 203, 204, 205, 206, 207, 208, 209, 210, 211, 212, 214, 237 Hidden Valley, 83, 84, 85, 86, 87 Himalayas, 11, 42, 77, 78, 90, 91, 146, 148, 151, 152, 154 Holocene, 8, 9, 14, 15, 59, 116, 119, 136, 139, 140, 141, 142, 143, 160, 161, 215 hPa, xxviii, 154, 163 Huascarán, xix, 6, 8, 9, 14, 15, 160 hydrogen, 4, 58, 65, 78 IAEA, xxvi, 23, 39, 54, 57, 58, 78, 118, 148 IC, xxvi, 22 Ice Core Paleoclimatology Research Group, xxvi, 4 ICPRG, xxvi, 4, 11 Illimani, 49, 58 India, 8, 12, 78, 91, 154, 155 Indian Peninsula, 155 infiltration, xxi, 43, 96 Inilchek Glacier, viii, 61, 62, 64, 66, 75, 76, 78, 240 International Atomic Energy Agency, xxvi, 21 International Research Institute for Climate Prediction, xxvi Intertropical Convergence Zone, 158, 238 ion-exchange chromatography, xxvi, 22 IRI, xxvi, 154 isotope ratios (įD), 65 Italy, 102, 104, 107, 197 Jostedalsbreen, 122, 125, 127, 128, 131, 132, 137, 141 Jungfrau, 96

Index Jungfraujoch, 42, 43, 44, 45, 46, 47, 105 Ka, xxviii Kangchung Glacier, 83 Kangwure Glacier, 83, 87, 88 Khumbu Himal, 82, 83, 90, 93 Kilimanjaro, xx, 6, 8, 10, 15 kilometer, xxviii km, xxviii, 43, 53, 72, 102, 104, 110, 113, 124, 127, 132, 133, 134, 135, 199, 201, 204, 217, 231 km2, xxviii, 10, 53, 121, 124, 133, 135, 164, 167 km3, xxviii, 210 Knife Point Glacier, 198 Krakatau, 195, 197, 199, 200, 204, 208, 210, 211, 214 Kyrghyzstan. See . See Laboratory of Ice Core and Cold Regions Environment, Cold and Arid Regions Environmental and Engineering Research Institute, xxvi, 6, 159 Laki, 128 Last Glacial Maximum, xxvi, 8, 161 LGM, xxvi, 8, 137 LIA, xxvi, 24, 25, 32, 121, 124, 127, 128, 130, 132, 137, 182, 200, 201, 202, 212, 221, 230, 232, 236 LICCRE, xxvi, 6 Little Ice Age, ix, xxvi, 24, 25, 35, 79, 121, 127, 140, 143, 179, 180, 182, 214, 221, 233, 236 m, xxviii, 10, 18, 20, 42, 43, 45, 47, 49, 51, 53, 54, 62, 64, 66, 67, 71, 72, 73, 76, 82, 83, 84, 85, 91, 95, 98, 100, 114, 116, 128, 129, 133, 137, 138, 152, 164, 165, 182, 186, 187, 188, 192, 193, 197, 198, 199,

Index 200, 201, 202, 204, 207, 210, 211, 212, 217, 220, 221, 230 m/y, 220 Makalu region, 83 marine, 3, 8, 11, 90, 94, 136, 157 masl, xxviii, 63, 84, 95, 96, 103, 124, 128, 129, 131, 146, 163, 164, 165, 217 massifs, 66 Medieval Warming, 10 Mediterranean, 13, 14, 75, 119 melt horizon, 42 meltwater, xvi, xxi, 10, 18, 20, 33, 40, 41, 44, 47, 48, 52, 181, 182, 184, 196, 198, 236 mercury, 24, 35, 182, 186, 189, 204, 205, 206, 209, 210, 211, 212, 213, 214, 215, 216, 221, 233, 234, 237 metals, xx, 24, 28, 108, 109, 116, 119, 187, 212, 216, 237 meteorological, 3, 40, 42, 46, 49, 53, 61, 75, 77, 96, 122, 124, 130, 131, 132, 135, 136, 138, 145, 155, 159, 163, 177, 193, 196, 220, 227, 239, 240 meter, xxviii, 7, 41, 143, 150, 185, 187, 207, 208, 209, 210, 221 meters above sea level, xxviii, 19, 63 meters per year, xxviii, 220 methane, 98, 99, 100, 110, 116, 237 mg/L, xxviii Microbiology, 28 micrograms per liter, xxviii micrometer, xxviii, 223 microohms, xxviii Middle East, 13 mil, xxviii, 186 milligrams per liter, xxviii millimeter, xxviii, 21, 185 mm, xxviii, 21, 51, 83, 164, 167, 171, 201 Mongolian Altai, 42

247 Monsoon, 8, 11, 14, 78, 160, 161 Monsoons, 3, 8, 151 Mont Blanc, 96, 97, 98, 102, 108, 112, 118, 119 Montana, 184 Monte Rosa, 57, 58, 96, 97, 98, 102, 110, 118 Mt. St. Helens, 197, 198, 208 Mt. Xixabangma, 83 mȍ, xxviii Nangpai Gosum Glacier, 83 nanograms per liter, xxviii NAO, xxvi, 129, 134, 136, 137, 155, 156 NASA, xxvi National Aeronautics and Space Administration, xxvi National Geographic Society, 159 National Ice Core Laboratory, xxvi, 19 National Oceanic and Atmospheric Administration, xxvi National Science Foundation, xxvi, 77, 159 Nepal, 12, 19, 77, 93, 94, 161 ng/L, xxviii, 24, 186, 187, 191, 202, 203, 208 NICL, xxvi, 19, 185 NOAA, xxvi, 154 Nordend, 96 North Atlantic, xxvi, 75, 129, 135, 136, 137, 139, 142, 154, 155, 156, 158, 160, 161 North Atlantic Oscillation, xxvi, 129, 142, 155, 156, 160, 161 Northern Hemisphere, 9, 25, 148, 149, 152, 158, 160 Norway, 121, 122, 125, 126, 127, 128, 129, 130, 131, 132, 133, 134, 136, 137, 138, 139, 140, 141, 142, 143

248 Norwegian Water Resources and Energy Directorate, xxvi, 125, 139, 141 NSF, xxvi, 159 NVE, xxvi, 125, 129, 131, 134, 138, 139, 141 OAR, xxvi o C, xxvii, 220, 224, 229, 230, 232, 236 OC, xxvi, 106, 107 ocean sediments, 33 Oceanic and Atmospheric Research, xxvi Øksfjordsjøkelen, 127, 138 one-thousandth of an inch, xxviii organic carbon, xxvi, 107 oxygen, xxviii, 4, 10, 21, 23, 28, 58, 85, 94, 129, 146, 148, 151, 160, 161, 165, 192, 193, 232 Patagonia, 11 Penitents, 41 percolation, xvi, 18, 20, 33, 43, 57, 95, 96, 181, 196, 236 Peru, xix, 5, 8, 10, 14, 15, 17, 19, 36, 49, 57, 59, 160, 179, 221, 232 Plaine Morte, 44 Polar Front, 129, 134, 137 pollution, xx, 90, 91, 92, 93, 95, 103, 108, 115, 117, 201, 214, 237, 239 post-depositional effects, 41, 42, 43, 44, 49, 53, 54 ppt, xxviii, 83, 208 precipitation, xxi, xxviii, 4, 8, 9, 11, 15, 17, 21, 22, 23, 25, 26, 31, 32, 36, 39, 42, 43, 44, 45, 46, 47, 49, 51, 53, 54, 55, 57, 58, 66, 72, 77, 78, 79, 82, 83, 87, 90, 91, 92, 93, 99, 103, 113, 115, 118, 128, 132,

Index 134, 135, 136, 145, 147, 148, 149, 150, 151, 152, 154, 155, 157, 158, 159, 160, 161, 163, 165, 167, 168, 169, 170, 171, 172, 173, 174, 175, 176, 177, 178, 179, 180, 192, 200, 201, 208, 209, 211, 214, 218, 226, 227, 229, 231, 232, 234, 237, 238, 239, 240 proxy, 3, 4, 5, 8, 13, 31, 61, 75, 96, 134, 136, 139, 151, 157, 172, 178, 181, 218, 228, 240 Puruogangri, 6, 7 QC, xxvi, 191 Qilian Mountains, 146 Qori Kalis, 10 quality control, xxvi, 187 Quaternary, xii, xxii, 15, 79, 138, 141, 142, 161 Quelccaya ice cap, xix, 3, 4, 5, 10, 15 radio-echo sounding, 29, 36 Radio-echo sounding, 220, 234 Radioisotopes, 25, 26 recrystallisation, 43 refreezing, 52 Rocky Mountains, 235 Russian, 5, 42 Saharan, 98, 116, 119 Sajama, 6, 8, 9, 49 sastrugi, 41, 42 Scanning Electron Microscopy, xxvi sea level pressure, xxvii, 133, 155, 156 sea surface temperature, xxvii, 178 seasonal cycles, 45, 56, 236 SEM, xxvi, 182, 186, 193, 198, 201 Seserjoch, 96 SLP, xxvii, 155, 156 SMOW, xxvii, 23, 65, 223

Index snow-depth sensor, 218, 220, 222, 227, 231 SOI, xxvii, 172, 176 solar, 19, 30, 40, 57, 63, 96, 117, 136, 140, 221 South Pole, 91, 94, 213 Southern Oscillation Index, xxvii, 172 Soviet Union, 75, 77 square kilometers, xxviii, 164 SST, xxvii Stable Isotope, 23, 58, 65, 77, 223 stable isotopes, 4, 39, 40, 41, 47, 56, 58, 78, 79, 148, 151, 159, 161, 216, 237 Standard Mean Ocean Water, xxvii, 23, 223 sublimation, 39, 41, 46, 49, 51, 53, 54, 56, 58, 236 Summit, Greenland, 91, 94, 216 Svartisen ice cap, 126, 140 Sweden, 34, 122, 125, 126, 127, 128, 129, 132, 134, 135, 140, 143 Swiss Rhone valley, 113 Swiss Topographic Institute, 45 Switzerland, xii, xiii, 27, 42, 58, 102, 104, 107, 110, 114, 117, 118 TA, 217, 229, 230, 231, 232, 236 Tambora, 128, 195, 197, 199, 200, 204, 208, 210, 211, 214, 215 Tanggula Mountains, 148, 155, 158, 169, 238 Tapado, 49 TC, xxvii, 106, 107 tephra, 128 The Ohio State University, iii, xi, xiii, 14, 19, 159, 160 thousands of years ago, xxviii Tibet, xx, 7, 19, 66, 77, 78, 94, 150, 155, 163, 164, 169, 177, 179, 180, 240

249 Tibetan Plateau, ix, 6, 8, 9, 10, 11, 14, 15, 17, 78, 79, 82, 89, 90, 94, 145, 146, 147, 148, 149, 150, 151, 154, 155, 156, 157, 158, 159, 161, 163, 164, 165, 170, 173, 238 Tien Shan mountains, 61 till, 138 total carbon, xxvii, 106, 107 transfer function, 218, 228, 229, 230, 231, 232, 236 tree rings, 3, 25, 33, 157 tree-ring, xvi, 136, 157 troposphere, 103, 158, 196 Tsast Ula, 53, 54 UFG, xxvii, 181, 182, 184, 185, 186, 187, 192, 193, 195, 198, 199, 200, 201, 202, 204, 205, 206, 207, 208, 209, 211, 212, 217, 218, 220, 221, 222, 223, 225, 226, 227, 228, 229, 231, 232, 236, 237 Ulan Bator, 54 UNESCO, xxvii, 77 United Nations Educational, Scientific, and Cultural, xxvii United States Department of Energy, xxv United States Geological Survey, xxvii University of Aarhus, 126 University of Bergen, 129 University of Cambridge, 125 University of Copenhagen, 159 University of Maine, xii, xix, 65, 77 University of Manchester, 126 University of New Hampshire, xiii, xix, 20, 65, 77, 83, 234 University of Pennsylvania, 65, 77 University of Washington, xxii, 77 Upper Fremont Glacier, ix, x, xxi, xxvii, 19, 24, 31, 35, 181, 184, 194, 197,

250 198, 203, 214, 217, 219, 220, 222, 224, 225, 227, 228, 230, 233, 236 USGS, xxvii, 5, 185 varves (lakes), 33 VEI, xxvii, 195, 196, 197, 198, 199, 214 Vienna Standard Mean Ocean Water, 23 volcanic eruptions, 22, 24, 33, 51, 98, 182, 195, 196, 204, 211, 221 Volcanic Explosivity Index, xxvii, 195, 198, 214 V-SMOW, 23

Index Wind River Range, 181, 196, 214, 220, 230, 233 Wisconsin District Mercury Research, 185 Wisconsin District Mercury Research Lab, xxvii WMO, xxvii, 39, 54, 57, 148 World Meteorological Organization, xxvii WS, xxvii, 221, 231 Wyoming, x, xxi, 19, 36, 181, 196, 214, 217, 219, 220, 221, 222, 224, 225, 227, 228, 230, 233, 236 Xixibangma, 66

water equivalent, xxviii, 41 water insoluble organic carbon, 106 WDMRL, xxvii, 185, 186, 191 weather station, xxvii, 67, 221, 222, 231 weq, xxviii, 41, 43, 44, 45, 51, 53, 83, 84, 85, 164, 166, 182

į18O, xxviii į18O ratios, 61 į2H, xxviii ǻC, 83, 84, 87, 89 µg/L, xxviii µm, xxviii, 223

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