The Quaternary Period in the United States, Volume 1 (Developments in Quaternary Sciences)

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Developments in Quaternary Science, 1 Series editor: Jim Rose

THE QUATERNARY PERIOD IN THE UNITED STATES

by

A.R. Gillespie and S.C. Porter University of Washington Seattle Washington USA and

B.F. Atwater U.S. Geological Survey Seattle Washington USA

2004

Amsterdam – Boston – Heidelberg – London – New York – Oxford – Paris San Diego – San Francisco – Singapore – Sydney – Tokyo

Developments in Quaternary Science Series editor: Jim Rose Volumes in this series 1.

The Quaternary Period in the United States Edited by A.R. Gillespie, S.C. Porter, B.F. Atwater 0-444-51470-8 (hardbound); 0-444-51471-6 (paperback) – 2004

2.

Quaternary Glaciations – Extent and Chronology Part I Europe; Part II North America; Part III South America, Asia, Africa, Australia, Antarctica Edited by J. Ehlers, P.L. Gibbard 0-444-51462-7 (hardbound + CD-ROM) – 2004

3.

Ice Age Southern Andes – A Chronicle of Paleoecological Events Authored by C.J. Heusser 0-444-51478-3 (hardbound) – 2003

4.

Spitsbergen Push Moraines – Including a translation of K. Gripp: Glaciologische und geologische Ergebnisse der Hamburgischen Spitzgbergen-Expedition 1927 Authored by J. van der Meer Forthcoming – 2004 0-444-51544-5

5.

Tropical West Africa – Marine and Continental Changes during the Late Quaternary Authored by P. Giresse Forthcoming – 2004

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DEVELOPMENTS IN QUATERNARY SCIENCE S E R I E S E D I T O R: J I M R O S E

1

THE QUATERNARY PERIOD IN THE UNITED STATES

ELSEVIER B.V. Sara Burgerhartstraat 25, P.O. Box 211, 1000 AE Amsterdam, The Netherlands © 2004 Elsevier Ltd. All rights reserved. This work is protected under copyright by Elsevier, and the following terms and conditions apply to its use: Photocopying Single photocopies of single chapters may be made for personal use as allowed by national copyright laws. Permission of the Publisher and payment of a fee is required for all other photocopying, including multiple or systematic copying, copying for advertising or promotional purposes, resale, and all forms of document delivery. Special rates are available for educational institutions that wish to make photocopies for non-profit educational classroom use. Permissions may be sought directly from Elsevier Global Rights Department, PO Box 800, Oxford OX5 1DX, UK; phone: (+44) 1865 843830, fax: (+44) 1865 853333, e-mail: [email protected]. You may also contact Global Rights directly through Elsevier’s home page (http://www.elsevier.com), by selecting ’Obtaining Permissions’. In the USA, users may clear permissions and make payments through the Copyright Clearance Center, Inc., 222 Rosewood Drive, Danvers, MA 01923, USA; phone: (+1) (978) 7508400, fax: (+1) (978) 7504744, and in the UK through the Copyright Licensing Agency Rapid Clearance Service (CLARCS), 90 Tottenham Court Road, London W1P 0LP, UK; phone: (+44) 207 631 5555; fax: (+44) 207 631 5500. Other countries may have a local reprographic rights agency for payments. Derivative Works Tables of contents may be reproduced for internal circulation, but permission of Elsevier is required for external resale or distribution of such material. Permission of the Publisher is required for all other derivative works, including compilations and translations. Electronic Storage or Usage Permission of the Publisher is required to store or use electronically any material contained in this work, including any chapter or part of a chapter. Except as outlined above, no part of this work may be reproduced, stored in a retrieval system or transmitted in any form or by any means, electronic, mechanical, photocopying, recording or otherwise, without prior written permission of the Publisher. Address permissions requests to: Elsevier Global Rights Department, at the mail, fax and e-mail addresses noted above. Notice No responsibility is assumed by the Publisher for any injury and/or damage to persons or property as a matter of products liability, negligence or otherwise, or from any use or operation of any methods, products, instructions or ideas contained in the material herein. Because of rapid advances in the medical sciences, in particular, independent verification of diagnoses and drug dosages should be made. First edition 2004 Library of Congress Cataloging in Publication Data A catalogue record from the British Library has been applied for. ISBN: 0-444-51470-8 (Hardbound) ISBN: 0-444-51471-6 (Paperback) ISSN: 1571-0866 (Series) ∞ The paper used in this publication meets the requirements of ANSI/NISO Z39.48-1992 (Permanence of Paper). Printed in The Netherlands.

Foreword

who wish to achieve a substantial treatment of major scientific concerns and a venue for those seeking the authority provided by such an approach. We are very proud indeed that The “Quaternary Period in the United States” edited by Alan Gillespie, Steve Porter and Brian Atwater will be the first issue of the series. This book has a distinguished lineage, and will be a benchmark publication for a part of the world that forms the basis of many models used to interpret Quaternary processes and patterns of environmental change throughout the globe. The underlying aim of the book is to update our understanding of the Quaternary of the United States over the 38 years since the 1965 INQUA. With this in mind, it is interesting to consider the scope of this text. In addition to the traditional concerns with topics like glacial, aeolian, coastal, fluvial, lake, permafrost and soil processes and changes in patterns of vegetation, insects and mammals including humans, a number of topics are given special attention for the first time. On the physical side, there are chapters on outburst floods, groundwater and speleothem, palaeoseismology and volcanism, while on the biological side chapters consider vegetation dynamics, fire history, as well as a much-enlarged examination of Quaternary insects. Perhaps it is most marked by the attention given to modelling throughout, and there are notable chapters on coupling of ice sheet and climate models for the simulation of former ice sheets and modelling palaeoclimate. The editors are to be congratulated on bringing together such a vast amount of up-to-date information together within such a short time, and to make available to the Quaternary community at large, a book that will act as a reference text for many years to come.

The “Quaternary Period in the United States” edited by Alan Gillespie, Steve Porter and Brian Atwater is the first volume in the Elsevier Book Series on “Developments in Quaternary Science.” This book is to be published in 2004 to coincide with the very successful XVI INQUA Congress held in Reno, Nevada, in July 2003, and follows the well-known and wellused volume, published for the 1965 VII INQUA Congress, on “The Quaternary of the United States” edited by H. E. Wright and D. G. Frey. Developments in Quaternary Science is a book series designed to bring together important texts within the continually expanding field of Quaternary science. This series is linked to the journals: Quaternary Science Reviews, Quaternary Research and Quaternary International, and provides an outlet for texts that require individual or special consideration, not achieved most effectively within an academic journal. In this respect the series will provide an outlet for topics that require substantial space, are linked to major scientific events, require special production facilities (i.e. use of interactive electronic methods), or are archival in character. The subjects covered by the series will consider Quaternary science across different parts of the Earth and with respect to the diverse range of Quaternary processes. The texts will cover the response to processes within the fields of geology, biology, geography, climatology archaeology and geochronology. Particular consideration will be given to issues such as the Quaternary development of specific regions, comprehensive treatments of specific topics such as global scale consideration of patterns of glaciation, and compendia on timely topics such as dating methodologies, environmental hazards and rapid climate changes. This series will provide an outlet for scientists

Jim Rose Series Editor

v

Contents

Foreword Jim Rose

v

Preface Alan R. Gillespie, Stephen C. Porter and Brian F. Atwater

ix

The southern Laurentide Ice Sheet David M. Mickelson and Patrick M. Colgan

1

The Cordilleran Ice Sheet Derek B. Booth, Kathy Goetz Troost, John J. Clague and Richard B. Waitt

17

Controls, history, outbursts, and impact of large late-Quaternary proglacial lakes in North America James T. Teller

45

Pleistocene glaciations of the Rocky Mountains Kenneth L. Pierce

63

Quaternary alpine glaciation in Alaska, the Pacific Northwest, Sierra Nevada, and Hawaii Darrell S. Kaufman, Stephen C. Porter and Alan R. Gillespie

77

Coupling ice-sheet and climate models for simulation of former ice sheets Shawn J. Marshall, David Pollard, Steven Hostetler and Peter U. Clark

105

Permafrost process research in the United States since 1960 Bernard Hallet, Jaakko Putkonen, Ronald S. Sletten and Noel Potter Jr.

127

Quaternary sea-level history of the United States Daniel R. Muhs, John F. Wehmiller, Kathleen R. Simmons, and Linda L. York

147

Western lakes Larry Benson

185

Isotopic records from ground-water and cave speleothem calcite in North America Jay Quade

205

Rivers and riverine landscapes David R. Montgomery and Ellen E. Wohl

221

Landscape evolution models Frank J. Pazzaglia

247

Eolian sediments Alan J. Busacca, James E. Begét, Helaine W. Markewich, Daniel R. Muhs, Nicholas Lancaster and Mark R. Sweeney

275

Soils and the Quaternary climate system Milan J. Pavich and Oliver A. Chadwick

311

Earthquake recurrence inferred from paleoseismology Brian F. Atwater, Martitia P. Tuttle, Eugene S. Schweig, Charles M. Rubin, David K. Yamaguchi and Eileen Hemphill-Haley

331

vii

viii Quaternary volcanism in the United States William E. Scott

351

Late-Quaternary vegetation history of the eastern United States Eric C. Grimm and George L. Jacobson Jr.

381

Quaternary vegetation and climate change in the western United States: Developments, perspectives, and prospects Robert S. Thompson, Sarah L. Shafer, Laura E. Strickland, Peter K. Van de Water and Katherine H. Anderson

403

Results and paleoclimate implications of 35 years of paleoecological research in Alaska Patricia M. Anderson, Mary E. Edwards and Linda B. Brubaker

427

Quaternary history from the U.S. tropics Sara Hotchkiss

441

Climatically forced vegetation dynamics in eastern North America during the late Quaternary Period Thompson Webb III, Bryan Shuman and John W. Williams

459

Holocene fire activity as a record of past environmental change Cathy Whitlock and Patrick J. Bartlein

479

Interannual to decadal climate and streamflow variability estimated from tree rings David W. Stahle, Falko K. Fye and Matthew D. Therrell

491

Quaternary Coleoptera of the United States and Canada Allan C. Ashworth

505

Vertebrate paleontology S. David Webb, Russell W. Graham, Anthony D. Barnosky, Christopher J. Bell, Richard Franz, Elizabeth A. Hadly, Ernest L. Lundelius Jr., H. Gregory McDonald, Robert A. Martin, Holmes A. Semken Jr. and David W. Steadman

519

Peopling of North America David J. Meltzer

539

Modeling paleoclimates Patrick J. Bartlein and Steven W. Hostetler

565

Preface

today. The book begins with sections on ice and water – as glaciers, permafrost, oceans, rivers, lakes, and aquifers. Six chapters are devoted to the high-latitude Pleistocene ice sheets, to mountain glaciations of the western United States, and to permafrost studies. Other chapters discuss ice-age lakes, caves, sea-level fluctuations, and riverine landscapes. With a chapter on landscape evolution models, the book turns to essays on geologic processes. Two chapters discuss soils and their responses to climate, and wind-blown sediments. Two more describe volcanoes and earthquakes, and the use of Quaternary geology to understand the hazards they pose. The next part of the book addresses plants and animals. Five chapters, organized mainly by region, consider the Quaternary history of vegetation in the United States. Other chapters treat forcing functions and vegetation response at different spatial and temporal scales, the role of fire as a catalyst of vegetation change during rapid climate shifts, and the use of tree rings in inferring age and past hydroclimatic conditions. Three chapters address vertebrate paleontology and the extinctions of large mammals at the end of the last glaciation, beetle assemblages and the inferences they permit about past conditions, and the peopling of North America. A final chapter addresses the numerical modeling of Quaternary climates, and the role paleoclimatic studies and climatic modeling has in predicting future response of the Earth’s climate system to the changes we have wrought. The resulting volume was made possible through the combined efforts of authors, reviewers, and a responsive production team. We are especially indebted to Karin StewartPerry. We hope the book stands alongside its predecessors as an enduring and useful resource for investigators of the Quaternary Period.

This book is devoted to advances in understanding of the past ca. two million years of Earth history – the Quaternary Period – in the United States. It is being published in the year of INQUA XVI, the first Congress of the International Union for Quaternary Research to meet in the U.S. since INQUA VII, in 1965. The intervening 38 years has brought enormous growth in Quaternary research, as illustrated by journals dedicated to it. In 1965, there was one such journal, Japan’s Daiyonki Kenkyu (Quaternary Research). Its U.S. namesake, Quaternary Research, appeared in 1970, followed by Quaternary Science Reviews in 1982, The Journal of Quaternary Science in 1985, Quaternary International in 1989, and The Holocene in 1990. “The Quaternary Period in the United States” has ample precedent. For INQUA VII, Herbert E. Wright, Jr. and David G. Frey edited a review volume on the Quaternary of the United States that has served as a standard reference for nearly four decades (Wright & Frey, 1965). Nearly 20 years later, Steve Porter and Herb Wright edited updated reviews on Late Quaternary Environments of the United States (Porter, 1983; Wright, 1983). Vladimir Sibrava, David Q. Bowen and Gerald M. Richmond then surveyed Quaternary glaciation in the Northern Hemisphere (Sibrava et al., 1986), and the Geological Society of America’s Decade of North American Geology (DNAG) series delved into geologic aspects of Quaternary geologic and geomorphic history in detail. “The Quaternary Period in the United States” is the first general overview since the DNAG series appeared in the 1980s. No review volume can provide comprehensive treatment for a topic as vast and rich as the Quaternary. We therefore had to limit this book to highlights from an admittedly incomplete selection of fields that are of widespread current interest Alan R. Gillespie, Stephen C. Porter & Brian F. Atwater

Seattle, March, 2003

References Wright, H.E., Jr. (ed.) (1983). The Late Quaternary of the United States, Vol. 2, The Holocene. Minneapolis, University of Minnesota Press, 277 pp. Wright, H.E., Jr. & Frey, D.G. (1965). The Quaternary of the United States. New Jersey, Princeton University Press, 922 pp.

Porter, S.C. (ed.) (1983). Late Quaternary environments of the United States, Vol. 1, The Late Pleistocene. Minneapolis, University of Minnesota Press, 407 pp. Sibrava, V., Bowen, D.Q. & Richmond, G.M. (eds) (1986). Quaternary glaciations in the northern hemisphere. Quaternary Science Reviews, 5, 534 pp.

ix

The southern Laurentide Ice Sheet David M. Mickelson1 and Patrick M. Colgan2 1

Department of Geology and Geophysics, University of Wisconsin Madison, Weeks Hall, 1215 West Dayton Street, Madison, WI 53706-1692, USA; [email protected] 2 Department of Geology, Grand Valley State University, 1 Campus Drive, Allendale, MI 49401, USA; [email protected] more than four glacial and interglacial episodes during the Pleistocene. Much work remains to be done to unravel this continental record of early glaciations and correlate it to the ocean record. There has been debate about the extent of the early Wisconsin Glaciation as well. In many areas, particularly in northern Illinois and southern Wisconsin, deposits thought to be of this age now appear to be older. Probably one reason for our rather poor understanding of pre-late Wisconsin glacial events is the decline of purely stratigraphic studies from the 1980s through the 1990s. State geologic surveys have in many cases reduced their staff and mapping has been displaced by topical studies and an emphasis on applied research related to groundwater and mineral extraction. In academia, traditional mapping and glacial stratigraphy have not been as common in the last 40 years as previously. Instead they have been replaced by, or combined with, studies of sediment genesis, glacial process, development of conceptual and quantitative models, and details of local chronology, and integrative studies of glacial deposits and other aspects of Quaternary history such as ice-marginal lakes, paleoclimatology, loess, soils, and the paleontologic record. In the 1990s the U.S. Geological Survey began to fund mapping projects in academia (EDMAP) and state geologic surveys (STATEMAP), and this has revitalized mapping of glacial deposits in the northern U.S. Since the late 1990s the Great Lakes Mapping Coalition, a joint effort of the U.S. Geological Survey and several state geologic surveys, has focused on detailed three-dimensional mapping that includes subsurface investigation. If this program continues to grow, it may re-stimulate interest in mapping glacial deposits. Much of our improved understanding of the southern LIS has come from studies of the new field of paleoglaciology. Reconstruction of ice-sheet surfaces, interpretation of former bed conditions and discussions of sliding vs. subglacial deforming beds, estimates of sediment fluxes, interpretations of the nature and distribution of subglacial meltwater, new interpretations of landform genesis, and modeling have all been major areas of research in the last 40 years. There have been several extensive compilations of the glacial record of the southern LIS since 1965 and we make no attempt to repeat these here. Instead, we highlight what we view as advances in our understanding of the southern LIS and its deposits since publication of the Quaternary of the United States in 1965. Events and processes along the southern margin of the LIS are closely tied to the behavior of the ice sheet in Canada, but much of this literature is not discussed here because of the scope of the book. Likewise, we refrain from discussing ancillary topics, like loess and

Changing Attitudes and Approaches The publication of the Quaternary of the United States (Wright & Frey, 1965) for the 1965 INQUA Congress in Denver was a milestone that summarized our knowledge of the Quaternary of the U.S. in a single volume. Glacial geology was a major component of the volume, and it contained 125 pages on the Laurentide Ice Sheet (LIS) in the U.S. In the present volume, almost 40 years later, many fewer pages are devoted to the same topic, indicating the vast increase in other aspects of Quaternary studies. In the U.S., glacial geology has expanded greatly in knowledge and interest, and now glacial geologists have a much richer field and variety of techniques with which to study Quaternary history. The other chapters in this book are a clear indication of the diversity of fields that now make up what traditionally was classified as glacial geology or did not exist before 1965. Radiocarbon dating remains the most important tool for determining the chronology of the last glaciation. Accelerator mass spectrometry (AMS) has allowed dating of smaller and somewhat older samples. Tree-ring calibration of the radiocarbon time scale has resulted in dating accuracy not possible 40 years ago. Newer dating methods such as thermoluminescence, amino-acid racemization, paleomagnetism, and cosmogenic-isotope methods have yielded mixed results, but have the potential to improve our interpretations of glacial chronology, especially those of the pre-late Wisconsin. There have been revolutionary changes in the way we study glacial sediments and reconstruct their depositional environments. Genetic classifications have been replaced by descriptive lithofacies approaches, which focus on modern-process analogs for interpretations of depositional environments. Correlations of till units from one area to another are now approached with more caution, and the use of facies models facilitates the understanding of complex glacial sequences. Geophysical techniques have been used to explore lake basins and subsurface stratigraphy. Models of glacial landform genesis are also driven by modern analogs and interpretations made in modern glacier settings. There have been major changes in our understanding of pre-late Wisconsin events since 1965. The terms “Nebraskan” and “Kansan” are no longer used, and now there is evidence of at least six pre-Illinoian glaciations in the continental record. Flint (1971), in a widely used textbook, hinted that there might be problems with correlations of what was called “Nebraskan Drift,” but nevertheless used the terminology accepted at that time. Based on the oxygen-isotope record from the oceans, we now know that there were clearly DEVELOPMENTS IN QUATERNARY SCIENCE VOLUME 1 ISSN 1571-0866 DOI:10.1016/S1571-0866(03)01001-7

© 2004 ELSEVIER B.V. ALL RIGHTS RESERVED 1

2

D.M. Mickelson & P.M. Colgan

the history of ice-marginal lakes, because these are covered elsewhere in the volume.

Mapping and Compilations of Glacial Geology and Geomorphology Since 1965 Since 1965 there have been several comprehensive reviews of the geology and geomorphology of the southern LIS. The “state-of-knowledge” was summarized by Flint in 1971 in part of his classic text. At this time, the deep marine record was only beginning to be discovered. Our understanding of modern ice-sheet dynamics in Greenland and Antarctica was in its infancy, and an understanding of surging glaciers and modern glacial environments was just emerging. Records of global climate change such as ice cores, pollen databases, loess records, and lake cores were fragmentary, few, and far between. Much of what was known was based on the incomplete terrestrial record of continental and mountain glaciation. Throughout the 1970s and 1980s numerous records of the last glaciation were collected, analyzed, and combined into

a global database (e.g. CLIMAP, 1976, 1984; COHMAP, 1988). For the INQUA meeting in Moscow, the U.S. INQUA Committee produced two volumes of edited papers on the Quaternary of the U.S. (Wright, 1983). These include a comprehensive review of the glacial record and a chronology of glacial and periglacial events during the late Wisconsin glaciation (Mickelson et al., 1983). Included were maps showing the nature of the glacier bed, moraines and other icemargin positions, and a generalized map of landform regions. Andrews (1987) summarized major issues in understanding the whole LIS: the thickness of the ice sheet, the extent of ice during the mid-Wisconsin (marine oxygen isotope stage (OIS) 3), the timing of the late-glacial maximum, and the chronology of deglaciation. The publication Quaternary Glaciations in the Northern Hemisphere (Sibrava et al., 1986), which contains several review papers on deposits of the LIS in the United States (Brown et al., 2001; Eschman & Mickelson, 1986; Fullerton, 1986; Fullerton & Colton, 1986; Hallberg, 1986; Hallberg & Kemmis, 1986; Johnson, 1986; Lasemi & Berg, 2001; Matsch & Schneider, 1986; Stone & Borns, 1986), is the most recently published compilation covering all of the area of the LIS in the United States. Dyke

Fig. 1. A shaded relief image created from digital elevation data (USGS ETOPO5) showing southern limits of selected ice advances, location of Great Lakes, and names of states (CT, Connecticut; IA, Iowa; IL, Illinois; IN, Indiana; KS, Kansas; MA, Massachusetts; ME, Maine; MI, Michigan; MN, Minnesota; MO, Missouri; ND, North Dakota; NE, Nebraska; NH, New Hampshire; NY, New York; OH, Ohio; PA, Pennsylvania; RI, Rhode Island; SD, South Dakota; WI, Wisconsin). MSL is location of isotope record of Dorale et al. (1998). Scale in km is given in Fig. 2.

The southern Laurentide Ice Sheet 3 & Prest (1987) and Dyke et al. (2002) summarize what is currently known about the extent and timing of the entire LIS during the last glacial maximum (LGM). A major contribution to our understanding of regional aspects of the glacial record is the compilation of Quaternary geologic maps at 1:1,000,000 scale for all of the area. Organized by G.M. Richmond and D.S. Fullerton, these maps, with many authors, are published as U.S. Geological Survey Miscellaneous Investigations Series I-1420. All of the southern LIS is covered by these maps, and although they were in most cases compiled from older mapping, they represent the most up-to-date regional maps available. Soller (1998) and Soller & Packard (1998) published a map of the thickness and character of surficial deposits across the area of the U.S. covered by the LIS. It portrays very thick deposits filling pre-late Wisconsin valleys and in interlobate areas. Many states have mapping programs that have added to our knowledge of glacial deposits and several different map scales are being used. Much of this mapping has been driven by the need for geologic information to help solve environmental problems or to help locate groundwater and mineral resources. Many states (Fig. 1) also have developed specific derivative maps (e.g. contamination potential, aggregate resources) based on maps of Quaternary deposits. Almost all states have glacial or Quaternary geology maps at 1:500,000 or smaller. A summary of recently active mapping programs is shown in Table 1. Massachusetts and Connecticut (Fig. 1) have been mapped mostly by the U.S. Geological Survey in cooperation with state agencies, and many of these maps are available at a scale of 1:24,000 as open file reports. Several states have had more limited recent mapping of glacial deposits (Vermont, New Hampshire, Michigan, Iowa, and South Dakota). The U.S. Geological Survey in cooperation with Ohio, Indiana, Illinois, and Michigan recently formed the Central Great Lakes Geologic Mapping Coalition. They are sharing resources to produce detailed (1:24,000) three-dimensional maps of Quaternary deposits in the southern Great Lakes area. This effort is reinvigorating interest in mapping of glacial deposits in the area, and in the future it may extend across all of the area covered by the LIS. Table 1. Typical map scales used in states with recent published mapping and a representative reference to each. Illinois Indiana Maine Minnesota New Jersey New York North Dakota Ohio Vermont Wisconsin

1:100,000 or 1:24,000 1:24,000 1:24,000 1:100,000 1:24,000 1:250,000 1:125,000 or 1:250,000 1:62,500 1:24,000 1:100,000

Curry et al. (1997), Grimley (2002) Brown & Jones (1999) Thompson (1999) Hobbs (1995) Stanford et al. (1998) Muller & Cadwell (1986) Harris & Luther (1991) Totten (1988) DeSimone & Dethier (1992) Clayton (2001)

New Data Sources During the last four decades the number of new data sources with which to study Quaternary landforms and sediments has increased greatly. These include better topographic maps, digital raster graphics, digital elevation models, remotely sensed data acquired by satellite-based systems, and improved drilling and geophysical techniques. Most of the advances in this area have been driven by advances in satellite technology, low-cost computing power, the increased availability of geographic information systems (GIS), and the efforts of government agencies in creating large spatial databases. A new series of 1:24,000 U.S.G.S topographic quadrangle maps has made it possible to compare landforms over all of the glaciated United States. These maps are now available as digital raster graphics (DRG) for the conterminous United States, Hawaii and portions of Alaska. By far the most important new data source is the digital elevation model (DEM). These arrays of elevations have revolutionized the way topography can be visualized, categorized, and analyzed. The first 3 arc-second DEMs (1 × 1 degree) were created from the processing of 1:250,000 scale topographic maps and, as a result, had relatively low spatial resolution (90 × 90 m), and contained numerous processing artifacts. The most recent series of 30- and 10-meter DEMs have improved on these problems. DEMs allow classification of slope, aspect, and many other parameters of interest to Quaternary scientists. Images derived from the raw DEMs allow the creation of visually stunning shaded-relief views, contour maps, and orthographic perspective views with overlying drapes of color-coded contour intervals or satellite images. Fairly subtle geomorphic features (such as small moraines and flutes) have been discovered from these DEMs. They have also allowed recognition of regional patterns by allowing areas much larger than a single 7.5-minute quadrangle to be viewed at once. An early example of a shaded relief image is the Digital Shaded Relief Map of the Conterminous United States (Thelien & Pike, 1989), and a smaller scale one is portrayed in Fig. 1. In the last 10 years numerous workers have used DEMs for mapping glacial features. Satellite images such as those from the Landsat Multispectral Scanner and Thematic Mapper have been used in a few studies (e.g. Boulton & Clark, 1990), but their use has been limited because of their relatively low spatial resolution (∼80 and 30 m respectively) compared to traditional low-altitude aerial photography (25 m) include multiple

22

Table 1. Sources of age data for Puget Lowland stratigraphic units. Type Section Location

Reference for Nomenclature

Reported Age (in 103 Years)

Type of Date

Location

Reference for Age

Comment

Sumas glaciation

Near Sumas, Canadian side

Armstrong (1957)

na

na

na

na

na

Easterbrook (1963)

11.3–10.0 and pre 11.9

14 C

yr B.P.

Clague et al. (1997)

New dates

11.5–10.0

14 C

yr B.P.

Aldergrove–Fort Langley and Chilliwack R. valley, B.C. Multiple locations, Fraser Lowland and Nooksack valley

Kovanen & Easterbrook (2001), Kovanen (2002)

Compilation and new dates; table of 69 dates on the Sumas interval

Armstrong et al. (1965)

na

na

na

na

na


13.0–11.0 13.0–11.5

14 C 14 C

yr B.P. yr B.P.

13.6–11.3

14 C

yr B.P.

Armstrong et al. (1965) Kovanen & Easterbrook (2001) Dethier et al. (1995)

New dates Compilation and new dates New dates

na

Includes the Kulshan glaciomarine drift, Deming Sand, and Bellingham glaciomarine drift

Sumas Drift

Sumas stade Everson Glaciomarine Drift, Everson interstade

Vashon till, Vashon glaciation

Upstream of Everson, on the Nooksack R.

Southeast of Cedarville on the Nooksack R.

Easterbrook (1963)

na

na

Type section Whidbey Is. to Campbell R. Northern Puget Lowland na

Vashon Island

Willis (1898)

>13.5

na

na

Rigg & Gould (1957)

Youngest limiting age


25.0–13.5

14 C

yr B.P.


New dates

18.0–13.0

14 C

yr B.P.

Multiple, Strait of Georgia to Lake Washington Fraser Lowland

Compilation

16.0–13.5

14 C

yr B.P.

Seattle, Bellevue, Issaquah

Kovanen & Easterbrook (2001) Porter & Swanson (1998)

Vashon Drift, Vashon stade

Compilation and new dates

Steilacoom Gravel

Steilacoom plains

Willis (1898), Bretz (1913), Walters & Kimmel (1968)

Younger than 13.5

14 C

yr B.P.

Ft. Lewis, Tacoma

Borden & Troost (2001)

Multiple, young, sub–Vashon dates

Esperance Sand Member of Vashon Drift

Fort Lawton, Seattle

Mullineaux et al. (1965)

15.0–13.5; 15.0–14.5

14 C

yr B.P.

Seattle; Issaquah

Mullineaux et al. (1965), Porter & Swanson (1998)

Limiting ages

Lawton Clay Member of Vashon Drift

Fort Lawton, Seattle


15.0–13.5; 15.0–14.5

14 C

yr B.P.

Seattle; Issaquah

Mullineaux et al. (1965), Porter & Swanson (1998)

Limiting ages

Derek B. Booth et al.

Name (Climatic Intervals in Italics)

Port Moody nonglacial deposits

Port Moody

Hicock et al. (1982)

23.0–21.0

14 C

yr B.P.

Port Moody interstade Coquitlam Drift

na Coquitlam–Port Moody

Coquitlam stade

Evans Creek Drift, Evans Creek stade

Olympia interglaciation

Carbon River valley, near mouth of Evans Creek

Fort Lawton

Olympia beds

Hicock & Armstrong (1981)

New dates

Hicock & Armstrong (1985) Hicock & Armstrong (1981)

Interstade informally introduced New dates

Compilation of 52 dates

Hicock (1976)

21.7–18.7

14 C

yr B.P.

Type section

Hicock & Armstrong (1985)

30.0–25.0

14 C

yr B.P.

Multiple locations

26.0–17.8

14 C

yr B.P.

Multiple locations

Hicock & Armstrong (1985) Clague (1980), Armstrong et al. (1985)

Crandell (1963)

25.0–15.0

14 C

yr B.P.

Type section


Armstrong et al. (1965) (Crandell)

na

na

na

na


35.0–15.0

14 C

yr B.P.

Fort Lawton and multiple locations in WA and BC

Armstrong et al. (1965), Troost (1999)

24.0–15.0

14 C

yr B.P.

Fort Lawton and West Seattle


Minard & Booth (1988)

>45–13.5

14 C

yr B.P.

Multiple locations around Seattle and Tacoma

Troost (1999), Borden & Troost (2001)

New dates

Equivalent to Evans Creek stade?

Alpine glaciation in Cascade Range na

Compilation and new dates; may be partly equivalent to Quadra sediments at Point Grey in Vancouver (Armstrong & Brown, 1953) New dates

Possession Point, Whidbey Island

Easterbrook et al. (1967)

80

Amino acid

Multiple locations

Easterbrook & Rutter (1981)

New dates

Whidbey Formation

Double Bluff, Whidbey Island


107–96, avg = 100 151–102

Amino acid

Multiple locations

Easterbrook & Rutter (1981) Easterbrook (1994)

New dates

Thermo– luminescence

The Cordilleran Ice Sheet 23

Possession Drift

24

Name (Climatic Intervals in Italics)

Type Section Location

Reference for Nomenclature

Reported Age (in 103 Years)

Type of Date

Location

Reference for Age

Comment

Double Bluff Drift

Double Bluff, Whidbey Island


250–150

Amino acid

Type section

New dates

178–111 291–177

Amino acid Thermo– luminescence

Easterbrook & Rutter (1982) Blunt et al. (1987) Easterbrook et al., 1992

na na

Salmon Springs Drift

Near Sumner

Crandell et al. (1958)

1000

Inferred, based on Lake Tapps

Type section

Easterbrook (1994)

Reversely magnetized (Easterbrook, 1986)

Lake Tapps Tephra

Near Sumner

Crandell (1963), Easterbrook & Briggs (1979)

840

Fission track

3 locations

Easterbrook & Briggs (1979)

1000

Fission track

Multiple locations

Westgate et al. (1987)

Correlation of other locations to type section based on chemistry na

Willis (1898)

1690–1640

Laser–argon

Type section


New date; reversely magnetized (Easterbrook, 1986)


na

na

na

na

na

Puyallup interglaciation, Puyallup Sand

Near Alderton

Puyallup Formation Stuck Drift

Near Alderton


Close to 1600

Based on bounding ages

Type section

Easterbrook (1994)


Alderton Formation

Near Alderton


2400–1000, avg = 1600

Laser–argon

Type section

Easterbrook (1994)


Orting Gravel, Orting Drift

Orting

Willis (1898), Crandell et al. (1958)

2000 (?)

Inferred

Type section

Easterbrook (1986), Easterbrook et al. (1988)



Table 1. (Continued )


Fig. 5. Locations of type sections for the recognized pre-Fraser stratigraphic units in the Puget Lowland. Locations of cross section of Fig. 6 and measured sections in Fig. 7 are also shown. The Olympia nonglacial interval was first defined by Armstrong et al. (1965) with its type section at Fort Lawton (Mullineaux et al., 1965). The Possession Drift, Whidbey Formation, and Double Bluff units were named and described by Easterbrook et al. (1967, 1981). The Salmon Springs and older drifts were first described by Willis (1898) and formally named by Crandell et al. (1958).

26


Fig. 6. East-west cross section through Commencement Bay near Tacoma, showing radiocarbon dates and topographic relief within the Olympia beds (unit Qob). Reversely magnetized nonglacial volcanic-rich deposits yield a zircon fission-track age of 1.1 × 106 years (modified from Troost et al., 2003). Unlabeled numbers are 14 C ages in 103 14 C yr B.P. tephra, lahar, peat, and diatomite layers (Troost et al., 2003). At least five discontinuous Olympia-age tephras and lahars have been identified near Tacoma, with source areas including Mt. St. Helens and Mt. Rainier. Freshwater diatomites and in situ tree roots reveal lacustrine and forested environments across the lowland. Mastodons, mammoths, and bison roamed the Puget and Fraser lowlands during this nonglacial interval (Barton, 2002; Harrington et al., 1996; Plouffe & Jette, 1997). The next-oldest Pleistocene sediment in the Puget Lowland is the Possession Drift, probably related to glaciation during MIS 4 (Easterbrook, 1994) (Fig. 7a). The ice sheet responsible for this drift may have been less extensive than during MIS 2, according to reconstructions of global temperature. Away from the type section on Whidbey Island, pre-Fraser glacial deposits cannot be uniquely correlated with Possession Drift without age control. Thermoluminesence dating may prove most useful in this age range (Easterbrook, 1994), with preliminary results suggesting localities of Possession-age outwash south of the type section (Easterbrook, 1994; Mahan et al., 2000). The Whidbey Formation and its counterpart in British Columbia, the Muir Point Formation, correlate with MIS

5, the youngest full interglacial interval of the Pleistocene record. Climate was similar to that of today, with sea level perhaps slightly above today’s level and as much as 60 m lower (Easterbrook, 1994; Easterbrook et al., 1967). At its type section (Easterbrook, 1994) (Fig. 7b), the Whidbey Formation includes silt, sand, gravel, ash, and diatomite. On Whidbey Island, extensive sand deposits may be deltaic in origin. Like Olympia nonglacial deposits, sedimentary layers surely vary in thickness and composition over short distances; relief on the upper surface of the Whidbey Formation probably resembles today’s landscape relief. Difficulties in dating sediments of this age, however, provide few constraints on the paleotopography from this time. Still older mid- and early-Pleistocene deposits in the Puget Lowland include the Double Bluff Drift (Easterbrook, 1994) (Fig. 7b) and various unnamed glacial and interglacial deposits in the interval from 250,000 to 780,000 years ago, the existence of which are anticipated from climatic fluctuations expressed by the marine isotope record. Recent chronological and stratigraphic correlation efforts have begun to identify deposits in this age range and to confirm the presence of pre-Fraser deposits at locations away from their

−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−→ Fig. 7. Measured sections at pre-Fraser localities on and near Whidbey Island and in the Puyallup River valley (reproduced from Easterbrook, 1994), and at Fort Lawton in Seattle. Fig. 7a depicts both the Possession Drift and the Whidbey Formation at Point Wilson. Fig. 7b shows the lithologies noted at the type locality of the Double Bluff Drift. Fig. 7c depicts the stratigraphic relationships between the Puyallup Formation, Stuck Drift, and Alderton Formation at the Alderton type locality; black dots depict reversely polarized samples. Fig. 7d shows the modern exposure at the type locality for deposits of the Olympia nonglacial interval, and for the Lawton Clay Member and Esperance Sand Member (the latter now generally mapped as Vashon advance outwash) of the Vashon Drift (Mullineaux et al., 1965).


Fig. 7.

28


type sections (Hagstrum et al., 2002; Mahan et al., 2000; Troost et al., 2003). The oldest pre-Fraser deposits, about 1 million years old and older, are the Salmon Springs Drift, Puyallup Formation, Stuck Drift, Alderton Formation, and Orting Drift (Crandell, 1963; Westgate et al., 1987) (Fig. 7c).

deposits. One such gravel was dated to between 200,000 and 400,000 Th/U yr ago and another to before 780,000 Th/U yr ago (Bjornstad et al., 2001).

Chronology of the Fraser Glaciation Eastern Washington Discontinuous drift extending beyond the limits of Fraser-age drift in the Pend Oreille, Columbia, and Little Spokane valleys has stones that are highly weathered or deeply penetrated by cracks, has a slightly argillic soil, and overlies granite and gneiss bedrock that is highly decayed, even to grus. These characteristics indicate that the drift is pre-Fraser in age. Direct dating of pre-Fraser sediments is poor, but radiocarbon dates in Canada have been interpreted as denying the existence of an ice sheet between 65,000 and 25,000 yr B.P. (Clague, 1980), consistent with nonglacial conditions west of the Cascade Range during this time. The weathering of the drift and surrounding bedrock in places is so strong as to suggest an age very much older than late Wisconsin – equivalent to MIS 6 (160,000–130,000 years ago) or older. In northeastern Washington and adjacent Idaho, however, there is no objective basis for Richmond’s (1986, Chart 1) assignment of any of these deposits to particular time intervals. Probably there were several pre-Wisconsin Cordilleran ice-sheet glaciations in eastern Washington and farther east in Idaho and Montana. Glacial Lake Missoula and great floods from it are possible only when the Purcell Trench lobe advances far enough south (to 48◦ 10′ N) to dam the Clark Fork of the Columbia. In southern Washington, deposits resembling Fraser-age Missoula-flood gravel bars but thickly capped by calcrete deeply underlie some of these Fraser Fig. 8. Growth of the Cordilleran Ice Sheet during the Fraser glaciation (from Clague, 1981).

The Cordilleran Ice Sheet most recently advanced out of the mountains of British Columbia about 25,000 14 C yr B.P. It flowed west onto the continental shelf, east into the intermontaine valleys of British Columbia where it probably merged with the western edge of the Laurentide Ice Sheet, and south into the lowlands of Washington State (Fig. 8, Table 1). In southern British Columbia and western Washington the Puget lobe filled the Fraser Lowland and the Puget Lowland between the Olympic Mountains and Cascade Range. The Juan de Fuca lobe extended east along the Strait of Juan de Fuca to terminate some 100 km west of Washington’s present coast. Several ice lobes east of the Cascade Range expanded south down the Okanogan Valley and down other valleys farther east. The Fraser-age ice-sheet maximum on both sides of the Cascade Range was broadly synchronous (Waitt & Thorson, 1983). It approximately coincided with the maximum advance of some parts of the Laurentide Ice Sheet in central North America at about 14,000 14 C yr ago. but lagged several thousand years behind the culminating advance of most of the Laurentide Ice Sheet (Lowell et al., 1999; Mickelson et al., 1983; Prest, 1969).

Northern Puget Lowland/Southern Fraser Lowlands The Fraser glaciation began about 25,000 14 C yr B.P. with the expansion of alpine glaciers in the Coast Mountains of British

The Cordilleran Ice Sheet 29 Columbia, the Olympic Mountains, and the Cascade Range of Washington. Glaciers in the Coast Mountains coalesced to form piedmont ice lobes that reached the Fraser Lowland of British Columbia about 21,000 14 C yr B.P. during the Coquitlam stade (Hicock & Armstrong, 1981). The Coquitlam stade correlates with the Evans Creek stade of Washington, an early alpine phase of the Fraser glaciation in the Cascade Range (Armstrong et al., 1965). The Coquitlam stade was followed by a period of climatic amelioration that lasted from about 19,000 to 18,000 14 C yr B.P. – the Port Moody interstade of Hicock & Armstrong (1985). The Port Moody interstade was in turn followed by the late Wisconsin advance of the Cordilleran Ice Sheet during the Vashon stade (Armstrong et al., 1965). The Puget lobe advanced into northern Washington about 17,000 yr B.P. (Clague, 1981; Easterbrook, 1986) and retreated rapidly from its maximum position around 14,000 yr B.P. (Clague, 1981; Easterbrook, 1986; Porter & Swanson, 1998). The Vashon stade was followed by a period of rapid and extensive glacier retreat (Everson interstade) that ended with a resurgence of the southwestern margin of the Cordilleran Ice Sheet in the Fraser Lowland about 12,000 14 C yr B.P. (Sumas stade) (Clague et al., 1997; Kovanen, 2002; Kovanen & Easterbrook, 2001). Several advances separated by brief periods of retreat apparently marked the Sumas stade. The final advance(s) occurred 11,000 14 C yr B.P. or shortly thereafter. Soon after 10,500 14 C yr B.P.,

the Cordilleran Ice Sheet rapidly disappeared from the lowlands.

Central Puget Lowland Rates of ice-sheet advance and retreat are well constrained in the central Puget Lowland. The Puget lobe advanced to the latitude of Seattle by about 14,500 14 C yr B.P. (17,590 cal yr B.P.) and to its maximum by 14,000 14 C yr B.P. (16,950 cal yr B.P.) (Porter & Swanson, 1998). The ice apparently remained near its maximum position only a few hundred years and then rapidly retreated. It retreated past Seattle by 13,600 14 C yr B.P. (16,575 cal yr B.P.) (Porter & Swanson, 1998) (Fig. 9). Glacial lakes, including Lake Russell, formed south of the retreating ice front, draining through a spillway to the Chehalis River (Bretz, 1913). The lakes coalesced into one lake, Lake Bretz (Lake Leland of Thorson, 1980), which enlarged northward as the ice front retreated until a northern spillway was uncovered. Further backwasting allowed sea water to enter the lowland from the Strait of Juan de Fuca. Glaciomarine drift and other marine deposits accumulated in the northern lowland where land had not yet rebounded from isostatic depression. This interstade – named the Everson by Armstrong et al. (1965) – ended about 12,000 14 C yr B.P. Isostatic rebound raised the glaciomarine and marine deposits above sea level between about 13,500 and 11,300 14 C yr B.P. (Dethier et al., 1995).

Fig. 9. Rates of Puget lobe advance and retreat in the Puget Lowland during the Vashon stade (modified from Porter & Swanson, 1998). Rapid advance and retreat are required to honor the limiting radiocarbon dates from Lake Carpenter, Seattle, Bellevue, and Issaquah. Maximum icesheet extent could have persisted at most a few hundred years.

30


Eastern Washington In contrast to the tight age constraints west of the Cascade Range, limits on the Fraser maximum east of the Cascades and Coast Mountains are broad. They include a date of 17,240 14 C yr B.P. for proglacial advance outwash, 100 km north of the ice limit, followed by advance to the glacier maximum, then a retreat of at least 80 km by 11,250 14 C yr B.P., judged partly on the distribution of Glacier Peak tephra layer G (Clague et al., 1980; Mehringer et al., 1984; Porter, 1978). Lake Missoula flood deposits, interbedded with varves of glacial Lake Columbia that contain detrital wood dated 14,490 14 C yr B.P., suggest that the Purcell Trench lobe blocked the Clark Fork for 2000–3000 yr and reached its maximum extent about 15,000 14 C yr B.P. (Atwater, 1986).

Sea-Level Record Changing sea levels greatly altered the shorelines of the Pacific Northwest. Variations in relative sea level, ranging from 200 m above present sea level to more than 100 m below, are the integrated result of eustasy, isostasy, and tectonism. These phenomena are difficult to assess separately, however, because eustasy and isostasy are interdependent and because the eustatic component has proven particularly difficult to quantify.

Eustasy Global Record Eustatic sea-level changes are global and are caused mainly by changes in volume of ocean water. Fluctuating continental glaciers are the most important cause of eustatic sea-level change on the time scale of concern here – sea level falls when ice sheets grow and rises when they shrink. Seawater also decreases in volume as it cools, which further lowers sea level during glaciations. The growth and decay of large ice sheets during the Pleistocene caused sea level to fluctuate by 120–140 m (Fairbanks, 1989; Lambeck et al., 2000, 2002; Peltier, 2002; Yokoyama et al., 2000). Estimates of sea-level lowering during the last glaciation (MIS 2) derive from fossil corals in Barbados, New Guinea, and Tahiti (Bard et al., 1990a, b, 1993, 1996; Chappell & Polach, 1991; Fairbanks, 1989) and from more recent sediment cores taken from the Sunda Shelf (Hanebuth et al., 2000) and Northwest Shelf of Australia (Yokoyama et al., 2000). Eustatic sea-level changes have also been estimated from variations in the oxygen-isotope composition of air in bubbles trapped in the Greenland and Antarctica ice sheets (Dansgaard et al., 1971; Epstein et al., 1970; Grootes et al., 1993; Johnsen et al., 1972; Jouzel et al., 2002; Lorius et al., 1985; Petit et al., 1999) and in foraminifera in ocean sediment (Chapman & Shackleton, 1999; Chappell & Shackleton, 1986; Lea et al., 2002; Shackleton, 1987; Waelbroeck et al., 2002). Numeric

Fig. 10. Eustatic sea-level curve based on dating of shallowwater corals at Barbados (after Fairbanks, 1989). modeling and geologic data (summaries in Clark & Mix, 2002) provide equivalent sea-level lowering of 118–130 m for the volume of ice locked in glaciers at the last glacial maximum. Eustatic sea level rose after about 18,000 14 C yr B.P. as ice sheets in the Northern Hemisphere began to decay. Sea-level rise accelerated after about 15,000 14 C yr B.P. and remained high until about 7000 14 C yr B.P. when the Laurentide Ice Sheet had largely disappeared (Fig. 10; Fairbanks, 1989). Rates of eustatic sea-level rise were exceptionally high between about 11,000 and 10,500 14 C yr B.P. and between 9000 and 8000 14 C yr B.P. After 7000 14 C yr B.P., the rate of eustatic sea-level rise sharply decreased, and by 4000 14 C yr B.P. sea level was within 5 m of the present datum.

Regional Expression It is difficult to disentangle the eustatic and glacio-isostatic components of the sea-level record in Washington and British Columbia. Isostatic depression and rebound dominate the late Pleistocene sea-level record in peripheral areas of the former Cordilleran Ice Sheet, but these effects decrease with distance beyond the ice margin. Estuaries in southwestern Washington record a mostly eustatic response, with the river valleys in this area drowned by rising sea level when ice sheets melted. In southwestern British Columbia and the northern Puget

The Cordilleran Ice Sheet 31 Lowland, in contrast, relative sea level during deglaciation was higher than at present because these areas were isostatically depressed during the last glaciation. Isostasy Global Record The growth and decay of ice sheets, and thus changes in global sea level, redistributed mass on the Earth’s surface. Ice sheets depressed the crust beneath them, but just beyond their margins the crust warped as a “forebulge” (Walcott, 1970). Melting ice sheets reversed the process: the forebulge migrated back towards the former center of loading to cause uplift there. Water transfer from oceans to ice sheets unloaded the seafloor; the opposite happened during deglaciation. These hydro-isostatic adjustments opposed the direction of glacio-isostatic adjustments. Continental shelves rose when seawater was removed and they subsided again when melting ice sheets returned water to the oceans. Regional Expression Expanding glaciers during the early part of the Fraser glaciation progressively depressed the land surface of southwestern British Columbia and northwestern Washington (Clague, 1983). This depression started beneath the Coast Mountains, where glaciers first grew. As glaciers continued to advance, the area of crustal subsidence expanded beneath coastal areas. Subsidence probably exceeded the eustatic fall in sea level as ice sheets grew between 25,000 and 15,000 14 C yr B.P. (Chappell et al., 1996; Lambeck et al., 2002; Shackleton, 1987; Waelbroeck et al., 2002). If so, relative sea level in the region rose during this period. The relative rise in sea level

controlled deposition of thick bodies of advance outwash (the Quadra Sand in British Columbia and the Esperance Sand in western Washington) on braided floodplains and deltas, and in littoral environments (Clague, 1976). As the Puget lobe reached its limit near the city of Olympia, the region to the north was isostatically depressed. The depression was greatest beneath the Strait of Georgia and Fraser Lowland and decreased south along the Puget Lowland. The height of the uppermost shorelines that formed during deglaciation gives some limits on isostatic depression. Marine deltas near Vancouver lie 200 m above present sea level (Clague et al., 1982). With eustatic sea level-100 m at the time the highest shorelines formed (Fairbanks, 1989), local glacio-isostatic depression must have exceeded 300 m. The depression was actually larger, because the Cordilleran Ice Sheet had thinned before the highest shorelines formed, and thus rebound had started already. The modern altitudes of the late-glacial marine limit display the variable isostatic influence of the Cordilleran Ice Sheet. The marine limit is highest around the Strait of Georgia and in the Canadian part of the Fraser Lowland, and it declines west and south (Clague et al., 1982; Dethier et al., 1995; Mathews et al., 1970). From about 125 to 150 m above sea level (asl) near Bellingham, it drops to 70 m asl west of Victoria, below 50 m asl on the west coast of Vancouver Island at Tofino, and probably below 50 m asl near the entrance to Juan de Fuca Strait. The marine limit decreases south of Bellingham to about 35 m asl at Everett. At the heads of the British Columbia mainland fiords to the north, the marine limit is fairly low because these areas remained ice-covered until isostatic rebound was well along (Clague & James, 2002; Friele & Clague, 2002). Isostatic uplift rates can be inferred from a variety of shoreline data. Proglacial lakes covered southern and central Puget Lowland during deglaciation (Fig. 11), the lakes

Fig. 11. Paleogeographic maps showing the maximum extents of Lake Russell and Lake Bretz (modified from Thorson, 1989, Fig. 2). Arrows show spillway locations controlling local and regional lake altitudes.

32


Fig. 12. North-south profile of shoreline features (delta tops) associated with Lake Russell, Lake Hood (confluent with Lake Russell), Lake Bretz, and the marine limit near Discovery Bay on the south side of the Strait of Juan de Fuca Strait (modified from Thorson, 1989, Fig. 5; with additional data from Dethier et al., 1995).

were dammed to the north by the retreating Puget lobe. The last lake drained when the Puget lobe retreated north of Port Townsend and marine waters entered Puget Sound. Differential isostatic rebound warped the shorelines of these proglacial lakes (Fig. 12) – shorelines of Lake Russell-Hood are tilted up to the north at 0.85 m/km; the tilt of Lake Bretz shorelines is 1.15 m/km (Thorson, 1989). Most uplift in the Fraser Lowland and on eastern Vancouver Island occurred in less than 1000 years (Clague et al., 1982; Mathews et al., 1970), as inferred from the shoreline tilt data and relative sea-level observations. These data underlie a postglacial rebound model of the Cordilleran Ice Sheet (Clague & James, 2002; James et al., 2000) that predicts low mantle viscosities (150 m across the northern basins of lakes McConnell, Agassiz, and Ojibway. Because the rate of rebound was greater during the early stages of deglaciation, older beaches have a steeper slope than do younger beaches along a specific isobase; this can be used as a relative age indicator in a proglacial lake basin.

48

James T. Teller

Fig. 3. Cartoons of bathtubs of water at times (T) 1, 2, 3, and 4, experiencing differential rebound and overflow through three different outlets. Strandlines are numbered in sequence of formation; for simplicity, beach curvature due to non-linear differential rebound is not shown (after Larsen, 1987; Teller, 2001). (A) Overflow from southern end of basin. Lake regresses everywhere through time. (B) Overflow from “middle” of basin. Through time, lake transgresses to south of outlet, regresses to north. (C) Overflow from northern end of basin. Lake transgresses upslope throughout basin. A Brief History of Large Proglacial Lakes in North America The largest proglacial lakes in North America during the last deglaciation were Lake McConnell, Lake Agassiz, Lake Ojibway, and the various lakes in the Great Lakes basins (Fig. 1). The record of older glaciations and proglacial lakes is incomplete, and their ages are not well constrained. Based on the magnetostratigraphic record, Barendregt & Irving (1998) concluded that the LIS was not as extensive during the early Pleistocene. If true, the LIS would not have provided a continuous barrier across the Hudson Bay Lowland, nor have expanded far enough west to impound drainage in the Mackenzie River Valley to the Arctic Ocean, and ancestors of glacial Lakes Agassiz, Ojibway and McConnell would not have formed. As a result, diversion of meltwater into the Mississippi River and Great Lakes from the vast region of Canada between the Rocky Mountains and Great Lakes would not have occurred, resulting in a dramatically different hydrological scenario from that of the last (and best known) glacial period,

as well as a notably different role of North American fresh waters in ocean systems and climate (see later section). The stratigraphic record of the last deglaciation is complex, and the chronology is not well controlled in some areas. Changes in proglacial lakes along the margin of the LIS were both frequent and rapid. Lakes periodically merged with others, sometimes forming bodies of water that spanned half a continent. At times, events in the upper end of the drainage system were responsible for re-directing millions of square kilometers of runoff from one lake basin to another. At other times, isostatic rebound or glacial advance deprived “downstream” basins of that water. Basin linkages were often abrupt, resulting in huge outbursts of overflow from one basin to the other. Now that we recognize this interrelationship and appreciate the magnitude and rapidity of the hydrological links, it is essential to examine the whole proglacial system along the LIS to understand how it worked. This is a challenge, because interpreted changes in one basin must be discussed (and tempered) with interpreted changes in hydrologically linked basins, even though these lakes may be

Controls, History, Outbursts, and Impact of Large Late-Quaternary Proglacial Lakes in North America 49

Fig. 4. Hypothetical development of lake levels through time as outlet changes from south to “middle” and back to south in basin. Sequence begins after southward drainage is abandoned, having formed as shown in Fig. 3A. Note that most large proglacial basins in North America experienced more outlet changes than shown here and that outlet erosion would also have impacted on lake level. (A) Overflow through southern outlet (S) ends after regressive lake levels (1, 2, 3) formed. Outlet then shifts to “middle” of basin (M1) and lake abruptly drops to level 4. No new beaches formed. (B) Overflow through outlet M1, with waters transgressing to south and regressing to north of outlet (levels 5, 6); beaches south of outlet reworked upslope; no distinct beaches formed north of outlet. (C) New lower outlet opens in “middle” of basin and overflow shifts to M2, with lake abruptly dropping to level 7; beach abandoned at transgressive maximum of level 6. (D) Lake transgresses to southern end of basin (level 8) and begins to flow out through outlet S, abandoning the now-higher outlet M2. As differential rebound continues, the lake will decline in level (regress) below level 8 throughout the basin. thousands of kilometers apart. Perhaps our biggest advance in recent decades has been this effort to understand the whole system, even though there are still many unresolved issues and, undoubtedly, will be many surprises ahead. Although the focus of this chapter is on large proglacial lakes, many smaller lakes developed along the ice margin as it retreated down slope (Teller, 1987). Many of these formed across the Canadian Prairies, commonly in north- or east-draining valleys and lowlands, and can be seen in Fig. 1 west of Lakes Agassiz and McConnell. Individual lakes typically were several thousand km2 in extent (although some were >10,000 km2 ) and survived only a century or two (Teller & Kehew, 1994); many drained catastrophically into Lake Agassiz (e.g. Kehew, 1982). These lakes were integral in the development of postglacial drainage. Their abrupt drainage scoured deep valleys, led to the domino-like drainage of other small lakes downstream, and deposited flood fans in those basins (Kehew & Lord, 1986). A number of syntheses of the history of North American proglacial lake systems have been published in the past few decades, many in (or as part of) books and special journal

issues, including Prest (1970), Clayton & Moran (1982), Mickelson et al. (1983), Teller & Clayton (1983), Karrow & Calkin (1985), Teller (1987, 1995b), Dyke & Prest (1987), Fulton (1989), Lewis & Anderson (1989), Teller & Kehew (1994), Licciardi et al. (1999), and Teller & Leverington (pers. comm., 2003); these syntheses provide hundreds of additional references. The following synthesis is taken from these and other papers published mainly in the past 20 years.

Speculation About the History of Lakes During Advance of the LIS Accumulation of continental ice over the upland areas of eastern Canada sometime after 110,000 years ago (marine oxygen isotope stage (MIS) 5e) probably was gradual, and did not initially result in the formation of large proglacial lakes. Not until ice thickened and expanded into the lower St. Lawrence Valley were major drainage basins impounded. Ice first advanced across southern Quebec into the St. Lawrence Lowland about 80,000–75,000 years ago

50

James T. Teller

(Lamothe et al., 1992; Occhietti, 1989), and may have extended into the Ontario basin and northern New England, eventually becoming extensive in the Great Lakes (see Clark, 1992). It is not known whether drainage west of the Great Lakes was impounded by this advance. Subaerial exposure and a warmer climate in the Hudson Bay Lowland followed. Extensive glaciolacustrine sediment, dated by thermoluminescence at 40,000 years in Manitoba, suggests that runoff to Hudson Bay was again dammed by readvancing Late Wisconsinan ice (Thorleifson et al., 1992); a final push of the LIS to its maximum extent across all proglacial lake basins occurred after 30,000 years ago (Dyke et al., 2002). As the Last Glacial Maximum approached, the LIS blocked a series of drainages (Fig. 2). The following speculative sequence of blockage during glacial advance illustrates the controls on glacial lakes (cf. Teller, 1995a):

(1) When ice first invaded the lower St. Lawrence Valley, east of Montreal, a lake would have formed in the western part of the lowland, re-routing overflow from the Great Lakes southward through the Lake Champlain and Hudson Valleys (CV and HV in Fig. 5; D in Fig. 1). (2) Further ice advance across the Lake ChamplainHudson outlet forced waters to deepen and expand westward into the Ontario basin, which overflowed through the Mohawk Valley (MV in Fig. 5; C in Fig. 1). (3) Expansion of the LIS over the eastern Great Lakes forced water levels throughout those basins to rise and seek outlets into the Mississippi River basin via the Wabash and Illinois River Valleys (B and A in Fig. 1). (4) Amalgamation of the Keewatin (western) and Labrador (eastern) ice centers eventually ponded waters in the Hudson Bay Lowland. Depending on the configuration of the ice margin, one or more ice-marginal lakes formed across the lowest terrain of Quebec, Ontario, Manitoba, and Saskatchewan, north of the continental divide (see Fig. 1); these were ancestors of Lakes Agassiz and Ojibway, which formed during the retreating phase of the LIS. Initially, overflow would have been from the Ojibway basin into the Ottawa River Valley (I and H in Fig. 1) but, as the eastern (Ojibway) and western (Agassiz) regions became separated by the ice, western overflow would have been diverted into the Superior basin (L and M in Fig. 1) and through the Clearwater outlet in the northwestern corner of the Agassiz basin (N in Fig. 1). Isostatic adjustment and outlet erosion would have combined with ice-margin location to dictate the extent and depth of waters. (5) The westward expansion of LIS to the Mackenzie Mts. of the Yukon Territory in northwestern Canada impounded drainage of the >1.5 million km2 Mackenzie River drainage basin, forming the predecessor of glacial Lake McConnell; overflow was west across the mountains to the Pacific Ocean (Lemmen et al., 1994).

(6) After ice blocked outlets L, M, and N from the Agassiz basin (Fig. 1), causing waters to rise further, overflow was directed south through the Minnesota River Valley (K in Fig. 1) into the Mississippi River. To the west, lakes overflowed across the divide into the Missouri River drainage basin. Eventually, all of these proglacial lake basins were completely overridden by the LIS, bringing this history of lakes during ice advance to an end. The subsequent northward retreat of ice is much better documented and, in a general way, followed a reverse pattern of lake evolution to that during the advancing stages, with lakes following on the heels of the retreating LIS, declining in level as new lower outlets were opened and amalgamating with (or overflowing into) other ice-marginal lakes as ice retreated down slope.

History and Interrelationship of Lakes Formed During Ice Retreat, Before 13,000 14 C yr B.P. The first lakes to form along the LIS margin during its retreat were in the southern parts of the Great Lakes’ basins. The ice margin fluctuated throughout this period, so lakes formed, were overridden, and reformed on numerous occasions. As new and lower outlets were opened, lake levels fell but immediately began to transgress across the newly emergent lake floor because of differential isostatic rebound. Figure 6 shows the fluctuating water level in the Huron basin interpreted by Lewis et al. (1994), and illustrates the frequent and rapid changes that occurred throughout the Great Lakes. Although these fluctuations were not always synchronous, because they were controlled by the cryodynamics of different ice lobes, there was an overall similarity of expansion and contraction during deglaciation. Ice-margin retreats (interstades) alternated with readvances (stades), resulting in an overall rise and fall pattern of lake levels. In later stages, climate may have played a role (Blasco, 2001; Lewis et al., 2002). As shown in Fig. 7A, the Erie and southern Huron basins contained ice-marginal lakes during the Erie Interstade, perhaps between about 16,000 and 15,000 14 C yr B.P.; ice retreated enough in the Ontario basin to allow overflow at this time through the Mohawk River Valley (M of Fig. 7A; MV of Fig. 5) to the Atlantic Ocean via the Hudson River Valley of New York state (C and D in Fig. 1). This outlet and these basins were overridden during the Port Bruce Stade (Fig. 7B). The subsequent retreat of ice led to a series of lakes in these basins and in the western Ontario basin during the Mackinaw Interstade (Fig. 7C, D, and E), as well as in the Michigan basin. Outflow from glacial Lake Maumee in the Erie basin first occurred through the Wabash River Valley of Indiana (W in Fig. 7C; B in Fig. 1) and included at least one catastrophic outburst (Fraser & Bleuer, 1988). When the ice margin retreated far enough, proglacial lakes in the Michigan basin received overflow from the Huron basin via the Imlay and Grand River Valley channels across the state of Michigan (Kehew, 1993) (Fig. 7D; E in Fig. 1). Retreat north of the isostatically depressed Fenelon Falls outlet (F in Fig. 7E; G in


Fig. 5. Topography of the St. Lawrence Valley Lowland and adjacent area (elevations in m and shaded above 100 m), showing the Mohawk Valley (MV), Lake Champlain Valley (CV), and Hudson Valley (HV) routes for overflow from the glacial Great Lakes (after Gadd, 1988, Fig. 1). Fig. 1) and the Ontario basin led to low lake levels in Erie and Huron basins, and overflow was again routed east through the Mohawk River Valley (M in Fig. 7E; also, see Fig. 6). These lake basins were not completely overridden during the next (Port Huron Stade) advance at 13,000 14 C yr B.P.

History and Interrelationship of Lakes Formed After 13,000 14 C yr B.P. The Port Huron readvance closed the eastern overflow route, displaced waters in the Huron and Ontario basins, and raised lake levels in the Erie basin (Fig. 7F). The subsequent ice retreat (Two Creeks Interstade) once more initiated overflow from the Huron, Erie, and Ontario basins through the Mohawk valley system (Fig. 7G), producing generally low lake levels (e.g. Hansel et al., 1985; Lewis et al., 1994). Waters in the Michigan basin, which had previously overflowed south into the Mississippi River basin, spilled into the Huron basin after ice retreated north of the Straits of Mackinac. Glacial Lake St. Lawrence formed in the western St. Lawrence Valley shortly after 12,000 14 C yr B.P. by the amalgamation of several smaller lakes (e.g. Rodrigues & Vilks, 1994). Other relatively small lakes formed as ice retreated from the northward-sloping valleys south of the St. Lawrence such as the Hudson and Connecticut River Valleys (Ashley, 1975; Connally & Sirkin, 1973; Fullerton, 1980).

Within a few hundred years after the next (Greatlakean) ice advance, which re-dammed waters in the Michigan basin and drowned the well-known Two Creeks forest bed, the ice margin again began to retreat. Overflow was progressively routed from the Huron basin through lower, more northern avenues, the Port Huron, Fenelon Falls, and North Bay outlets (F, G, and H in Fig. 1); the northern outlet carried water to the Ottawa River Valley and the Champlain Sea. Because of isostatic depression in the area of their outlets, lake levels in the Michigan, Huron, and Erie basins generally remained low for much of the next 5000 years (e.g. Figs 6 and 8) (Lewis et al., 1994). During this time, ice retreated far enough to allow glacial Lake St. Lawrence to drain and the Champlain Sea to penetrate westward in the St. Lawrence Valley to the Ontario basin (Anderson & Lewis, 1985), bringing the history of ice-marginal lakes in the St. Lawrence Valley to an end (Rodrigues & Vilks, 1994). The Superior basin was first deglaciated between 11,000 and 10,000 14 C yr B.P., but ice invaded the basin again during the Marquette advance at around 10,000 14 C yr B.P. (Lowell et al., 1999). The Agassiz basin remained largely filled by ice until after 11,700 14 C yr B.P. (Fenton et al., 1983). Overflow from the early stages of lakes in the Superior and Agassiz basins was south into the upper Mississippi River drainage system. Lake Barlow (Fig. 8), and its successor Lake Ojibway, trapped between the retreating LIS and the continental divide, expanded into the eastern Hudson Bay Lowland (Fig. 1).

52 James T. Teller Fig. 6. Lake level fluctuations in the Huron basin, showing the relationship to inflow (mainly meltwater and Lake Agassiz) and the outlets used (see Fig. 1 for geographic locations) (Lewis et al., 1994, Fig. 14). Names of stades and interstades, and various Great Lakes’ stages are shown; ages are in 14 C years.


Fig. 7. Maps showing 7 snapshots of water bodies in the eastern glacial Great Lakes, resulting from advance and retreat of the LIS (hachured line) and differential isostatic rebound (after Lewis et al., 1994, Figs 3–5). Arrows show routing of overflow from lakes. Ages in 14 C yr B.P. (A) Maximum extent of lakes during the Erie Interstade at 15,500 14 C yr B.P. (B) Extent of ice during the Port Bruce Stade about 14,800 14 C yr B.P. (C) Highest phase of Lake Maumee about 14,100 14 C yr B.P. at the start of the Mackinaw Interstade, which overflowed southwest via the Wabash River (W) in the Ohio River basin. (D) Continued ice retreat led to lake expansion and overflow west via the Grand River Valley (GRA) into the Michigan basin, which then spilled south into the Mississippi River basin via the Chicago outlet. (E) Continued ice retreat during Mackinaw Interstade opened the Fenelon Falls outlet (F) from the Huron basin and the Mohawk Valley route (M) from the Ontario basin, lowering lake levels and isolating these basins from one another about 13,200 14 C yr B.P. (F) Reduction in extent of Great Lakes during the Port Huron Stade readvance about 13,000 14 C yr B.P., when overflow was forced back into the Mississippi River basin. (G) Two Creeks interstade, showing low levels of the Huron, Erie, and Michigan basins with deeper water in the Ontario basin and in confluent glacial Lake St. Lawrence about 11,900 14 C yr B.P.; overflow was east through the Mohawk Valley.

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Fig. 7. (Continued ) As the LIS continued to retreat, other topographic thresholds (Fig. 2) came into play to control depth, distribution, and interconnection of North America’s largest proglacial lakes. Differential isostatic rebound continued to cause southward transgression of these water bodies everywhere south of their outlets; the lake level changes of all proglacial lakes reflect the interplay of these factors and resulted in a saw-toothed lake-level history, like that shown for Lake Agassiz in Fig. 9. Although there were several aborted starts to Lake Agassiz in the southern end of the Red River Valley, terminated when ice re-advanced to displace the lake (e.g. Clayton & Moran, 1982; Fenton et al., 1983), early overflow was south, mainly though the Minnesota River Valley and into the Mississippi River at Minneapolis (K in Fig. 1) until ice retreated several hundred kilometers north, where a low point on the basin rim led east into the Superior basin. The level of Lake Agassiz abruptly dropped by about 110 m (Leverington et al., 2000) when this eastern outlet was opened around 10,900 14 C yr B.P., producing a low-water period (the Moorhead phase; Fig. 9) with an extensive subaerially

exposed and vegetated lake floor that was progressively inundated by transgressing waters as differential isostatic rebound uplifted the outlet (e.g. Bajc et al., 2000; Fenton et al., 1983). This eastward diversion of Agassiz overflow generally corresponds to a relatively high lake level in the Huron basin (Fig. 6). Lake Agassiz eventually transgressed to the southern outlet again, but waters soon fell as a different outlet in the northwestern corner of the basin was opened (N in Fig. 1). Although details of events during the period from 10,100 to 9400 14 C yr B.P. have been debated (see Teller, 2001), most overflow from Lake Agassiz, which was the world’s largest proglacial lake at this time, was through the northwestern outlet to the Mackenzie River Valley and to the Arctic Ocean (Fisher & Smith, 1994; Smith & Fisher, 1993); a readvance about 9900 14 C yr B.P. severed the link between Lake Agassiz and the northwestern outlet, forcing the lake to rise and overflow briefly through the southern outlet (Fig. 9) (Teller, 2001). This coincided with the Marquette re-advance into the Superior basin, which forced that lake to overflow west into the Agassiz basin and deposit a distinctive series of 24 red varves in northwestern Ontario (Zoltai, 1961).


Fig. 8. Proglacial lakes along the LIS about 9500 14 C yr B.P., just before the eastern outlets of Lake Agassiz were reopened, showing the lower levels of the Great Lakes resulting from their isostatically depressed outlets and the loss of overflow from the Agassiz basin (Teller, 1987, Fig. 20). The fluctuating level of Lake Agassiz produced a sawtooth pattern (Fig. 9), dropping abruptly when new outlets opened and rising more gradually afterward (south of the outlet) as a result of differential rebound. The changing bathymetry and extent of Lake Agassiz, including the volume of flood outbursts that occurred when lower outlets were deglaciated, is shown and discussed by Leverington et al. (2000, 2002) and Teller & Leverington (in press). Overflow and catastrophic outbursts from Lake Agassiz have been correlated with lake level rises and changes in the sedimentary and isotopic records in the Superior, Michigan, Huron, and Erie basins (e.g. Colman et al., 1994; Farrand & Drexler, 1985; Lewis et al., 1994; Rea et al., 1994; Teller, 1985; Teller & Mahnic, 1988; Tinkler et al., 1992; Tinkler & Pengelly, 1995). It was suggested by Kehew (1993) that the influx of Agassiz waters during its Moorhead phase (ca. 10,900–10,100 14 C yr B.P.) may have forced the level of Lake Michigan to rise and overflow through the Chicago outlet. Lewis et al. (1994) suggest a linkage of Agassiz overflow to increasing lake level in the Huron basin (see Fig. 6) and a correlation of enriched (higher) 18 O values in ostracode shells to the dilution of Huron basin waters by Agassiz waters that had a high precipitation/meltwater ratio; Rea et al. (1994) weave this into a history of the paleohydrology of the Great Lakes system. Colman et al. (1994), however, correlate lake level rise and depleted δ18 O values in the Michigan basin with the influx of Lake Agassiz overflow. Regardless, it is clear that

the addition of the >2 million km2 Agassiz drainage basin would have impacted the Great Lakes, especially at times when catastrophic outbursts through its eastern outlets introduced thousands of km3 into the upsteam end of the Great Lakes in only a few months or years (see later discussion). When Agassiz flow was cut off, such as during the Marquette readvance into the Superior basin around 10,000 14 C yr B.P., lake levels in the Great Lakes declined (Figs 6 and 8). Following that, meltwater continued to enter the Great Lakes from the Agassiz basin and from remnant ice until about 8000 14 C yr B.P., when the LIS retreated far enough north of the divide with the Hudson Bay basin to allow waters to overflow from the eastern (Ojibway) part of this vast proglacial lake (I of Fig. 1). During the period between about 10,000 and 9000 14 C yr B.P., another great proglacial lake formed along the northwestern side of the LIS, Lake McConnell (Fig. 1). Trapped in the isostatically depressed area of the Mackenzie River basin, between the LIS and the Mackenzie Mountains, this lake changed its depth and configuration in response to ice marginal fluctuations, outlet change and incision, and isostatic rebound. Lemmen et al. (1994) and Smith (1994) summarize the history of this giant and complex lake, which eventually covered an area of 215,000 km2 . During its life, it received inflow from the Lake Agassiz basin when that lake overflowed through its northwestern outlet (N of Fig. 1) between about 10,100 and 9400 14 C yr B.P. (Fig. 9). The

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Fig. 9. Schematic of relative changes in level of Lake Agassiz south of its outlet from 11,700 to 7700 14 C yr B.P., showing rising (transgressing) stages that resulted from differential isostatic rebound after each abrupt decline in lake level when new lower outlets were opened. Variable routing of overflow to oceans shown by bars. Names of selected beaches and the five main stages of the lake are shown. eventual disappearance of glacial Lake McConnell, after reaching its maximum extent 10,500–10,000 14 C yr B.P. (Lemmen et al., 1994), occurred gradually, as differential rebound raised the depressed basin, causing it to drain through the Mackenzie River into the Arctic Ocean. Some of North America’s largest modern lakes, Lakes Athabasca, Great Slave, and Great Bear, are remnants on the floor of this lake. After about 9400 14 C yr B.P., when ice had retreated from the Superior basin, overflow from Lake Agassiz was directed eastward through a series of 17 progressively lower channels into the Nipigon-Superior basin (Leverington & Teller, 2003; Teller & Thorleifson, 1983) until about 8000 14 C yr B.P., producing the stair-step series of beaches seen in Fig. 9. When ice receded north of the Great LakesHudson Bay divide, Agassiz overflow began to by-pass the Great Lakes, entering glacial Lake Ojibway, which had been expanding along the retreating LIS in the eastern Hudson Bay Lowland (Fig. 1); overflow from this giant ice-marginal lake was through the Kinojévis outlet into the Ottawa River-St. Lawrence Valley system (near I in Fig. 1). Lakes Barlow (Fig. 8) and Ojibway (Fig. 1) formed north of the Great Lakes as ice retreated down slope into the Hudson Bay basin, overflowing through the Ottawa River Valley and into the Champlain Sea (St. Lawrence Valley,

Fig. 8) (Veillette, 1994; Vincent & Hardy, 1979). Northward expansion was interrupted by several glacial surges into the lake (the Cochrane surges). After the merger with Lake Agassiz about 8000 14 C yr B.P., this lake survived another few centuries, until ice in Hudson Bay thinned enough so it no longer provided a barrier to outflow to Hudson Strait and the North Atlantic Ocean. The catastrophic final drainage of this giant lake appears to have had an impact on oceans (see later discussion) (Barber et al., 1999; Teller et al., 2002). After about 7700 14 C yr B.P., the water level in the Huron basin fell below its outlet as climate warmed, possibly resulting in a closed basin and more saline environment for several hundred years (Blasco, 2001) (see Fig. 6). In general, throughout this late-glacial period, lakes in the Great Lakes’ basins transgressed southward, although where the outlet lay on an isobase that extended through the “middle” of the basin rather than along its northern side, such as in the Superior basin, waters transgressed on to the southern shore, but receded from the northern side (see Fig. 3B). The Nipissing transgression brought all the western Great Lakes to a common level about 6500 14 C yr B.P., and waters eventually reached the Chicago outlet, diverting overflow southward into the Mississippi River system (e.g. Eschman & Karrow, 1985; Hansel et al., 1985; Larsen, 1987). This transgression

Controls, History, Outbursts, and Impact of Large Late-Quaternary Proglacial Lakes in North America 57 also reached the Port Huron outlet from the Huron basin to the Erie basin, which led to erosion and eventual capture of overflow from the western Great Lakes by the eastern Great Lakes system by about 4700 14 C yr B.P. (Lewis et al., 1994). The complex history of the Great Lakes is summarized in Karrow & Calkin (1985), Larsen (1987), Teller (1987), Lewis & Anderson (1989), and Lewis et al. (1994).

The Impact of Large Glacial Lakes on Climate and Oceans In recent years there has been an increasing recognition of the role that proglacial lakes played in influencing climate and in bringing about climate change. One aspect relates to the thermal capacity of water, which makes a lake slow to respond to seasonal warming and cooling; lake temperatures fluctuate less than do those on land and tend to moderate climate and increase cloudiness and precipitation in the region around it. As a result, ecosystems around large lakes tend to be different. The early stages of most North America proglacial lakes probably were turbid and cold, and lake productivity (and organic content in the sediment) was low. In their review, Lewis et al. (1994) note that waters in the ice-marginal Great Lakes were cold, with a mean temperature no higher than 3.6 ◦ C, based on 112 taxa of aquatic and terrestrial mollusks; these authors and Moore et al. (2000) used variations in the 18 O of ostracodes to identify the relative contributions of meltwater and precipitation runoff during the history of proglacial waters in the Huron basin. The “lake effect” on land adjacent to large proglacial lakes produced contrasting climate, ecosystems, and isotopic signatures to those a few tens of kilometers away (e.g. Norton & Boisenga, 1993). Hu et al. (1997) found a shift to more depleted 18 O values in a small lake in northern Minnesota from 10,200 to 9100 14 C yr B.P., which coincided with the time when Lake Agassiz was expanding. Even though there was no increase in Picea at this time, which typically reflects cooler conditions, they interpreted that the isotopic depletion was due to decreased summer temperatures and the addition of precipitation derived from the isotopically light waters of nearby Lake Agassiz. Dean et al. (2002) suggest that the shift in climate signal about 7300 14 C yr B.P. in the varved sequence of a small lake just south of the Agassiz basin reflects a change in atmospheric circulation that correlates with the final drainage of Agassiz. Hostetler et al. (2000) used an atmospheric GCM to model the likely effect of Lake Agassiz on the region, concluding that, during an expansive lake stage, there were cooler temperatures and less precipitation, because there was less evaporation from the cooler lake than from the land, and because the cool lake induced anticyclonic winds that blocked moisture-bearing winds from outside the region. During expanded lake stages they conclude that the suppressed precipitation would have contributed to retreat of the LIS. Whenever new lower outlets were deglaciated and proglacial lake waters were drawn down, the influence of

those lakes on regional climate would have been reduced. However, the climate around basins receiving these waters may have been influenced. Although the impact of small influxes of new water probably had a minimal effect on the depth of a lake and its overall nature, large influxes, such as the catastrophic outbursts from Lake Agassiz into the Great Lakes, may have briefly raised lake levels by several tens of meters and affected their sedimentary record, as noted before. Lewis & Anderson (1992) even suggested that the rise in Picea and shift in 18 O and 13 C around the Erie basin about 11,000–10,500 14 C yr B.P. reflects cooling related to the first influx of Lake Agassiz overflow during that period. Rodrigues & Vilks (1994) identified an ostracode and isotopic signal in sediments of the Champlain Sea in the western St. Lawrence Valley that resulted from the diversion of Agassiz overflow into the Great Lakes. The abrupt transfers of water from one lake basin to another had significant hydrological impacts on lakes, rivers, and oceans, as well as on climate. Of course there also was a morphological impact on many valleys receiving additional runoff or draw down from a lake basin; Teller & Thorleifson (1983) and Fisher & Smith (1994) describe some of the dramatic geomorphic and sedimentological changes that occurred as a result of outbursts from Lake Agassiz. The re-direction of overflow from one basin to another and the simple opening of a lower overflow route may produce such changes. In both cases there would have been a large initial, short-term increase in outflow (producing a flood hydrograph). The arrows in Fig. 1 show many of the new channels and routes for lake overflow around the LIS, which may have experienced these hydrological events. As previously discussed, the diversions of overflow were related to isostatic rebound, glacial advance, and glacial retreat. Some of the diversions of lake overflow were permanent, some lasted for centuries, and some last for only a few years. The record of some drainage diversions in glaciated parts of North America is found in sediments of the Gulf of Mexico, and Leventer et al. (1982), Broecker et al. (1985), and others used this to suggest how the routing of North American meltwater might have impacted on ocean circulation. The potential role of Lake Agassiz overflow on ocean circulation and related climate was proposed by Broecker et al. (1989), who suggested that the diversion of Agassiz overflow from the Gulf of Mexico to the North Atlantic about 11,000 14 C yr B.P. may have led to the Younger Dryas cooling. Although the impact of these fresh waters on ocean circulation continues to be debated, many GCM models indicate that a relatively small addition of fresh water to specific places in the North Atlantic Ocean, particularly at certain “transitional” stages of circulation, can shut down the production of North Atlantic Deep Water (NADW) (e.g. Fanning & Weaver, 1997; Manabe & Stouffer, 1997; Rind et al., 2001); as discussed by Teller et al. (2002), several outbursts from Lake Agassiz exceeded those modeled values. As a result, thermohaline circulation and the transfer of heat into high latitudes of the North Atlantic were reduced, leading to climate cooling. The flux

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of freshwaters needed to alter NADW is related to the length of time over which this flux continued and to the specific nature of ocean circulation at the time of the freshwater addition. The potential impact of fresh water additions to the North Atlantic Ocean and the close chronological relationship of major late-glacial climate changes to both short outbursts of overflow from Lake Agassiz and the diversion of North American drainage from one ocean to another have led Clark et al. (2001) and Teller et al. (2002) to conclude that those freshwater fluxes triggered or contributed to global climate change during the last period of deglaciation. In turn, these global climate changes are recorded in some pollen and isotopic records in small lakes in the Great Lakes and Agassiz region (Yu & Wright, 2001), although the degree to which these are related to global climate forcing from the North Atlantic region or reflect the control on climate by the proglacial lake system itself remains uncertain. Four major late-glacial climate changes have been related to the influx of freshwaters from North America; these were superposed on the overall global warming, and are: (1) The Older Dryas cold period. This cooling occurred between about 16,500 and 13,000 14 C yr B.P. It spans the Erie Interstade in North America, in which the first large eastward overflow of freshwater from the Great Lakes to the North Atlantic Ocean (the Heinrich 1 event of ice rafting in the North Atlantic) was initiated. It also includes the Mackinaw Interstade, which led to the second period of Great Lakes outflow to the North Atlantic. This suggests a link between cooling and outflow from the Great Lakes and glacial re-advance, represented by the Port Huron Stade and by iceberg production (Clark et al., 2001). (2) The Younger Dryas cold period. This abrupt cooling at about 11,000–10,000 14 C yr B.P. coincides with the largest outburst of water from Lake Agassiz and the accompanying 800-year diversion of its overflow into the Great Lakes and North Atlantic. (3) The Preboreal Oscillation. Shortly after the end of the Younger Dryas, a relatively small cooling in the North Atlantic region occurred, which coincides in time with an outburst from Lake Agassiz to the Arctic Ocean (Fisher et al., 2002; Teller et al., 2002) and the draining of the Baltic Ice Lake into the North Atlantic (Björck et al., 1996). (4) The 8200 cal yr B.P. cold event. Following the draining of 163,000 km3 of fresh water from Lake AgassizOjibway about 7700 14 C yr B.P. (8400 cal yr B.P.), a small cooling was recorded in the isotopic signal of the Greenland ice sheet, suggesting a link between these events (Barber et al., 1999; Teller et al., 2002). The 200-year difference between the age of this drainage, estimated by Barber et al. (1999), and the cooling may reflect (a) a two-step final draw down of Lake Agassiz, where complete drainage of the western part of this huge lake was delayed by 200 years (Leverington et al., 2002), or it may indicate that the ocean car-

bonate reservoir effect used by Barber et al. (1999) to estimate the age was too small (Teller & Leverington, pers. comm., 2003). Only because the global ocean was in a more stable interglacial mode did this large freshwater flux not have a larger impact on climate. Thus, not only did large proglacial lakes respond to climate and the bounding ice sheet, but they also influenced the ice and climate, both on a regional scale and, via the oceans, on a global scale. The chronological coincidence of large changes in proglacial lakes, the climate record in oceans and on adjacent continents, and the Greenland isotopic record supports a cause-effect linkage. Specifically, large fluxes of water from Lake Agassiz, along with changes in the routing of overflow from lakes around the LIS, slowed thermohaline circulation in the North Atlantic Ocean which, in turn, led to climate cooling.

Acknowledgments This synthesis is the product of years of research by many people. My own research has been generously supported by the Natural Sciences and Engineering Research Council of Canada. My thanks to the many Quaternarists who have shared their expertise with me over the years. Reviews of this chapter by Mike Lewis, Tim Fisher, and Alan Kehew have been helpful in improving the final product. Thanks to Allen Patterson for his help in preparing the illustrations.

References Anderson, T.W. & Lewis, C.F.M. (1985). Postglacial waterlevel history of the Ontario basin. In: Karrow, P.F. & Calkin, P.E. (Eds), Quaternary evolution of the Great Lakes. Geological Association of Canada Special Paper 30, 231–253. Ashley, G.M. (1975). Rhythmic sedimentation in glacial Lake Hitchcock, Massachusetts-Connecticut. In: Jopling, A.V. & McDonald, B.C. (Eds), Glaciofluvial and glaciolacustrine sedimentation. Society of Economic Paleontologists and Mineralogists Special Publication 23, 304–320. Bajc, A.F., Schwert, D.P., Warner, B.G. & Williams, N.E. (2000). A reconstruction of Moorhead and Emerson Phase environments along the eastern margin of glacial Lake Agassiz. Rainy River basin, northwestern Ontario. Canadian Journal of Earth Sciences, 37, 1335–1353. Barber, D.C., Dyke, A., Hillaire-Marcel, C., Jennings, A.E., Andrews, J.T., Kerwin, M.W., Bilodeau, G., McNeely, R., Southon, J., Morehead, M.D. & Gagnon, J.-M. (1999). Forcing of the cold event of 8200 years ago by catastrophic drainage of Laurentide lakes. Nature, 400, 344–348. Barendregt, R. & Irving, E. (1998). Changes in the extent of North American ice sheets during the late Cenozoic. Canadian Journal of Earth Sciences, 35, 504–509. Björck, S., Kromer, B., Johnsen, S., Bennike, O., Hammarlund, D., Lemdahl, G., Possnert, G., Rasmussen, T.L., Hammer, C.U. & Spurk, M. (1996). Synchronized

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Kehew, A.E. (1993). Glacial lake outburst erosion of the Grand Valley, Michigan, and impacts on glacial lakes in the Lake Michigan basin. Quaternary Research, 39, 36–44. Kehew, A.E. & Lord, M.L. (1986). Origin and large-scale erosional features of glacial-lake spillways in the northern Great Plains. Geological Society of America Bulletin, 97, 162–177. Kehew, A.E. & Teller, J.T. (1994). History of late glacial runoff along the southwestern margin of the Laurentide ice sheet. Quaternary Science Reviews, 13, 859–877. Lamothe, M., Parent, M. & Shilts, W.W. (1992). Sangamonian and early Wisconsinan events in the St. Lawrence Lowland and Appalachians of southern Quebec. In: Clark, P.U. & Lea, P.D. (Eds), The last interglacial-glacial transition in North America. Geological Society of America, Special Paper 270, 171–184. Larsen, C.E. (1987). Geological history of glacial Lake Algonquin and the Upper Great Lakes. U.S. Geological Survey Bulletin B1801, 36 pp. Last, W.M. & Smol, J.P. (Eds) (2001a). Basin analysis, coring, and chronological techniques. Vol. 1 of “Tracking environmental change using lake sediments”. Kluwer Academic Publishers, Dordrecht, Netherlands, 548 pp. Last, W.M. & Smol, J.P. (Eds) (2001b). Physical and geochemical methods. Vol. 2 of “Tracking environmental change using lake sediments”, Kluwer Academic Publishers, Dordrecht, Netherlands, 504 pp. Lemmen, D.S., Duk-Radkin, A. & Bednarski, J.M. (1994). Late glacial drainage systems along the northwestern margin of the Laurentide Ice Sheet. Quaternary Science Reviews, 13, 805–825. Leventer, A., Williams, D.F. & Kennett, J.P. (1982). Dynamics of the Laurentide Ice Sheet during the last deglaciation: evidence from the Gulf of Mexico. Earth and Planetary Science Letters, 59, 11–17. Leverington, D.W., Mann, J.D. & Teller, J.T. (2000). Changes in the bathymetry and volume of glacial Lake Agassiz between 11,000 and 9300 14 C yr B.P. Quaternary Research, 54, 174–181. Leverington, D.W., Mann, J.D. & Teller, J.T. (2002). Changes in the bathymetry and volume of glacial Lake Agassiz between 9200 and 7600 14 C yr B.P. Quaternary Research, 57, 244–252. Leverington, D.W. & Teller, J.T. (2003). Paleotopographic reconstructions of the eastern outlets of glacial Lake Agassiz. Canadian Journal of Earth Sciences, 40, 1259–1278. Lewis, C.F.M. & Anderson, T.W. (1989). Oscillations of levels and cool phases of the Laurentian Great Lakes caused by inflows from glacial Lakes Agassiz and Barlow-Ojibway. Journal of Paleolimnology, 2, 99–146. Lewis, C.F.M. & Anderson, T.W. (1992). Stable isotope (O and C) and pollen trends in eastern Lake Erie, evidence for a locally induced climatic reversal of Younger Dryas age in the Great Lakes basin. Climate Dynamics, 6, 241–250. Lewis, C.F.M, Blasco, S.M. & Coakley, J.P. (2002). Severe dry climate impact on the Laurentian Great Lakes indicated by early to middle Holocene lake closure. Internat. Assoc.

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Controls, History, Outbursts, and Impact of Large Late-Quaternary Proglacial Lakes in North America 61 Shaw, J. (1996). A meltwater model for Laurentide subglacial landscapes. In: McCann, S.B. & Ford, D.C. (Eds), Geomorphology sans frontière. J. Wiley, Chichister, 182–226. Shaw, J. (2002). The meltwater hypothesis for subglacial bedforms. Quaternary International, 90, 5–22. Shoemaker, E.M. (1991). On the formation of large subglacial lakes. Canadian Journal Earth Sciences, 28, 1975– 1981. Shoemaker, E.M. (1992). Water sheet outburst floods from the Laurentide Ice Sheet. Canadian Journal Earth Sciences, 29, 1250–1264. Smith, D.G. (1994). Glacial Lake McConnell: paleogeography, age, duration, and associated river deltas, Mackenzie River basin, western Canada. Quaternary Science Reviews, 13, 829–843. Smith, D.G. & Fisher, T.G. (1993). Glacial Lake Agassiz: the northwestern outlet and paleoflood. Geology, 21, 9–12. Smol, J.P., Birks, H.J.B., Last, W.M. (Eds) (2001a). Terrestrial, algal, and siliceous indicators. Vol. 3 of “Tracking environmental change using lake sediments”. Kluwer Academic Publishers, Dordrecht, Netherlands, 371 pp. Smol, J.P., Birks, H.J.B. & Last, W.M. (Eds) (2001b). Zoological indicators. Vol. 4 of “Tracking environmental change using lake sediments”, Kluwer Academic Publishers, Dordrecht, Netherlands, 220 pp. Teller, J.T. (1985). Lake Agassiz and its influence on the Great Lakes. In: Karrow, P. & Calkin, P.E. (Eds), Quaternary evolution of the Great Lakes. Geological Association of Canada, Special Paper 30, 1–16. Teller, J.T. (1987). Proglacial lakes and the southern margin of the Laurentide Ice Sheet. In: Ruddiman, W.F. & Wright, H.E. (Eds), North America and adjacent oceans during the last deglaciation. Geological Survey of America, Decade of North American Geology Vol. K-3, 39–69. Teller, J.T. (1995a). The impact of large ice sheets on continental paleohydrology. In: Gregory, K., Baker, V. & Starkel, L. (Eds), Global Continental Paleohydrology. J. Wiley & Sons, New York, pp. 109–129. Teller, J.T. (1995b). History and drainage of large ice-dammed lakes along the Laurentide Ice Sheet. Quaternary International, 28, 83–92. Teller, J.T. (2001). Formation of large beaches in an area of rapid differential isostatic rebound: the three-outlet control of Lake Agassiz. Quaternary Science Reviews, 20, 1649– 1659. Teller, J.T. & Clayton, L. (1983). An introduction to glacial Lake Agassiz. In: Teller, J.T. & Clayton, L. (Eds), Glacial Lake Agassiz. Geological Association of Canada, Special Paper 26, 3–5.

Teller, J.T. & Kehew, A.E. (Eds) (1994). Late glacial history of proglacial lakes and meltwater runoff along the Laurentide Ice Sheet. Quaternary Science Reviews, 13, 795–981. Teller, J.T. & Leverington, D.W. (in press). Glacial Lake Agassiz: a 5000-year history of change and its relationship to the isotopic record of Greenland. Geological Society of American Bulletin. Teller, J.T. & Mahnic, P. (1988). History of sedimentation in the northwestern Lake Superior basin and its relation to Lake Agassiz overflow. Canadian Journal of Earth Sciences, 25, 1660–1673. Teller, J.T. & Thorleifson, L.H. (1983). The Lake Agassiz – Lake Superior connection. In: Teller, J.T. & Clayton, L. (Eds), Glacial Lake Agassiz. Geological Association of Canada, Special Paper 26, 261–290. Teller, J.T., Leverington, D.W. & Mann, J.D. (2002). Freshwater outbursts to the oceans from glacial Lake Agassiz and their role in climate change during the last deglaciation. Quaternary Science Reviews, 21, 879–887. Thorleifson, L.H., Wyatt, P.H., Shilts, W.W. & Nielsen, E. (1992). Hudson Bay lowland Quaternary stratigraphy: evidence for early Wisconsinan glaciation centered in Quebec. In: Clark, P.U. & Lea, P.D. (Eds), The last interglacialglacial transition in North America. Geological Society of America, Special Paper 270, 207–221. Tinkler, K.J. & Pengelly, J.W. (1995). Great Lakes response to catastrophic inflows from Lake Agassiz: some simulations of a hydraulic geometry for chained lake systems. Journal of Paleolimnology, 13, 251–266. Tinkler, K.J., Pengelly, J.W., Parkins, W.G. & Terasmae, J. (1992). Evidence for high water levels in the Erie basin during the Younger Dryas chronozone. Journal of Paleolimnology, 7, 215–234. Veillette, J.J. (1994). Evolution and paleohydrology of glacial lakes Barlow and Ojibway. Quaternary Science Reviews, 13, 945–971. Vincent, J.-S. & Hardy, L. (1979). The Evolution of Glacial Lakes Barlow and Ojibway, Quebec and Ontario. Geological Survey of Canada Bulletin, 316, 18 pp. Warner, B.G. (Ed.) (1990). Methods in Quaternary Ecology. Geoscience Canada Reprint Series 5, Geological Association of Canada, 170 pp. Yu, Z. & Wright, H.E. (2001). Response of interior North America to abrupt climate oscillations in the North Atlantic region during the last deglaciation. Earth Science Reviews, 52, 333–369. Zoltai, S.C. (1961). Glacial history of part of northwestern Ontario. Proceedings Geological Association of Canada, 13, 61–83.

Pleistocene glaciations of the Rocky Mountains Kenneth L. Pierce U.S. Geological Survey, Northern Rocky Mountains Science Center, Montana State University, Bozeman, Montana, 59717-3492, USA; [email protected] early, middle, and late stades (Richmond, 1965). Moraines of the early and middle stade were distinguished at and near the terminus, whereas the late stade moraines were 25–75% farther upvalley from the terminal moraines to the valley heads. In 1965, Rocky Mountain glacial subdivisions and correlations were closely linked with those of the mid-continent. The Bull Lake and Pinedale glaciations were correlated with the early and late Wisconsin respectively (Richmond, 1965). Curiously, Leverett (1917) had correlated the Bull Lake with the Illinoian Glaciation of the mid-continent that he defined in 1899, but this correlation was not accepted in 1965. Three pre-Bull Lake glaciations were correlated with the Illinoian, Kansan, and Nebraskan Glaciations (Richmond, 1965). Also in 1965, erratic boulders and diamictons well beyond or above moraines of Pinedale and Bull Lake age had been noted at many sites in the Rocky Mountains; these were attributed to an older glaciation vastly more extensive than the Bull Lake or Pinedale. Soil development was then emerging as the primary relative-age method (Birkeland, 1964; Morrison, 1965; Richmond, 1962). Pinedale deposits have an “immature zonal soil” with “B-horizons 0.3–0.6 m thick, that display very little illuviation and weak to moderate structural development” (Richmond, 1965). In contrast, Bull Lake deposits have “mature zonal soils” with “B-horizons 0.3–1.2 m thick, with sufficient illuvial clay to be slightly plastic . . . and moderately developed subangular blocky structure.” The 1965 synthesis was primarily descriptive; glacialgeologic sequences were correlated according to succession and the general appearance of moraines. Pleistocene ELA’s were estimated to be about 1000 m lower than present ELAs and both ELAs and moraines decreased in altitude northward through the Rocky Mountains (Richmond, 1965).

Introduction This chapter presents the status of Rocky Mountain glacial studies in 1965 and progress from that time to the present. The Rocky Mountains and the adjacent Basin and Range of the United States consist of about 100 ranges distributed in a northwest trending belt 2000 km long and 200–800 km wide. Glaciation created much of the grandeur of the high parts of these ranges. Figure 1 shows the extent of late Pleistocene glaciers as well as a measure of snowline altitude across the western U.S. Early in the 20th century, some noteworthy regional Rocky Mountain glacial studies included: the Uinta and Wasatch Mountains of Utah (Atwood, 1909), western Wyoming (Blackwelder, 1915), the San Juan Mountains of Colorado (Atwood & Mather, 1932), western Montana (Alden, 1953), and southern Rocky Mountains (Ray, 1940). For simplicity in communication, we use the following regional terms and their probable correlations with marine oxygen-isotope stages (MIS): Pinedale (Late Wisconsin, MIS 2), early Wisconsin (MIS 4), Bull Lake (now thought to be largely MIS 6 and perhaps 5d), and pre-Bull Lake (pre-MIS 6).

Status in 1965 Blackwelder (1915) named the Bull Lake and Pinedale glaciations for moraines on the east and west sides of the Wind River Range. In glaciated valleys of the Rocky Mountains, researchers generally distinguished a younger set of moraines (Pinedale) from an older set 5–10% further downvalley (Bull Lake). In 1965, G.M. Richmond summarized the status of glacial studies in the Rocky Mountains in two publications: (1) an INQUA chapter on the glacial geology (Richmond, 1965); and (2) a guidebook and 16-day INQUA field trip through much of the Northern and Middle Rocky Mountains (Richmond et al., 1965). Also in 1965, Crandall (1967) mapped and described the glacial sequence on the Wallowa Mountains of eastern Oregon. In 1965, non-quantitative morphology was the prime basis for distinguishing and correlating Bull Lake and Pinedale moraines. Richmond (1965) noted Bull Lake moraines were bulky with smooth slopes, did not retain lakes, and were less bouldery than Pinedale moraines. Bull Lake moraines were commonly subdivided into early and late stades. In contrast, Pinedale moraines were described as steep, irregular, having kettles commonly with ponds, and studded with numerous relatively unweathered boulders. Pinedale moraines were commonly subdivided into

Advances since 1965 The next major synthesis of Rocky Mountain glaciations was by Porter et al. (1983). It included chronology, ELA patterns, basal shear stress, types of glaciers based on valley slope and glacier thickness, and contrasting mass balance between ranges based on the modern snowpack at the Pleistocene ELA. This study was followed by Richmond (1986a), who summarized glacial extents, stratigraphy, and chronology for glaciated ranges in the Rocky Mountains. Additionally Richmond (1986b) presented a detailed chronology of Yellowstone. Madole (1976) summarized the Colorado Front Range. For the Great Basin west of the Rockies,

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Fig. 1. Map showing extent of Late Wisconsin glaciers in the Rocky Mountains and other mountains in the western United. States (from Porter et al., 1983). Contours in meters are on the lowest cirque floors. This measure of late Pleistocene snowline increases in altitude both inland from the Pacific Ocean as moisture decreases, and southward as temperature increases (Hammond, 1965).

Pleistocene Glaciations of the Rocky Mountains 65 Wayne (1984) made a detailed study of the Ruby-Humboldt area, and Osborn & Bevis (2001) achieved an extensive overview including mapping and sequence descriptions and mapping for many ranges.

Quaternary Geochronologic Studies Dating and correlation of glacial and other Quaternary deposits depend on a variety of dating techniques, as recently described in Noller et al. (2000). Dating techniques may be classified (Colman & Pierce, 2000) as: (1) relative age (chemical, biologic, geomorphic methods that include soils and weathering rinds, some of which can be refined to calibrated age; (2) numerical age (sidereal, isotopic, and radiogenic); and (3) correlated-age methods.

Selected Relative-Age Methods Relative-age methods are broadly applicable and are useful in distinguishing deposits in a sequence. Correlation between sequences is better established if variables such as climate and lithology can be controlled. Methods may be combined; for example, Miller (1979) measured six relative age parameters on a glacial moraine sequence and applied statistical clustering techniques to define two ages of Pinedale moraines, two ages of Bull Lake moraines, and one age of pre-Bull Lake moraine. Soil development. After 1965, soils became a primary basis for distinguishing glacial and other deposits in the Rocky Mountains, and other western mountains, where Pete Birkeland, his students and his text (1974, 1999) played a central role. Studies of his students include: (1) soils along a Rocky Mountain glacial transect (Shroba, 1977; Shroba & Birkeland, 1983); moraine-terrace transects in Colorado (Nelson & Shroba, 1998; Netoff, 1977); (2) the moraine sequence in the Wallowa Lake area, Oregon (Fig. 1, northeast of Strawberry Mountains) as well as part of the Sierra Nevada (Burke, 1979; Burke & Birkeland, 1979); (3) a terrace sequence near the Wyoming-Montana border (Reheis, 1987); (4) a toposequence in central Idaho from moraine crest to foot slope on Bull Lake and Pinedale glacial moraines (Berry, 1987). Some other important contributions were the formulation of a soil development index by Harden (1982), and quantitative studies of soils in the type areas of the Bull Lake and Pinedale glaciations (Dahms, 1991; Hall & Shroba, 1993; Swanson, 1985). Weathering rinds and carbonate coats. Motivated by the results of Porter (1975), Colman & Pierce (1981) measured weathering rinds on basalt and andesite in the western U.S. and found that rind thicknesses systematically increased with stratigraphic sequence of moraines. They calibrated this method based on dating of the West Yellowstone Bull Lake as MIS 6. For the glacial succession near McCall, Idaho, weathering rinds show a clear difference between moraines correlated with MIS 2, 4, and 6. But for the corresponding soils, clay increase expressed as grams/cm2 shows little

change between moraines correlated with MIS 4 and MIS 6 based on weathering rinds. In areas of calcic soils (pedocals), carbonate coats on the undersides of stones increase in thickness with time. Based on Uranium-series ages, carbonate coats in central Idaho accumulate at ∼0.6 mm/103 yr (Pierce, 1985; Pierce & Scott, 1982). Numerical age methods. Few numerical ages existed in 1965 for glacial deposits in the Rocky Mountains. The establishment of marine oxygen isotope stages provided a reference standard for the large number of late Cenozoic glacial-interglacial oscillations and potentially associated glacial advances. The AGU Handbook (Noller et al., 2000) describes many numerical-age techniques useful to dating glacial deposits, including radiocarbon and luminescence that can be applied to glacially related deposits such as bogs and loess. For Rocky Mountain and other glacial sequences, cosmogenic dating (Gosse & Phillips, 2001) is producing a numerical chronology that is based on the accumulation of isotopic changes in morainal boulders based on the duration of surface exposure to cosmic rays, provided one can account for surface erosion of the boulders and possible exhumation of the boulders. It has yielded ages that suggest a younger age than indicated by radiocarbon dating for the Pinedale glaciation and a possible MIS 5d correlation for type Bull Lake deposits.

Pre-Bull Lake Glacial Deposits We use the designation pre-Bull Lake for glacial deposits that are older than MIS 6 (or 8?) and include deposits as old as the late Pliocene. Most deposits once attributed to very large pre-Bull Lake glaciations are now interpreted to have non-glacial origins. Primary reasons are lack of till fabric and glacial striations, and their satisfactory explanation as deeply weathered fluvial gravel. In addition, many of these deposits are located well outside Pinedale and Bull Lake glaciated areas; this requires an earlier, much more extensive glaciation generated during a much more severe glacial climate. However, the marine isotope record shows no glacial intervals much more severe than MIS 2 or 6, and suggests that pre-Bull Lake glaciations were not much more extensive than those of the Bull Lake or Pinedale. Richmond (1965 and references therein) defined three pre-Bull Lake glacial deposits on the east side of the Wind River Range, from oldest to youngest: Washakie Point, Cedar Ridge, and Sacagawea Ridge. Later work establishes that only the Sacagawea Ridge has a glacial origin. At the type section for the Cedar Ridge and Washakie Point deposits, Hall & Jaworowski (1999) conclude that the only pre-Bull Lake glacial till present is the Sacagawea Ridge till. Chadwick et al. (1997) also did not confirm the presence of Cedar Ridge or Washakie Point glacial till. Moraines and associated outwash terraces of the Sacagawea Ridge glaciation (Richmond & Murphy, 1989) are somewhat older than the Lava Creek Ash dated at 640,000 ± 2000 yr (Christiansen, 2001), and probably correlate with MIS 16 (Chadwick et al., 1997).

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Three extensive pre-Bull Lake glaciations were also named and defined from diamictons of inferred glacial origin in the La Sal Mountains of Utah (Richmond, 1962). Shroder & Sewell (1985) concluded that these diamictons are extensive mass-movement deposits. They found the total area glaciated was only one twentieth (5%) the area mapped as glaciated in pre-Bull Lake time by Richmond (1962). Madole (1982) concluded that upland bouldery deposits on the Colorado Front Range previously considered to be glacial were Tertiary deposits of non-glacial origin. For the southern Rocky Mountains, Scott (1975) concluded that all pre-Bull Lake glacial deposits are close to Bull Lake and Pinedale moraines, and that bouldery deposits once attributed to an extensive icecap glaciation are weathered Cenozoic fluvial gravels. In Jackson Hole, deposits of the “ghost” glaciation (Love, 1977) are either outlying deposits of no necessary glacial origin, or are part of the Munger (Bull Lake) glaciation of Pierce & Good (1992). The best evidence for multiple pre-Bull Lake glaciations is exposed in sections just east of the Rocky Mountains near the US-Canada border. A succession of pre-Bull Lake diamictons identified as glacial by striations, till fabrics, stone shapes, and erratics from the Rocky Mountains are separated by soils but do not require a much more extensive glaciation than the Bull Lake (Karlstrom, 2000 and references therein; Fullerton et al., 2003). Karlstrom (2000) identifies, from youngest to oldest, the following: two glaciations early in the Brunhes normal Chron (0–0.78 myr ago), at least three glacial events during the Matuyama reversed Chron (0.78–2.6 myr ago), and two events with normal polarity either in the Gauss Chron (2.6–3.6 myr ago), or possibly the Reunion or Olduvai events (2.23–2.20 or 1.93–1.76 myr ago). Terrace sequences, which include pre-Bull Lake terraces, represent climatically modulated cycles of erosion followed by lateral planation and deposition. However, few terraces are tied directly to pre-Bull Lake glacial moraines. With age control on one or more terraces, the ages of other terraces can be approximated by incision rates (Reheis, 1987; R.C. Palmquist, written comm., 1989; Chadwick et al., 1997). Locally such terrace sequences have age control provided by one or more volcanic ashes, such as the 640,000-yr-old Lava Creek ash in Verdos Alluvium roughly 100 m above drainage in the Colorado Piedmont (Scott, 1975).

Bull Lake Glaciation

Thickness of hydration rinds on glacial pressure cracks in obsidian clasts from the Bull Lake moraines at West Yellowstone are calibrated by hydration thicknesses on cooling cracks of dated rhyolite flows. A plot through time shows that the Bull Lake glacial cracking is ∼30,000 yr older than the 122,000-yr-old West Yellowstone flow, and is ∼40,000 yr younger than the 183,000 ± 3000 yr old Obsidian Cliff flow. Thus the age is about 150,000–140,000 yr (Pierce et al., 1976) and correlates with the later part of MIS 6 (190,000–130,000 yr ago, Martinson et al., 1987). A younger, less extensive glacial margin is indicated by the unusual embayed and perlitic eastern margin of the West Yellowstone rhyolite flow (Christiansen, 2001, p. G44). This recessional(?) glacial margin is about 15 km east of the Bull Lake terminus. Christiansen (2001, p. G46) estimates the K-Ar ages have a geologic uncertainty of ±10,000 yr. Thus, the age of the West Yellowstone flow may be from ∼135,000 to 110,000 yr, and the glacier related to the embayed flow margin may date from late MIS 6 time, during recession from Bull Lake moraines, or a separate advance during MIS 5d time. Wave-cut bluffs eroded into the Bull Lake end moraines at The Narrows of Hebgen Lake near West Yellowstone expose a 10-cm ash bed between imbricate thrusts of till. The ash had an apparent K-Ar age of 481,000 yr (Obradovich, 1992; Richmond, 1986b). Richmond (1986b) defined two glaciations based on position relative to the ash: the 610,000–481,000-yr-old “Till of Horse Butte” below the ash and the 481,000–399,000-yr-old “Till of Hebgen Lake” above the ash. J.D. Obradovich (spoken comm., 2002) and I conclude that contamination of the feldspar concentrate is likely. The ash chemistry is similar to the 162,000 ± 2000 yr old tuff of Bluff Point (G.A. Izett, written comm., 1990; Obradovich, 1992). I interpret the till thrust above the 162,000 yr old ash to be the same as the surface Bull Lake till and consistent with late MIS 6 time, and the till beneath the glacially thrusted ash to possibly be older than 162,000 yr and perhaps of an early MIS 6 age.

Cosmogenic Exposure Dating of Bull Lake Deposits At the type Bull Lake on the east side of the Wind River Range, boulder exposure ages (36 Cl supplemented by 10 Be; Phillips et al., 1997) from a sequence of 15 Bull Lake moraines (Chadwick et al., 1997) yield the following ages and suggested MIS correlations:

Combined K-Ar and Obsidian Hydration Dating In 1965, Bull Lake moraines were widely considered to be early Wisconsin in age (Richmond, 1965). However, the Bull Lake moraines near West Yellowstone are clearly older than the West Yellowstone rhyolite flow (Pierce et al., 1976; Richmond, 1986b; Waldrop, 1975), which is best dated as 122,300 ± 2200 yr old (Obradovich, 1992, and spoken comm., Sept. 2002, based on the three older sanidine ages selected because complete degassing is a primary concern).

Moraine Group & No.

Moraine Dated

Cosmogenic Age (103 yr)

MIS

D D C B A

XIII XII IX (not dated) II–III

95–120 95–120 100–130 >130 >130

5d 5d 5d 6 6

XII–VX IX–XI IV–VIII I–III

Pleistocene Glaciations of the Rocky Mountains 67 Phillips et al. (1997) caution that in addition to laboratory analysis, additional uncertainty is about 10–15%. Contrasting cosmogenic ages on boulders on the same Bull Lake moraines on the west side of the Wind River Range are illustrated and described in Fig. 2. Gosse & Phillips (2001) show that both 10 Be and 26 Al ages increase with increasing resistance to erosion (Fig. 2A), and conclude that these Bull Lake moraines are ∼150,000 yr old. For the same moraines (Fig. 2B), Phillips et al. (1997) find an age of 120,000–100,000 yr, an age 20–35% younger than the 10 Be and 26 Al ages, but similar to the 36 Cl ages on Group C and D across the range at Bull Lake. In light of the U-series dating discussed next, it seems likely that the 36 Cl ages are too young, on both sides of the range. U-series dating of terraces along the Wind River in the type area of the Bull Lake glaciation (Blackwelder, 1915) produces ages older than the cosmogenic ages of the nearby correlative moraines. Multiple analyses determining 230 Th/U ages on micro-stratigraphic layering of the carbonate coats on stones from soils on the terraces yield the following ages (Sharp et al., 2003): Terrace

Age (103 yr)

MIS

Glacial Unit

WR2 WR3 WR4

55 ± 8.6 150 ± 8.3 167 ± 6.4

4 6 6

Early Pinedale correlative Late Bull Lake Early Bull Lake

Based on these ages, the early and late Bull Lake terraces at the type area correlate with early and late MIS 6. In the Colorado Front Range, 10 Be and 26 Al ages on only 2(?) boulders in Bull Lake moraines are 122,000 ± 26,000 yr and those from a Bull Lake terrace are 136,000 ± 28,000 yr, both with no correction for erosion of boulder surface (Schildgen et al., 2002; Dethier et al., 2003).

Discussion of Age and Correlation of the Bull Lake Glaciation In the West Yellowstone area using combined K-Ar and obsidian hydration dating, Pierce et al. (1976) determined a MIS-6 age (Illinoian) for the Bull Lake Glaciation, revising the widely accepted early Wisconsin correlation (MIS 4?). Next, for Bull Lake moraines of the Wind River Range primarily based on 36 Cl ages, Phillips et al. (1997) dated Groups A and B as MIS 6 and Groups C and D as MIS 5d. However, dating the firmest boulders, Gosse & Phillips (2001) determined a mid-MIS-6 10 Be age on the same moraine that Phillips et al. (1997) determined a MIS 5d 36 Cl age. Finally, U-series ages by Sharp et al. (2003) date the type terrace of the Bull Lake glaciation (WR3) as late MIS 6, and an older Bull Lake terrace (WR4) as early MIS 6, arguing against a MIS 5d age for the Bull Lake.

Fig. 2. Plots showing differences between two cosmogenic-dating studies on boulders from the same Bull Lake moraine on west side of Wind River Range near Fremont Lake. These differences illustrate the current difficulty in using exposure ages to distinguishing MIS 5d from MIS 6. (a) Plot of ages in order of increasing resistance to erosion and difficulty in sampling, from (1) weathered plagioclase porphyroblastic granodiorite to (4) unweathered granite. Both 10 Be and 26 Al exposure ages increase with increasing resistance of boulders to erosion. Gosse & Phillips (2001) conclude these moraines are ∼150,000 yr old (MIS 6). (b) Plot of 36 Cl ages supplemented by 36 Cl/10 Be age for the same moraine as in Fig. 2A (Phillips et al., 1997). Phillips et al. (1997) conclude “the distribution of 36 Cl ages as quite similar to that for the Bull Lake IX moraine at Bull Lake, giving limits of 120 to 100 × 103 yr.” This age range would suggest correlation with MIS 5d, although the left side of this figure as well as U-series dating (see text) indicate a MIS 6 age.

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The ice-contact West Yellowstone flow is either recessional Bull Lake or a younger advance. Its age of 122,300 ± 2200 yr favors correlation with MIS 5d, although recession from the Bull Lake moraines (MIS 6) is permissible with the time scale of Martinson et al. (1987). However, if Winograd et al.’s (1997) dates at Devils Hole in southern Nevada are correct, and the last interglaciation (MIS 5e) in the western U.S. began 142,000 yr ago and lasted until 120,000 yr ago, then it follows that correlation of the ice-contact flow with MIS 5d is implied. An extensive mountain glaciation during MIS 5d appears plausible in that a major minimum in solar insolation for the Northern Hemisphere culminated about 110,000 yr ago (Berger, 1978), and cool summers combined with enhanced precipitation from warm oceans then could have fostered the expansion of mountain glaciers. For Mono Basin moraines in Bloody Canyon of the Sierra Nevada, Phillips et al. (1990) advocate a MIS 5d age, although this age is appears to be out of stratigraphic sequence (see Kaufman et al., this volume). Glaciation during MIS 5d at high latitudes is advocated for the Lake Baikal area, Siberia (Karabanov et al., 1998). One implication of the assignment of some Bull Lake advances to MIS 5d is that glaciers more extensive than the Pinedale glaciers advanced and retreated in only ∼12,000 yr (the MIS5d age-span of Martinson et al., 1987). Studies of ice cores in Antarctica and Greenland indicate it has taken ∼10,000 yr or longer to accumulate the upper 1 km of ice in these areas of relatively slow accumulation. The Pinedale icecap on the Yellowstone Plateau built up to a thickness of a kilometer after an advance from the adjacent mountains (Good & Pierce, 1996). Also, the Wind River Pinedale icecap on the east side of the range locally was 600 m thick (William Locke, Earth Sciences web site at Montana State University, 2002, http://www.homepage.montana.edu/∼ueswl/winds.html). Thus, there is a question whether adequate time exists between warm periods MIS 5c and 5e to build up icecaps thicker than the Pinedale, and then to deposit multiple, large-volume moraines. If this did occur in the 12,000-yr span of MIS 5d, it demonstrates the rapidity of icecap buildup, moraine deposition, and recession. Lake-sediment records from the U.S. have not been interpreted to show a major glaciation in MIS 5d comparable in magnitude with MIS 6 or 2. Such records have been studied from Lake Bonneville (Oviatt et al., 1999), Owens Lake, California (Bischoff et al., 1997, using the timescale of Bischoff & Cummings, 2001; Litwin et al., 1999), or from the mid-continent (Zhu & Baker, 1995). For Clear Lake, California, Adam (1988) shows that MIS 5d consisted of three cold intervals, each lasting only 1400–2400 yr. They were separated by two warm intervals of 3100–4700-yr duration. The oxygen-isotope record from Devils Hole, Nevada (Winograd et al., 1997) also shows MIS 5d as shorter than MIS 2, 4, and 6. Whitlock et al. (2000; and spoken comm., 2002) concluded from pollen studies of cores from Carp Lake, western Washington, that MIS 5d was cool and humid, and not as severe than MIS 2. The MIS-5d landscape was covered by open forest. In conclusion, early and late Bull Lake moraines have been distinguished in many Rocky Mountain areas (Richmond,

1965), locally on the basis of an intervening soil. Distinction as early and late Bull Lake may correlate either with: (1) MIS 6 and 5d respectively, as suggested for the Wind Rivers (Phillips et al., 1997; Chadwick et al., 1997); or (2) early and late MIS 6 (190,000–170,000 and 150,000–130,000 yr ago) as indicated by: (a) Sharp et al. (2003) for the Wind River areas (b) Fullerton et al. (2003) for early and late Bull Lake in the Glacier Park area, and (c) this paper for the West Yellowstone area. For different places, such contrasting correlations may be valid and show that different areas have contrasting surviving successions of Bull Lake-like moraines, a topic extensively developed in Gillespie & Molnar (1995). Nevertheless, I consider surviving Bull Lake moraines of MIS 5d yet to be established, where whereas those of MIS 6 are quite credible.

MIS 4 or Early Wisconsinan Glaciation Based on weathering-rind thickness, Colman & Pierce (1981, 1986, 1992) found three successions in the western U.S. had early Wisconsin moraines (∼60,000–70,000 yr old), including the Rocky Mountain sequence at McCall, Idaho. Other information also favors glaciation during MIS 4. On the east side of the Wind River Range, an outwash(?) terrace (WR2) between the Pinedale (WR1) and late Bull Lake (WR3) has a 230 Th/U age on carbonate coats of 55,000 ± 8600 yr (Sharp et al., 2003), suggesting an early Wisconsin glacial advance, but with moraines subsequently overridden by younger glaciers. For Owens Lake, eastern California, abundant rock flour signifies Sierran glaciation from 78,000 to 66,000 yr ago (Bischoff & Cummings, 2001). A relatively deep lake in the Bonneville Basin, Utah dates 59,000 ± 5000 yr, and probably represents cooler conditions near the MIS 4/3 boundary (Kaufman et al., 2001). A high stand of Summer Lake, southeastern Oregon dates between 89,000 and 50,000 yr, a period that includes MIS 4 (Cohen et al., 2000). A loess-buried soil section in southern Jackson Hole records glacial(?) loess deposition ∼65,000–75,000 yr ago based on TL ages, 10 Be accumulation, soil development, and minimum 14 C ages (K.L. Pierce, unpub. data). At Carp Lake, Washington, a coolhumid interval from 58,000 to 43,000 yr ago and cool-dry interval from 72,000 to 58,000 yr ago may represent glaciation in MIS 4 and early MIS 3 time, bracketed by warmer intervals (Whitlock et al., 2000). Thus, an early Wisconsin (MIS 4) glacial advance probably has occurred in the Rocky Mountains, but surviving moraines are not recognized (see Gillespie & Molnar, 1995), excepting for the McCall, Idaho area.

Numerical Ages of Pinedale Deposits Ages of the Pinedale (last) glaciation based on calibrated (cal) radiocarbon and obsidian-hydration dating are greater than new ages based on cosmogenic methods (Fig. 3). Fig. 3B shows the extent of Pinedale glaciation based on a 1983 compilation of 14 C and obsidian-hydration ages (Porter et al., 1983, Fig. 4–28). Between Fig. 3A and B, the correction of 14 C age (right) to calibrated age (left) is show by the sloping

Pleistocene Glaciations of the Rocky Mountains 69

Fig. 3. Development of chronologies of the last (Pinedale) glaciation of the Rocky Mountains showing age vs. percent full-glacial length. A. Current cosmogenic dating for Wind River Range (Gosse et al., 1995a, b) and the Yellowstone icecap (Licciardi et al., 2001). EWR and WWR-east and west Wind River Range with WWR-a, original terminal moraine age of Goss and WWRb, age calculated by J.M. Licciardi (written comm., 2003) using current scaled production rates; Y-Yellowstone. B. Plots on radiocarbon time scale showing 1965 model and 1983 compilation (both from Porter et al., 1983). For 1983 compilation of mountain glaciers and icecaps, control points are plotted and indexed to three tables (Porter et al., 1983, Figs 4–28, Tables 4–6 to 4–8). The center column between A and B shows the increase in age between radiocarbon ages (right) to calibrated ages (left); the shaded bands highlight this age offset. The cosmogenic chronology dates recession from near the terminal moraine position about 15,000 yr ago, whereas the calibrated 14 C compilation indicates an age nearer 20,000 cal yr B.P. lines and is based on Stuiver & Reimer (1993, Calib. program) and Bard et al. (1998). Figure 3B shows the Pinedale glacial maximum occurred 25,000 to 20,000 14 C yr B.P. (∼23,400 to 29,200 cal yr B.P.) and was followed by rapid recession of ice-caps by 15,000 14 C yr B.P. (∼18,000 cal yr B.P.) and recession of valley glaciers to near cirque positions (shown as a small readvance) by 12,000 14 C cal yr B.P. The 1983 model shows near complete deglaciation about 5000 yr earlier than the 1965 model (Fig. 3B, dotted line). More recent 14 C ages also support the plot shown in Fig. 3B, with relatively old ages for the Pinedale maximum, and rapid recession to the cirques. For two glacially dammed lakes in the Colorado Front Range, Rosenbaum & Larson (1983) report the four oldest ages on finely disseminated plant debris were between 22,000 and 23,200 (± ∼1000) 14 C yr B.P. (27,000 to 25,700 cal yr B.P.). These newer 14 C ages are consistent with the antiquity of 14 C ages from the Colorado Front Range that were used in plotting Fig. 3B: (1) organic matter concentrated from sediments in a morainedammed lake dates to 22,400 14 C yr B.P. (Madole, 1980), and (2) a section of interbedded till and peat indicates that the Pinedale glaciation lasted from >30,000 to 13,700 14 C yr B.P. (Nelson et al., 1979). In the Yellowstone-Grand Teton area, Whitlock (1993) obtained basal ages from cores in sediments deposited during the Pinedale recession. These ages were as old as 14,580 ± 150, 17,160 ± 210, 16,040 ± 220, 15,640 ± 150 and 14,490 ± 700 14 C yr B.P., which would indicate glacial recession there by 20,000–19,000 cal yr B.P.

Cosmogenic Surface Exposure Ages of Pinedale Deposits Cosmogenic ages for the Wind Rivers and Yellowstone produce ages for the Pinedale glacial maximum and recession that are younger than the 14 C based time-distance plot (Fig. 3). For the type Pinedale moraines on the west side of the Wind River Range, Gosse et al. (1995a) determined a terminal moraine age of 21,700 ± 700 10 Be yr. They also found that moraine building continued near the terminus for ∼6000 yr, or until ∼15,800 10 Be yr ago. Similarly, on the east side of the Wind River Range, Phillips et al. (1997) measured 36 Cl ages supplemented by 10 Be ages for boulders on the three outer Pinedale moraines, and concluded the moraines are between 23 and 16,000 yr old. From the values determined by Gosse et al. (1995a), recalculation of the mean of each moraine group using current scaled production rates yields the following ages for the type Pinedale moraines (J.M. Licciardi, written comm., 2003): Moraine or Group

Age (×103 10 Be yr)

Terminal Pinedale Recessional group Difference

20.1 ± 1.0 17.6 ± 0.8 ∼2.5

In contrast, for the Pinedale sequence of the northern Yellowstone outlet glacier, Licciardi et al. (2001) analyzed a

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Kenneth L. Pierce

goodly number of boulders on each deposits and determined the following younger exposure ages:

had occurred by about 12,000 14 C yr B.P. (14,000 cal yr B.P., Carrara, 1995), also suggesting rapid retreat.

Deposit (Number of Boulders Dated)

Cosmogenic Ages (×103 yr)

Discussion of Pinedale Ages

Eightmile terminal moraines (8) Eightmile terminal moraines (9) Chico recessional moraines (8) Deckard Flats readjustment (10) (Yellowstone Plateau icecap no longer contributing) Late glacial outburst flood (6)

16.5 ± 0.4 3 He 16.2 ± 0.3 10 Be 15.7 ± 0.5 10 Be 14.0 ± 0.4 10 Be 13.7 ± 0.5 10 Be

These cosmogenic ages for the Eightmile (Pinedale) terminal moraines of the Yellowstone outlet glacier are about 3000–4000 10 Be yr younger than the 20,100 10 Be yr ages on outermost Pinedale terminal moraines on both sides of the Wind River Range (Fig. 3). This is probably a real age difference that can be explained by either: (1) the time interval required for the progressive buildup of the Yellowstone icecap; or (2) a climatic difference related to nearness to the Laurentide ice sheet (J.M. Licciardi, written comm., 2003). The 3000–4000 10 Be yr contrast between the Yellowstone and the Wind Rivers in Pinedale terminal moraines shows that subdivision into early and middle stades does not necessarily indicate correlation from place to place. For the Wallowa Range in eastern Oregon, the following are averages of cosmogenic ages on boulders on close-spaced moraines in the Wallowa Lake area (TTO, TTY, WTO) and in the glacial cirques (GL) (Licciardi, 2000; written comm., 2003): Moraine (Number of Boulders Dated)

Age (×103 10 Be yr)

TTO & some TTY (9) WTO & some TTY (9) GL (4), near cirques

21.1 ± 0.4 17.0 ± 0.3 10.2 ± 0.6

These ages are quite similar to those of the Wind Rivers. From a Pinedale moraine site in Colorado Front Range, a few boulders yield an exposure age of 16,800 ± 3400 10 Be and 26 Al yr B.P. (Dethier et al., 2003). This age is also more than 5000 yr younger than corrected 14 C ages for moraine-dammed lakes in the same area. The rate of retreat from the terminal-moraines areas to the cirques was remarkably fast (Fig. 3). For the Wind Rivers, glacial retreat of 33 km occurred at an average rate of 7.4 m/yr to moraines dated 12,100 ± 500 10 Be yr in the cirque basins (Gosse et al., 1995b). Retreat of the northern Yellowstone outlet glacier also occurred at the rapid rate of ∼30 m/yr, given the 15,700 10 Be yr for Chico recessional moraines and 13,700 10 Be yr for flood deposits that originated more than 60 km upvalley (Licciardi et al., 2001). For the Glacier National Park area (Carrara, 1987), much of deglaciation

The cosmogenic ages for moraines of the northern Yellowstone outlet glacier are thousands of years younger than the Yellowstone chronology based on 14 C, obsidian-hydration (Pierce et al., 1976), and U-series ages (Sturchio et al., 1994), although none of non-cosmogenic methods used directly dated moraine deposition. As discussed previously, the calibrated 14 C ages for deglaciation of the Yellowstone Plateau are as old as 20,000–19,000 cal yr B.P. (Millspaugh et al., 2000; Porter et al., 1983; Whitlock, 1993; see discussion in Licciardi et al., 2001). Either the cosmogenic ages are too young, or the carbon samples contained a large fraction of old carbon. Many of the carbon samples dating deglaciation were of finely disseminated organic carbon, which could be contaminated by older carbon. But some samples such as a peaty mud dating ∼15,800 cal yr B.P. (13,140 ± 700 14 C yr B.P.; Porter et al., 1983, W-2285) from near the center of the Yellowstone Plateau ice cap consisted of recognizable plant fragments and is inconsistent with the 14,000 ± 400 10 Be yr age of the Deckard Flats readjustment which predates the complete deglaciation of the Yellowstone Plateau. Obsidian-Hydration Age of Pinedale Glaciation The obsidian-hydration dating for Yellowstone assumes that under a constant temperature the square of hydration thickness is a linear function of age and that the Yellowstone obsidians sampled hydrate at the same rate (Pierce et al., 1976). Hydration was measured on pressure cracks resulting from glacial grinding up to the time of moraine deposition. Age calibration for obsidian in Bull Lake moraines is based on their hydration thickness compared to that on bracketing 122,000- and 183,000-yr-old rhyolite flows (Pierce et al., 1976). But no bracketing rhyolite flows exists for the Pinedale moraines, and age estimation is particularly sensitive to both the difference in soil temperature between late Pleistocene and Holocene time and the present soil temperature difference between the Yellowstone Plateau and the West Yellowstone Basin (Pierce et al., 1976). An age younger than the ∼30,000–25,000 yr obsidian-hydration age calculated for the Pinedale moraines near West Yellowstone (Pierce et al., 1976) would result either if (1) the Pleistocene-Holocene temperature difference was larger than the estimated 6 ◦ C or (2) the temperature difference between the Yellowstone Plateau and West Yellowstone basin was larger than the estimated 0.5 ◦ C, or (3) both. Moraines Attributed to Younger Dryas The younger Dryas (YD) was a dramatic cooling episode between 11,000 and 10,000 14 C yr B.P. (12,800 ± 200 and

Pleistocene Glaciations of the Rocky Mountains 71 11,500 ± 300 cal yr B.P.) in Europe. It has been attributed to a dramatic southward extension of cold North Atlantic water. At several localities in the Rocky Mountains, dating studies correlate a minor readvance in or near high cirques with the YD. In the core of the Wind River Range, the Inner Titcomb Lakes moraine dates 13,800 ± 600 to 11,400 ± 500 10 Be yr (Gosse et al., 1995b; see Fig. 3). Calculation of the mean age using the latest scaled production rates yields a mean age of 12, 600 ± 500 10 Be yr (J.M. Licciardi, written comm., 2003). The Titcomb Lakes moraines correlate with the nearby Temple Lake moraines dated by sediment changes in nearby lake sediments as between 13,800 ± 900 and 11,800 ± 700 cal yr B.P. (Davis et al., 1998). In the Colorado Front Range, the Triple Lakes cirque moraines of Benedict (1985) have minimum 14 C ages just over and under ∼10,000 14 C yr B.P. (11,200 cal yr B.P.) indicating a latest Pleistocene age and suggesting they are candidates for the YD event (Davis, 1987). Also in the Colorado Front Range, lake sediments reflecting a nearby glacier advance date between 13,200 and 11,100 cal yr B.P. (Menounos & Reasoner, 1997). In Colorado Mountains, pollen spectra suggest a YD cooling event between 13,500 and 12,900 cal yr B.P. in Black Mountain Lake and 13,600 and 12,900 cal yr B.P. in Sky Pond (Reasoner & Jodry, 2000). In the southern Sangre de Cristo Range of New Mexico, lake sediments date an YD advance about 11,800 to 11,00 cal yr B.P. (Armour et al., 2002).

Other Rocky-Mountain Glaciation Studies Modeling Glacial Flow Basal shear stress, mass balance, and glacial flow have been reconstructed for some late Pleistocene glaciers in the Rocky Mountains. For the glacial geologic reconstruction of the northern Yellowstone outlet glacier, basal shear stress averaged 1.2 bars for strongly extending reaches (converging flow lines), 1.0 bars for 11 uniform reaches, and 0.8 bars for 16 strongly compressing reaches (Pierce, 1979). Assuming precipitation similar to present, best estimates of mass balance yielded an annual accumulation of 2.8 km3 above the ELA of 2850 m and annual loss of 3.3 km3 below the ELA. Based on these parameters for a cross-section downvalley from the ELA, annual ice discharge was 2.7 km3 , of which 10% is modeled by flow within the glacier and the remaining 90% is attributed to basal sliding (Pierce, 1979). On the mountains of northwestern Montana, a large ice cap built up on a complex mountain topography that includes several ranges. Based on terminal moraine positions, divide crossings, and nunataks, Locke (1995) modeled this ice sheet assuming basal shear stress values near 1 bar to produce a contour map of the ice surface. His model revealed much about the sources and non-sources of the multiple glacial lobes that radiate outward from this glacial source area. For some canyons draining the east side of Wind River Range icecap, W. Locke (written comm., Earth Sciences web site at Montana

State University, 2002) calculates that the basal shear stress was as high as 8 bars. Possible factors for such high values are: (1) the need to funnel accumulation from a central icecap into narrow, steep canyons; and (2) low subglacial water pressure due to the draining of the glacial-bed water into permeable limestones high above the base level of the Wind River Basin.

Sedimentology of Glacial Deposits On the southern margin of the San Juan Mountains, Colorado, glacial system, Johnson & Gillam (1995) found that end moraines are primarily of debris-flow sediment interbedded with sandy stream sediment. They conclude that “existing moraines were built rapidly (in 10 yr to a few tens of years)” but that outwash deposits indicate that the glacier stood at its terminus for a long time.

Pleistocene Climates The contrast between present and glacial climates and departures of Pleistocene climate patterns from present can be made by comparison of late Pleistocene ELA’s with modern snowpack patterns. For the Great Basin, Zielinski & McCoy (1987) located anomalies in the distribution of modern snowpack at late Pleistocene ELA’s, in particular an anomaly in the NW Great Basin which appears to have been considerably wetter in Pleistocene times than present; this difference was also noted by Porter et al. (1983). Leonard (1984) found that contours on late Pleistocene ELA’s for the San Juan Mountains in southern Colorado closely resembled modern snowpack pattern, suggesting little change in atmospheric circulation patterns. For the Sawatch Range in central Colorado, late Pleistocene mean summer temperatures are estimated to have been 7–9 ◦ C colder (Brugger & Goldstein, 1999). Using modern climatic conditions at Pleistocene ELAs, Leonard (1989) found that Pleistocene glaciers could be maintained with either: (1) no change in total precipitation and 8.5 ◦ C colder in summer, or (2) about 10–13 ◦ C colder using Mears (1987) estimate of regional cooling and at least a 44% reduction in fall-through-spring precipitation. Locke (1990) used Mears (1987) estimate of regional cooling of 10 ◦ C to infer an ∼25% decrease in last glacial precipitation in Montana; Murray & Locke (1989) used ice-flow theory to independently estimate a similar drying. Hostetler & Clark (1997), using a nested modeling strategy, determined that at 18,000 yr ago climatic conditions varied across the western U.S. Glaciers in the southern Rockies were sustained through decreases of 9–12 ◦ C with little change in precipitation, whereas glaciers in the northern Rockies existed under relatively cold-summer and dry-winter conditions resulting from easterly anticyclonic wind flow off the Laurentide Ice Sheet. However, no evidence for this easterly flow was found by Muhs & Bettis (2000) based on late Wisconsin loess distribution, or by Locke (1990) or Gillespie (1991) based on changes in glacial ELA

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patterns in the northern Rocky Mountains and Sierra Nevada respectively. Modeling of the climatic response in Western North America to the changes associated with sudden lowering of the Laurentide Ice Sheet in a Heinrich Event show a complex response that in different areas and times reinforce, cancel, or reverse local changes and thus make difficult simple correlations in the region (Hostetler & Bartlein, 1999).

Possible Uplift and Subsidence Belts of uplift and subsidence may explain different patterns in the relative extent of Bull Lake and Pinedale glaciations near Yellowstone. The normal pattern is that Pinedale glaciers were ∼90% as long as Bull Lake glaciers. But for both the Greater Yellowstone glacial system and nearby independent valley glaciers on the western and southern sides of Yellowstone, Pinedale glaciers are less than 80% and locally 60% the length of Bull Lake glaciers; whereas, on the northern and eastern sides of Yellowstone, Pindeale glaciers were >100% the length of Bull Lake glaciers and overran Bull Lake moraines (Pierce & Morgan, 1992). An explanation of this pattern is that the terrain to the northeast is rising and that to the southwest is subsiding. These areas appear to be on the currently uplifting (NE) and subsiding (SW) slopes of the “bow wave” of the Yellowstone hot spot. Compared to Bull Lake time, areas to the northeast became relatively higher in Pinedale time whereas those to the southwest became relatively lower (Pierce & Morgan, 1992).

Conclusions and Recommendations for Future Study In their provocative paper “Asynchronous maximum advances of mountain and continental glaciers” Gillespie & Molnar (1995) cite much evidence that mountain glaciation did not proceed in lock step with either continental glaciation or its proxy, the marine oxygen isotope record. More dating may establish such variation for the Rocky Mountains. Global climate models (Kutzbach & Guetter, 1986) suggest that glacial-anticyclonic circulation weakened westerly flow and resulted in air cooler and drier than present, particularly for the northern Rocky Mountains. But Locke (1990) observed that the gradient of maximum glacial ELAs exactly parallels that of modern ELAs, indicating a similar rather than contrasting climatic pattern. In particular Locke noted no maximum glacial lowering of ELAs due to upslope precipitation on the east side of ranges that would be associated with such easterly flow. Assuming the Kutzbach & Guetter model results would suggest that differences in timing and relative magnitude of late Pleistocene glaciation is likely southward through the Rocky Mountains with greater distance from the continental ice sheet. Glacial culminations might follow storm tracks located at some distance south of the expanding or contracting continental ice sheet, a topic addressed by Licciardi (2000; written comm., 2003). More precisely dated glacial and lacustrine records may reveal

patterns in such non-parallelism from south to north (colder) or east to west (wetter) throughout the Rocky Mountains. Cosmogenic dating of Pinedale glaciers on both sides of the Wind River Range suggest a Pinedale culmination near the last glacial maximum 21,000 yr ago, but fail to reveal surviving older moraines of MIS 2, 3, or 4 age. Although imprecise, weathering rind thicknesses, some radiocarbon ages, and obsidian hydration suggest Wisconsin glacial advances in the 50,000–35,000 and 70,000–60,000 yr range (Colman & Pierce, 1981; Pierce et al., 1976; Porter et al., 1983). A 55,000-yr-old Wind River terrace (Sharp et al., 2003) probably reflects early Wisconsin (MIS 4) glaciation. Records of glacial flour in lake sediments of the Rocky Mountains similar to that for the Sierra Nevada (Bischoff & Cummings, 2001) and the Cascades (Rosenbaum & Reynolds, 2003) could reveal much about the glacial record both preserved or subsequently overridden in the end moraine record, including helping resolve the relative magnitude of glacial activity in MIS 3, 4, 5d, and 6. For the many glacial cycles documented by the MIS record older than MIS 6, only the Sacagawea Ridge moraines of MIS 16 are well correlated with a particular MIS. The potential distinction of moraines of MIS 4, 5d and MIS 6 may present interesting contrasts. The conference on the last interglaciation (Kukla et al., 2002 and references therein) presents current knowledge about MIS 5d and its contrasts with MIS 5e and 5c. A warm ocean, a major low in northern Hemisphere solar insulation, and glacial expansion particularly at high latitudes, accompanied MIS 5d. What were the glacial and climatic conditions along a north-south transect through the Rockies? In MIS 2, climate was cold and probably drier in the northern Rockies (Whitlock et al., 2000), and MIS 6 was probably similar. Was MIS 5d a different character of glaciation than MIS 2 and 6 in the Rockies? If MIS 6 was similar to MIS 2, why are Bull Lake moraines (MIS 6 or MIS 5e) more bulky? The cross-feed between data for Pleistocene glaciations across the western mountains and climatic models such as Hostetler & Clark (1997) and Hostetler & Bartlein (1999) will enhance understanding in both subject areas. Cosmogenic dating can benefit by refinement in calibration and changes in cosmic-ray flux through time. For boulders on deposits ∼100,000 yr and older, the history of the boulder needs to be better understood including erosion of the boulder surface, emergence of the boulder from the eroding moraine, deposition and deflation of loess, and burial by snow.

Acknowledgments I thank Pete Birkeland, Bill Locke, Joe Licciardi, Linda Pierce, and Nicole Davis for helpful reviews and editing. In the preparation of this paper, I benefited from discussions with Peter Clark, Bill Locke, Alan Gillespie, Steve Colman, Dave Fullerton, Joe Licciardi, Warren Sharp, Darrell Kaufman, Dave Dethier, Cathy Whitlock, Pete Birkeland, Dennis Dahms, Ralph Shroba, Eric Leonard, John Good, Don Easterbrook, John Obradovich, and Joe Rosenbaum.

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Ph.D. dissertation, University of Colorado, Boulder, 169 pp. Noller, J.S., Sowers, J.M. & Lettis, W.R. (Eds) (2000). Quaternary geochronology, methods and applications. AGU Reference Shelf, 4, 582 pp. Obradovich, J.D. (1992). Geochronology of the Late Cenozoic Volcanism of Yellowstone National Park and adjoining areas, Wyoming and Idaho. U.S. Geological Survey Open File Report 92–408, 45 pp. Osborn, G. & Bevis, K. (2001). Glaciation of the Great Basin of the western United States. Quaternary Science Reviews, 20, 1377–1410. Oviatt, C.G., Thompson, R.S., Kaufman, D.S., Bright, J. & Forester, R.M. (1999). Reinterpretation of the Burmeister Core, Bonneville Basin, Utah. Quaternary Research, 52, 180–184. Phillips, F.M., Zreda, M.G., Gosse, J.C., Klein, J., Evenson, E.B., Hall, R.D., Chadwick, O.A. & Sharma, P. (1997). Cosmogenic 36 Cl and 10 Be ages of Quaternary glacial and fluvial deposits of the Wind River Range, Wyoming. Geological Society of America Bulletin, 109, 1453–1463. Phillips, F.M., Zreda, M.G., Smith, S.S., Elmore, D., Kubik, P.W. & Sharma, P. (1990). Cosmogenic Chlorine-36 chronology for glacial deposits at Bloody Canyon, eastern Sierra Nevada. Science, 248, 1529–1532. Pierce, K.L. (1979). History and dynamics of glaciation in the northern Yellowstone National Park area. U.S. Geological Survey Professional Paper 729 F, 91 pp. Pierce, K.L. (1985). Quaternary history of movement on the Arco segment of the Lost River fault, central Idaho. In: Stein, R.S. & Bucknam, R.C. (Eds), Proceedings of Workshop XXVII on the Borah Peak, Idaho, earthquake. U.S. Geological Survey Open-File Report 85–290, pp. 195– 206. Pierce, K.L. & Good, J.D. (1992). Field guide to the Quaternary geology of Jackson Hole, Wyoming. U.S. Geological Survey Open-File Report 92–504, 49 pp. Pierce, K.L. & Morgan, L.A. (1992). The track of the Yellowstone hot spot–volcanism, faulting and uplift. In: Link, P.K., Kuntz, M.A. & Platt, L.W. (Eds), Regional geology of eastern Idaho and western Wyoming. Geological Society of America Memoir, 179, 1–53. Pierce, K.L., Obradovich, J.D. & Friedman, I. (1976). Obsidian hydration dating and correlation of Bull Lake and Pinedale glaciations near West Yellowstone, Montana. Geological Society of America Bulletin, 87, 703–710. Pierce, K.L. & Scott, W.E. (1982). Pleistocene episodes of alluvial-gravel deposition, southeastern Idaho. In: Bonnichsen, B. & Breckenridge, R.M. (Eds), Cenozoic geology of Idaho. Idaho Bureau of Mines and Geology Bulletin 26, 685–702. Porter, S.C. (1975). Weathering rinds as a relative-age criterion: Application to sub-division of glacial deposits in the Cascade Range. Geology, 3, 101–104. Porter, S.C., Pierce, K.L. & Hamilton, T.D. (1983). Late Pleistocene glaciation in the Western United States. In: Porter, S.C. (Ed.), The Late Pleistocene, Vol. 1, of: Wright, H.E., Jr. (Ed.), Late Quaternary Environments of the United

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Quaternary alpine glaciation in Alaska, the Pacific Northwest, Sierra Nevada, and Hawaii Darrell S. Kaufman1 , Stephen C. Porter2 and Alan R. Gillespie2 1 2

Department of Geology, Northern Arizona University, Flagstaff, AZ 86011, USA; [email protected] Quaternary Research Center, University of Washington, Seattle, WA 98195, USA; [email protected], [email protected]

most mountain fronts. Evidence for multiple glacier fluctuations is also preserved in successions of glacially influenced deposits of lacustrine (e.g. Lake Atna: Ferrians, 1963; Lake Noatak: Hamilton, 2001), marine (e.g. Yakataga Formation: Plafker & Addicott, 1976; Hagemeister Island: Kaufman et al., 2001a), fluvial (e.g. Epigurak Bluff: Hamilton et al., 1993), and eolian systems (loess and sandsheets: Begét, 2001; Lea & Waythomas, 1990; and dunes: Carter, 1981; Mann et al., 2002). The glacial history is often better dated in these depositional settings where volcanic products and organic material are more commonly preserved. The available age control shows that Alaska has a glacial record as long as that of anywhere else in the Northern Hemisphere. The geochronology also shows that glaciers fluctuated on time scales ranging from tens of thousands of years to decades, consistent with other records of global climatic changes. The ages of Holocene (e.g. Calkin et al., 2001) and late Wisconsin (e.g. Porter et al., 1983) glacier fluctuations are best documented. The ages and correlations of glacier deposits of the next-older (penultimate) ice advance are beyond the range of 14 C dating and have remained controversial. Considering the vastness of the glaciated area in Alaska, the diversity of its glacial systems, and the widespread impact of glaciers on the non-glaciated regions, a single review cannot include a full account of Alaskan glacial geology. Instead, this section presents a brief summary of the major developments in Alaskan glacial geology prior to 1990, then highlights some of the progress in the last 10 years, since the last state-wide summaries of the Pliocene-Pleistocene (Hamilton, 1994) and Holocene (Calkin, 1988) glacial chronologies.

Introduction This chapter deals with mountain glaciation in the North American Cordillera and Hawaii, exclusive of the Rocky Mountains (Pierce, this volume) and Canada. We focus on a few critical areas in each of these regions where research since 1965 has produced significant new results that advance our understanding of the extent, chronology, and dynamics of mountain glaciers, and enhance the paleoclimatic inferences that can be drawn from them.

Alaska Unlike other high-latitude areas of North America, much of Alaska was never glaciated (Fig. 1). Even more land area lay exposed to the arid Pleistocene climate during intervals when sea level was lower than present. Despite the vastness of its unglaciated area, Alaska’s mountainous terrain generated a mass of glacier ice on a par with all the rest of the western United States combined. The largest expanse of glaciers comprised the coalescent ice caps and piedmont lobes that extended from the Alaska Range to the Gulf of Alaska and from the southeastern panhandle to the Aleutian Islands. This amalgamation formed the western extension of the North American Cordilleran Ice Sheet, and it contained most of the glacier ice in Alaska. The ice caps that grew in the Brooks Range, a northern extension of the Rocky Mountain system in northern Alaska, and the Ahklun Mountains in southwestern Alaska were the only other major centers of ice accumulation in the state. Lower uplands across the state supported hundreds of smaller valley glaciers; most notable are the small ranges of Seward Peninsula and the Yukon-Tanana Upland. In all, glaciers once covered about 1,200,000 km2 of Alaska and its adjacent continental shelf; during late Wisconsin time, the area was 727,800 km2 (Manley & Kaufman, 2002). Presently, 74,700 km2 of Alaska is covered by ice, or 4.9% of the state; most of the present volume of glacier ice is in the coastal ranges proximal to moisture sources around the Gulf of Alaska. Because most of Alaska was not glaciated, mountain glaciers were free to expand onto adjacent lowlands where they left a rich record of moraines and morphostratigraphically related glacial-geologic features. Evidence for the extent of glaciers around the Gulf of Alaska is now submerged and obscured, but a succession of moraines is preserved along

Status in 1965 Prior to the widespread use of helicopters, glacial-geologic research in Alaska was focused in the south-central part of the state, particularly in areas accessible by road (Péwé, 1965). The first attempt at a state-wide summary of Pleistocene glacier extents was by Péwé (1953). This effort led to the statewide surficial map (Karlstrom et al., 1964) and its derivative map of glacier extents (Coulter et al., 1965), both published about the time of the VII INQUA Congress. By the mid-1960’s, the overall distribution of Quaternary glaciers in Alaska was known generally. Fifteen local Quaternary glacial sequences had been studied and the major


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Fig. 1. Pleistocene maximum, late Wisconsin, and modern glacier extent across Alaska. Image from the Alaska Paleoglacier Atlas website (Manley & Kaufman, 2002).

advances in each sequence were tentatively correlated (Péwé et al., 1965). Drift was subdivided into three principal units: Wisconsin, Illinoian, and pre-Illinoian, based mainly on semiquantitative relative-weathering criteria and comparison with the mid-continent region. About half of the local sequences included a two-fold subdivision of the Wisconsin glaciation (early and late) and many authors recognized evidence of multiple advances during the early Holocene. Drift interstratified with marine deposits around Nome, Kotzebue (Hopkins et al., 1965), and Cook Inlet (Karlstrom, 1964), provided additional age control based on the now-discredited Th/U dating of molluscs, and a few 14 C ages had been determined; these were some of the first 14 C analyses (ca. 1953) ever made. The dramatic difference between the limited extent of late Wisconsin glaciers and the vast extent of pre-Wisconsin ice was identified in most areas, and the exceptionally long record of glacial-marine sedimentation dating back to the latest Miocene had been recognized around the Gulf of Alaska (Miller, 1957).

Glacial-Geologic Research Between 1965 and 1990 Péwé (1975) reviewed the Quaternary geology of Alaska. His comprehensive synthesis included the first and only statewide compilation of glaciation thresholds. By the 1980s widespread application of 14 C dating had greatly refined the chronology of late Wisconsin glacier fluctuations; in Alaska; the technique was then, and has been since, most extensively applied to deposits of the Cordilleran Ice Sheet (Hamilton & Thorson, 1983) and in the Brooks and northern Alaska ranges (Porter et al., 1983). These studies demonstrated that the most recent glaciation in Alaska occurred between 24,000 and 11,500 years ago and was therefore broadly synchronous with late Wisconsin glaciation elsewhere in North America. Hamilton et al. (1986) complied the most complete collection of glacial-geologic studies in Alaska yet published. The volume includes detailed reports on ten different glaciated regions of the state, including: central Brooks Range (Hamilton, 1986), Seward Peninsula (Kaufman &

Quaternary Alpine Glaciation in Alaska, the Pacific Northwest, Sierra Nevada, and Hawaii 79 Hopkins, 1986), Yukon-Tanana Upland (Weber, 1986), Nenana River valley (Thorson, 1986), Beaver Mountains and northwestern Alaska Range (Kline & Bundtzen, 1986), Alaska Peninsula (Detterman, 1986), Aleutian Islands (Thorson & Hamilton, 1986), Cook Inlet (Schmoll & Yehle, 1986), Gulf of Alaska (Molnia, 1986), and southeast Alaska (Mann, 1986). Readers interested in the glacial geology of these regions are referred to this volume. The regional studies employed a wide range of approaches, but nearly all included 14 C ages or other geochronology, descriptive accounts of relative-weathering features, and paleoclimatological inferences. The authors recognized new complexities in the glacial sequences resulting from their detailed stratigraphic and geochronological studies; these were commensurate with their emerging understanding of Quaternary global climate changes provided by marine oxygen isotopes. About the same time, evidence for Holocene glaciation in Alaska was reviewed by Calkin (1988) and modern glaciers were discussed by Krimmel & Meier (1989). Hamilton’s (1994) review of Pleistocene glaciation of Alaska is the most up-to-date and complete synthesis. Based largely on the 1986 regional summaries, and updated with information from 109 additional papers published after 1986. Hamilton integrated evidence for late Cenozoic glaciation in 15 regions of the state and correlated glacial deposits within six broad age categories. Although the extent of glaciers during late Wisconsin time was relatively well known by the time of the review, the ages and correlations of glacier deposits of the next-older (penultimate) ice advance were controversial. Previous studies generally inferred that the penultimate advance was younger than the last interglaciation (i.e. early Wisconsin s.l.), but tephrostratigraphic, paleoecologic, pedogenic, and thermoluminescence evidence from the southern and central parts of Alaska suggested that the penultimate drift might predate the last interglaciation. Recognizing the likelihood that the drift might be of different ages in different places, Hamilton (1994) avoided the term “early Wisconsin” in favor of “penultimate” for glacier advances beyond the range of 14 C dating regardless of whether they predated or postdated the last interglaciation, and he assigned the next older drift, known to predate the last interglaciation, to the middle Pleistocene. Several new studies (see below) have amassed geochronological evidence favoring an early Wisconsin (s.l.) age for the penultimate drift.

Progress During the Last Decade With the recent retirement and death of several prominent Alaskan Quaternary geologists, research into the glacial geology of Alaska has slowed during the last decade. In the recently published volume on the Quaternary Paleoenvironments of Beringia (Elias & Brigham-Grette, 2001), for example, only five of 35 chapters are devoted to the glacial geology of Alaska. Nonetheless, significant progress has been made as the research has shifted from the firstgeneration, regional mapping (mainly by the U.S. Geological Survey) to technique- and hypothesis-driven investigations

at smaller scales (mainly by university scientists and their students). Areas of active glacial-geologic research include the Brooks Range, Ahklun Mountains, and Pacific coastal mountains.

Brooks Range Building on decades of systematic, surficial-geologic mapping across the central Brooks Range, Hamilton (2002) recently completed a detailed glacial-geologic study of the Itkillik-Sagavanirktok River area. He identified six distinct late Pleistocene moraine sets, the most complete subdivision of the Ikillik (Wisconsin) glaciation yet discovered in the Brooks Range. The glacial history of the western and eastern sectors of the Brooks Range remains uncertain, however. In the west, Hamilton’s (2001) recent study of the stratigraphic record exposed in the Noatak River basin has revealed evidence for multiple expansions of glaciers from the DeLong Mountains that dammed a succession of proglacial lakes. New photo-interpretive mapping and 14 C dating (T.D. Hamilton, unpub. data) indicate that glaciers were considerably less extensive (by an order of magnitude) in the DeLong Mountains during the late Wisconsin than has been depicted in previous state-wide compilations. Restricted ice in the DeLong Mountains during the late Wisconsin is consistent with the near absence of glaciers in the Baird Mountains south of the Noatak River (D.S. Kaufman, unpubl. data).

Ahklun Mountains The Ahklun Mountains of southwestern Alaska supported the largest center of Pleistocene glaciers outside the Cordilleran Ice Sheet and the Brooks Range. Research during the last decade has clarified the age and extent of multiple Quaternary glacier advances. Amino acid analysis of mollusc shells from glaciomarine sediment in coastal exposures of northeastern Bristol Bay, combined with 40 Ar/39 Ar dating of lava from a tuya eruption, indicates that extensive piedmont glaciers emanating from an ice cap centered over the Ahklun Mountains advanced south across the present coast as many as four times during the middle Pleistocene (Kaufman et al., 2001a). The youngest glaciers to reach the coast formed ice-thrust ridges containing glacially influenced marine sediment in Nushagak Bay area (Lea, 1990) and ice-contact stratified drift in the Togiak Bay area (Kaufman et al., 2001b). Luminescence, amino acid, paleoecologic, and tephrostratigraphic evidence shows that the drift is younger than the last interglaciation (Kaufman et al., 1996, 2001a, b; Manley et al., 2001). The glacial-geologic evidence for thin, low-gradient glacier ice indicates a limited glacial-isostatic effect; this, combined with the available geochronology and the long distance to the edge of the shallow continental shelf, suggests that the glaciomarine sediment was deposited during periods of high eustatic sea level. Similar evidence from elsewhere in central Beringia (Brigham-Grette et al., 2001; Huston

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et al., 1990; Pushkar et al., 1999) indicates that glaciers attained their maximum Pleistocene extent prior to the buildup of Northern Hemisphere ice sheets. The transitions between the interglaciations of marine oxygen-isotope stages (MIS) 11 and 5 and the subsequent glacial intervals are likely times when high sea level, warm sea-surface temperatures, and decreasing summer insolation conspired to generate the largest volumes of glacier ice in Alaska. The expansion of ice over high-latitude landmasses may have had an important positive feedback in the climate system during the onset of global glaciations. The first published exposure ages for moraines in Alaska (Briner & Kaufman, 2000; Briner et al., 2001) and 14 C ages from lake-sediment cores (Kaufman et al., in press) clarify the ages of late Pleistocene glaciations in the Ahklun Mountains. No evidence for an extensive ice advance during MIS 6 has yet been discovered. Instead, outlet glaciers in the southwestern part of the mountain range advanced beyond the present coast and reached their maximum late Pleistocene extent ∼60,000 years ago. Glaciers attained their maximum late Wisconsin extent ∼24,000–20,000 cal yr B.P. when they terminated more than 60 km upvalley from their early Wisconsin limits. They then experienced a series of fluctuations as summer insolation increased, sea level rose, and ocean-atmospheric circulation shifted to its interglacial mode. The most dramatic readvance culminated at the end of the Younger Dryas interval (Briner et al., 2002), consistent with emerging paleoenvironmental evidence for cooling around the state at that time (e.g. Bigelow & Edwards, 2001; Hu et al., 2002; Mann et al., 2001). In the Brooks Range (Hamilton, 1986) and Cook Inlet region (Reger & Pinney, 1996), however, prominent readvances occurred ∼1500 to 1000 years before the Younger Dryas, and no Younger Dryas advances have been recognized in these two regions. Glaciers retreated during an interval of early Holocene warmth, then reformed in the highest elevations of the Ahklun Mountains beginning ca. 3400 cal yr B.P. (Levy et al., in press).

The North Pacific Coast Motivated by continued interest in biotic exchanges between the old and new worlds, Mann & Hamilton (1995) recently reviewed the paleogeography of the North Pacific coast since the last glacial maximum. They summarized evidence from around the southern margin of the Cordilleran Ice Sheet for time-transgressive glacier fluctuations and for several major climatic transitions between about 26,000 and 10,000 cal yr B.P. The understanding of the glacial history of the Cook Inlet region has been improved recently with work by Schmoll et al. (1999), Reger & Pinney (1996), and Reger et al. (1995). On the Alaska Peninsula, recent work by Wilson & Weber (2001) has attempted to correlate multiple Pleistocene drift units, and to understand the interaction of volcanic and glacier activity. These efforts have been frustrated, however, by differential glacier response of different source areas (Stilwell & Kaufman, 1996), and by the recognition that many of the moraines surrounding

Bristol Bay are glacially tectonized ridges that do not necessarily record climatically significant ice-marginal positions (Kaufman & Thompson, 1998). Recent studies of the dendrochronology, lichenometry, and moraine geomorphology of the coastal mountains rimming the northern Gulf of Alaska (Barclay et al., 2001; Calkin et al., 2001; Wiles & Calkin, 1994; Wiles et al., 1999) and the Wrangell Mountains (Wiles et al., 2002) provide the most detailed and geographically most extensive record of late Holocene glaciation in the state. Neoglacial expansions of many glaciers took place by about 4000–3500 cal yr B.P. Glaciers retreated by ∼2000 cal yr B.P. before expanding again during the Little Ice Age advances of the 13th, 15th, middle 17th, and second half of the 19th centuries A.D.

On-Going Efforts Three decades after the last Alaska-wide compilation of glacial geology (Coulter et al., 1965), a collaborative effort has produced a new synthesis of reconstructed Pleistocene glacier extents (Manley & Kaufman, 2002). The Alaska PaleoGlacier (APG) Atlas integrates the results of glacialgeologic studies from 26 publications and 42 source maps into a Geographic Information System (GIS) targeted for a scale of 1:1,000,000. Maps and several GIS layers are available online (http://instaar.colorado.edu/QGISL/ak paleoglacier atlas). The APG Atlas is part of a larger effort led by the INQUA Commission on Glaciation to create a global GIS database of Pleistocene glacier extents. The atlas depicts several glaciated massifs that were not previously recognized by Coulter et al. (1965), mainly in the Yukon-Koyokuk region and the Kuskokwim Mountains. On-going spatial analysis based on more detailed digital mapping is aimed at reconstructing equilbribum-line altitudes (ELAs) and their paleoclimatic forcing using Pleistocene valley and cirque glaciers across Alaska (e.g. Manley & Kaufman, 1999). Despite substantial progress since 1965, our understanding of the Alaskan glacial record is hindered by major gaps in ground-based mapping and geochronologic control. Recently, for example, interpretations made from satellite images led to the inference that a major Pleistocene ice sheet covered much of central Beringia (the “Beringian Ice Sheet”; Grosswald, 1998). New research in far eastern Russia (e.g. Gualtieri et al., 2000), and previous studies (summarized by Brigham-Grette, 2001), refute the existence of this former ice sheet. New research is needed to address: (1) the age of the penultimate glaciation; (2) teleconnections to the global record of rapid climatic changes; (3) relation of the Alaskan glacial record to that in adjacent regions of the Yukon Territory (e.g. Froese et al., 2000; Westgate et al., 2001) and western Beringia (e.g. Glushkova, 2001; Heiser & Rousch, 2001); and (4) the role of sea level, atmospheric moisture, continentality, and other physiographic effects in controlling regional-scale response of glaciers to climate forcing.

Quaternary Alpine Glaciation in Alaska, the Pacific Northwest, Sierra Nevada, and Hawaii 81 Fig. 2. (a) Shaded relief map of Olympic Mountains and Cascade Range of Washington and Oregon showing localities mentioned in the text. (b) Shaded relief map of western Washington showing localities mentioned in the text.

Cascade Range and Olympic Mountains, Washington and Oregon The glaciated Cascade Range and Olympic Mountains of Washington and Oregon (Fig. 2) contain a wealth of data bearing on the Quaternary climatic and environmental history of the Pacific Northwest. The record of Pleistocene alpine glaciation is juxtaposed to a long record of ice-sheet glaciation found in adjacent lowlands, and terrain adjacent to hundreds of modern glaciers contains evidence of Holocene glacier variations. During their greatest Pleistocene advance, alpine glaciers in the Washington Cascade Range and Olympic Mountains terminated as much as 70–80 km from their sources. During the last glaciation, the largest glaciers were only half as long. In the Oregon Cascades, glacier tongues terminated 10–30 km from ice fields that mantled the range crest. The glaciated region encompasses a wide range of environments resulting from strong longitudinal and altitudinal climatic gradients. Cool, moist climates of western

Washington and high snowfall zones on the western flank and crest of the Cascades and Olympics contrast with rainshadow conditions and drier climate farther east. Not surprisingly, radiocarbon dating of events during the last glaciation and the Holocene is largely restricted to the wetter flanks and crests of the mountains. Only recently have other dating methods become available that have the potential of developing a chronological framework spanning a significant part of the glacial history of these ranges.

Status of Alpine Glacial Studies in 1965 Pleistocene Glaciation in the Cascade Range In his review of the Quaternary glacial record of the Puget Lobe of the Cordilleran Ice Sheet and the adjacent Cascade Range, Crandell (1965) recognized four Pleistocene drift units. The two youngest, Salmon Springs and Fraser, he

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Fig. 2. (Continued)

Fig. 3. Map of glaciated region of upper Yakima River drainage basin showing extent and inferred ages of Pleistocene drifts.

Quaternary Alpine Glaciation in Alaska, the Pacific Northwest, Sierra Nevada, and Hawaii 83 inferred to be of early to middle Wisconsin age and late Wisconsin age, respectively. Crandell divided the alpine glacial record into pre-Salmon Springs (?), Salmon Springs (?), Fraser, and Neoglacial drifts. Evidence of multiple glaciations had been found in valleys of the Entiat, upper Wenatchee, upper Yakima, Puyallup, Carbon, and White rivers, and at Mount Mazama (Fig. 2). Along much of the western range front in Washington, alpine limits were overlapped and obscured by drift of the Puget Lobe. In 1965, the general advance-retreat chronology of the Puget Lobe was based on constraining 14 C dates, but age limits for alpine chronologies were based mainly on dated Holocene tephra layers. Crandell (1965) inferred that Evans Creek drift near Mount Rainier (ca. 20,000 14 C years old) predated the maximum (Vashon) advance of the Puget Lobe (ca. 15,000–13,000 14 C yr B.P.). According to prevailing opinion, by the time of the Vashon advance, the largest alpine glaciers had greatly shrunk in size or even disappeared. Crandell (1965) summarized previous studies of the Oregon Cascades, noting that ice had covered the High Cascades at least once in late Pleistocene time. He also reported evidence of more-extensive glaciation in the west-draining valleys, as well as in the North Santiam River basin.

Pleistocene Glaciation on the Olympic Peninsula Crandell (1964, 1965) focused his reconnaissance study of the glacial record on the southwestern part of the peninsula where glaciers spread across low terrain and advanced toward the Pacific coast. He recognized three drifts of likely Wisconsin age (early, middle, and late), as well as one or more of preWisconsin age, distinguishing them on the basis of extent and contrasts in weathering.

Holocene Glaciation in the Cascades In 1965, Crandell & Miller (1965) proposed a chronology for post-Hypsithermal (middle Holocene) glacier advances on Mt. Rainier based on limiting tephra ages and dendrochronology. Moraines of the Burroughs Mountain advance are overlain by tephra layers Wn (AD 1480) and C (ca. 2200 14 C yr B.P.), but are younger than layer Yn (ca. 3300 14 C yr B.P.). Moraines of Garda age were deposited during the Little Ice Age and were dated by tree-ring measurements (Sigafoos & Hendricks, 1972). An initial advance began in the late 12th or early 13th century and culminated variously in the mid-14th century to the mid-19th century.

Table 1. Post-1965 glacial-geologic studies in the Cascade Range and Olympic Mountains.a Washington Cascades Western Drainages Mt. Baker Nooksack Skagit Glacier Peak and vicinity North Cascades Tolt Middle Fork Snoqualmie South Fork Snoqualmie Cedar Mt. Rainier and vicinity

White Pass Toutle Lewis Eastern Drainages Mt. Adams and vicinity Yakima Wenatchee Ingalls Icicle

Chumstick Upper Wenatchee Entiat

Thomas et al. (2000) Kovanen & Easterbrook (2001) Heller (1980) Begét (1981), Davis & Osborn (1987) Miller (1969), Porter (1978) Knoll (1967) Williams (1971) Porter (1976) Hirsch (1975) Crandell & Miller (1965, 1974), Burbank (1981), Heine (1998) Clayton (1983) Hyde (1973) Mundorf (1984), Hammond (1987) Hopkins (1976) Porter (1976) Hopkins (1966) Waitt et al. (1982), Swanson & Porter (1997, 2000) Merrill (1966) Nimick (1977) Thorp (1985)

Olympic Mountains Wynoochee Quinnault Queets Hoh

Carson (1970) Moore (1965) Thackray (2001) Thackray (2001)

Oregon Cascades Metolius Mt. Jefferson and vicinity Mountain Lakes

Scott (1977) Scott (1977) Carver (1973)

a Numerous unpublished manuscripts by William A. Long on

various aspects of Quaternary glaciation in the Cascades and Olympics are housed in the library of the Quaternary Research Center, University of Washington.

Post-1965 Studies of Pleistocene Glaciation Washington Cascades Mapping and Relative-Age Control. During the past 30 years, most of the major Cascade valleys have been mapped and attempts made to derive a chronology of glaciation based

largely on relative-age criteria (Fig. 2A and B and Table 1). As a result, limits of the last glaciation are well-delineated in major drainages. In most valleys evidence of two or more glaciations has been recognized, and in the most-thoroughly studied examples, evidence of at least nine ice advances has been documented.

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Earliest-Dated Cascade Glaciation. The earliest evidence of glaciation that is at least partially controlled by radiometric ages is found southeast of Mt. Rainier in the southern Cascades, where Clayton (1983) found an alpine till near Penoyer Creek on the Tumac Plateau beneath a basalt flow with a K/Ar age of 1.75 ± 0.35 myr. Another till along the South Fork of Clear Creek, lies beneath a basaltic-andesite flow having a K/Ar age of 0.65 ± 0.08 myr. The regional extent of these tills is unknown. Glaciation of Mt. Rainier Volcano. In their exhaustive study of Quaternary glaciation on and near Mt. Rainier volcano, Crandell & Miller (1974) recognized the deposits of four Pleistocene glaciations and two Holocene ice advances. At least one early glaciation is inferred from an intracanyon lava flow dating between 600,000 ± 60 K/Ar yr (for a mineral separate) and 325,000 K/Ar yr (whole-rock sample) that flowed down a glaciated valley. The largest glacier of the Wingate Hill glaciation extended 105 km from the mountain. Wingate Hill drift was differentiated from the subsequent Hayden Creek drift on the basis of weathering characteristics. The last glaciation (Fraser) is represented by the Evans Fig. 4. Map of Leavenworth in the Wentachee River drainage basin showing extent and ages of Pleistocene moraines of Icicle Creek glacier.

Creek till, deposited by glaciers as much as 64 km long. A late-Fraser advance (McNeeley) left moraines far upvalley from their Evans Creek limit. No radiometric dates are available for these drifts, but Crandell and Miller inferred that the Evans Creek advance preceded the Vashon advance of the Puget Lobe. The McNeeley advance occurred prior to ca. 8850 14 C years ago, the age of Rainier tephra layer R that mantles McNeeley moraines. Glaciation of Upper Yakima River Drainage. The glacier system that occupied the upper Yakima River drainage was one of the longest in the Cascades and its record one of the most detailed (Fig. 3). Porter (1976) mapped the extent of eight glacier advances and determined their relative age based on various weathering parameters. 36 Cl ages for moraines of the Domerie advance, which impound three large lakes in the major tributary valleys, cluster in two groups, averaging ca. 23,200 ± 1000 and 16,300 ± 1600 years. Only a few preliminary ages are available for the next-older (preDomerie) Ronald and Bullfrog moraines, but both apparently are older than 50,000 36 Cl years (T.W. Swanson, unpublished data). Three still-older drifts (Indian John, Swauk Prairie, and

Quaternary Alpine Glaciation in Alaska, the Pacific Northwest, Sierra Nevada, and Hawaii 85

Fig. 5. Cosmogenic isotope (36 Cl) ages of moraine systems in Icicle Creek valley compared with the standard marine oxygenisotope record (Martinson et al., 1987). The 36 Cl curves represent the collective probability distribution of individual samples (±1). Lookout Mountain) have not been dated, but likely are earlylate to middle Pleistocene in age. Glaciation of Icicle Creek Drainage. Right-lateral moraines of the Icicle Creek glacier, which occupied the next major valley north of the Yakima River, record seven advances that reached to or beyond the town of Leavenworth (Fig. 4). Two additional late-glacial moraines are found in tributary valleys upriver. The original sequence proposed by Page (1939) has been expanded with the discovery of five additional moraine systems. More than 100 36 Cl ages have been obtained for moraine boulders and provide a chronology of late Pleistocene glacier variations for four of these moraines (Figs 5 and 6) (Swanson & Porter, 1997, and unpublished data). Prior to this study, the only age control was the presence of Mazama tephra (ca. 6800 14 C yr B.P.) on late-glacial moraines of the Rat Creek advance (= Stuart glaciation of Page, 1939). Two moraines of the last glaciation (Leavenworth I and II) have mean 36 Cl ages of 20,000–18,000 and 18,000–15,000 years, respectively. Ages for the next older drift (Mountain Home) range from 77,000 to 71,000 years. Nine boulders on the Peshastin-age moraines range from 112,000 to 103,000 years and average 107,700 years. The oldest, strongly weathered drift (Boundary Butte) has not yet been dated. In addition, two late-glacial (late-Leavenworth) moraines have been dated (see below).

Time of Maximum Late Pleistocene Advance. Moraines marking the greatest expansion of ice in the mountains of western conterminous United States have often been inferred to predate the last interglaciation and correlate with MIS 6 or earlier isotope stages. The 36 Cl ages for Peshastin drift in the Icicle Creek drainage, however, imply an early last-glaciation age (i.e. equivalent to MIS 5d; Swanson & Porter, 2000). If this chronology is correct, it raises a question as to whether moraines marking maximum ice extent in other mountain ranges may also have formed early in the last glacial cycle. The difficulty is that surface exposure ages for moraines this old can be seriously affected by boulder weathering and moraine degradation, therefore requiring careful sample selection and interpretation. Nevertheless, these old moraines, heretofore largely “undatable” except by using relative-age criteria, may now be amenable to dating in carefully selected areas. Late-Glacial Ice Advances. Was there an advance of Cascade glaciers during the Younger Dryas Stade (11,000–10,000 14 C yr B.P.; ca. 12,900–11,600 cal yr B.P.)? This question has motivated several investigations during the past two decades. Moraines of presumed late-glacial age have been found in almost every glaciated valley investigated, but dating them has proved difficult. Often the only age control is the age of overlying Mazama tephra. Three recent studies

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Fig. 6. Large granodiorite boulder of Rat Creek II moraine being sampled for provide closer-limiting ages based on radiocarbon and 36 Cl surface-exposure dating. Late-glacial moraines in the Yakima River and Icicle Creek valleys have been dated by the 36 Cl method. In the former valley, two post-Domerie (i.e. late-glacial) moraines at Snoqualmie Pass have mean 36 Cl ages of 14,100 ± 500 and 12,700 ± 800 yr. The age of the younger moraine is consistent with a minimum limiting radiocarbon date from a bog inside the moraine of 11,050 ± 50 14 C yr B.P. (ca. 12,600 cal yr B.P.) (Porter, 1976). In the Icicle Creek drainage, two late-glacial moraines (Rat Creek I and Rat Creek II) have mean 36 Cl boulder ages of 14,000–13,000 and 13,000–12,000 years, respectively. In each of these drainages the paired moraines are closely nested. Kovanen & Easterbrook (2001) reported that three forks of the Nooksack valley in the western North Cascades contained long glacier systems during late-glacial time. Logs in a lateral moraine of the Nooksak Middle Fork glacier, lying 3 km beyond the terminus of the Derning Glacier have ages of 10,680 ± 70 and 10,500 ± 70 14 C years. However, outwash heading at a moraine in the lower North Fork valley, some 50 km from the glacier source area at Mt. Shuksan (2782 m) and Mt. Baker (3285 m), overlies glacialmarine drift dated 11,910 ± 80 yr and contains charcoal dated 10,605 ± 70 and 10,790 ± 80 14 C yr B.P., implying an ice tongue of exceptional length during Younger Dryas time. The proposed exceptional extent of the Nooksack Valley glaciers at a time when glaciers along much of the crest of the Cascades were confined to valley heads and cirques is puzzling. Only on the north side of the Stuart Range, the

36 Cl

dating.

source area of the type Rat Creek glacier, were late-glacial valley glaciers unusually long (ca. 8–10 km). Presumably this was because of their high accumulation area (1800–2200 m), as well as the potential for enhanced nourishment related to avalanching from steep valley slopes onto glacier surfaces. Mount Rainier, which is more than 1000 m higher than Mt. Baker and receives comparable record-high precipitation, supported long outlet glaciers during full-glacial time, but late-glacial ice was restricted in extent (see below). A record of late-glacial glacier variations in Mt. Rainier National Park is based on mapping of moraines and radiocarbon dating of lake and bog sediments associated with the moraines (Heine, 1998). The outer of two moraines (McNeeley 1) in cirques and valley heads was deposited before 11,320 ± 60 14 C yr B.P., and therefore predates the Younger Dryas interval. Moraines of a subsequent readvance (McNeeley 2) are overlain by layer R tephra (8850 14 C yr B.P.). Basal organic matter at one site gives a minimum age of 9140 ± 100 14 C yr B.P. for this advance. Detrital sediment layers in peat bogs that are inferred to be outwash from McNeeley 2 moraines are bracketed by ages of 10,080 ± 60 and 9120 ± 80 14 C yr B.P. at one site, and between 9850 ± 60 and 8990 ± 60 14 C yr B.P. at another site. These ages imply a post-Younger Dryas age for the advance. Heine therefore inferred that glaciers in this area were less extensive during the Younger Dryas interval than either before or after, possibly due to colder temperatures and diminished precipitation at that time. These contrasting local moraine chronologies within the same mountain range, reflecting: (1) moderate advance; (2) exceptional advance; and (3) retreat during Younger Dryas

Quaternary Alpine Glaciation in Alaska, the Pacific Northwest, Sierra Nevada, and Hawaii 87 time, raise questions about interpretations of the geology, the chronology, and the paleoclimatic environment, that will require additional research to resolve. A recent study in southern British Columbia 40,000 14 C yr B.P.) is found in reworked beach samples (EB and FrnB in Figs 16 and 19) as well as at beach and shelf sites near the North Carolina-South Carolina border (Fig. 19). The age of this aminozone is enigmatic, as it appears younger than 80,000 yr when compared with VLPG data from late Pleistocene emergent sites, but it is definitely

older than ∼40,000 14 C yr B.P. (Wehmiller & York, 2001; Wehmiller et al., 1995; York & Wehmiller, 1992a; York et al., 2001). This same aminozone is inferred from the results of Mitterer (1975) for samples collected from an emergent outcrop near Bon Terra, Florida (R.M. Mitterer, pers. comm. to J.F. Wehmiller, 1995) (BT, Fig. 16). The Bon Terra collection is particularly important, as it most likely represents the Anastasia Formation, widely considered as of “last interglacial” age. Other data from northern Florida localities (Mitterer, 1975) are consistent with this four-fold amino acid zonation for the southern portion of the Sea Islands, although specific locality information for Mitterer’s collections is not available. Gulf Coast North and west of the carbonate-rich records of central and southern Florida (discussed above), other coastal records from the Gulf Coast include those from the Florida panhandle extending west to southern Mississippi and those along the Louisiana-Texas coast (Fig. 3). Geochronological data for Pleistocene units in this entire region are limited, particularly for emergent sequences, the focus of this review. Seismic exploration of the thick Quaternary and pre-Quaternary records in the northern Gulf of Mexico, combined with extensive marine biostratigraphy, allow high-resolution chronology for portions of the offshore record, especially on the Texas-Louisiana shelf. DuBar et al. (1991) reviewed the characteristics and nomenclature of the onshore and offshore Quaternary record of the entire region. By analogy with the Atlantic Coastal Plain, it is quite likely that Gulf Coast Quaternary deposits are equally complex and variable. In the panhandle region of northwest Florida (Fig. 4), the number of preserved Pleistocene shoreline features has been the subject of considerable debate (Donoghue & Tanner, 1992; Donoghue et al., 1998; Otvos, 1992, 1995), as the location of “last interglacial” features remains in doubt. Additionally, the potential age range (late Pleistocene or most of the Pleistocene) of preserved coastal features is questioned. Mollusk samples from shallow excavations in Franklin County in the Florida panhandle have yielded racemization results that imply a late Pleistocene age when compared with other data from Florida and Georgia, but correlation of these results to a specific sea level transgression is not possible (Otvos & Wehmiller, unpublished data; Wehmiller et al., 1996). Pleistocene coastal features on the Texas coast include the Ingleside Barrier (interpreted as broadly of “last interglacial” age by Otvos & Howat, 1996), also referred to as the “Ingleside surface.” The Ingleside surface is underlain by the Beaumont Clay, a complex unit that probably includes multiple ages of fluvial and nearshore units, perhaps representing the past 400,000 years of glacial-interglacial incision and valley filling (Blum & Price, 1998). Records on the Texas shelf provide an important perspective on the origin and age of both submergent and emergent landforms in the region (Anderson et al., 1996; Morton & Price, 1987; Rodriguez et al., 2000; Suter et al., 1987).

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D.R. Muhs et al.

Valley incision, infilling, and delta construction are all recognized in various seismic studies, reinforced with core data. In one of the more provocative studies of the implications of these records for Pleistocene sea-level history, Rodriquez et al. (2000) identified a submerged escarpment and associated barrier features that indicate a shoreline at a depth (corrected for subsidence) of about −15 m. Based on seismic stratigraphic relations to other sequence boundaries in the record (interpreted to represent ∼140,000 and ∼20,000 yr, or “transitional MIS 6/5e” and “MIS 2” transgressive surfaces), this paleoshoreline is interpreted to be between 35,000 and 50,000 yr old (∼MIS 3). Its depth is significantly shallower than would be predicted from most marine isotope or dated coral reef records (e.g. Lambeck & Chappell, 2001). Finite radiocarbon ages are consistent with this estimate (Rodriguez et al., 2000), although ages in this range are near the limit of reliability. Racemization data for mollusks from these shelf sequences and from cores beneath the Ingleside Barrier are all internally consistent with the stratigraphic position of the samples (Wehmiller et al., 2000a). Kinetic modeling is also consistent with an age of between 40,000 and 80,000 yr for the −15 m paleoshoreline (Wehmiller et al., 2000a). Elsewhere on the Gulf of Mexico shelf, Schroeder et al. (1995) used radiocarbon dating of oyster shells to produce a rough curve of postglacial sea-level rise and a more uncertain record for the time interval between about 27,000 and 40,000 14 C yr B.P. Although many shell radiocarbon ages in this range can be challenged (Bloom, 1983a), some paired AAR/radiocarbon (AMS) analyses of massive and well-preserved shells from shelf/beach environments indicate that ages in this range may be valid (Wehmiller & York, 2001; Wehmiller et al., 1995). Because all of the dated shells of Schroeder et al. (1995) were transported to some extent, a precise sea level curve cannot be constructed.

Summary: Atlantic and Gulf Coasts

the Texas coast also appears to represent a complex of multiple units representing several transgressive-regressive cycles during the middle- to late Pleistocene.

Last Interglacial and Last Glacial Periods Atlantic coast sites with U-series ages representing early MIS 5 (∼125,000 yr) are rare north of the Florida Keys; where found, these units are at roughly the same elevation (all within ca. 6–8 m of present sea level) as those dated at 80,000 yr. These observations have been the focus of debate about the Atlantic Coastal Plain geochronology for nearly three decades. Nevertheless, a high sea level at ∼80,000 yr is consistent with similar data from Bermuda (Muhs et al., 2002b) and some data from the U.S. Pacific Coast (Fig. 13). Emergent shoreline features are found throughout the Gulf and Atlantic coastal plains, and in the absence of numerical geochronologic data, it is common to assign “last interglacial” ages to features found at elevations up to approximately +8 m. The current elevations of emergent Atlantic coast paleoshorelines or shallow marine units U-series dated to late MIS 5 (80,000 yr) are higher than expected from marine isotope curves or dated tropical coral terrace records (Lambeck & Chappell, 2001; Shackleton, 1987). Records on the New Jersey, Maryland, and Texas shelves (Sheridan et al., 2000; Rodriguez et al., 2000; Toscano & York, 1992) also indicate sea levels during either late MIS 5 or MIS 3 that were higher than predicted from marine isotope curves. These observations suggest that major sections of the Atlantic and Gulf coasts, although considered as passive margin regions, may have been influenced by a combination of postglacial isostatic adjustment (through multiple glacial cycles), hydroisostasy, and shelf sediment loading.

Records of Holocene Sea-Level Rise, Atlantic and Gulf Coasts

Early Pleistocene to Last Interglacial Period Age assignments for Atlantic coast marine units rely upon either biostratigraphic, radiometric, or suitably calibrated racemization modeling. All of these methods have significant uncertainties, particularly for the earlier Pleistocene, and racemization models are quite subjective, especially when used to correlate over broad latitude ranges. Nevertheless, for the Atlantic Coastal Plain, early to middle Pleistocene depositional records appear to be preserved throughout the region. Notable examples include the older parts of the Albemarle Embayment section in North Carolina and the Waccamaw formation in northeastern South Carolina. Late-middle Pleistocene, or pre-last interglacial units, include paleochannel fill in the Chesapeake Bay region, and U-series and racemizationdated units from the Albemarle and Cape Fear Arch regions. Determining the origin(s) and age(s) of the Suffolk Scarp would provide major insights into the history of sea level change on the Atlantic Coastal Plain. The Beaumont Clay on

Preserved records on continental shelves provide a general insight into the potential complexity of emergent records on any low gradient coastal plain. The youngest portions of these sequences represent the processes that occurred during last glacial low-stands and the rise in sea level since the last glacial maximum. In many cases, these are depositional records, preserved in areas of relatively large sediment supply. Examples include, but are not limited to, the inner shelves of New Jersey (Sheridan et al., 2000), Maryland (Toscano, 1992; Toscano & York, 1992), North Carolina (Boss et al., 2002; Riggs et al., 1995), large estuaries (Colman et al., 1992) and thick sections on the Texas shelf (Anderson et al., 1996; Rodriguez et al., 2000). Where sediment supply is more limited (such as the shelves of South Carolina and Georgia), extensive recycling of shelf sediments during multiple Quaternary transgressions is inferred from local stratigraphic sections where 1–2 m of preserved section may represent the entire Quaternary (Gayes et al., 1992; Harris et al., 2000; Heron

Quaternary Sea-Level History of the United States 173 et al., 1984; Hine & Snyder, 1985; Macintyre et al., 1978; Pilkey et al., 1981). Records of Holocene sea-level rise on the Atlantic and Gulf coasts are numerous, having been obtained in order to understand local or regional histories of coastal evolution or geophysical models of postglacial isostatic adjustment (e.g. Peltier, 1999). Major processes affecting these records include isostatic rebound in glaciated areas (Kelley et al., 1992), steady subsidence associated with forebulge collapse in regions proximal to glacier advance (Kraft et al., 1987), and, occasionally, rises of sea level to near present level during the middle Holocene in regions farther from glacial influence (Blum & Carter, 2000). Holocene depositional records that provide good models for Pleistocene sequences in a particular region are especially important. Good examples include channel erosion and filling in Chesapeake Bay (Colman et al., 1990, 1992) and in Albemarle Sound (Riggs et al., 1992, 1995), and Holocene coastal environments of the Florida peninsula (Davis et al., 1992). In each of these regions, good analogies between preserved Pleistocene and Holocene records can be recognized.

Unresolved Issues on the Atlantic Coastal Plain The following are three major unresolved geochronological issues related to the Pleistocene of the Atlantic Coastal Plain. Quite likely similar issues will arise for Gulf Coast sites as additional chronologic data (or samples for dating) become available for this region, but limited results prevent more specific conclusions. (1) Identification of 125,000 yr erosional and depositional records using geomorphic and geochronologic analysis. Detailed mapping and dating in selected regions (such as the Cape Fear Arch and Albemarle Embayment) where both surface landforms and subsurface units can be clearly mapped is a high priority for establishing confidence in any of the Atlantic or Gulf coast chronologic records. (2) Reconciling the elevation of the 80,000-yr-old units with the sea-level record from marine isotope and tropical coral reef records. The similar elevation of these units, and of possible 125,000-yr units (all within 6 m) from Virginia to Georgia (and also Bermuda), must be interpreted in the context of glacial-margin tectonics and hydroisostatic effects. (3) Understanding the age-resolution capabilities of amino acid racemization methods, including assessing the reality and significance of the “offshore aminozone” found between North Carolina and Georgia. Inconsistencies in correlation of racemization data from region to region appear to contradict the stratigraphic coherence of racemizatoin data in local regions with common temperature and depositional histories. Continued refinement of racemization kinetic modeling by comparison with rigorous independent calibration is needed.

Summary Much progress has been made in the past three decades in our understanding of the Quaternary sea-level history of U.S. coastlines. Two complimentary dating methods, U-series analyses of fossil corals and amino acid racemization of fossil mollusks, have generated hundreds of numerical and correlated ages. These data have allowed lateral correlation of marine deposits on the Pacific and Atlantic coasts and the Hawaiian Islands, as well as on the Bering Sea and Arctic Ocean coasts of Alaska. Many of the sea-level high stands of the Quaternary are recorded in the reef record of the tectonically stable Florida Keys. Stratigraphic studies show that deposits of pre-lastinterglacial high stands are present, although dating has yet to establish the precise timing of these deposits. Nevertheless, preliminary ages suggest that two high sea stands of the mid-Pleistocene are recorded on the Florida Keys, perhaps on the order of ∼300,000–340,000 and 220,000–230,000 yr. Corals in reefs of two high sea stands of the last interglacial complex, the ∼80,000 and ∼120,000 yr stands, are present in this island chain. Offshore, Holocene reefs of the Florida Keys and mainland extend the Barbados record of sea-level rise since the last glacial period up to present sea level. Reefs and coral-bearing marine deposits, both emergent and submergent, have been identified, mapped and dated in the Hawaiian Islands. On the island of Oahu, the Waimanalo Limestone was deposited during the peak of the last interglacial complex. Because Oahu is uplifting only very slowly, the range of U-series ages for this deposit is a good estimate of the duration of the peak of the last interglacial period. Results of recent high-precision dating indicate that the peak of the last interglacial period could have lasted at least 15,000 yr and possibly longer. Waimanalo Limestone deposits contain a significant number of extralimital southern mollusks, indicating warmer-than-present waters during the last interglacial period. The “big island” of Hawaii, unlike Oahu, is undergoing long-term Quaternary subsidence, due to volcanic loading. As a result, submerged reefs off Hawaii mark deglacial periods and yield a record of such events for more than 400,000 yr. The island of Lanai has long been the object of study because of high-elevation marine deposits. A novel hypothesis that these deposits were formed by a “giant wave” generated by a submarine landslide was proposed as a challenge to a decades-old concept that sea level was much higher than present. This hypothesis has in turn been challenged by new mapping and dating that suggest that the deposits were elevated due to lithospheric flexure, in response to the subsidence of the island of Hawaii, in turn induced by volcanic loading. Multiple marine terraces on the Pacific Coast record a long history of sea-level high stands superimposed on longterm crustal uplift. Pre-last-interglacial high sea stands are recorded on the coast of California and several fossil-bearing localities in southern California hold promise for unraveling middle and early Pleistocene sea-level history. A high terrace on San Nicolas Island dates to ∼500,000–600,000 yr ago and suggests that much of the sea-level history of the middle

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Pleistocene may be recorded in a series of lower terraces. U-series analyses show that marine terraces from both the ∼80,000- and ∼120,000-yr-old sea-level high stands of the last interglacial complex, long recognized on tropical coasts, are also present on the Pacific Coast. The ∼120,000 yr high sea stand lasted at least 8000–10,000 yr on the Pacific Coast, based on the range of individual coral ages. It was a major interglacial period that is distinguished by the presence of extralimital southern mollusks, indicating eastern Pacific Ocean water temperatures warmer than present. The ∼80,000 yr high sea stand could also have been on the order of 8000–10,000 yr long, but it is distinguished by the presence of extralimital northern mollusks, indicating water temperatures cooler than present. Reworking of ∼120,000 yr fossils during the ∼100,000 yr high stand occurred on the Pacific Coast, based on U-series dating of individual corals. This reworking suggests that sea level at ∼100,000 yr ago may have been close to that of today. “Capture” of the ∼120,000 yr high sea stand during the ∼100,000 yr high stand may also explain the mixture of extralimital southern and extralimital northern fossils in some deposits. As with the coasts of California, Hawaii, and Florida, the peak of the last interglacial period is recorded on the Bering Sea and Arctic Ocean coasts of Alaska. Although corals have not been found in these deposits, amino acid ratios, non-finite radiocarbon ages, and terrace elevations all indicate that the “Pelukian” high sea stand corresponds to the peak of the last interglacial period. Similar to low-latitude and mid-latitude marine deposits of this high sea stand, Pelukian deposits of Alaska contain a number of extralimital southern species of mollusks, indicating warmer-than-modern water temperatures. In addition, an older high sea stand, that is probably on the order of ∼400,000 yr old, may record another major interglacial period, also with warmer-than-modern water temperatures. The Atlantic Coastal Plain of the eastern United States has been one of the most challenging regions for studies of past sea levels. Because this coast is on a passive continental margin, sea-level records are complex: long-term uplift does not “isolate” individual high stands of sea as distinct terraces in the manner that is typical for the Pacific Coast. Many transgressive-regressive cycles are preserved at roughly the same elevations, resulting in a highly complex stratigraphic record. Nevertheless, amino acid racemization studies have shown that there is a rich record of Quaternary sea-level fluctuations, at least as far back as the middle Pleistocene and more likely the early Pleistocene. An unexpected result of recent studies on the Atlantic Coastal Plain is that, unlike the other coasts of the U.S., records of the peak last-interglacial high sea stand of ∼120,000 yr ago have been remarkably difficult to find. The few U-series ages that have been reported for this high sea stand are equivocal. In contrast, the ∼80,000 yr high sea stand is recorded as low-elevation (but emergent) marine deposits in Virginia, North Carolina, South Carolina and Georgia. Individual coral ages of this high sea stand are very similar to those on the Pacific Coast and indicate that the ∼80,000 yr high stand could have lasted as long as ∼10,000 yr. The presence

of ∼80,000 yr deposits at or even above sea level on a tectonically stable coast is a major challenge to widely held views that this sea stand was significantly lower than present sea level. The longer-term Quaternary sea-level record of the Atlantic Coastal Plain is apparent in thick stratigraphic sequences, multiple aminozones ranging back to the early Pleistocene, and some independent chronologic control for units that date to >750,000 yr. Shore-parallel correlation of these deposits is also possible, as has been done on the Pacific Coast. The actual length of the Quaternary sea-level record in this region, as shown by amino acid racemization data, depends on the choice of calibration samples and on either a “short” or “long” time constant for racemization. Consequently, age estimates based on racemization results might vary by 100% or more. The “short” chronology suggests that the Atlantic Coastal Plain record spans only ∼500,000 yr, whereas the “long” chronology would suggest a considerably greater part of the Quaternary is represented. Certain results of sea-level studies on the coasts of the United States challenge the Milankovitch (1941) theory of climate change, as interpreted from the deep-sea oxygen isotope record. The duration of the peak of the last interglacial period (MIS 5e), when sea level was as high or higher than present, is interpreted from the oxygen isotope record to be on the order of only a few thousand years (Imbrie et al., 1984; Martinson et al., 1987). In contrast, the coral record from Oahu, Hawaii indicates that sea level as high or higher than present could have lasted 15,000–20,000 yr. The U-series ages from both Oahu and California also suggest that sea level was relatively high about 115,000 yr ago, a time that the oxygen isotope record would suggest that sea level was relatively low. U-series ages of ∼80,000 yr for corals from the tectonically stable Atlantic Coastal Plain and the slowly uplifting Pacific Coast suggest that sea level at that time was near present, whereas the oxygen isotope record suggests that sea level was then well below present. The reasons for the discrepencies between the coastal sea-level record and the deep-sea oxygen isotope record are not understood, but provide an important challenge to future investigations on the coasts of the United States. Acknowledgments Work on sea-level history by Muhs and Simmons was supported by the Earth Surface Dynamics Program of the U.S. Geological Survey and is a contribution to the LITE (Last Interglacial: Timing and Environment) Project. Studies by Wehmiller and York were supported by the National Science Foundation. We thank Robert S. Thompson, Richard Z. Poore, and two anonymous referees for helpful reviews of an earlier version of this paper. References Addicott, W.O. (1966). Late Pleistocene marine paleoecology and zoogeography in central California. U.S. Geological Survey Professional Paper 523-C, pp. C1–C21.

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Western lakes Larry Benson U.S. Geological Survey, 3215 Marine St., Boulder, CO 80303, USA of those who attempted to quantify the elements of climate change will be featured. Records that extend beyond the Last Interglaciation (e.g. Adam et al., 1989; Cohen, 1996; Davis & Moutoux, 1998; Kowalewska & Cohen, 1998; Thompson, 1996), records that are salt-dominated (e.g. Jannik et al., 1991; Li et al., 1996; Lowenstein et al., 1999; Smith, 1984), records that are discontinuous (e.g. Hooke, 1999; Wells et al., 1987), and most outcrop-based records will not be discussed in this chapter. Studies of western lakes and their sediments have improved in several ways since 1965, including: (1) recovery of continuous sediment records from the deepest areas of lake basins by coring; (2) improvements in age control; and (3) introduction and application of new proxies of climate change.

Status of Western Lake Studies in 1965 Wright & Frey (1965) contained a chapter entitled the “Quaternary Geology of the Great Basin” in which Roger Morrison (1965) emphasized the use of soils as stratigraphic markers. He used these markers to correlate the rise and fall of Great Basin lakes (Fig. 1) with the advance and retreat of North American alpine and continental glaciers. This study was entirely outcrop-based and the correlations were made under the assumption that soils had formed synchronously across North America during relatively dry and warm periods. Although not reflected in the Morrison (1965) chapter, new concepts in climate forcing and methods of age control had already emerged. Antevs (1948), nearly two decades previously, had suggested that maximum levels of Great Basin lakes were linked to the presence of a permanent ice sheet over North America. He hypothesized that the size of the ice sheet, combined with a permanent high-pressure area located over it, caused storm tracks associated with the polar jet stream (PJS) to be pushed south of their present-day average path over the northern Great Basin. In addition, Emiliani & Geiss (1957) had suggested that variation in insolation strongly influenced Pleistocene glacial-interglacial cycles. Broecker and his colleagues (Broecker & Kaufman, 1965; Broecker & Orr, 1958; Broecker & Walton, 1959; Kaufman & Broecker, 1965) had already applied the relatively new radiocarbon (14 C) and uranium-series methods of age control to studies of carbonates precipitated from Lake Bonneville and Lake Lahontan (Fig. 1), and Flint & Gale (1958) and Stuiver (1964) had obtained 14 C ages of organic materials in sediment cores from Searles Lake, California (Fig. 1). These studies demonstrated that the last large lake cycle in each of these basins occurred between ∼25,000 and ∼10,000 14 C yr B.P.

Sediment Coring Coring of a few lake basins had occurred prior to 1965; e.g. Great Salt Lake, Utah (Eardley & Gvosdetsky, 1960), and Searles lake, California (Smith, 1962); however, the core sites were few and some were chosen for their economic rather than their climatic value. During the past two decades, scientists have sought to obtain continuous or nearly continuous records of climate change by coring the deepest areas of lake basins. Cores have been recovered from several presently dry lake basins including Owens Lake basin, California (Lund, 1996; Smith & Bischoff, 1997), Summer Lake basin, Oregon (Cohen et al., 2000), and Estancia basin, New Mexico (Allen & Anderson, 2000) (Fig. 1). Numerous perennial lakes also have been cored, including Great Salt Lake, Utah (Spencer et al., 1984), Clear Lake, California (Sims et al., 1988), Carp Lake, Washington (Whitlock & Bartlein, 1997), and Pyramid Lake, Nevada (Benson et al., 1997a, 2002) (Fig. 1).

New Approaches and Technologies Methods of Age Control Because of constraints on the length of this chapter, I will emphasize high-resolution studies of sediment cores taken from lakes located within the western United States as defined by the boundaries of Fig. 1. High-resolution will be taken to mean a resolution of 500 to 2000 yr for records ≤150,000 yr in length, a resolution of 100 to 500 yr for records ≤50,000 yr in length, a resolution of 5 to 100 yr for records ≤10,000 yr in length, and a resolution of ≤5 yr for records ≤1000 yr in length. Sediment-based records that illustrate evolving concepts of climate change (e.g. solar forcing) will be emphasized. The work of those who broke new ground as well as the work

Methods of age control have been expanded and in many cases greatly improved since 1965. Radiocarbon (14 C). In 1965, ∼5 g of C was needed for 14 C dating. With the advent of accelerator mass spectrometry (AMS) in the late 1970s, very precise 14 C analyses could be performed on as little as 1 mg of C (Southon et al., 1982). This enhancement in sensitivity has allowed the dating of terrestrial macrofossils, lacustrine shells, humic acids (e.g. Reasoner & Jodry (2000), pollen (Long et al., 1992; Mensing & Southon, 1999), and the total inorganic or organic carbon fraction of lake sediments (e.g. Benson et al., 1997a, b).


PUBLISHED BY ELSEVIER B.V. 185

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Larry Benson Fig. 1. Location map of lake and marine core sites cited in this paper. This map defines the boundary of the western United States for the purposes of this paper. Dotted line indicates the boundary of the Great Basin. Solid black areas indicate existing water bodies and gray areas indicate the extent of late-Pleistocene pluvial lakes.

Terrestrial macrofossils deposited in lake sediments generally yield reliable 14 C dates unless the time they have spent in the terrestrial environment is long compared to the time they have spent in the lacustrine environment. Unfortunately, deep-water sediments of large lakes such as Walker Lake, Pyramid Lake, and Great Salt Lake, usually lack terrestrial macrofossils. Other carbon sources often yield 14 C ages older than their date of deposition because of reworking of “old” forms of detrital carbon to the core site (e.g. pollen, Benson et al., 2002) or because the carbon source was formed from a lake whose dissolved inorganic carbon (DIC) reservoir was not in equilibrium with atmospheric 14 C (reservoir effect) (Broecker & Walton, 1959). Radiocarbon dates of organic and inorganic carbonbearing materials that are now subaerially exposed offer a means of estimating lake-level histories of Great Basin paleolake systems. Such data have been obtained for Lake Bonneville, Utah (Oviatt et al., 1992), Lake Lahontan, Nevada (Benson et al., 1995, and references therein), and Mono Lake, California (Benson et al., 1998a, and references therein) (Fig. 1). Two of these paleolake systems reached their highest levels between ∼16,000 and ∼14,000 cal yr B.P. (Fig. 2). Lake Bonneville’s well-documented earlier decline at 17,000 cal yr B.P. resulted from the catastrophic downcutting of its basin margin and not from climate change. If the downcutting had not occurred, Lake Bonneville may have achieved its highstand at ∼15,000 cal yr B.P. (dashed line in Fig. 2A). Calibration of 14 C ages. Not only has the 14 C method been improved in terms of greater sensitivity and smaller sample size, there are now procedures for conversion of 14 C dates to calibrated ages. Stuiver et al. (1998) obtained 14 C ages of dendrochronologically dated wood samples (mostly 10-yr segments) back to 11,620 cal yr B.P. Hughen

et al. (1998) used varved sediments from the Cariaco Basin to provide a new 14 C calibration data set for the period 14,500 to 10,000 cal yr B.P., and Bard et al. (1998) have used high-precision 14 C and 230 Th-234 U ages of marine corals to calibrate the 14 C time scale to ∼23,500 cal yr B.P. In addition, Kitagawa & van der Plicht (1998) used 250 14 C dates of terrestrial macrofossils from annually laminated sediments from Lake Suigetsu, Japan, to provide an atmospheric calibration for the past 45,000 cal yr B.P. Most recently, Schramm et al. (2000) have calibrated the 14 C time scale to 40,000 cal yr B.P. using high-precision 14 C and 230 Th-234 U ages of aragonites found in laminated sediments of Lake Lisan, Israel. Voelker et al. (1998) have correlated a 14 C-dated plankton-based marine 18 O record with the GISP2 18 O record, providing a means of linking terrestrial 14 C records with the GISP2 record. In the present paper conversions from 14 C to calibrated ages have been done using, in order of preference, the work of Stuiver et al. (1998), Hughen et al. (1998), and Bard et al. (1998). Uranium series. Uranium-series methods, have been greatly improved since the seminal study of Kaufman & Broecker (1965). For example, the work of Bischoff & Fitzpatrick (1991) and Luo & Ku (1991) have demonstrated that the correction for contamination by a single source of detrital Th is best accomplished using total sample dissolution and an isochron technique. Much of the unreliability in uranium-series age estimates may be due to multiple and often poorly understood sources of initial Th (Bischoff & Fitzpatrick, 1991; Lin et al., 1996; Szabo et al., 1996). If there are more than two sources of Th, the total dissolution and isochron approach may not succeed. Uranium-series methods have been applied with some success in attempts to date carbonates from Lake Lahontan

Western Lakes 187

Fig. 2. Lake-level histories based on surficial materials for: (A) the Bonneville basin (Oviatt et al., 1992), (B) the Lahontan basin (Benson et al., 1995), and (C) the Mono Lake basin (Benson et al., 1998a). (Bischoff & Fitzpatrick, 1991; Lao & Benson, 1988; Lin et al., 1996; Szabo et al., 1996) and to date salts and carbonates from Searles and Manly lakes (Bischoff et al., 1985; Ku et al., 1998; Lin et al., 1998; Peng et al., 1978) (Fig. 1). Uranium-series age estimates of salts and/or clays and organic materials contained within salts appear to yield reasonable values after correction for a single source of detrital 230 Th (Bischoff et al., 1985; Ku et al., 1998; Lin et al., 1998; Peng et al., 1978); i.e. most of the ages are in stratigraphic order. Tephrochronology. Wilcox (1965) was one of the first to suggest the usefulness of volcanic ashes as stratigraphic

markers in the western United States. However, it wasn’t until the 1970s that analytical tools such as the electron microprobe (EMP) were available for chemical analysis of individual glass shards. Tephrochronology was applied to California lakes by Sarna-Wojcicki et al. (1988) and Adam et al. (1989) and to Great Basin lakes by Davis (1978, 1983). A recent review of tephrochronologic methods can be found in SarnaWojcicki (2000), and a recent application of tephrochronologic methods to lake-sediment age control can be found in Whitlock et al. (2000). New lake-sediment-based ages of two of the most widespread Great Basin tephras (Wono and Trego Hot Springs tephras) can be found in Benson et al. (1997a). Paleomagnetic secular variation (PSV). At the time of the publication of the Wright & Frey (1965) volume, geophysicists were beginning studies of paleomagnetic reversals (e.g. Cox et al., 1965) and had not developed detailed PSV (inclination, declination, intensity) records of the Earth’s field. Since then, studies of lake sediments and archeological materials have helped create detailed PSV records (Lund & Banerjee, 1979, 1985a, b; Verosub, 1979; Verosub & Mehringer, 1984). Creer & Tucholka (1982, 1983) were among the first to suggest that type PSV curves could be applied to the dating of lake sediments and Lund (1996) was the first to demonstrate that distinctive field signatures in inclination and declination could be traced across North America without significant change in pattern. During the last 25 yr, secular variation has been used in studies of several Great Basin lakes, including Mono Lake, California (Liddicoat & Coe, 1979; Lund et al., 1988), Fish Lake, Oregon (Verosub et al., 1986), Pyramid Lake, Nevada (Benson et al., 2002), pluvial Lake Lahontan (Liddicoat, 1992, 1996; Liddicoat & Coe, 1997), Owens Lake, California (Li et al., 2000), pluvial Lake Bonneville (Liddicoat & Coe, 1998), and Summer Lake, Oregon (Negrini et al., 1984, 2000). Recently, Benson et al. (2003a) have shown that the geomagnetic feature, termed the Mono Lake excursion, has a date of 28,620 ± 300 14 C yr B.P. (∼32,400 GISP2 yr using the Voelker et al. (1998) conversion). Methods for age control of young sediments. Various methods for age control of sediments 40 ka by 234 U-230 Th dating of Lake Lisan sediments (last glacial Dead Sea. Earth and Planetary Science Letters, 175, 27–40. Scott, W.E., McCoy, W.D., Shroba, R.R. & Rubin, M. (1983). Reinterpretation of the exposed record of the last two cycles of Lake Bonneville, Western United States. Quaternary Research, 20, 185–261. Sims, J.D., Rymer, M.J. & Perkins, J.A. (1988). Late Quaternary deposits beneath Clear Lake, California; physical stratigraphy, age, and paleogeographic implications. In: Sims, J.D. (Ed.), Late Quaternary Climate, Tectonism, and Sedimentation in Clear Lake, Northern California Coast Ranges. Geological Society of America Special paper 214, 21–44. Smith, G.I. (1962). Subsurface stratigraphy of late Quaternary deposits, Searles Lake, California, a summary.U.S. Geological Survey Professional Paper, 450-C, 65–69. Smith, G.I. (1984). Paleohydrologic regimes in the southwestern Great Basin, 0–3.2 my Ago, compared with other long records of “global” climate. Quaternary Research, 22, 1–17. Smith, G.I. & Bischoff, J.L. (1997). Core OL-92 from Owens Lake: project rationale, geologic setting, drilling procedures, and summary. In: Smith, G.I. & Bischoff, J.L. (Eds), An 800,000 Year Paleoclimatic Record from Core OL-92, Owens Lake, Southeast California Geological Society of America Special Paper 317, 1–8. Snyder, C.T. & Langbein, W.B. (1962). The Pleistocene lake in Spring Valley, Nevada, and its climatic implications. Journal of Geophysical Research, 67, 2385–2394. Southon, J.R., Nelson, D.E., Korteling, R., Nowikow, I., Hammaren, E., McKay, J. & Burke, D. (1982). Techniques for the direct measurement of natural beryllium-10 and carbon-14 with a tandem accelerator. In: Currie, L.A. (Ed.), Nuclear and Chemical Dating Techniques. American Chemical Society, Washington, DC, 75–88. Spencer, R.J., Baedecker, M.J., Eugster, H.P., Forester, R.M., Goldhaber, M.B., Jones, B.F., Kelts, K., McKenzie, J., Madsen, D.B., Rettig, S.L., Rubin, M. & Bowser, C.J. (1984). Great Salt Lake and precursors, Utah; the last 30,000 years. Contributions to Mineralogy Petrology, 86, 321–334. Stine, S. (1990). Late Holocene fluctuations of Mono Lake, eastern California. In: Meyers, P.A. & Benson, L.V. (Eds), Special Issue Paleoclimates: the record form lakes,

ocean and land. Palaeogeography, Palaeoclimatology, Palaeoecology, 78, 333–381. Stine, S. (1994). Extreme and persistent drought in California and Patagonia during mediaeval time. Nature, 369, 546–549. Stoner, J.S., Channell, J.E.T. & Hillaire-Marcel (1998). A 200 ka geomagnetic chronostratigraphy for the Labrador Sea: Indirect correlation of the sediment record to SPECMAP. Earth and Science Planetary Letters, 159, 165–181. Stoner, J.S., Channell, J.E.T., Hillaire-Marcel & Kissel, C. (2000). Geomagnetic palaeointensity and environmental record from Labrador Sea core MD95-2024: global marine sediment and ice core chronostratigraphy for the last 110 kyr. Earth and Science Planetary Letters, 183, 1651–177. Stuiver, M. (1964). Carbon isotopic distribution and correlated chronology of Searles Lake sediments. American Journal of Science, 262, 377–392. Stuiver, M., Reimer, P.J. & Braziunas, T.F. (1998). Highprecision radiocarbon age calibration for terrestrial and marine samples. Radiocarbon, 40, 1127–1151. Szabo, B.J., Bush, C.A. & Benson, L.V. (1996). Uraniumseries dating of carbonate (tufa) deposits associated with Quaternary fluctuations of Pyramid Lake, Nevada. Quaternary Research, 45, 271–281. Thompson, R.S. (1996). Pliocene and early Pleistocene environments and climates of the western Snake River Plain, Idaho. Marine Micropaleontology, 27, 141–156. Verosub, K.L. & Mehringer, P.J. (1984). Congruent paleomagnetic and archeomagnetic records from the western United States: A.D. 750 to 1450. Science, 224, 387–389. Verosub, K.L., Mehringer, P.J. & Waterstraat, P. (1986). Holocene secular variation in western North America: paleomagnetic record from Fish Lake, Harney County, Oregon. Journal of Geophysical Research, 91, 3609–3623. Voelker, A.H., Sarnthein, M., Grootes, P.M., Erlenkeuser, H., Laj, C., Mazaud, A., Nadeau, M. & Schleicher, M. (1998). Correlation of marine 14 C ages from the Nordic seas with the GISP2 isotope record: implications for 14 C.calibration beyond 25 ka BP. Radiocarbon, 40, 517–534. Verosub, K.L. (1979). Paleomagnetism of carved sediments from western New England: variability of the paleomagnetic recorder. Geophysical Research Letters, 6, 241–244. Wagner, G., Beer, J., Laj, C., Kissel, C., Masarik, J., Muschele, R. & Synal, H. (2000). Chlorine-36 evidence for the Mono Lake event in the Summit GRIP ice core. Earth and Planetary Science Letters, 181, 1–6. Weimer, M.B. (1997). An examination of solute evolution and lake level fluctuations of Mono Lake, California, using minor element and oxygen isotope chemistries of the ostracod limnocythere ceriotuberosa. Unpublished Master of Science thesis, University of Southern Colorado. Wells, S.G., McFadden, L.D. & Dohrenwend, J.C. (1987). Influence of late Quaternary climatic changes on geomorphic and pedogenic processes on a desert piedmont, Eastern Mojave Desert, California. Quaternary Research, 27, 130–146. Whitlock, C. & Bartlein, P.J. (1997). Vegetation and climate change in northwest America during the past 125 kyr. Nature, 388, 57–61.

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Isotopic records from ground-water and cave speleothem calcite in North America Jay Quade Department of Geosciences, University of Arizona, Tuscon, AZ 85721, USA

All the situations under review here involve dissolution of carbonate-rich rocks somewhere along the ground-water flow path. Calcite dissolution is initiated in the vadose (recharge) zone in the presence of elevated soil pCO2 (>350 ppmV) produced by plant respiration and decay. This dissolution process can continue below the vadose zone and into the saturated zone until significant levels of supersaturation with respect to calcite are reached, whereupon calcite precipitation occurs. Supersaturation also develops through: (1) off-gassing of dissolved CO2 due to a reduction in ambient pCO2 (as in caves and springs) or due to increased T; (2) evaporation, as in some caves and springs; and/or (3) through reduction of pressure, as with ascending ground water, perhaps such as at Devils Hole. Soil solutions should undergo little or no change in temperature or pressure between the deep soil zone and caves at depths of generally ten’s of meters. Therefore, soil solutions entering a cave must lose water or CO2 or both in order for calcite to form. This in turn requires some advective or diffusive gas leakage to outside of the cave (Wigley & Brown, 1976). Most caves in temperate regions are at or near 100% humidity except in ventilated areas such as around entrances. Thus, CO2 off-gassing is thought to be the cause of speleothem formation in caves from temperate regions. Humidity is typically 1 km.

can be developed as follows: d{c-w } 18 T + (18 Odw ) Oc = dT

18 Odw = 18 Omw (3) 18 d{ Omw } T + (18 Osw ) (4) (18 Omw ) = dT substituting Eqs (3) and (4) into Eq. (2) yields: 18 d{c-w } d{ Omw } 18 Oc = T + T dT dT + (18 Osw )

The ␦18 O Record Speleothems In this section we will review how the paleoclimatic interpretation of the 18 O record from speleothems has evolved in the last three decades in North America, with particular emphasis on the much studied speleothem records in Iowa and Missouri. The seminal works of Hendy & Wilson (1968), Hendy (1971), Schwarcz et al. (1976), and Harmon et al. (1978, 1979a) framed the basic approach of most studies prior to 1990 to reconstructing paleoclimate from speleothem isotope records. Three factors were and are seen to potentially control the 18 O value of speleothem calcite (henceforth 18 Oc ): (1) cave and air temperature (2) the 18 O value of meteoric water (henceforth 18 Omw ), and (3) the kinetic effects of evaporation and rapid CO2 off-gassing on the 18 O value of soil and cave waters. In temperate caves, the effects of evaporation in soils on 18 Omw are generally viewed as negligible. This assumption and the strong correlation between the 18 Omw and the 18 O of cave drip water (henceforth 18 Odw ) were confirmed in such studies as Schwarcz et al. (1976), Harmon (1979), Goede et al. (1982), and Yonge et al. (1985). The kinetic effects of evaporation and rapid CO2 offgassing can fatally compromise a paleoclimate 18 O record from speleothems. These processes can lead to significant enrichment in 18 O in the case of evaporation, and in both 13 C and 18 O in the case of rapid CO off-gassing. Hendy 2 (1971) devised several tests for these effects that have since been widely applied in speleothem research. Water will both evaporate and lose CO2 as it drips from a hanging stalactite and then splashes onto and flows as a thin film down a stalagmite or flowstone. If pCO2 is much lower in the cave than in the drip, off gassing of CO2 can be so rapid that 13 C and 18 O are preferentially (kinetically) lost with little or no isotopic back-reaction with the solution. This process can be identified as a gradual enrichment (and covariance between) in 13 C and 18 O in 18 Oc along a single growth layer. The effects of evaporation can be identified as an increase in 18 Oc alone along a growth layer. In addition to these screening criteria, sampling from poorly ventilated parts of the cave where pCO2 and humidity are high will reduce the risk of – but are no absolute insurance against – kinetic effects. The general equation used by most in describing the controls on change () in the 18 Oc value in speleothems

(2)

(5)

where (1) d{c-w }/dT is the temperature-dependent fractionation (Kim & O’Neil, 1997) between calcite and water, or 18.03 × 103 c-w = 1000ln ␣c-w = − 32.42 T (6) which is the same as approximately −0.2‰/◦ C at 0 ◦ C and −0.24‰/◦ C at 30 ◦ C. Temperature (in ◦ K) in d{c-w }/dT is assumed to be cave temperature, which is close to mean annual temperature (±1 ◦ K) in most caves provided that the cave is: (1) not so shallow that dripwater does not reach thermal equilibrium with the host rock; (2) so deep (>100 m) that rock temperature exceeds mean annual air temperature; or (3) sampled too closely to the cave entrance. (2) (␦18 Osw ) = changes in the ␦18 O value of seawater. This amounts to −0.08‰/10 m of sea-level change, or ∼−1.0 ± 0.1‰ across the last glacial/interglacial transition (Shrag et al., 1996). This can be factored into evaluations of ␦18 Oc provided that the age of the speleothem is known. (3) d{␦18 Omw }/dT = the dependence of ␦18 Omw on local air temperature variations. This dependence is a major imponderable in Eq. (5) and requires some careful consideration here. The most widely cited value for d{␦18 Omw }/dT in the speleothem literature is ∼+0.7‰/◦ C from Dansgaard (1964) (Fig. 1). Closer inspection of more extensive compilations of rainfall/T data shows that ␦18 Omw is a non-linear function of temperature (Yurtsever & Gat, 1981). Arctic stations display higher d{␦18 Omw }/dT values of +0.67–0.90‰/◦ C, whereas low-latitude sites show lower values ( 0

d␦18 Oc /dt < 0

d␦18 Oc /d␦18 Ov > 0

Cascade Cave (Gascoyne et al., 1981)

Gardner’s Gut and Waipuna Caves (Hendy & Wilson, 1968) Norman Bone/Grapevine Cave (Thompson et al., 1976) Caves in Bermuda, Kentucky, and W. Virginia (Harmon et al., 1978) Cold Water Cave (Harmon et al., 1979a) Unnamed cave (Wilson et al., 1979) Ease Gill Caverns (Gascoyne, 1992)

West Cave (Frumkin et al., 1999)

Lynds Cave (Goede & Hitchman, 1983) Little Trimmer Cave (Goede et al., 1986) Francome Cave (Goede et al., 1990) Cold Water Cave (Dorale et al., 1992) Onondaga Cave (Denniston et al., 2001) Okshola and Stordalsgrotta Caves (Lauritzen, 1995) Lithofagus Cave (Lauritzen & Onac, 1999) Spannegal Cave (Spötl et al., 2002)

Crag Cave (McDermott et al., 2001) Cold Water Cave (Denniston et al., 1999) Hulu Cave (Wang et al., 2001) Burns et al. (2001)

Ingleborough Cave (Gascoyne, 1992) Lost John’s Pothole (Gascoyne, 1992) Victoria Cave (Gascoyne, 1992) Exhaleair and Nettlebed Caves (Hellstrom et al., 1998) Soreq Cave (Bar-Mathews et al., 1999)

For these reasons, little paleo-temperature estimation based on fluid inclusion analysis has been published since 1985. Although too imprecise for paleo-temperature reconstruction, the ␦Dfiw analyses have shown that the ␦Dmw was ∼ 12‰ (or very roughly −1.5‰ in ␦18 Omw ) lower in the fullglacial period than today. Moreover, all but a few studies prior to 1990 found a negative relation between coexisting ␦Dfiw and ␦18 Oc values through a stalagmite time series (Table 1). Since ␦Dmw is positively correlated with temperature, so must temperature be negatively correlated with ␦18 Oc . This relation was further confirmed in most speleothem data sets by the fact that ␦18 Oc values are lower in cold (full glacial) times compared to values in modern calcite (Gascoyne, 1992). Dorale et al. (1992) revisited the paleo-temperature issue at Cold Water Cave (first examined by Harmon et al., 1979b) by studying a record from deep in the cave spanning much of the Holocene. Rather than relying in the ␦D value from fluid inclusions, Dorale et al. (1992) made a simplifying assumption to get at paleo-temperature (after Hendy & Wilson, 1968), that all variations in the ␦18 Omw in the past are T dependent and can be approximated as d{␦18 Omw }/dT = 0.695‰/◦ C (from Dansgaard, 1964). Equation (5) can be solved iteratively for T by inserting this term in the equation and by assuming that (␦18 Osw ) is negligible for the Holocene. Using this approach, Dorale et al. (1992) interpreted the increase of nearly 2‰ in ␦18 Oc during the mid-Holocene (Fig. 2b) to indicate a 2–3 ◦ C increase in temperature. This approach is also reflected in the interpretation of ␦18 Oc results from Crevice Cave (Dorale et al., 1998) and Onondoga Cave (Denniston et al., 2001) in nearby Missouri. In both caves, decreases in ␦18 Oc of 2‰ during the Younger Dryas (Denniston et al., 2001) and 1.5‰ between 55,000

and 41,000 yr ago (Dorale et al., 1998) were interpreted to result from temperature decreases. The conundrum here is that most other caves studies up until 1990 (Gascoyne, 1992), as described previously and including Cold Water Cave (Harmon et al., 1978, 1979a), concluded just the opposite, that increases in ␦18 Oc values reflected a temperature decrease. Differing estimates of d{␦18 Omw }/dT contributed to the differing interpretations (Fig. 1). Dorale et al. (1992, 1998) and Denniston et al. (2001) assumed the rather high, global average value for d{␦18 Omw }/dT from Dansgaard (1964), high enough that d{␦18 Omw }/dT > (d{D c-w }/dT)T + ␦D 18 Osw in Eq. (5). Harmon et al. (1978) and Schwarcz & Yonge (1983), using constraints from ␦Ddw , concluded that d{␦18 Omw }/dT must have been smaller, on the order of 0.3‰/◦ C and as a consequence that (d{c-w }/dT)T + (␦18 Osw ) > (d{␦18 Osw }/dT)T in Eq. (5) (Fig. 1). Denniston et al. (1999) obtained the most detailed Holocene ␦18 Oc records from Cold Water Cave yet (Fig. 2b), and their interpretations potentially add a further complication to paleoenvironmental interpretation of speleothem ␦18 Oc values. Three speleothems were analyzed from three widely spaced sites deep in the cave. One of the speleothems (1S) was the same as that analyzed by Dorale et al. (1992) in less detail, and the same clear increase in ␦18 Oc was observed during the mid-Holocene. The three separate speleothem records show poor agreement (Fig. 2b), especially with respect to the strongly opposed shifts in ␦18 Oc values in speleothem 1S and 2SS during the mid-Holocene, and the comparative lack of response in ␦18 Oc values from speleothem 3L. Denniston et al. (1999) attributed the large discrepancies between records to the differing effects of

Isotopic Records from Ground-Water and Cave Speleothem Calcite in North America 209 Fig. 2. Time (103 yr) vs. a. δ13 C c values and b. δ18 O c values from Cold Water Cave, Iowa (from Denniston et al., 1999). Estimated percent C4 biomass shown in a. uses the soil-diffusion model of Cerling (1984) and Quade et al. (1989). The percent C4 biomass (summer grasses) increases to near 50% in the mid-Holocene “prairie” period. Denniston et al. (1999) interpret the δ18 O c values from CWC 2SS as the least effected by evaporation and therefore the most representative of changes in δ18 O mw through time. They view the decrease in δ18 O mw values in CWC 2SS during the mid-Holocene as indicating an increase in the fraction of moisture derived from Pacific vs. the Gulf of Mexico.

evaporation on recharge waters above each of the speleothem sites. For example, speleothem 1S shows the largest enrichments in 18 O during the mid-Holocene of the three records. And the 1S site underlies a south-facing slope most subject to evaporative enrichment of 18 O in soil water. In contrast, speleothem 2SS underlies a north-facing slope and was therefore deemed to be the most useful for paleoenvironmental reconstruction due to a lack of evaporation effects. The results from 2SS show a ∼1.8‰ increase in ␦18 Oc from the early to mid-Holocene (Fig. 2b). Using the same assumptions as Dorale et al. (1992) did previously for interpretation of 1S, this 1.8‰ shift would imply 3–4 ◦ C cooling during the mid-Holocene. However, no

such cooling is indicated by other, independent records from the region. To explain the decrease in ␦18 Oc , Denniston et al. (1999) invoked an increase in the proportion of winter (with lower ␦18 Omw values) compared to summer rainfall (with higher ␦18 Omw values) during the mid-Holocene. In essence, Denniston et al. (1999), along with other recent publications described below, add another variable (␦18 Ov ) to Eq. (5): 18

␦ Oc =

18 d{␦ Omw } d{c-w } T + T dT dT

+ (␦18 Osw ) + (␦18 Ov )

(9)

210

J. Quade

where ␦18 Ov denotes the ␦18 O value of water vapor that arrives at a site, and which is determined by a variety of factors not directly dependant on temperature, such as storm trajectory, seasonality of rainfall, and rainfall amount (but not sea-water ␦18 O changes). In summary, the five studies over the past 25 years of the ␦18 Oc records from Iowa and Missouri have yielded up three very different paleoclimatic interpretations. Harmon et al. (1978, 1979a) combined ␦18 Oc and ␦18 Oiw results to interpret the higher ␦18 Oc values during full-glacial period as indicating a colder cave temperatures during speleothem formation. From the same cave, Dorale et al. (1992) took higher ␦18 Oc values to indicate warmer conditions in the mid-Holocene. Denniston et al. (1999), acting from a larger data set, invoked evaporative effects and changes in ␦18 Ov responding to changes in storm track, not temperature, as the main determinants of ␦18 Oc . More fundamentally, the interpretations of Denniston et al. (1999) challenge a key assumption of nearly all previous speleothem studies, that in temperate caves ␦18 Omw = ␦18 Odw . For temperate caves this assumption was seemingly satisfied by the work of Thompson et al. (1976), Goede et al. (1982), Yonge et al. (1985), and others, although in drier regions such as the Judean Hills (mean annual precipitation ∼0.5 m/yr) evaporative enrichment in 18 O of all drip waters in the vadose zone has been convincingly argued (Ayalon et al., 1998; Bar-Mathews et al., 1995). In the end, the causes for the lack of agreement in the ␦18 Oc values between stalagmites at Cold Water Cave could be explained by the evaporation effects suggested by Denniston et al. (1999). Clearly, there is a need to rigorously test this explanation in a number of caves by systematically sampling temperate cave drip waters fed by recharge on slopes of different aspect. Some Results From Outside North America Other recent results from around the world reflect the same evolution in thinking in North America on how ␦18 Oc results should be interpreted in the absence of independent constraints on ␦18 Omw (once thought to be provided by fluid-inclusion analysis). As one example, the record from Hulu Cave in eastern China spans 10,000–70,000 yr and shows a long-term oscillation in ␦18 Oc values of 4–5‰ but also shorter term shifts of 1–2‰ that seem to correlate with Heinrich Events in the Greenland ice records (Wang et al., 2001). ␦18 Oc is positive during cold periods such as marine oxygen isotope stage (OIS) 2, the Heinrich Events, and the Younger Dryas. Whereas the short-term shifts in ␦18 Oc might be explainable by changes in cave temperature alone, the longer term 4‰ shift between ∼16,000 and 10,000 yr ago is not (Fig. 1). Barring any unforeseen kinetic fractionation effects in the cave, a major, temperature-independent shift in ␦18 Omw seems called for. Wang et al. (2001) invoke insolation-related oscillations in the strength of the East Asian summer (low ␦18 Omw ) and winter (high ␦18 Omw ) monsoon to explain the record, and down-play the role of temperature change. Burns et al. (2001) similarly interpreted the ␦18 Oc speleothem record from Hoti Cave in Oman in terms of local

changes in the strength of the Indian Monsoon. Both these instances fit with the strong correlation between rainfall amount and ␦18 Omw values at low latitudes (Hong Kong, r 2 = 0.71; New Delhi, r 2 = 0.45, Rozanski et al., 1993). Other recent cases for large, temperature-independent shifts in the ␦18 Ov come from West Cave near Jerusalem studied by Frumkin et al. (1999) and Crag Cave in Ireland (McDermott et al., 2001). The West Cave record extends almost continuously for the last 170,000 yr, and shows a very similar coarse structure to ␦18 O results from ice and marine records for the same period. Glacial-age (OIS 2 and 6) ␦18 Oc values are 3.5–5‰ more positive than interglacial values. This would appear to place the West Cave results in the same category as many other speleothem records, where (d{c-w }/dT)T + ␦18 Osw dominates over the (d{␦18 Omw }/dT)T term in Eq. (5). However, even assuming a very low d{␦18 Omw }/dT = 0.3‰/◦ C would only produce a ∼1‰ increase in glacial-age ␦18 Oc (see Fig. 1, T = −8 ◦ C line), well short of the observed 3.5–5‰ shift. Frumkin et al. (1999) therefore invoke a major increase in local glacial ␦18 Osw , due to evaporative enrichment of the Mediterranean during glacial lowstands to largely explain the very low glacial-age ␦18 Oc values in West Cave. In an analogous manner, McDermott et al. (2001) interpreted decreases in ␦18 Oc values from Crag Cave in Ireland in terms of temperature decreases during the Holocene enhanced by decreases in ␦18 Ov caused by North Atlantic meltwater surges. As in North America, however, some researchers outside the U.S. continue to view changes in temperature, not in ␦18 Ov , as the main determinant of ␦18 Oc . One recent example comes from Soreq Cave, only 330 m higher and 10 km away from West Cave, which yielded a ␦18 Oc record remarkably similar to that of West Cave and yet the paleoclimatic interpretation of Bar-Mathews et al. (1997, 1999) is totally different. The Soreq Cave record displays the same 3–3.5‰ decrease in ␦18 Oc during the last deglaciation as in the West Cave record. However, Bar-Mathews et al. (1999) interpreted the shift largely in terms of temperature change, not changes in ␦18 Ov . In summary, most speleothem studies prior to 1990 both within and outside North America documented higher ␦18 Oc values during cold phases such as the last glacial period compared to values from modern speleothems. This response was viewed to indicate that the combined (d{c-w }/dT)T + ␦18 Osw terms dominated over a relatively low (d{␦18 Omw }/dT)T term for most continental settings. In contrast, most speleothem records documented in the last ten years show lower ␦18 Oc values during cold phases. The reason for this contradiction remains unclear and only partly can reside in the improved quality of U-series dating using TIMS (thermal ionization mass spectrometry) compared to the less precise alpha-count dating of pre-1990. The more recent results from North America are consistent with the recognition that d{␦18 Omw }/dT is much larger than once thought for inland continental settings. More recent studies also are recognizing that Eq. (5) is an unrealistically simplistic view of the determinants of ␦18 Oc values in most caves, and that in many caves the former promise of speleothems

Isotopic Records from Ground-Water and Cave Speleothem Calcite in North America 211 in yielding paleotemperature estimates appears to be out of the question.

Devils Hole Three important isotope records from ground-water calcites, two from Devils Hole, and one from nearby Death Valley, have come out of the Mojave Desert. The Death Valley record involved measurement of the ␦D values of inclusive water in ground-water calcites exhumed by recent tectonism along the east side of Death Valley (Winograd et al., 1985). The record spans the last 2 million years and shows a large increase of 40‰ (much larger than the uncertainty on the ␦Dfiw measurements) in ␦Dfiw values. Winograd et al. (1985) attributed this to a decrease in the average value of ␦18 Ov as the Sierra Nevada mountains were uplifted to the west. The best-known ground-water calcite records in the region come from two cores obtained by SCUBA divers at Devils Hole. DH–2 spans 50,000 to 310,000 yr ago (Winograd et al., 1988) and DH–11 50,000–550,000 yr ago (Fig. 3; Winograd et al., 1992), based on multiple alpha and TIMS U-series dates from both cores. Neither record has any apparent depositional hiatuses. Further records filling in much of the last 80,000–19,000 yr have been obtained but are largely unpublished (Winograd et al., 1996). Dates from the cores lag climatic events by the time required for water to travel from the recharge area to Devils Hole. Thus, dates from Devils Hole represent minimum ages (by KF) Q > KF > OW = PF = H = P Q > KF > OW = PF = H = P Q > KF > OW = PF = H = P Q > KF > PF = OW = H = P (Pye and Johnson show KF ≫ PF) N.D.

N.D.

N.D.

Locally 10–30 if thick and deeply buried or erosional remnant

MI ≫ V

Q > KF > OW = PF = P

N.D.

N.D.

N.D.

25% pine (Pinus), 5% beech (Fagus), 15% prairie forb (Asteraceae, Chenopodiaceae/Amaranthaceae, and Artemisia), 10% ragweed (Ambrosia), and 10% hemlock (Tsuga) pollen (from Shuman et al., 2002b).

assemblages (Fig. 9) (Levesque et al., 1993; Walker et al., 1991; Williams et al., 2002). Both types of data can be controlled by factors other than temperature, but in certain settings, such as those considered here, each can yield useful paleotemperature estimates. We also map past moisture balance trends derived from lake-level data (Fig. 3) (e.g. Harrison, 1989; Shuman, 2001; Shuman et al., 2002b; R.S.

Webb et al., 1993). Striking regional similarities among the lake-level histories of multiple lakes studied by the Digerfeldt (1986) method likely reflect long-term changes in regional moisture-balance (e.g. among New England and Quebec lakes studied by Almquist et al., 2001; Lavoie & Richard, 2000; Newby et al., 2000; Shuman, 2001; Shuman et al., 2001). The maps also include: (1) evidence for drier-than-modern

Climatically Forced Vegetation Dynamics in Eastern North America During the late Quaternary Period 463

Fig. 4. Maps of individual taxa, combinations of three taxa, and biomes for selected dates during the past 21,000 years. Upper three rows show the distributions of individual taxa: spruce (Picea), birch (Betula), and ash (Fraxinus). Three progressively darker shades of green represent progressively higher relative abundance. Grey indicates the absence of the taxa at the mapped abundance level, and white indicates regions with no data. The fourth through sixth rows contain maps of three combinations of three taxa: spruce, sedge (Cyperaceae), and pine (Pinus); ash, hornbeam (Ostrya-type), and elm (Ulmus); pine, hemlock (Tsuga) and beech (Fagus). Individual taxa are plotted, where they grow alone, as either red, yellow, or blue, and combinations of the taxa are plotted as combinations of these primary colors. The bottom row shows maps of biome distribution. Biomes include cool deciduous forest (CDEC, light blue), taiga (TAIG, dark blue), cool conifer forest (CCON, light green), cool mixed forest (CLMX, yellow), temperate deciduous forest (TDEC, pale blue), warm mixed forest (WMMX, bright green), xerophytic woodland (XERO, orange), mixed parkland (MXPA, red), spruce parkland (SPPA, gray), conifer woodland (CEOD, pink), steppe (STEP, orange), desert (DESE, bright orange), and tundra (TUND, purple). Maps are accurate within an envelope of 500 years about the time assigned to them.

464

T. Webb, B. Shuman & J.W. Williams

Fig. 4. (Continued )

conditions from sedimentary hiatuses in lake cores (Webb & Webb, 1988); and (2) qualitative estimates of lake-level change based upon sediment type, aquatic macrofossils, and other indicators (e.g. Harrison, 1989; Webb, 1990). General circulation model experiments have provided another source of independently derived information about past climates at broad scales. The experiments show possible climate responses to known changes in climatic boundary conditions (Fig. 1), such as the long-term changes in inso-

lation (Berger, 1978), glacial extent (Dyke & Prest, 1987), and atmospheric carbon dioxide concentrations (Indermühle et al., 1999; Monnin et al., 2001). Model simulations may not be accurate at all scales or for each climatic variable, but previous data-model comparisons support the general simulation of glacial-to-interglacial climate change (Webb, 1998; Whitlock et al., 2001; Wright et al., 1993) and of certain millennial-scale climate changes (Rind et al., 1986; Rutter et al., 2000).


Fig. 5. Tree cover maps for all woody taxa, needle-leaved taxa, and broad-leaved taxa. Tree coverages are expressed as percentages of total area within each 50 × 50 km grid cell (from Williams, 2003). Results North American Climates Over the Past 21,000 Years In the time since the last glacial maximum (LGM), 21,000 years ago, North American climates have changed significantly at both orbital and millennial time scales. Between 21,000 and 6000 years ago, climates warmed and full-glacial conditions gave way to interglacial conditions (Bartlein et al., 1998; COHMAP, 1988). Changes in seasonal insolation, ice-sheet extent, and atmospheric carbon dioxide and dust concentration are the main controls for these orbital-scale changes (Kutzbach et al., 1993, 1998) (Fig. 1). The glacialinterglacial transition was punctuated, however, by fast changes in global climate controls (Fig. 1), such as: (1) the reorganization of oceanic heat transport ca. 14,600, 12,900, and 11,600 cal yr B.P. at the beginning and end of the Bølling/Allerød and Younger Dryas chronozones (Broecker et al., 1985; Clark et al., 2002; Ruddiman & McIntyre, 1981; Rühlemann et al., 1999), which is tracked by changes in atmospheric 14 C concentration (Clark et al., 2001; Stuiver et al., 1995); (2) the rapid increases in atmospheric carbon dioxide concentration ca. 15,000 and 11,000 cal yr B.P. (Monnin et al., 2001); and (3) the final collapse of the Laurentide ice sheet ca. 8200 cal yr B.P. (Barber et al., 1999). At orbital time scales, data and model syntheses by COHMAP (1988; Webb, 1998; Wright et al., 1993) show how: (1) ice-sheet retreat; and (2) the shift from low to high to low seasonality in insolation (Fig. 1) created spatially

variable patterns of climate change across North America and the globe. In North America, the southern branch of the jet and winter storm track moved north, and the glacial anticyclone lessened and then disappeared as continental ice sheets retreated. From 16,000 to 9000 cal yr B.P., the combination of ice sheet coverage, CO2 levels, and increased insolation seasonality (Fig. 1) was different enough from today to induce the NCAR CCM1 to simulate climates unlike any today in the North American mid-continent (Fig. 2). These simulated climates without modern analogs were characterized by warmer-than-present summers, colderthan-present winters, and low precipitation levels relative to present (Kutzbach et al., 1998; Webb et al., 1998). The timing of maximum warmth, cooling, dryness, or wetness varied geographically, however, in response to the spatial variations in: (a) forcing; and (b) atmospheric circulation changes, which differentially altered the advection of heat and moisture (Bartlein et al., 1998). For example, during the early Holocene, higher-than-present seasonality of insolation and the retreating ice sheet altered radiation and temperature gradients. As a result, circulation patterns changed. The maps of lake-level variations (Fig. 3) in eastern North America illustrate the resulting spatial patterns of moisture-balance change by showing how the Midwest dried out while the Southeast and then the Northeast became wetter. Part of this gradual change was punctuated by an abrupt change in atmospheric circulation at 8200 cal yr B.P. induced by the final collapse of the Laurentide ice sheet (Shuman et al., 2002b).

466


Fig. 6. Vegetation anomaly maps for the last glacial maximum and mid-Holocene, expressed in terms of biomes, oak, spruce, pine, and prairie forbs. All anomalies are expressed as differences between the past time period and pre-settlement vegetation (21,000–500 cal yr B.P.; 6000–500 cal yr B.P.). Blank areas in the biome maps indicate no data or no change in biome type. In the top panels, colored areas show the biome assignments for the past interval in grid cells that have changed between the present and past. Biome abbreviations same as in Fig. 4. Green colors in the individual taxon anomaly maps indicate that a taxon was locally more abundant in the past; browns indicate areas where a taxon is more abundant at present.


Fig. 7. Maps of square-chord distance between fossil pollen spectra at each mapped 50 × 50 km grid point and pollen spectra at the same grid point 1000, 3000, and 5000 years later. The difference between modern (0 yr ago) and just before European settlement (500 yr ago) is also shown. Large differences between intervals are estimated by large square-chord distances, and plotted as dark gray or black. Distances greater than 0.20 are greater than the modern difference among biomes (see Overpeck et al., 1985). The contour intervals were chosen because they allow the maps to illustrate the contrast in the spatial patterns between the 1000-yr difference maps and the 3000-yr and 5000-yr difference maps.

Millennial-scale climate variability also resulted in distinct spatial patterns of change. Changes during the Younger Dryas chronozone (12,900–11,600 cal yr B.P.) provide an illustrative example. Sea-surface temperatures (SST’s) in the North Atlantic were colder then than during the previous two millennia (Ruddiman & McIntyre, 1981) and directly cooled proximate regions of North America (Mott et al., 1986; Peteet et al., 1990; Walker et al., 1991). These patterns of climate change during the Younger Dryas chronozone resulted from reorganized ocean and atmospheric heat transport, which

also led to seasonally warmer climates elsewhere in North America (Kneller & Peteet, 1999; Shuman et al., 2002a) and in other regions of the world (Bluiner et al., 1998; Clark et al., 2002; Rühlemann et al., 1999). Globally, Younger Dryas climate patterns were unique within the past 21,000 years because the changes in ocean circulation were embedded within a context of radiation forcing, ice-sheet extent, and carbon dioxide concentrations that was strikingly different from other times (Figs 1 and 4) (see also figures in Whitlock et al., 2001).

468


Fig. 8. Summary of New England climate and vegetation history. Hydrogen isotope ratios and lake-level estimates from Crooked Pond (Huang et al., 2002; Shuman et al., 2001) in southeastern MA are compared with pollen percentages from North Pond (Whitehead & Crisman, 1978) in western MA 206 km to the northwest of Crooked Pond. Inferred climate phases are given at top (from Shuman et al., 2003).

Individualistic Responses by Taxa to Orbital-Scale Climate Change Our maps match those of Jackson et al. (2000) in showing how much the vegetation patterns at the LGM differed from those today (Figs 4 and 6). Conifer parklands grew south of the ice sheet grading into pine woodlands along the southeast coast into Florida (Fig. 4). Deciduous trees and shrubs were not numerous, but grasses and sedges were relatively abundant implying an openness to the vegetation (Figs 4 and 5).

The full-glacial climatic conditions produced a distribution of biomes that differed markedly from today. Several modern biomes (e.g. taiga, cool conifer forest, and temperate deciduous forest) were not evident or were much restricted in range (Figs 4 and 6). With the beginning of the long-term ice-sheet retreat and temperature increase (Fig. 1), both the range boundaries and the regions of peak abundance for many taxa shifted northward as well as east or west (Fig. 4). For example, spruce was abundant from 21,000 to 13,000 cal yr B.P. in the east and its area of high abundance then rapidly shifted into western Canada by 11,000 cal yr B.P. By the early Holocene, the range of spruce covered much of boreal Canada as it does today, but was restricted by the ice sheet. Following 7000 years of low abundance in the east, however, the peak of spruce abundance shifted eastward from western to eastern Canada after 6000 cal yr B.P. The maximum abundance of many taxa also increased as ranges shifted. Birch abundance began increasing in Beringia after 18,000 cal yr B.P. and south of the ice sheet only after 16,000 cal yr B.P. In eastern North America, hemlock only began increasing as it spread north through the Appalachians after 10,000 cal yr B.P. (Figs 3 and 4). Other taxa (e.g. ash and elm) reached greater-than-modern abundance and range extent during the late Pleistocene and early Holocene from 16,000 to 10,000 cal yr B.P. (Fig. 4). As a consequence of these varying independent movements among the taxa in terms of their geographic range and the location of their abundance maxima, new assemblages emerged and then disappeared. For example, spruce, sedge, and ash overlapped in their area of abundance from 16,000 to 12,000 cal yr B.P. to create a mixed parkland biome that does not exist today (Fig. 4). The mixed forest that contains both hemlock and beech today only formed after the ice sheet collapsed ca. 8000 cal yr B.P. (Figs 3 and 4). The associations of plant taxa and functional types that were unlike any growing today in North America (i.e. noanalog vegetation) appear to have developed under climates that were also dissimilar from today (Fig. 2). From 16,000 to 11,000 cal yr B.P., the NCAR CCM1 simulated climates without modern analogs just where the pollen data indicate that plant assemblages without modern equivalents grew (Fig. 2) (Williams et al., 2001). The observed match between the pollen data and model simulations supports earlier hypotheses that the no-analog plant associations formed as a consequence of plant taxa responding individualistically to multivariate changes in climate (Cushing, 1965; Overpeck et al., 1992). Besides changes in temperature and the seasonality of insolation, variations in moisture balance also played a large role in altering vegetation patterns. For example, within the early Holocene, lake-level maps show the Midwest becoming dry between 10,000 and 7000 cal yr B.P. and the vegetation responding with the eastward expansion of the prairie (Fig. 3). In the Northeast, conditions were dry at 10,000 cal yr B.P., and then became wetter as mesic taxa like birch, hemlock, maple, and beech replaced the white pine populations that had been dominant there (Fig. 3; Davis & Jacobson, 1985; Jackson et al., 1997; Newby et al., 2000; Shuman et al., 2001, 2002b). The Southeast was

Climatically Forced Vegetation Dynamics in Eastern North America During the late Quaternary Period 469 Fig. 9. (a) Summer temperatures at Splan Pond, New Brunswick, spanning the Younger Dryas chronozone, inferred from fossil chironomid evidence (Levesque et al., 1993) and compared to pollen data from the same site (Mayle & Cwynar, 1995). Arboreal pollen types shown are spruce (Picea), aspen (Populus), cedars (Cupressaceae), alder (Alnus), birch (Betula), and total pines (Pinus) with white pine (P. strobus-type) shown in gray; nonarboreal pollen types are willow (Salix) and sedges (Cyperaceae). (b) A histogram of the time lags (positive if pollen data are lagging) associated with the most significant (i.e. lowest p-value) cross-correlations between the climate proxy record and pollen principal components done down-core for each of eleven sites in Maritime Canada and Europe (Williams et al., 2002).

also drier than today before 10,000 cal yr B.P. and then its moisture balance became more positive as southern pines replaced oaks there after 8000 cal yr B.P. These maps show how closely the vegetation changes match those in moisture balance and how well taxa favoring wet or dry conditions respond to the appropriate changes in moisture (Fig. 3). The climate response surfaces for these pollen taxa support these interpretations (T. Webb et al., 1993, 1998). The adjustments of individual plant taxa to climate change also led to new arrangements of vegetation cover, density, and functional composition. Biome maps (Figs 4 and 6) show large changes in the position and area of biomes between the last glacial maximum and present, with a major reorganization between 14,000 and 9000 cal yr B.P. Compared to the individual taxa, the reconstructed biomes did not move much, however. Instead of migrating large distances as climate changed, biomes grew and shrank in places where novel associations of taxa and functional types appeared and disappeared (Fig. 4). Biomes characteristic of the Holocene (e.g. temperate deciduous forest, taiga) were absent or occurred only in limited areas during the late Pleistocene, whereas the mixed parkland biome largely disappeared after 11,000 cal yr B.P. (Fig. 4). In some cases, shifts in biome position and type can be traced

to the distribution of individual plant taxa – for example, the northward spread of spruce populations caused the end of its association with sedge, ash, and hornbeam (the end of the mixed parkland) and the beginning of an association with alder and birch (the rise of taiga) (Fig. 4). Here, not only range shifts but the increases in maximum abundances of alder and birch after 16,000 cal yr B.P. also played a key role in the development of the new biome. In other cases, variations in the distribution of individual taxa are subsumed as internal variations in biome composition (Williams et al., 2003). As categorical representations of the vegetation, the biome maps limit apparent vegetational change to ecotonal regions between biomes (see the 6000 cal yr B.P. to present differences in Fig. 6). Continuous measures of the vegetation variation such as shown on isopoll maps (Fig. 4) and difference maps (Fig. 6) demonstrate that this emphasis on ecotones is an artifact of the biome categories; actual change was far more widespread (Figs 4–7). Describing the relative density of various plant life forms avoids limitations inherent to categorical classifications of the vegetation while still providing an estimate of vegetation structure (DeFries et al., 2000). Reconstructions of % tree cover based on calibrating pollen records to AVHRR

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Fig. 10. Pine and spruce pollen percentages track abrupt range shifts during the Younger Dryas chronozone (12,900–11,600 cal yr B.P.) Records of pine (Pinus) (A) and spruce (Picea) (B) pollen percentages are plotted with time on a vertical axis to show changes in their geographic distributions. Arrows indicate the east-west range shifts of pine (A) and the north-south range shifts of spruce (B). Each site is labeled by state as in Shuman et al. (2002a). The Stotzel-Leis site in Ohio (Shane, 1987) is shown in gray in inset A, and superimposed upon the stratigraphy from Pretty Lake in Indiana (Williams, 1974) (from Shuman et al., 2002a). (Advanced Very High Resolution Radiometer) observations of the modern vegetation indicate that total tree cover has increased since the last glacial maximum (Fig. 5) (Williams, 2003; Williams & Jackson, 2003). Total tree cover in eastern North America at the last glacial maximum was generally 10% (from Shuman et al., 2002a).

When we zoom in to specific time series in the Northeast (Shuman et al., 2003), stable isotope ratios and lake-level estimates show six phases of climate history in southern New England during the past 15,000 years, corresponding with six regional pollen zones (Fig. 8). The different arrival times of the dominant taxa match well with the development of favorable conditions for their growth and do not appear to be controlled by dispersal limitations (Johnson & Webb, 1989). The new climate records support Deevey’s (1939) original climatic interpretation of the sequence of pollen zones and add to the climate history in Davis et al. (1980). Highly negative hydrogen isotope ratios and lower-thanmodern lake levels (Huang et al., 2002; Newby et al., 2000; Shuman et al., 2001) indicate that New England climates were colder and drier than present before 13,000 cal yr B.P., when cold-tolerant spruce and pine populations were common. An expansion of spruce populations and a decline of pine populations during the Younger Dryas chronozone followed a cooling indicated by lowered stable isotope ratios (Huang et al., 2002; Shemesh & Peteet, 1998). After 11,600 cal yr B.P., a rapid positive shift in hydrogen isotope

Climatically Forced Vegetation Dynamics in Eastern North America During the late Quaternary Period 471 ratios and a decrease in lake levels indicate a rapid shift to warmer- and drier-than-earlier conditions coincident with the decimation of spruce populations and expansion of white pines (Huang et al., 2002; Lavoie & Richard, 2000; Newby et al., 2000; Shuman et al., 2001; R.S. Webb et al., 1993). The pines remained abundant until the populations of mesic types like hemlock and beech expanded. This change coincided with a rise of lake levels ca. 8000 cal yr B.P. (Shuman et al., 2002b). Hickory populations did not expand in New England until more than a thousand years later when the hydrogen isotope ratios indicate that conditions became at least as warm as modern. Chestnut populations, which prefer higher soil moisture levels, expanded after 3000 cal yr B.P. when lake-levels indicate that conditions became as wet as modern (Shuman et al., 2003). Warm and dry conditions from 5400 to 3000 cal yr B.P. likely contributed to the decline of hemlock populations at 5400 cal yr B.P. by contributing to moisture stress in the trees and making them more susceptible to disease (Bhiry & Filion, 1996; Davis, 1981a, b; Haas & McAndrews, 2000; Newby et al., 2000; Shuman et al., 2001, 2003; Yu et al., 1997). The independent climate records clarify many of the issues that Davis (1965) critically evaluated.

Individualistic Responses by Taxa to Millennial-Scale Climate Change Paired records of fossil pollen data and independent climate estimates at individual sites indicate rapid vegetation responses to the abrupt climate changes associated with the Younger Dryas chronozone (Fig. 9). Many North American sites show a shift to cold-tolerant vegetation between 12,900 and 11,600 cal yr B.P. (e.g. Mayle & Cwynar, 1995; Mott et al., 1986; Peteet et al., 1990; Shane, 1987; Shuman et al., 2002a). High-resolution records reveal that such vegetation changes were nearly as fast as the climate changes that caused them. For example, cross-correlation analysis between pollen and chironomid data from Splan Pond, New Brunswick, shows that vegetation lagged climate by less than 100 years (Williams et al., 2002) (Fig. 9). Similarly rapid responses (3000 years, did not register as millennial-scale differences (Fig. 7). Likewise, slow orbitally forced climate changes probably contributed little to the large changes calculated for the beginning of the Younger Dryas chronozone (1000-yr distance maps in Fig. 7). Much of the difference from 13,000 to 12,000 and from even 12,000 to 11,000 cal yr B.P. resulted from century-scale changes in vegetation composition that were synchronous at many sites ca. 12,900 and 11,600 cal yr B.P. (Fig. 10) (Shuman et al., 2002a). The vegetation changes across 3000- and 5000-year intervals (Fig. 7) and across the past 21,000 years (Fig. 6), however, dwarfed even the large millennial-scale differences associated with the Younger Dryas climate changes. The glacial-to-interglacial climate warming was far larger than the reorganization of climate patterns during rapid climate changes caused by ocean and atmospheric circulation changes. Maps of the long-term changes in insolation (Berger,

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1978; Whitlock et al., 2001) and net radiation (Whitlock et al., 2001) show that the mid-latitudes underwent a large longterm increase in summer radiative forcing between 21,000 and 11,000 cal yr B.P. (as well as between LGM and modern). Consequently, even a large reorganization of the thermohaline circulation and oceanic heat transport could not widely alter climatic patterns of ca. 12,000 cal yr B.P. to be like those of 21,000 cal yr B.P. In some regions, cooling was significant at the onset of the Younger Dryas (e.g. Greenland, Stuiver et al., 1995) (Fig. 1) and local vegetation patterns changed significantly (Figs 9–11), but the continental-scale climate gradients (and vegetation patterns) immediately before, during, and after the Younger Dryas remained more like each other than those even 3000 years later (Figs 4–7). The longterm northward migration of taxa, such as spruce (Fig. 4), was altered only subtly because the latitudinal band where climate was appropriate for them remained north of their LGM positions and well south of their Holocene positions. Comparison between pollen diagrams and maps shows that large local changes are not always indicative of large continental-scale changes. On local scales, orbital- to centennial-scale climate changes had a large impact on ecosystem composition (Figs 8–10), but only the larger magnitude, long-term changes appear important in continentalscale maps of individual taxa and assemblages (Fig. 4). The geographic displacement of vegetation types associated with century-scale climate changes (even when those that involved range shifts of hundreds of kilometers, Figs 10 and 11) resulted in large local changes in vegetation composition, but was small compared to the range shifts of thousands of kilometers from 17,000 to 6000 cal yr B.P. (Fig. 4). On the continental scale, rapid change is only evident (e.g. 13,000–12,000 and 12,000–11,000 cal yr B.P.) when many individual sites show rapid change simultaneously (Fig. 10). Many pollen diagrams show several rapid (century-scale) changes during the past 21,000 years, but most of these changes resulted from long-term time-transgressive changes that rapidly impacted individual sites while gradually shifting species’ ranges (see spruce decline discussion in Bernabo & Webb, 1977). The inference that large local changes reflect global climate changes is “iffy” at best and can only be sustained when maps show how the local changes link to other changes over a large area.

Conclusions The emerging view from independently inferred paleoclimate estimates and patterns is allowing new assessment of climate-induced vegetation dynamics. Forcing factors at both orbital and millennial scales cause climate to change which in turn affects the pattern of major vegetation changes across the continent and at individual sites. The data show that the vegetation has changed when climate changed both at continental scales and at individual sites. The rates of migration, abundance change, and assemblage change have all varied at the recorded rate of climate change. Different recorders of the climate along with climate modeling results are allowing: (a) the influence of temperature changes to be separated from that

of moisture change (Huang et al., 2002; Shuman et al., 2002b, 2003); and (b) the impact of the changing seasonality to be assessed (Williams et al., 2001). The multivariate nature of climate change combined with multivariate response surfaces that are unique for each plant taxon allow for widely varying responses among taxa to similar changes in climate (Bartlein et al., 1986; T. Webb et al., 1993). Differential migration paths and rates among taxa are expected as the result of such multivariate climate forcing and responsiveness (Prentice et al., 1991), and taxa today that grow together should have in general arrived or at least expanded in abundance at different times (Shuman et al., 2002a, b, 2003; Williams et al., 2003). Individualistic plant behavior scales up to cause the emergence and disappearance of plant assemblages, of combinations of plant functional types, and hence of biomes. Changes at these higher levels of ecological organization ultimately result in feedbacks to the climate system (Kutzbach et al., 1996; Street-Perrott et al., 1990). The individual taxa move farther and more dramatically, however, than the biomes that they comprise. Changes by individual taxa combine and scale up to become differences in vegetation structure and composition not captured by biome classification (Williams, 2003; Williams et al., 2003). Multiple depictions and visualizations of vegetation change are therefore needed to portray its full diversity of changes. As new data sets and increasingly sophisticated vegetation and climate models come on-line to improve past climate estimates, we foresee many new insights into vegetation dynamics that will include how: (1) vegetation-climate feedbacks (Kutzbach et al., 1996); (2) changing atmospheric concentrations of carbon dioxide (Cowling & Sykes, 1999; Davis, 1991); and (3) evolution and extinction of taxa (Davis & Shaw, 2001; Jackson & Weng, 1999) enhance or inhibit the shifting of taxon abundances and distributions as climate changes. A current fascination with understanding the impact of global warming on ecosystems, landscapes, and human livelihood has placed a premium on studying decade-to-millennial scale climate changes. We are hopeful that the reductionism of this approach will in time give way to a more holistic approach to understanding vegetation and climate dynamics. Such will be necessary if we are to understand the roles of millennial and orbital scale climate and vegetation dynamics in evolution and speciation. Studies of the genetic make-up of fossil plants are just beginning and should open many exciting lines of inquiry (Davis & Shaw, 2001). Such studies will allow us to understand how Hutchinson’s (1965) evolutionary play unfolds in the continuously changing ecological theater (Webb, 1995).

Acknowledgments Grants from the Earth System History Program at NSF supported this research. Earlier support came from the Office of Energy Research at DOE. B. Shuman was supported during the writing by a NOAA Climate and Global Change Postdoctoral Fellowship and the University of Oregon. J. Williams was supported by the National Center for

Climatically Forced Vegetation Dynamics in Eastern North America During the late Quaternary Period 475 Ecological Analysis and Synthesis, a Center funded by NSF (Grant #DEB-0072909), the University of California, and the Santa Barbara campus. We thank P. Leduc and P. Newby for technical support and P.J. Bartlein for cartographic advice and paleogeographic information. Our mapping relied on excellent work by E. Grimm and J. Keltner in developing the North American Pollen Database (NAPD). We thank J.T. Overpeck and R.S. Webb for their work at NGDC at NOAA to support the development of this database. We also thank numerous palynologists for making their data available to the NAPD and Pierre Richard for giving access to data from Base de Donnees polliniques et Macrofossiles du Québec.

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Holocene fire activity as a record of past environmental change Cathy Whitlock and Patrick J. Bartlein Department of Geography, University of Oregon, Eugene, OR 97403, USA; [email protected]; [email protected] above-background levels of charcoal particles. The size and exact location of fires cannot be resolved with the specificity of dendrochronological studies, but charcoal records have the advantage of providing a fire reconstruction that spans thousands of years and encompasses periods of major climate change and vegetation reorganization. Annually resolved fire reconstructions are possible, but in most cases fire history is described in terms of fire episodes (one or more fires) during a time span of years to a few decades. Fire was recognized as a past and present link between climate change and vegetation response in one chapter (Davis, 1965) of the review volume for the VII Congress of the International Association for Quaternary Research (Wright & Frey, 1965). Since 1965, research in fire history has undergone a renaissance that has improved the use of fire data as both a paleoclimatic and paleoecologic tool. Recent studies consider fire as a proximal cause of vegetation changes and also recognize that vegetation patterns (both spatially and temporally) help shape fire regimes. The role of climate as the ultimate control of both vegetation composition and fire regimes is also widely recognized. In this chapter, we discuss some of the recent advances, including efforts to: (1) better understand the processes that introduce charcoal into lakes and wetlands; (2) refine the methods to interpret these deposits; and (3) evaluate the response of fire to climate and vegetation controls operating on different time scales based on paleoecological evidence, paleoclimate simulations, and modern assessments. We focus this review on research in North America.

Introduction Fire is the dominant form of natural disturbance in temperate forests, and, as such, it serves as a process that modulates forest susceptibility to climate change, disease, and other forms of disturbance. Fire has been identified as an important catalyst of vegetation change during rapid climate shifts in the past (e.g. T. Clark et al., 1996; Swetnam & Betancourt, 1998), and it has been implicated as the primary agent of ecosystem change in the future (e.g. Overpeck et al., 1990, 2003; Watson et al., 2000). At the global scale, biomass burning is considered an important but poorly understood process in the global carbon cycle, one that releases greenhouse gases, aerosols, and particulates to the atmosphere but also sequesters carbon as inert charred matter and ash (Cofer et al., 1997; Watson et al., 2000). At the regional scale, fire plays an essential role in maintaining the integrity of forest ecosystems (MacNeil, 2000; Mills & Lugo, 2001; Nature, 2000). Because of fire’s importance as an ecosystem process at large and small scales, it is necessary to understand: (1) the response of fires to past, present, and future climate change for global change assessments; and (2) the role of fire in maintaining forest health and promoting ecosystem change for better forest management. Like many types of paleoenvironmental data, information on past fires can be interpreted in climatic terms as well as used as an indicator of how particular ecosystems respond to known climate changes. The ultimate objective of paleoenvironmental research is to do both – understand the cause and ecological consequences of climate change. Two sources of paleoecological data provide information on fire-climate interactions. One source, dendrochronological data, includes records of fire-scarred tree-rings and maps showing the distribution of forest-stand ages following fire (see Agee, 1993; Johnson & Gutsell, 1994, for information on methods). Dendrochronological methods provide highly resolved spatial reconstructions of past fire activity, but they are limited by the age of living trees, which spans only the last few centuries in most places. This relatively short duration makes it difficult to examine the role of fire during periods of major climate change. Moreover, tree-ring records are best suited to reconstruct low-intensity ground fires that do not kill trees and often offer little information on the frequency of stand-replacing crown fires, which have become more widespread in western forests in the last two decades. The second data source is the record of particulate charcoal deposited in lakes and wetlands during and shortly after a fire (see Whitlock & Anderson, 2003; Whitlock & Larsen, 2002, for information on methods). Fire occurrence is identified by sedimentary layers with abundant or

Refinements in Charcoal Analysis The use of charcoal data to reconstruct fire history in North America began in the late 1960s and early 1970s when microscopic charcoal particles (generally 100 ␮m in diameter) recovered in sieved residues (e.g. Carcaillet et al., 2001b; Long et al., 1998) and on petrographic thin-sections (e.g. Anderson & Smith, 1997; Clark, 1988b, 1990) from contiguous core samples. Studies that use macroscopic charcoal emphasize the fire history of the local watershed, and regional reconstructions are based on networks of local records. The attention on local fire reconstructions has necessitated a better understanding of charcoal taphonomy, i.e. the processes that introduce and deposit charcoal to a lake or wetland. The principles of particle motion physics have been used to describe the transport of charcoal particles from a point source, suggesting that particles >1000 ␮m in diameter are deposited near a fire, particles 102 years), large-scale changes in the climate system caused by variations in the seasonal cycle of insolation, atmospheric composition, and atmosphere-ocean interactions emerge as important controls of the fire regime. Variations in the seasonal cycle of insolation on orbital time scales (103 –105 years) govern the slowly varying components of the climate system, and these apparently shape the long-term fire regime and vegetation. High-resolution charcoal records provide an opportunity to examine changes in fire activity in response to changes in the large-scale controls. For example, a record from Cygnet Lake in Yellowstone National Park was examined to reconstruct the fire history of the last 17,000 years, and compare it with changes in summer insolation that varied the intensity of

←−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−− Fig. 2. A–F. Cygnet Lake, Yellowstone National Park (Millspaugh et al., 2000) provides an example of the charcoal decomposition approach of Long et al. (1998). A. Charcoal concentration calculated in contiguous samples; B. Data re-expressed as charcoal accumulation rates; C–D. Data decomposed into background trend (smooth curve) and charcoal peaks above background (Peaks are inferred to be fire episodes); E. Locally weighted frequency of charcoal peaks per 1000 years; F. Summer insolation anomaly. (Comparison of E and F suggests that time of maximum fires and the broad rise and fall of fire frequency is governed by the influences of the insolation anomaly.) G–L. Climate patterns shown in global climate model (GCM) simulations and modern NCEP data provide an explanation for high-fire activity noted at Cygnet Lake. G–I. 11,000 cal yr B.P.-minus-present anomalies of July 500 mb heights, vertical velocity, sea-level pressure, as simulated by a GCM (Kutzbach et al., 1998; see also Bartlein et al., 1998). J–L. Composite anomalies for same variables in NCEP data (Kalnay et al., 1996; Kistler et al., 2001) for recent years of large-area burned in Montana.

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drought in the region (Millspaugh et al., 2000) (Fig. 2). The record suggested that fire frequency in the early Holocene (between 10,000 and 11,000 cal yr B.P.) reached a maximum of 10–15 episodes/1000 years, in contrast to frequencies of 600 gms in mass (e.g. Frison & Bradley, 1999), requiring quarried masses of stone of even greater size. Such high-quality stone was less failure-prone, and more easily re-worked as supplies dwindled, both of which would have been important to mobile hunter-gatherers who could not predict when they would next be near a stone source (Goodyear, 1979). Because the stone can often be identified to the outcrop from which it was obtained, it is apparent Clovis groups were routinely moving hundreds of kilometers across the landscape (Tankersley, 1991). The scale of their mobility varied by region, depending on the nature and density of resources being exploited, but in general Clovis groups had far more extensive ranges than groups in later prehistoric times. There have been several efforts to determine more precisely the size of those ranges

Peopling of North America 549 and the populations occupying them (e.g. Anderson, 1995; Fiedel, 2000), and although interesting, these are necessarily speculative equations with many uncontrolled variables. The high mobility of Clovis groups is reflected in their sites (Kelly & Todd, 1988; Surovell, 2000). For the most part these are small, lack site furniture (items that go with a place and not with the persons occupying the place, Binford, 1979), and rarely include storage pits or evidence of habitation structures. Those structures that do occur are insubstantial. All of which indicates these groups were not staying for long periods of time at particular places, nor returning to specific places repeatedly (the exception here are their stone sources, which were often returned to repeatedly [e.g. Gardner, 2002]). This, in turn, suggests they were exploiting resources that were widely available and not place-restricted: such is the advantage of colonists on a landscape without other people, a landscape without social boundaries. Clovis artifacts and sites have been found across the continent in a variety of environments, from the rich grasslands of the western Plains to the complex boreal/deciduous forests of the American southeast (papers in Bonnichsen & Turnmire [Eds], 1991). No subsequent North American occupation was so widespread or occupied such diverse habitats. It is important to add, however, many areas appear only sparsely occupied, including the Great Basin, the Columbia and Colorado Plateaus, northern Great Plains, northern Rockies, and the uppermost and lowermost reaches of the Mississippi Valley (e.g. Anderson & Faught, 2000). This spotty distribution is undoubtedly biased by differences in the ages of exposed geomorphic surfaces, contemporary land use patterns, and the amount of archaeological and/or collector activity (Shott, 2002). Still, a more representative sample would likely only change the details, and not the essential fact that the Clovis presence on the landscape was broad, not deep – the manifestation of highly mobile people at low population densities. The geographic spread of Clovis across the continent was often envisioned as a more-or-less uniform diffusion across space, perhaps in an expanding wave front (e.g. Martin, 1973; Young & Bettinger, 1995). More recent GIS-based studies using the details of North American topography and terrain, the inferred viability of different Pleistocene habitats, and the continent-wide distribution of Clovis sites and isolated fluted points, predictably paint a more complicated picture (e.g. Anderson & Gillam, 2000; Steele et al., 1996, 1998; cf. Fiedel, 2000). Anderson & Gillam (2000), for example, propose that expansion threaded across the continent through areas of relatively low topography, along major river valleys, in proximity to glacial and pluvial lakes and, in eastern North America, along the coastal margin. Expansion could have proceeded in either a “string of pearls” or a “leap-frog” mode, the former a spatially continuous expansion, the latter rapid jumps across large stretches of landscape, possibly with pauses in “staging areas” along the way (Anderson & Gillam, 2000; also Anderson, 1995). A “leap-frog” model seems to better fit the dense but widely separated clusters of Clovis and Clovis-age materials across the continent, but must have entailed strategies to maintain a critical mass of

population and/or interaction with distant kin, so as to offset the demographic danger of living in small numbers over large areas devoid of other people (Anderson, 1995; Anderson & Gillam, 2001; Moore, 2001; Moore & Moseley, 2001). The chronology of the Clovis occupation varies across the continent and, as was the case in 1965, remains better known for some areas than others. The oldest sites are still those on the Great Plains and in the Southwest, which range in age from 11,570 to 10,900 14 C yr B.P. (Holliday, 2000; Stanford, 1999; Taylor, 2000). Significantly, the earliest appearance of Clovis continues to coincide with the opening of a viable ice-free corridor – granting that the timing of the opening is made fuzzy by uncertainty about the timing of deglaciation, draining, revegetation, etc. of the emerging corridor (Mandryk, 2001; Mandryk et al., 2001). There are now reliable radiocarbon dates on a dozen or so Clovis and Clovis-like fluted point sites in the eastern United States, but despite the continued suspicion that Clovis technology emerged in this area (Stanford, 1991), no eastern fluted point site yet pre-dates 11,500 14 C yr B.P. (Anderson et al., 2002). A very few approach that antiquity (e.g. Shawnee-Minisink, Pennsylvania; Dent, 2002]), but the majority fall between 10,600 and 10,200 14 C B.P. – a period contemporaneous with the Folsom (post-Clovis) occupation on the Great Plains (Bonnichsen & Will, 1999; Haynes et al., 1984; Lepper, 1999; Meltzer, 1988). Clovis-like materials occur in the Great Basin and far west, perhaps as early as 11,500 14 C yr B.P., although their ages remain uncertain, and the cultural chronology is confused by the possibly contemporaneous occurrence of large unfluted stemmed points (Beck & Jones, 1997; Bryan & Tuohy, 1999; Grayson, 1993). The radiocarbon record supports the long-held suspicion these groups radiated rapidly across the continent, the process taking perhaps no more than 500 radiocarbon years. The latter part of this episode, of course, overlaps the Younger Dryas (YD), which includes several radiocarbon plateaus that distort radiocarbon ages (Hughen et al., 2000; Kitagawa & van der Plicht, 1998; Peteet, 2000; Taylor et al., 1996). Although calibration may ultimately change the apparent speed of dispersal (perhaps “slowing” it to, say, approaching 1000 calendar years – see Dincauze, 2002; cf. Fiedel, 2000), it will nonetheless remain one of the fastest expansions of an archaeological complex known in prehistory (among the few cases faster were the expansion of prehistoric Thule across northern Canada in the centuries after 900 A.D. [Meltzer, 2002, discusses similarities and differences between the Thule and Clovis dispersals], and the dispersal of Lapita groups throughout Near and Remote Oceania beginning some 3000 years ago [Kirch, 1997]). It seems reasonable to suppose, given human population density was lower at this time than at any subsequent period in American prehistory, and that these groups were occupying a relatively rich landscape, that demographic pressure was not fueling that dispersal. These groups traveled much farther and faster than they had to if they were just looking for new land to siphon off burgeoning populations that would have otherwise put a strain on local resources (Kelly, 1996; Mandryk, 1993; Meltzer, 1995). But why (or how) did they move so far, so fast?

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Human Hunting and Pleistocene Mammalian Extinctions The traditional explanation is that Clovis people were specialized hunters in pursuit of wide-ranging big-game, notably the now-extinct Pleistocene megafauna, which were themselves able to override ecological boundaries (Kelly & Todd, 1988; Mason, 1962; Martin, 1973). Clovis colonizers could compensate for their lack of knowledge about the landscape by exploiting the same prey-species niche through all the habitats they traversed (Kelly, 1996). Harvesting the same food resources in new locations using traditional weaponry, hunting skills, and tactics would allow efficient and rapid dispersal (Keegan & Diamond, 1987; Kelly & Todd, 1988). Such a subsistence strategy and the rapid dispersal it could permit would buttress the claim humans were responsible for the extinction of the megafauna at the end of the Pleistocene in North and South America (Martin, 1973). By this scenario the herbivores met their demise directly as a result of intensive hunting, while carnivore population temporarily boomed with the sudden increase in scavengeable carcasses on the landscape, then went bust when the supply ran out. If hunters were killing off the megafauna at rates proposed by the overkill model (Martin, 1973; Mosimann & Martin, 1975), that would be incentive enough for rapid expansion, for their prey were always in front of them, not behind them. However, there are many reasons to doubt that scenario. First, recent studies of faunal remains from Clovis sites, as well as of the isotope geochemistry of rare human bone from this period, show that these late Pleistocene groups exploited a greater variety of animal and plant resources than traditionally supposed (Green et al., 1998; Johnson, 1991; Meltzer, 1993; Spiess et al., 1985; Stanford, 1991, 1999). Second, it is unlikely hunters could convert search and processing strategies and tactics successful against one prey species for use on another (Frison, 1991, 1999). Third, ethnographic and archaeological evidence demonstrates that specialized big-game hunting – let alone the hunting of a continent of animals to extinction – was rare among hunter-gatherers, and linked to particular habitat types (Binford, 2001; Hofman & Todd, 2001; Meltzer, 1993). This is not surprising: although models of foraging theory as well as empirical evidence suggest animals of large body size are high-ranked prey, that is not necessarily true of animals of the largest body sizes (risk comes into play); and, on productive, game-rich landscapes essentially devoid of other people – late Pleistocene North America – foragers would likely abandon a patch before extinction of the local fauna (Broughton, 1994; Grayson, 2001; Kaplan & Hill, 1992; Kelly, 1995). Winterhalder & Lu (1997) model circumstances under which depletion and extinctions are quite plausible, but it is doubtful those circumstances characterized this place and time, given the richness of the late Pleistocene landscape. Finally, and most telling, in spite of four decades of “time and more concerted effort” (Williams & Stoltman, 1965, p. 674), few additional Clovis big-game kill sites have been found, and most of those are still in western North America.

Clovis kills continue to be conspicuously scarce in eastern North America, despite the rich record there of terminal Pleistocene fossil localities. In fact, it is only a slight exaggeration to say that the overall tally of kill sites may even be less today than it was in 1965. For we now have greater knowledge of the various natural processes that can fracture or disarticulate skeletal remains in ways that mimic human activity. We can employ more stringent criteria to differentiate claims of association between artifacts and extinct faunal remains that are compelling, from claims that are not; and we can better differentiate evidence of hunting from evidence for other behaviors – such as scavenging (Binford, 1981; Grayson & Meltzer, 2002; G. Haynes, 1991; Haynes & Stanford, 1984; Lyman, 1994). Even recently, for example, Laub re-evaluated the Hiscock (NY) mastodon “kill” and concluded there is no evidence to support that interpretation. Instead, he now argues the site was a quarry where Clovis groups obtained mastodon bone and ivory from geological deposits (Laub, 2002). Applying such criteria to all purported Clovis or Clovisage megafaunal “kill” sites – of which there are more than 75 – leaves only 14 sites in North America (Fig. 1) for which there is secure and unambiguous evidence of human hunting (Grayson & Meltzer, 2003). When measured against the archaeological record of the hunting of extinct mammals of, for example, Upper Paleolithic France, that is a remarkably thin record (Grayson & Meltzer, 2002). Of those 14 sites, twelve contained mammoth, and the other two mastodon. There are no unequivocal kill sites for any of the other 33 genera of North American large mammals that went extinct at the end of the Pleistocene (Table 2a), again despite their abundance in the late Pleistocene paleontological record (Grayson, 2001). That remains of a few of those other genera are occasionally found in small numbers in archaeological sites is intriguing, but proves little more than their contemporaneity with Clovis people on the late Pleistocene landscape (Grayson & Meltzer, 2002). Turning the matter around, bison and other large mammals (Table 2b and c) were hunted in North America beginning as early as ∼10,900 14 C yr B.P. In the case of bison, there is abundant archaeological evidence of planned hunts, bone beds containing hundreds of slaughtered animals, impactfractured projectile points and skinning and butchering tools (e.g. Frison, 1991). Such intensive predation was often highly wasteful: of the 200 bison stampeded by hunters into an arroyo at the Early Holocene-age Olsen-Chubbuck site (Colorado), some 25% of the animals at the bottom of the carcass pile were left to rot untouched (Wheat, 1972). Nearly 11,000 years of human predation culminated in widespread slaughter by buffalo hide hunters in the late 19th century. Yet, despite being hunted for millennia bison (and, for that matter, the taxa in Table 2c) failed to go extinct; while 35 genera of animals that were not hunted at all or very little (mammoths and mastodons) did go extinct. Of course, as noted above, it has not been demonstrated that all 35 of those now extinct genera were contemporaries of humans, or even lasted until the terminal Pleistocene – let alone, that all went extinct simultaneously. Although we often assume as much (Martin, 1984), only 14 of the 35

Peopling of North America 551 Table 2. North American late Pleistocene mammals. Taxa that are in bold are ones for which there is secure archaeological evidence of human predation. Order

Family

Genus & Species

2a. North American Genera that went extinct in the Late Pleistocene Xenartha Dasypodidae Pampatherium sp. Holmesina septentrionalis Glyptodontidae Glyptotherium floridanus Megalonyhiae Megalonyx jeffersonii Megatheriidae Eremotherium rusconii Nothrotheriops shastensis Mylontidae Glossotherium harlani Carnivora Mustelidae Brachyprotoma obtusata Canidae Cuon alpinusa Ursidae Tremarctos floridanusa Arctodus simus Felidae Smilodon fatalis Homotherium serum Miracinonyx trumani Rodentia Castoridae Castoroides ohioensis Hydrochoeridae Hydrochoerus holmesia Neochoerus pinckneyi Lagomorpha Leporidae Aztlanolagus Perissodactyla Equidae Equus spp.a Tapiridae Tapirus spp.a Artiodactyla Tayussuidae Mylohyus nasutus Platygonus compressus Camelidae Camelops hesternus Hemiauchenia macrocephala Paleolama mirifica Cervidae Navahoceros fricki Cervalces scotti Antilocapridae Capromeryx minor Tetrameryx shuleri Stockoceros spp. Bovidae Saiga tataricaa Euceratherium collinum Bootherium bombifrons Proboscidea Mammutidae Mammut americanum Elephantidae Mammuthus spp.

Common Name Southern pampathere Northern pampathere Simpson’s glyptodont Jefferson’s ground sloth Ruscon’s ground sloth Shasta ground sloth Harlan’s ground sloth Short-faced skunk Dhole Spectacled bear Giant short-faced bear Sabertooth cat Scimitar cat American cheetah Giant beaver Holmes’s capybara Pinckney’s capybara Aztlan rabbit horses tapirs Long-nosed peccary Flat-headed peccary Western camel Large-headed llama Stout-legged llama Mountain deer Stag-moose Diminutive pronghorn Shuler’s pronghorn Pronghorns Saiga Shrub ox Harlan’s muskox American mastodon Mammoth

2b. North American species that went extinct, while other members of the same genus survived in North America Xenartha Dasypodidae Dasypus bellus Beautiful armadillo Carnivora Canidae Canis dirus Dire wolf Ursidae Temarctos floridanus Spectacled bear Felidae Panthera leo atrox American lion Artiodactyla Bovidae Oreamnos harringtonii Harrington’s mountain goat Bison antiquus Bison 2c. Select North American large mammal genera/species that survived Artiodactyla Cervidae Alces alces Cervus elaphus Odocoileus spp. Rangifer tarandus Antilocapridae Antilocapra americana Bovidae Ovibos moschatus Ovis spp.

Moose Elk Deer Caribou Pronghorn Musk ox Mountain sheep

Sources: Taxonomic data from Anderson (1984), Grayson (1991), Kurten & Anderson (1980). Information on archaeological occurrences from Frison (1991), Grayson & Meltzer (2002), and papers in Damas [Ed.] (1984), Gerlach & Murray [Eds] (2001), Helm [Ed.] (1981). a Others members of the same genus survived outside of North America.

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genera have reliable 14 C ages indicating survival past 12,000 years ago, the other 21 do not (FAUNMAP, 1994; Grayson, 2001, using criteria developed in Meltzer & Mead, 1983). In fact, the youngest reliable ages we have on some of those other 21 genera predate the LGM (Grayson, 2001). Even though that opens the possibility that some extinctions took place long prior to the appearance of Clovis in an area – and possibly on the continent (Goodyear, 1999; Grayson, 1991, 2001) – confirming that possibility will require additional evidence. The reason, simply, is that many of those same genera lacking terminal Pleistocene ages are also relatively rare in the fossil record, and the number of radiocarbon ages we have for a particular genus is strongly determined by how many fossils of that genus have been found (plotting the number of fossil occurrences against the number of radiocarbon ages [by genera] yields a highly significant correlation, r2 = 0.903 [data from FAUNMAP, 1994; Meltzer & Mead, 1985]). Until we get more radiocarbon dates, and can better discern the timing of their disappearance, we cannot conclude all genera disappeared simultaneously, or gradually, or let assumptions about the timing of extinctions be marshaled in support of arguments about its cause. Recent years have seen renewed efforts to bolster the case for human overkill: Alroy (2001) provides an elegant simulation model which he believes proves extinctions were an “unavoidable” consequence of the arrival of human hunters, while G. Haynes (2002) argues overkill was an inevitable outcome of late Pleistocene environmental change. However, the test of any simulation model, Alroy’s included, is not whether it can show a process could have occurred, but how it fares against the empirical evidence it purports to explain. In the absence of kill sites, the model fails. Haynes’ argument depends on there having been a “near continental drought” in late Clovis times (C. Haynes, 1991) which caused the megafauna to crowd together at water-holes where they were easy-picking for human hunters (Haynes, 2002). Unfortunately, there is no evidence of a drought of this magnitude or extent in Clovis times (Holliday, 2000), of crowding at the supposed megafaunal oases, or of slaughter thereat. The matter returns, as it must, to the empirical record, and there the facts are clear: very few kill sites have been found – and then only of mammoth and mastodon – this in spite of decades of intensive searching, and a rich paleontological record of many of these animals (we do not lack for fossils of this age). All of which makes it hard to avoid hearing the echo of Griffin’s (1965) conclusion: if human hunters had a role in killing off the megafauna they must have used magic, and then carefully hid the evidence. Living and Learning on a New Landscape Since specialized big-game hunting was not a significant component in Clovis subsistence, we still face the question of why (or how) Clovis groups moved so far, so fast. Haynes (1987) proposed the engaging scenario that curiosity, a charismatic leader(s) with the urge to see what was over the next hill or around the bend, and a landscape teeming with megafauna,

lured Clovis groups across the continent. Although curiosity and charisma assuredly played a role in individual cases, it does not provide a robust model for expansion across an entire continent over many centuries. Others attribute the fast dispersal to the changing climates and environments of Late Glacial North America, including the shift from patchy to zonal environments, a decline in faunal biomass, extinction of megafauna, and a change from equable to more continental regimes (Fiedel, 2000; Kelly, 1996, 1999; Kelly & Todd, 1988). Clovis groups had to move long distances and quickly, as local game populations declined precipitously in response those changes. Yet, the scale of those changes was on the order of centuries. Huntergatherers respond to the local weather – primarily on a daily and weekly basis – but also as it varied seasonally, annually, or over the course of their lifetimes. How or whether they respond depends on whether those changes triggered prey population fluctuations, reduced surface water, or otherwise restructured resource availability in ways that would have been detectable to and directly impacted their foraging activities. Long-term patterns of low frequency climatic variation over centuries (and many human generations) of the colonization process may not have been detectable on a human scale, and thus not directly relevant. Efforts have been made to link more rapid late Pleistocene climatic excursions – the Younger Dryas most prominently (Severinghaus & Brook, 1999; Steig, 2001) – to patterns in the contemporary archaeological record (e.g. Fiedel, 1999, 2000; C. Haynes, 1991). The YD has even been invoked by geneticists to explain the high frequency and reduced diversity of mtDNA haplogroup A2 among Na-Dene and Eskimo groups, on the assumption that harsh YD climates forced an occupational hiatus and thus a population bottleneck (Forster et al., 1996). However, evidence of a YD impact on humans in the high Arctic, where there is little doubt YD climate change was rapid and potentially significant, is equivocal (but see Mann et al., 2001). It is not apparent the YD triggered an occupational hiatus, or that changes in the distribution and abundance of sites during this period are real (as opposed to a vagary of sampling in this still little known region) (Bigelow & Powers, 2001). Yet, the Arctic may be the most likely region to see a YD impact on humans, if one is to be seen. So far, there is little evidence climate and landscape changes were as dramatic in mid-latitude, temperate North America (Grimm et al., 1993; Peteet, 2000), or had any impact on Clovis and later Pleistocene groups (Holliday, 2000; cf. C. Haynes, 1991). For my part, I have attributed the rapid and widespread movement of Clovis groups in part to their unfamiliarity with the landscape (Meltzer, 2002, 2003, 2004). Anthony (1990:901) observes that immigrants are not likely to move into areas about which they have no secure prior knowledge. Although the point is well taken, at some time in the North American past there was little choice in the matter. If, in fact, Clovis groups were colonizing a diverse and unfamiliar new continent, there was likely strong selective pressure to learn their landscapes (Meltzer, 2002). Landscape learning has at least three elements (entailments of which are discussed in Meltzer, 2003): wayfinding, tracking weather and

Peopling of North America 553 climate, and mapping resources (of all kinds: food, water, stone, etc.). Landscape learning would be especially important early in the colonization process, when environmental uncertainty was high, environments were patchy and varied temporally and at large scales (as they would relative to colonizers on a new landscape), and when human population numbers were low and groups were most vulnerable to extinction (Kaplan & Hill, 1992; also Kelly, 1995; Moore, 2001; Stephens & Krebs, 1986). Under these circumstances, selection would favor rapid and extensive exploration in order to reduce environmental uncertainty and forager risk, and provide foragers with the knowledge that would enable rapid niche shifts. There are demographic costs to moving that far that fast, for on a continent the size of North America populations would have been stretched thinly across the landscape (evidence of which may appear in the genetics of their descendants [Malhi et al., 2002]). To avoid inbreeding or, worse, extinction, groups would have to maintain a “critical mass” of population and an accessible source of potential mates, by participating in a larger effective gene pool (cf. Surovell, 2000). This would have been more or less difficult depending on the local group’s size, population growth rates, kin structure, age and sex composition, as well as how rapidly it was moving away from its geographic homeland and/or from other groups, and on environmental constraints on group size and population densities (Moore, 2001). Demography and landscape learning are tightly linked, as the decision to stay in a patch or move onto the next is in part based on the suitability of a new patch relative to the current one, after factoring in the costs of moving (Kelly, 1995). Those foragers who can better calculate those costs increase their chances of success and survival. By gaining information about a landscape one potentially reduces risk and mortality, and thus can increase population growth and recruitment rates. Arguably, then, the colonization process on a new landscape involved trade-offs between multiple competing demands (Meltzer, 2002): maintaining resource returns, or keeping food on the table, particularly as preferred or high-ranked resources declined, and in the face of limited knowledge of the landscape; maximizing mobility, to learn as much as possible, as quickly as possible about the landscape and its resources (in order to reduce environmental uncertainty in space and time), while maximizing residence time in resource-rich habitats to enhance knowledge of specific changes in resource abundance and distribution; minimizing group size, to buffer environmental uncertainty or risk on an unknown landscape; and, finally and most critically, maximizing the effective gene pool by maintaining contact between dispersed groups, in order to sustain information flow, social relations and, most especially, demographic viability. Colonizers had to balance the equation of moving to learn and explore, and staying to observe. Under this model we expect to see among colonizers large scale exploration to map the landscape (which, arguably, might be marked by stone tool caches); periodic aggregations of widely dispersed groups, to exchange mates, resources, and information; and extensive mating networks, in which

spouses can be drawn from distant groups. Central to making all this work would be high settlement mobility to maintain contacts with distant groups, map the landscape, and monitor resources and environmental conditions beyond the social and geographic boundaries of the local group; and open social networks, to enable individuals to move easily between and be readily integrated within distant groups. Highly territorial behavior would be decidedly disadvantageous. Although certain of those expectations are met (tool caches indeed occur in Clovis [Fig. 1], but not later Paleoindian times), the model has not been fully put to the test, largely for lack of sufficient data with the requisite temporal resolution (Meltzer, 2004).

The End of the Era One element common to many models of colonization (in North America and elsewhere), is that colonizers on a landscape with few other people not only had to be able to track great distances to find mates and exchange information and resources, they also had to be able to get along with near and distantly related groups they encountered (Kirch, 1997; Lourandos, 1997). Having large and open social networks based on flexible and fluid social and kin relations, fewer languages, the easy integration of individuals and groups, and sometimes long-distance exchange and alliance networks – all combine to diminish differences among peoples who need to be able to readily renew ties under geographic circumstances that might keep them apart for years at a time. One way these open social systems are manifest in the archaeological realm is by the widespread distribution, use, and exchange of instantly recognizable, and sometimes highly symbolic artifacts – such as unique styles of projectile points or, in the case of prehistoric Oceania, ceramic vessels (Kirch, 1997; Whallon, 1989). These forms served as a “currency” (a term not to be taken too literally) for social and ritual functions, and over long spans and large areas served to maintain recognition and alliances. Early in the Paleoindian period Clovis points are broadly similar stylistically, technologically, and typologically across a vast area of North America. The extent of Clovis distribution is likely a by-product of the size of the dispersal, but their similarity across that range may well reflect common symbols of an extensive social and mating network, which helped to check the attenuating effects of distance. But those effects were inevitable, cultural drift becoming more pronounced. Although the timing varies by area, new stylistic variants begin appearing sometime after 10,900 14 C yr B.P. in the central and western portions of the continent, and after 10,600 14 C yr B.P. in eastern North America. By 10,500 14 C yr B.P. the once pan-North American form is replaced by a variety regionally-distinctive point forms (Anderson, 1995; Anderson & Faught, 2000; Meltzer, 2002). Archaeologists have learned in recent years not to place undue weight on style and stylistic change in projectile points, for these may not be telling us about on-theground groups, dispersals, or adaptations (Dillehay, 2000;

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David J. Meltzer Fig. 4. Folsom projectile point in situ between the ribs of a Bison antiquus at the type site near Folsom, New Mexico. Like Clovis points, Folsom points are also fluted, but these points are routinely smaller, thinner, and more finely made, sometimes showing fine pressure flaking along the margins. This specimen, which is 4.54 cm in length, and is made of Flattop chert, which outcrops ∼450 km north of the site. Photograph taken ca. September, 1927, and it marks the first occasion in which artifacts were found in unequivocal association with an extinct species – although in this instance, the genus survived. See the text. Photograph courtesy of Denver Museum of Nature & Science, Denver, Colorado. All rights reserved, Image Archives, Denver Museum of Nature & Science.

Pluciennik, 1996). That said, these more-regionally specific styles that appear in mid-latitude North America starting after 11,000 14 C yr B.P. do seem to correspond with distinct adaptive strategies. Unlike Clovis, these later forms are more restricted geographically, have new and sometime preyor region-specific foraging strategies which occasionally involved new technologies (the two often co-occur [Binford, 2001]), and in places relied more on locally available stone indicating more restricted mobility. Folsom occupations, for example, appear on the Plains and Rocky Mountains of western North America, with a subsistence strategy tied to exploitation of Bison antiquus (Fig. 4) which, owing to a combination of competitive release following megafaunal extinction and the postglacial expansion of C4 grasslands, exploded in numbers. In other areas, late Paleoindian adaptations are different or sometimes less well known (Anderson & Sassaman [Eds], 1996; Beck & Jones, 1997; Frison, 1991). Assume, for the moment, these stylistic forms mark cultural groups (however defined), and that different forms

mark different groups in time and space (leaving aside how or whether these may have been isomorphic with languages or demes). Viewed that way, the shift from a single broad and relatively homogeneous form to multiple regional forms in the span from 11,500 to 10,500 14 C yr B.P., can be seen as the settling of colonizers in specific areas; a relaxation in the pressure to maintain contact with distant kin, and thus a reduction in the spatial scale and openness of the social systems (Meltzer, 2002). Stepping further out onto this speculative limb, one could attribute these patterns to an overall, continent-wide increase in population which reduced the need to maintain large and open social systems critical to insuring access to resources, information, and mates. Thus, once descendant populations were no longer demographically vulnerable and had little incentive to sustain long distance mating networks, the very vastness of North America and its topographic and geographic barriers would have conspired to impede interaction. The isolation of populations that resulted might

Peopling of North America 555 help explain the apparently anomalous human skeletons we see in early Holocene times at places like Kennewick, and the variability evident in the genetics of their descendants (Malhi et al., 2002; Powell & Neves, 1999). In effect, these early remains may not resemble contemporary Native Americans, but are ancestral nonetheless (Powell & Neves, 1999). The challenge, of course, is to devise ways of testing such hypotheses.

Where Do We Go Next? – Unresolved Issues in the Peopling of North America The last forty years, and especially the last half dozen, have seen great changes in our understanding of the first Americans. Much of what we knew or thought we knew about when the process of colonization began, how often migrations occurred, and how they played out on the late Pleistocene landscape, has been turned on its head. The conventional view of fast moving big-game hunters exploding on to the continent as the ice-free corridor zipped opened in terminal Pleistocene times is no longer a tenable scenario for the initial peopling of North America. Yet, though we now believe the peopling process to have been far more complex – beginning much earlier, possibly involving multiple migratory pulses, and involving a different entry route – fundamental questions about the antiquity, number, timing, and adaptive strategies of the Pleistocene peoples of North America remain. In the scramble to address those questions many ideas – some controversial – are being tried on for size. Time will tell which ones fit. In the meantime, there is much to do (cf. Fiedel, 2000; Meltzer, 2002). At the most basic level, we need more archaeological data. Although we may never detect the very first archaeological “footprints” of people on the North American continent, for reasons already noted, the lesson of Monte Verde is there must be unequivocal pre-Clovis sites here. The search for these will force the methodological question of whether we are (or have been) looking in the right places and in the right ways for early sites (Butzer, 1991; Collins, 1991; Dillehay & Meltzer, 1991), the answer to which will likely demand a greater contribution of geological and geophysical tools and techniques than now employed in archaeological field programs. Doing so should help fill in the temporal and spatial gaps in the archaeological record between the currently oldest evidence from Monte Verde, and the later and better known Clovis record. To resolve the question of entry routes, whether via the coast or interior, and if along the coast whether on foot or by boat or some combination therein, we will need direct archaeological evidence: sites, of course, but one can always hope for the discovery of a late-Pleistocene boat. Useful as well would be a firmer handle on the Quaternary geology and environmental history along the possible routes. There has been an increased effort in recent years to find sites marking a coastal entry using coring, remote sensing, bathymetry, and sampling of the submerged paleo-terrain (Fedje & Christensen, 1999; Fedje & Josenhans, 2000). None older

10,500 14 C yr B.P.

than have so far been found, but this is an effort with great potential. So too are efforts to model landscape evolution on the mainland, to help find uplifted Pleistocene coast. Although heavy vegetation here and on off shore islands (where remnant coastline may also be preserved) limits surface visibility, these are obvious areas to search. Ironically, we also still need more data on Clovis. That Clovis groups were not big game hunters begs the question of what they were eating. There is much to be gained by applying intensive recovery methods to Clovis age sites to capture organic remains of their diet (Dent, 2002; Ferring, 2001), by isotopic analysis of recovered human bone, and by searching for habitation and camp sites that can round out the picture of their adaptations. A critical part of that effort will be the derivation of finergrained data on the climates and environments of terminal Pleistocene times. What were the available resources, and how they were being exploited by these groups? As a subset of this point, we must address the question of whether late Pleistocene climatic and environmental change was sufficiently rapid and severe that it would have been detectable to and had an impact on human foragers. Ultimately, we need to close the gap between the archaeological and paleoecological temporal scales (recording change on a scale of decades or centuries), and our real-time (daily, weekly, yearly) models of human colonization processes. With finer-grained climatic and ecological data in hand, we are then poised to apply models of foraging theory to understand the adaptive strategies of hunter-gatherer populations as they may have responded to changes in resource abundance and distribution while moving across space and through time. Although these models were built for real-time activities of hunter-gatherer groups, properly scaled and employing archaeologically measurable variables (Grayson & Cannon, 1999), they could have great value here. For these models can give us insight – testable insight – into Clovis subsistence patterns, as well as the closely linked issue of how foragers on a people-free and unfamiliar landscape grapple with incomplete information and uncertainty. Developing and testing models of subsistence and landscape learning should ultimately answer the still-looming larger question of how and why Clovis groups expanded so far and so fast across late Pleistocene North America. Finally, that expansion and the potential processes of drift – cultural and genetic – have clear implications for both the archaeological record and the genetics of the descendant populations. It would be fruitful to model how population processes and social mechanisms of initial colonizers (wide-ranging mobility, the maintenance of long-distance mating networks, etc.) may have played out on the late Pleistocene North American landscape, and how these and possible later migrations might be manifest archaeologically, and what might be reflected in the genetic diversity of their descendants. Once the gaps in the archaeological record are filled in, and we possess a more extensive and detailed record of mtDNA and NRY diversity, it would then be especially interesting to determine whether (or at what points) the archaeological and genetic results converge.

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One obvious place to seek convergence would be in ancient DNA from late Pleistocene North American human skeletal remains. So far, however, none has been recovered, and the effort to do so will face problems of sample size, the difficulty of amplifying damaged DNA, the potential for contamination with modern DNA – not to mention concerns over scientific access to ancient human skeletal remains (Kolman & Tuross, 2000; cf. Merriwether, 2000). Even so, if (when) ancient DNA is recovered, it will provide the opportunity to test hypotheses derived from DNA of living populations, and reveal whether there were genetic lineages that entered the New World but subsequently went extinct (Meltzer, 1989; Merriwether, 2002). And that might go a long way toward explaining why patterns in the archaeological record and the genetic evidence from modern populations at the moment give different conclusions about the antiquity, number, and timing of migrations to the Americas.

Acknowledgments I am most grateful to Michael Cannon, Donald Grayson, Vance Holliday, Richard Klein, and Daniel Mann, for their valuable and constructive comments on a draft of this paper. Susan Bruning and Joanne Dickenson supplied important information; Judy Cooper provided much needed expertise in creating the figures. I thank Stephen Porter and Alan Gillespie for the opportunity to write this chapter, and especially their editorial patience and pointers. If all this collective expertise could not save me from error, the fault must surely be mine.

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Modeling paleoclimates Patrick J. Bartlein1 and Steven W. Hostetler2 2

1 Department of Geography, University of Oregon, Eugene, OR 97403-1251, USA Geological Survey at Geosciences Department, Oregon State University, Corvallis, OR 97331, USA

both global and local scales, using as input only the records of the external controls of climate (i.e. an “Earth-system model”), is perhaps a decade away. The indeterminacy of the data and the present limitations of the models thus dictate a synergistic approach for understanding climate variations that relies on integrating paleodata with paleoclimate model simulations. In this chapter we review the process of climate-system modeling and present a taxonomy of the models recently applied in the study of Quaternary climate change and variation. We also briefly trace the development of climate modeling since the 1965 INQUA volume and its companions were published. A synopsis of climate-modeling results for North America is provided, and we conclude with a discussion of some the emerging issues in the application of models for understanding climatic variations.

Introduction Data describe, models explain. Both are required to document and understand the past variations of Earth’s climate, and to help address the present problem of assessing climate change that may result from human activities. Models (for the most part conceptual as opposed to numerical) have long been applied for understanding climate variations during the Quaternary. Indeed, over a century ago, in a set of papers that contributed to the foundation of scientific method in the geosciences (multiple working hypotheses), T.C. Chamberlin (1897, 1899) provided a comprehensive conceptual model for explaining long-term climatic changes that is remarkably modern in some of its elements. What is regarded as a “climate model” today is generally a computerized numerical representation of the physical processes involved in the climate system, but conceptual models still play an important role in paleoclimate research. Whenever any kind of paleoclimatic data is interpreted, either quantitatively or qualitatively, some kind of model is invoked. Paleoclimatic data (of the kinds reviewed in this volume) and climate models play a complimentary role in understanding climate change. The data record how climate has changed, but data alone cannot provide an unambiguous explanation of why a particular climate state occurred or changed. This situation arises because most climatic variations recorded geologically have multiple, hierarchial causes (e.g. there is more than one way to create drought in a region) and because environmental subsystems display generally nonlinear responses to climate. Consequently, multiple cause-and-effect pathways can produce the same response in a paleoclimatic indicator. This indeterminacy of the “climate signal” is mitigated somewhat by considering networks of paleoclimatic data and by examining multiple indicators at individual sites, but such “multi-proxy mapping” cannot in itself eliminate the indeterminacy. Models based on physical principles (or widely accepted empirical representations of those physical principles) do have the potential to provide mechanistic explanations of past climatic variations, provided they are known to work, are applied in an appropriately designed experiment, and (perhaps most importantly) explicitly account for all of the components of the climate system that are involved in a particular climate change. Although comprehensive models of the climate system and its individual components (the atmosphere, oceans, biosphere, hydrosphere, and cryosphere) are evolving rapidly, the development of a comprehensive model that can simulate the temporal and spatial variations of climate on

Climate-System Modeling Many conceptual and numerical models that describe the workings of the climate system and its components have been developed, and there probably are as many taxonomies of those models as there are reviews of them. Primary reviews of climate modeling in general include Trenberth (1992), McGuffie & Henderson-Sellers (1997), and Randall (2000). Substantial information for paleoclimate modeling in particular can be found in chapters by Crowley & North (1991, Chaps 1 and 2), Kutzbach (1992), and Peteet (2001). Saltzman (2002) provides a coherent framework for understanding past, present and future climatic variations. Earlier modeling reviews include those by Schneider & Dickinson (1974), NRC (1974), Hecht (1985) and Kutzbach (1985). Because no two sources classify climate models or modeling studies in the same way, the task of providing an overview of the field is complicated. One way to organize a discussion of climate models and their application (climate modeling) is to consider first the nature of the climate system and what controls its variations through time, and then to describe a few large classes of climate models and their applications. Traditional definitions of climate are typically couched in statistics. For example, climate can be thought of as “. . . a set of averaged quantities completed with higher moment statistics (such as variances, covariances, correlations, etc.) that characterize the structure and behavior of the atmosphere, hydrosphere, and cryosphere over a period of time” (Piexoto & Oort, 1992), or, less explicitly, as “. . . the synthesis of


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weather in a particular region.” (Hartmann, 1994). Such statistics-based definitions are being replaced in practice (e.g. IPCC, 2001) by one in which climate is regarded as the collection of individual environmental components (jointly the climate system), and the record of their interactions and variations through time. Although the number of major components of the climate system is relatively small, the number of variables that describe these components is quite large, making a full cataloging of the climate variables that might be represented by models tedious and not very informative. However, the many variables involved in fully describing the climate system generally fall into one of three categories (Saltzman, 2002): those that describe the external forcing of the system (boundary conditions), those that describe the slowly varying aspects of the system (e.g. the size of ice sheets) that have traditionally been the focus of Quaternary paleoclimatology (slow-response variables), and those that describe the internal variables that are ordinarily thought of as weather (fast-response variables). A fourth category of variables, which we call subsystem variables, describes the state and function of the many environmental subsystems that are governed by climate and which provide paleoclimatic evidence or “paleodata” (Fig. 1).

Climate-System Variables Boundary Conditions In theory, the external controls of climate (or boundary conditions, a term borrowed from numerical analysis) include variables beyond the influence of climate. Such boundary conditions include: (1) the latitudinal and seasonal distribution of insolation (incoming solar radiation), as determined by variations in solar output and the elements of Earth’s orbit; (2) the configuration of continents and ocean basins including their topography and bathymetry, and the location of mountain chains and gateways between ocean basins; (3) the abiotic component of atmospheric composition, as determined by volcanic emissions; (4) a small flux of geothermal heat; and (5) human activities not controlled by climate. In practice, what is regarded as an external control as opposed to an internal response depends on both the experimental design of a modeling study and the timescale that it focuses on. Ice sheets, the terrestrial biosphere, and ocean temperature and salinity, for example, are most appropriately regarded as variables internal to the climate system. The areal extent and volume of the major ice sheets are ultimately controlled by

Fig. 1. The climate system (after Fig. 5–3 in Saltzman, 2002). The variables that describe the external forcing of the system (or boundary conditions) directly or indirectly influence the slow-response, fast-response, and environmental-subsystem variables, which in turn influence each other and determine the state of the climate system as a function of longitude, latitude, elevation and time. The arrows labeled “fields” indicate that one set of variables influences the other through patterns of atmospheric circulation, moisture and heat, while those labeled “fluxes” indicate that one set of variables influences another through the transfer of mass and energy. The dashed arrows indicate that the influence of the fast-response variable on environmentalsubsystem variables is currently unidirectional in climate models, but that eventually environmental-subsystem variables will interact with the fast-response variables as climate models develop.

Modeling Paleoclimates 567 external forcing and (over long time spans) can be thought of as an index of the internal state of the climate system. On century-to-annual time scales, however, the sizes of the ice sheets are comparatively constant. In simulations of particular times (e.g. the Last Glacial Maximum) or of sequences of times that are century in length or shorter, the ice sheets can therefore be regarded as one of the external controls that must be “prescribed” (or specified ahead of time) in those simulations. The same is also true for the temperature of the deep ocean and the components of atmospheric composition.

The slow-response variables of the climate system include ice sheets, large ice shelves, crust and mantle deformable by ice sheets, sea level, temperature and salinity of the deep ocean and the long-term state of its thermohaline and horizontal circulation, and slowly varying reservoirs involved in biogeochemical cycling that determine atmospheric composition. These slowly varying components are most often visualized as time series (as opposed to sequences of maps), and together describe the state of the climate system on century and longer time scales. As is the case for the boundary conditions, the particular role of these variables in climate models can be ambiguous and depends on the specific model and experimental design.

and feedback between the atmosphere and the surface (and might therefore be thought of as fast-response variables), and some that are dependent on climate but do not feed back to the climate system except in limited ways. There is no common mode for the visualization of these subsystem variables. Some variables are not easily categorized. Vegetation, for example, plays a key role in the instantaneous coupling of the atmospheric boundary layer and land surface by controlling the exchanges of energy and moisture. The rates of these exchanges depend on the structure of the vegetation and on the states of the atmosphere and underlying soil (including atmospheric humidity, wind, net radiation at the surface and soil-moisture availability) that together influence plant physiology (e.g. stomatal conductance). It was formerly thought that vegetation structure responds slowly to climate changes (on the order of hundreds to thousands of years), placing it in the category of slow-response variables. It is now clear that vegetation structure responds rapidly to climate changes over time spans of years to decades (Tinner & Lotter, 2001; Webb et al., this volume). Soils are dependent on climate and vegetation, but also have strongly expressed geological and geomorphic controls. Key attributes of the soil such as waterholding capacity (WHC) may be dominated by parent material (as in arenaceous soils), and so WHC might be regarded as a boundary condition; in other situations WHC is dominantly controlled by soil morphology, and hence acts like a slowresponse variable. The particular category a variable falls into is thus largely dependent on context, location and scale.

Fast-Response Variables

The Climate-Modeling Problem

The three-dimensional states of the atmosphere, land and ocean surfaces are represented by the fast-response variables. For the atmosphere, the key properties that are either observed or simulated include the distribution of temperature, pressure, wind and moisture (including clouds), and its trace-gas and mineral-aerosol composition. At the surface, the fastresponse variables include seasonally varying sea-ice extent, soil-moisture content, vegetation cover (including evapotranspiration rate, rooting depth, and albedo), and the temperature, depth and other physical and biological characteristics of the mixed layer of the ocean. On the shortest of time scales (hours-to-days) the fast-response variables can be thought of as describing the weather. In most instances, the fast-response variables are visualized using maps or sequences of maps such as those commonly used in weather forecasts.

The ultimate goal of climate modeling is to consider simultaneously the first three groups of variables listed above, and, as necessary, also to treat the other environmental systems that depend on climate (e.g. those described by the fourth group, the subsystem variables) – all in order to provide both a description and an explanation of the variations of climate through time. The result of modeling may be a single map or a series of maps or one or more time series. One can decompose the basic problem of climate modeling into a sequence of tasks: (1) use the record of boundary conditions to simulate the time history of the slow-response variables, (2) use the boundary conditions and state of the slow-response variables to simulate the fast-response variables, and (3) use the fast-response variables to understand and simulate the environmental subsystem variables.

Subsystem Variables

Climate Models

A large number of environmental systems and processes respond to variations of climate, and many of these, like those included in the terrestrial or marine biospheres, or surficial hydrologic systems, also provide the principal lines of paleoclimatic evidence that are used to reconstruct past climatic changes (Bradley, 1999). These systems have many variables, including some that play a role in the interaction

Climate models can be classified by describing the particular applications to which they are put (e.g. simulating the variations of the second-through-fourth set of variables described above), or by describing the comprehensiveness of different models – the number of processes and major components of the climate system they include and the (temporal and spatial) resolution at which those processes and components are

Slow-Response Variables

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represented. These classification schemes provide different ways of organizing the various models and are applicable to both conceptual and numerical models. Here we emphasize numerical models, but we do not underestimate the necessity, utility, and power of conceptual models.

Model Applications Most applications of numerical climate models can be categorized as having one of three general goals: (1) simulating the evolution over time of the climate system or one of its major components; (2) simulating the spatial patterns of climate-system components; and (3) simulating the detailed function of a single component or process. However, the present trends in climate-model development, which are leading to more comprehensive models of the climate system (see below), are in fact aimed at blurring the distinction among these applications.

Time-Evolution Applications Numerical climate models have been developed to simulate the evolution of climate over a range of time scales from geological (i.e. those that treat Cenozoic cooling or the onset of Quaternary glaciation) to inter- (and intra-) annual variations, as well as the trends in climate over the past several millennia (see Crowley & North, 1991, Chap. 1). The main goal of time-evolution modeling experiments is to simulate some macro-scale feature of the climate system, such as global ice volume or hemispheric-average temperature, using the record of external controls of climate as input (e.g. Birchfield et al., 1994). Models used in simulating time evolution commonly represent components of the climate-system at low resolution, as in box models or energy-balance models that include representations of continental and marine reservoirs or active layers without being spatially explicit (e.g. Harvey & Huang, 2001).

Spatial-Pattern Applications Models that focus on simulating the spatial patterns of climate include general circulation models (GCMs) and spatially resolved energy-balance models (EBMs) that include realistic geography (Crowley & North, 1991, Chap. 1). General circulation models, which simulate the three-dimensional structure of the atmosphere, ocean and land surface on time steps of minutes to hours, were originally developed for operational weather forecasting. The models continue to evolve and the operational models being run, for example, by the National Oceanic and Atmospheric Administration and the European Center for Medium-Range Weather Forecasts are now also being used for “reanalysis” projects (Kistler et al., 2001) an approach in which observed parameters of global weather (e.g. atmospheric soundings), sea-surface temperatures (SSTs), and ice cover are incorporated into

GCMs, which then provide simulations of the 3-D structure of the atmosphere and state of the land surface. The acronym “GCM” is often being taken to mean “global climate model.” A true global model, however, would be one in which all “internal” climate system components (i.e. variable groups 2 through 4 above) are explicitly represented (as opposed to being prescribed). Global climate models now exist in preliminary form as “EMICs” (Earth-system Models of Intermediate Complexity; Claussen et al., 2002). In practice, the extent to which other components of the climate system are included in a GCM (in addition to the atmosphere and land and ocean surfaces) is represented by additions to the GCM acronym. For example, AOGCMs include explicit representation of a three-dimensional ocean, while AVGCMs not only represent the regular physiological variations that must be included in any GCM, but also allow for variations in vegetation structure (Sellers et al., 1997). As more interactive components are added to GCMs, they will gradually evolve toward full Earth-system Models (ESMs). An important aspect of both GCMs and EBMs is their potential for simulating climate variables that may be crucial for understanding the response of particular paleoclimatic indicators to climatic variations. For example, many terrestrial subsystems (vegetation, lakes) are governed directly by the surface water and energy balances, or by “bioclimatic variables” (e.g. Prentice et al., 1992), as opposed to standard climate variables like temperature and precipitation. The former, mechanistic, variables are not commonly observed, or may indeed be unobservable, especially over regional scales. The potential for GCMs and EBMs, along with process models, to simulate presently unobservable variables also contributes to understanding the mechanistic controls of the variations in paleoclimatic indicators. Simulation of the time-evolution of the climate system can also be approached by conducting a series of spatialpattern simulations, or “snap-shots,” wherein the boundary conditions at a number of key times are established from the geologic record and are used to initialize models that produce “equilibrium” simulations for those times (e.g. COHMAP Members, 1988; Valdes, 2000; see also Peteet, 2001, for a discussion of the distinction between “snap-shot” and “sensitivity-test” experimental designs).

Subsystem (Process-Model) Applications Models that attempt to represent the variables that describe individual environmental subsystems are often called “process models” inasmuch as they are designed to incorporate, either explicitly or implicitly, the actual climatic mechanisms that govern a subsystem such as a watershed, lake, or plant. We exclude from this category statistical relationships that are developed by screening a large number of potential predictors of the distribution or variation of some paleoclimatic indicator, but which do include statistical relationships between predictors and mechanistically related responses (i.e. using mean July temperature or growing-degree days as a proxy for the heat and energy requirements of plants). Many

Modeling Paleoclimates 569 process models exist and we do not attempt an exhaustive review here. Process models range in scope (and scale of application) from those that represent the dynamics of ice sheets (e.g. Marshall, 2002) or global vegetation patterns to those that simulate the responses to climate variations within individual watersheds, lakes, or forest stands. Many subsystem process models are run in a “stand alone” or “off line” mode in which the there is no feedback with the climate model and input for the models is derived from the output of climate simulations. Process models are generally applied to individual points, such as a particular lake or forest stand, but they can also be run over a global network of sites or grid points, thereby producing simulations that are comparable with the spatial-pattern applications of global or regional climate models. Examples include equilibrium vegetation models (e.g. Harrison et al., 1998; Prentice et al., 1992) and ice-sheet mass-balance models (Pollard et al., 2000), applied over a grid, that use the results of a paleoclimatic simulation as input. The individual model grid points do not communicate with one another as they do in an AGCM or AOGCM.

Model Comprehensiveness A second way to classify climate models is by their comprehensiveness or scope, as was done by Claussen (2001) and Claussen et al. (2002) (Fig. 2). Claussen (2001) considers three attributes or axes: (1) the degree of model integration, or the number of interacting components of the climate system that are coupled within the model; (2) the number of processes explicitly simulated in the model, which can also be thought of as the cumulative dimension of the model; and (3) the detail of description in both time and space, commonly thought of as model resolution (Fig. 2). Highly integrated models account for the interactions and feedbacks among multiple components of the climate system, like the atmosphere, ocean, and terrestrial biosphere. In contrast, less integrated models explicitly represent one component (e.g. the atmosphere in an AGCM), and “parameterize” (or represent by very simple, sometimes empirically based relationships) the behavior of others (like the ocean and land surfaces in an AGCM). Models that represent many processes in a spatially explicit way (e.g. multiple-layer soil-moisture storage simulation) have higher cumulative dimensionality than others that include fewer processes in a more generalized way (e.g. the single-layer or “bucket” approach to soil-moisture storage). The third dimension of Claussen’s scheme, resolution, which has always been limited by computing resources, is easiest to envision. When individual models are placed in Claussen’s three-dimensional framework, several clusters emerge: (1) conceptual models that are of low spatial resolution (treating, for example, the whole Earth or greatly generalized continents and oceans) and in process dimensionality; (2) elemental (or low-dimensional) models of an individual or small number of the major components of the climate system; (3) EMICs, which are spatially explicit, and often represent multiple components, but generally at low spatial resolution (Claussen et al., 2002); and (4) comprehensive models,

Fig. 2. Classes of climate models plotted in a threedimensional space that describes model comprehensiveness (after Claussen, 2001; Claussen et al., 2002). The box is defined by the degree of model integration, or the number of interacting components of the climate system that are coupled together in the model, the number of processes explicitly simulated in the model, and the detail of description in time and space, commonly thought of as the resolution of the model. The gray shadows on the wall of the box represent the projections of each cluster of models. ESMs are shown using dashed lines because they currently do not exist. such as high-resolution “coupled” GCMs (Grassl, 2000). (Claussen et al. (2002) recognized clusters 1, 3, and 4, but we found enough differentiation among models to add cluster 2.) We also show a fifth cluster in Fig. 2 (represented by dashed lines) to illustrate the position of a full Earth-system model (ESM). In a progression of models across the three-dimensional space that defines comprehensiveness, the complexity of the models is measured to a large extent by their resolution and by the number of individual climate-system components described. This progression, however, is not a measure of the worthiness or value of the models. For example, conceptual models, which could be viewed as simple pencil-and-paper “thought experiments,” are actually among the more sophisticated of models in use in Quaternary science, whereas GCMs, which appear to be exotic and computationally demanding, are actually familiar to us as the source of day-to-day weather forecasts.

Conceptual Models Conceptual models include ideas and facts that we know are true (or that we are pretty sure are true, e.g. that Earth has experienced repeated glacial/interglacial variations during the Quaternary, related in some way to orbital variations), but

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they also include basic statements about how things work, often phrased in terms of hypotheses. Conceptual models are the oldest and most frequently applied kind of model used for understanding past climatic variations. The models are routinely applied by Quaternary scientists when looking at “raw” paleoclimatic data, no matter what kind (geologic, geomorphic, sedimentologic, paleoecologic, isotopic, etc.). Inasmuch as they often consider the Earth system as a whole, conceptual models cover the entire range of integration in the climate-model comprehensiveness space. As will be illustrated in the next section, conceptual models can take the straightforward form of an annotated list of potential controls of climatic variations (e.g. Mitchell, 1965), but most consist of a description of processes and interactions usually applied to a common sequence of events (e.g. Ruddiman, 2003). The common medium for reporting these models is a commentary in Science or Nature (e.g. Clark et al., 2002) or the summary chapter of a proceedings volume (e.g. Alley et al., 1999; Stocker et al., 2001). These latter discussions are most always supported by a number of less conceptual, more explicitly numerical modeling studies. Applications of conceptual models span the full range of temporal and spatial scales in Quaternary paleoclimatology and address issues that range from the question of why there are ice ages at all, to what controls individual wiggles in a single time series. Some conceptual models are expressed numerically. Examples include the Imbrie & Imbrie (1980) or Paillard (1998) models of glacial/interglacial variations in which global ice volume is related to insolation through simple differential equations, with the specific form of the relation determined by the state of the climate system (as represented by ice volume) in a geologically reasonable fashion.

Elemental (or Low-Dimensional) Models The second cluster includes models that numerically represent one or more components of the climate system in some kind of elemental or single-component fashion, and have lower spatial resolution, integration, or process dimensionality compared with models that are more comprehensive. This simplification results in lower computational demands (again relative to more comprehensive models). The number of models that potentially can be assigned to this cluster is quite large. The models range from highly spatially aggregated energy-balance models – like those of Budyko & Sellers (Budyko, 1982), which attempt to simulate global-average temperature from first principles and can be implemented with spreadsheet software – to spatially explicit energy-balance models like that of Crowley & North (1991). This cluster also includes simulating the slowly varying components of the climate system that must be viewed over relatively long time spans (Saltzman, 2002). So-called “box models” (e.g. Toggweiler, 1999) in which a few very large-scale geochemical reservoirs (and the flows among them) are simulated are also included in this cluster, as are models that feature coupling between components of the climate system such as the atmosphere and ocean, but for

only part of the globe (e.g. the tropics, as in Clement et al., 2001). Elemental or low-dimensional models have also been used extensively to assess potential future climate changes (e.g. Houghton et al., 1997), and to estimate of the sensitivity of global climate to carbon dioxide variations (e.g. Berger et al., 1998; Harvey & Huang, 2001).

EMICs (Earth-System Models of Intermediate Complexity) The cluster of models described as EMICs occupies an important position in the continuum of climate-model comprehensiveness because it offers a bridge between simple, low-resolution models and more comprehensive spatially explicit models (Claussen et al., 2002). In some ways, EMICs are very low-resolution AOVGCMs; in other ways, they are simple, low-dimensional models (like EBMs) to which some kind of simplified depiction of atmospheric circulation dynamics has been added. Fundamentally, EMICs simulate the interactions among more components of interest than can be done with the current generation of AGCMs or AOGCMs. The low spatial resolution of the EMICs, and the “parameterization” of many processes explicitly represented by GCMs, permits very long simulations of the temporal evolution of climate – an advantage over “snap-shot” simulations. Similarly, when used in “sensitivity-test” mode, EMICs allow a large number of combinations of inputs or parameters to be explored.

Comprehensive Models – GCMs The fourth cluster of climate models includes those that attempt to simulate the three-dimensional structure of the climate system and its variation over time. This cluster is exemplified by general circulation models (Randall, 2000), which in the current generation, include coupling among one or more climate-system components (although not as in as many combinations as the EMICs). In the most common form of application to paleoclimate, GCM experiments are designed by providing a set of boundary conditions of interest (such as insolation, atmospheric composition or the distribution of ice sheets), and the resulting simulations are analyzed as a sequence of gridded observations of a set of climate variables. In practice, GCMs are less integrated than the EMICs, owing to the smaller number of climate-system components that are directly coupled. They do, however, define the cluster of models that currently have the highest resolution and greatest number of explicitly represented physical processes among the several that we have described. GCMs play a major role in the assessment of potential future climate changes because they are able to provide spatially explicit simulations of climate under different scenarios of atmospheric trace gas and aerosol composition. GCMs have been applied in sequences of “time-slice” or snapshot simulations (i.e. Charbit et al., 2002; COHMAP Members, 1988; Valdes, 2000), with the goal of revealing the

Modeling Paleoclimates 571 mechanisms responsible for the regional patterns of climate change. Suites of models have been run under the same set of boundary conditions (i.e. Joussaume et al., 1999) to understand the role (if any) that the structures of individual models play in adequately simulating past climates. Both kinds of studies have featured comparisons with synthesis of paleoclimatic data (Kohlfeld & Harrison, 2000). Such comparisons have implicated the absence of feedbacks among components of the climate system in the mismatches between simulations and observations (Harrison et al., 2002). Regional climate models (RCMs) can be viewed as a subset of the GCMs, and have a relationship to regional or “fine-mesh” weather forecasting models that parallels that of GCMs to global forecasting models. Regional models require “lateral boundary conditions” – a temporal sequence of three-dimensional fields of variables describing the atmosphere and surface (i.e. SSTs) that are usually provided by a GCM simulation. Regional models can be considered to be “nested” within the lower-resolution global model in that the driving GCM fields are ingested by the RCM along the model’s boundary. Compared with GCMs, RCMs allow more spatially explicit simulation of the sensitivity of regional climate and subsystems to large-scale atmospheric controls (Hostetler & Bartlein, 1999). GCMs further allow evaluation of surface-atmosphere feedbacks, such as those associated with large lakes (Hostetler et al., 2000), that are not resolved in coarser-resolution GCMs.

and controls that by 1974 resulted in the National Research Council report Understanding Climatic Change, which is essentially modern in its scope and outlook. The principal materials related to climate modeling from the seventh congress include reviews by Broecker (1965) and Mitchell (1965) in the Wright & Frey (1965) “INQUA volume,” and a 1968 volume in the Meteorological Monographs series (Causes of Climatic Change; Vol. 8, Number 3), which was edited by Mitchell. Several other proceedings volumes were also published, but the majority of papers or chapters with climate modeling content are found in the volumes edited by Wright, Frey, and Mitchell. The Wright & Frey (1965) volume, The Quaternary of the United States, focused on the paleoclimate of the United States, and as a consequence it does not offer a comprehensive review of the fields of paleoclimatology or climate modeling of the time. It can be supplemented, however, by nearly contemporaneous books by Lamb (1966) and Budyko (1982, a summarization of his earlier work). With the organization of the Intergovernmental Panel on Climate Change in 1990, the publication of proceedings from NATO Advanced Study Institutes (e.g. Berger, 1980), and joint U.S./USSR syntheses (e.g. Porter, 1983; Velichko et al., 1984; Wright, 1983) and reports (MacCracken et al., 1990), the study of global climate change was shown to be truly global in perspective and participation.

Climate Modeling in the VII INQUA Congress Materials Comprehensive Models – ESMs GCMs are evolving toward greater comprehensiveness through the development of coupled models that include, for example, explicit representation circulation of the ocean (i.e. AOGCMs), and the terrestrial biosphere (AVGCMs). Coupled models will eventually fill in the region in model continuum now occupied mainly by the EMICs, and will extend into a presently unoccupied region of highly integrated models that incorporate many processes on a highresolution grid. However, because of computational limits on comprehensive-model simulations, models that feature coupling between a small number of climate-system components will be the rule for the near term. Ultimately, however, “superGCMs” (Saltzman, 2002), coupled with models of slowly varying components of the climate system, will form a true climate-system model (CSM), or Earth-system model (ESM).

VII INQUA Congress (1965) The seventh congress of the International Association for Quaternary Research marked a stage in understanding of climate variations in general, and paleoclimate modeling in particular. The congress took place at an interesting time, because both plate-tectonic theory and the astronomical (Milankovitch) theory of climate change emerged in their present forms and evolved over the following decade, as did the fuller depiction of the climate system and its variations

With the exception of EMICs, it is possible to see the same classes of models we described above both in the INQUA publications, and in journal articles not part of the formal proceedings but related to the scientific threads of the meeting.

Conceptual Models Mitchell’s review of the causes of climate change in the chapter “Theoretical Paleoclimatology” in the Wright & Frey (1965) volume represents a comprehensive listing of the conceptual models of Quaternary climate variations that were current at the time (see Mitchell’s Tables 1 and 2). The particular “causative factors” and the mechanisms through which they control climate reviewed include: (1) Autovariation, or internal variations of the climate system stemming from its highly nonlinear nature. (2) Air-sea interaction, including the role of the thermohaline circulation in transporting heat throughout the climate system. (3) Continental drift, which Mitchell viewed as an indirect cause of Quaternary glaciation. (4) Orogeny and continental uplift, and their potential effects on large-scale atmospheric circulation patterns. (5) Carbon-cycle variations, in which the potential for human action to have a significant impact on climate was discussed.

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(6) Volcanism, and the effect of dust and aerosols on incoming solar radiation. (7) Solar variability, including long-term and periodic variations of the solar constant. (8) Orbital variations, including the potential role of landsurface feedback in amplifying insolation changes. (9) Feedbacks, in which Mitchell reviewed a number of hypotheses that attempt to explain glacial/interglacial variations from combinations of internal variations and external forcing, including the one advanced by Chamberlin in 1899.

AGCMs might be used to investigate paleoclimatic questions. Although not explicitly part of the INQUA materials, the importance of contemporaneous work by Smagorinsky (1963), and Smagorinsky et al. (1965) to later work with GCMs is evident in subsequent publications. In the decade following the VII INQUA Congress, routine application of GCMs in paleoclimatic studies emerged (e.g. Gates, 1976; Williams et al., 1974).

Chamberlin’s hypothesis, which we would regard today as one expressed in terms of biogeochemical cycles, is particularly interesting in its consideration of the above controls (although not necessarily in modern terms), as well as the interactions among them. Interactions were referred to by Chamberlin as “intercurrent agencies,” an idea now usually described as “coupling” among systems. The only large gap in Mitchell’s list, though filled implicitly, is the potential role of land cover in controlling climate. Changes in land cover, on both the Quaternary and historical time scales, have the potential to influence significantly the emission of dust and mineral aerosols to the atmosphere (Harrison et al., 2001; Mahowald et al., 1999) and to change surface energy balances (DeFries et al., 2002) and consequently other components of the climate system (Chase et al., 2001).

It is apparent that many of the questions and issues that were raised in 1965 the INQUA Congress are still relevant today. Moreover, the specific contributions of the congress and its proceedings contributed to the foundation of the U.S. National Research Council (1974) report Understanding Climatic Change and its successors. These reports initiated research agendas for the study of global change (e.g. Malone et al., 1985) that remain relevant today (National Research Council, 1999).

Elemental Models Several elemental or low-resolution models were discussed in the INQUA proceedings. In the Wright & Frey (1965) volume, Broecker (1965) used a number of elemental models while reviewing the isotopic record of paleoclimatic variations. The chapter by Schumm (1965) included discussion of several elemental geomorphic and hydrologic models to examine the impact of climate changes while Meier (1965) presented an analysis of the response of glacier mass-balance and flow to variations in climatic controls. In the Meteorological Monographs volume, elemental models were used in the discussion of the thermohaline circulation (Weyl, 1968), and the surface energy balance (Ericksson, 1968). Saltzman (1968) considered the surface forcing of atmospheric circulation, and Kutzbach et al. (1968) examined the effects of changes in the latitudinal temperature gradient on atmospheric circulation. Both these latter studies examined with simple models the sensitivity of one component of global climate (atmospheric circulation) to changes in forcing in a manner that anticipates those authors’ later work.

Comprehensive Models GCMs appeared in a chapter in the Meteorological Monographs volume by Mintz (1968), which suggested how

Subsequent Developments

Synopsis of Results from Modeling Quaternary Paleoclimates of North America Climate models have been applied to advance understanding of many of the aspects of Quaternary climate changes in North America. The presence of the Laurentide Ice Sheet (LIS) makes climatic variations over North America a key component of the general description of long-term climatic variations. As a spatially heterogeneous region subject to the influence of the major northern hemisphere atmospheric circulation mechanisms, the patterns of regional climate changes across the continent have also been of interest. Paleoclimate modeling studies that have focused on North America fall into two general groups: those that focus on the slow-response components of the climate system like the LIS, and those that focus on the spatial patterns of the fast-response components at key times.

Temporal Variations of Climate Studies of the temporal variations of the climate system have addressed the onset of glaciation over the Cenozoic, as well as the nature of individual glacial cycles, and have also been used to examine the potential controls of glacial/interglacial variations and the genesis of millennial-scale variability. The long-term cooling during the Cenozoic, which ultimately led to Quaternary glacial/interglacial variations, and the higher-frequency variations superimposed on them, present several features or “targets” for which explanations have been attempted using various classes of models. These targets include: (1) the cooling itself, and the reorganization of the ocean, atmosphere, and cryosphere that is implied.

Modeling Paleoclimates 573 (2) the non-reversing steps toward more extensive glaciation, such as those around 35 myr ago, 12 myr ago; and within the past 5 myr. (3) the onset of extensive northern hemisphere glaciation around 2.8 myr ago. (4) the changes in periodicity and amplitude of global ice volume variations during the last 3 myr. (5) the sequence of global climate changes during a single glacial cycle. (6) “sub-millennial”-scale variations in climate.

Cenozoic Cooling and the Quaternary Ice Age The long-term cooling of the Cenozoic, leading ultimately to the onset of extensive glaciation in the northern hemisphere around 2.8 myr ago, has been examined using a variety of approaches, which have generally featured conceptual models supported by syntheses of data, elemental models of particular components of the climate system, or more comprehensive models used to simulate key times or to explore particular combinations of controls (using “snap-shot” simulations). Examples of the first application include the examinations of isotopic records by Miller et al. (1987) and Zachos et al. (2001) for the entire Cenozoic, or by Driscoll & Haug (1998) and Haug & Tiedemann (1998) for the past 5 myr. The transition to a more glacial state described by the latter two studies was also examined in simulations with Saltzman’s model of paleoclimatic dynamics (Saltzman, 2002; Saltzman & Verbitsky, 1993) and with the LLN 2-D model (Li et al., 1998), both of which are sophisticated elemental models. In another example of a general conceptual model supported by simulations with a more comprehensive model, sensitivity tests by Kutzbach et al. (1997) contributed to the evaluation of the tectonic hypotheses of Cenozoic climate change (Ruddiman, 1997). In a similar fashion, the role of the changes in paleogeography from Cretaceous times to present have been explored with coupled AOGCMs (Huber & Sloan, 2001; Otto-Bliesner et al., 2002), which account for the effects of changes in ocean basins and gateways on global climates. Further applications of GCMs and related models to pre-Quaternary climates are described by Parrish (1998). A number of studies have focused on the onset and maintenance of glacial/interglacial variations, again using a combination of modeling approaches. These include the conceptual (but mechanistic in character) model of Imbrie et al. (1992, 1993), and the aforementioned models of Imbrie & Imbrie (1980) and Paillard (1998). In these latter two applications, data analyses or relatively simple numerical models are used to illustrate the features of “thought experiments” that attempt to explain, for example, the features of the oxygen isotopic record. The inception of a single glaciation, as occurred around 115,000 yr ago, has been examined in several GCM-focused studies. Rind et al. (1989) found that the insolation changes between the time of the northern hemisphere summer maximum around 126,000 yr ago, and the relative minimum around 115,000 yr ago were insufficient to initiate permanent

snow cover in northeastern North America in their model. In contrast, subsequent simulations by Dong & Valdes (1995), Gallimore & Kutzbach (1996), and deNoblet et al. (1996) were able to simulate the accumulation of permanent snowfields, particularly if the models included feedback from climate-induced changes in land cover.

Millennial-Scale Variations Millennial-scale climate variations have also been examined with combinations of conceptual models, data analyses, and comprehensive models, in particular time-evolving elementary models and EMICs. The conceptual models, which include those described by Alley et al. (1999), Stocker et al. (2001) and Clark et al. (2002), have developed or attempted to test hypotheses for millennial-scale variations that generally involve reorganization of the circulation of the atmosphere and ocean, including the thermohaline circulation, and the global transmission or propagation of climate variations in the North Atlantic. The kinds of modeling studies used in the development and testing of those hypotheses span the entire continuum of model comprehensiveness. Saltzman (2002) (see also Saltzman & Verbitsky, 1995) showed how millennial-scale variability, like that associated with the Heinrich events, emerged from a dynamical model of the slowly varying components of the climate system. Similar variability emerges in simulations using EMICs (e.g. Crucifix et al., 2002; Ganopolski & Rahmstorf, 2001), which add some spatial specificity to the simulated climate variations. Simulations with GCMs and RCMs that examine the sensitivity of the climate and subsystems at the LGM to imposed changes in North Atlantic sea-surface temperatures reveal further details of the spatial patterns of millennial-scale variations (Hostetler & Bartlein, 1999; Hostetler et al., 1999). The integrated modeling studies and data analyses of the temporal variations of climate across the different timescales described above support make several generalizations about the temporal variations of the climate of North America leading up to and during the Quaternary: (1) changes in paleogeography, including changes in mountain belts and oceanic gateways, explain much of the pattern of climate change over the Cenozoic, if the synergistic effects of changes in atmospheric composition and ocean heat transport are considered. (2) the ice sheets are active components of the climate system, and no realistic account of the temporal and spatial patterns of Quaternary climate change can be made without considering them. (3) the thermohaline circulation of the ocean seems involved in climate variations across all time scales. (4) feedback from changes in the land surface and in ocean circulation appear to be involved in amplifying or attenuating the climatic effects of changing boundary conditions such as insolation and the arrangement of continents.

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Spatial Patterns of Fast-Response Variables – LGM to Present The LGM-to-Present “Natural Experiment” Various modeling studies have focused on the interval between the LGM (Last Glacial Maximum, 21,000 yr ago) and present. During this interval, nature performed experiments with the climate system (Webb & Kutzbach, 1998), and recorded the results in paleoclimatic data sets like those reviewed in this volume and elsewhere (Kohlfeld & Harrison, 2000). Both the nature of the boundary-condition changes over the interval and our knowledge of them has facilitated numerous application of elemental models, EMICs, and, particularly, GCMs. Comparing the climates of the LGM and 6000 yr ago with present provides an optimal experimental design in which only a few controls are changed from their present settings. At the LGM, there were extensive continental ice sheets, low concentrations of greenhouse gasses in the atmosphere, high aerosol loadings, relatively cold sea-surface temperatures, and land-cover characteristics that featured reduced areas of forest, but the latitudinal and seasonal distribution of insolation was similar to that at present. After the LGM, the amplitude of the seasonal cycle of insolation increased, reaching a maximum around 11,000 yr ago, so that at 6000 yr ago, insolation during the northern summer was greater than at present, while the remainder of the boundary conditions were close to their present (or pre-industrial) values. Because of the elegance of this natural experiment, many simulations have been done for 21,000 and 6000 yr ago, including those in PMIP (Harrison et al., 2002; Joussaume et al., 1999; Palaeoclimate Modeling Intercomparison Project). These two times have also been the focus for simulations with coupled AOGCMs (Braconnot et al., 2000; Harrison et al., 2002; Hewitt & Mitchell, 1998; Hewitt et al., 2001; Shin et al., 2002). Relatively few sequences of simulations with GCMs over this interval have been able to exploit the full natural experiment. The published sequences of experiments include simulations conducted with an early version of the NCAR CCM (National Center for Atmospheric Research, Community Climate Model) (COHMAP Members, 1988; Wright et al., 1993), a subsequent version of the CCM (CCM 1, Webb & Kutzbach, 1998), and with the UGAMP GCM (U.K. Universities Global Atmospheric Modelling Programme General Circulation Model; Valdes, 2000), and the LMD5.3 model (Laboratoire del Météorolog`ıe Dynamique; Charbit et al., 2002), models similar to CCM 1 in the degree of coupling among systems. These sequences of simulations, along with the suites of simulations for 6000 and 21,000 yr ago, jointly show that much of the variation in global and regional climate over this interval can broadly be explained by the influence of the ice sheets on atmospheric circulation and the influence of insolation on circulation and surface water- and energybalances. Comparisons of model simulations with paleodata demonstrate that, as is the practice of the Intergovernmental

Panel on Climate Change (Houghton, 2001), it is indeed feasible to simulate climates different from that at present using the kinds of models reviewed here. The detailed data-model comparisons organized by PMIP do show, however, that the present generation of global models may underestimate the magnitude of the responses of the climate system to changes in its controls (Harrison et al., 2002).

LGM-to-Present Simulations for North America The sequence of simulations conducted with CCM 1 (Fig. 3), although obsolete by today’s standards in terms of model resolution and interactivity of the ocean, still provide the only complete sequence of simulations performed using a coherent experimental design that have been extensively analyzed for North America (see Bartlein et al., 1998; Webb et al., 1998). Figure 3 shows the simulated sequence for several variables for January and July, expressed both as the actual values for each time in the sequence and as anomalies, or differences between each “paleo” simulation and the present-day or “control” simulation. Table 1 summarizes some of the main features in Fig. 3 that were discussed by Bartlein et al. (1998) and Webb et al. (1998). The variables in Fig. 3 appear roughly in a sequence that represents cause-and-effect. The top row of maps shows insolation (and anomalies) over the sequence of simulations, along with the outlines of the Laurentide ice sheet in the model – the two major controls of the sequence of regional climatic changes. Net radiation and surface air temperature illustrate the direct effects of insolation and ice-sheet size on the surface energy balance and temperature. The latitudinal and continental-marine contrasts in temperature, in concert with the topographic effects of the ice sheet, influence atmospheric circulation as represented by 500 mb horizontal wind speed and vertical velocity, and sea-level pressure. In turn, atmospheric circulation, in particular the large-scale patterns of vertical motions, and moisture availability (determined mainly by temperature) jointly influence the patterns of precipitation and thus precipitation minus evaporation. The principal features of the simulated climate over North America include: (1) displacement by the Laurentide Ice Sheet of the band of fast upper-level winds to the south of its present location in both winter and summer during the LGM and afterward. (2) development of a “glacial anticyclone” over the ice sheet in eastern North America, and consequent generation of large-scale sinking motions in the eastern and southern quadrants of the ice sheet. (3) existence of generally drier-than-present conditions during glacial times (when it was colder than present), giving way to wetter-than-present conditions as the continent warmed. (4) changes in the strength of surface atmospheric circulation features that follow the trends in the boundary conditions: weakening of the Aleutian low

Modeling Paleoclimates 575

Fig. 3. Sequence of CCM 1 simulations for North America from the LGM to present, for January (top) and July (bottom) (Bartlein et al., 1998; Kutzbach et al., 1998). The simulated values for each experiment are shown on the left-hand side of the figure, while the anomalies (“paleo” experiment minus present “control”) are shown on the right. For 500 mb vertical velocity, positive values (orange) on the left-hand panels indicate large-scale sinking motions, while negative values (blue) indicate large-scale rising motions. For 500 mb vertical velocity anomalies, negative values (blue) indicate more rising or less sinking than present, while positive values (orange) indicate more sinking or less rising than present.


Fig. 3. (Continued )

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Modeling Paleoclimates 577 Table 1. Features of the simulated climate of North America, LGM to present. Feature

CCM 1 Simulationa

Last Glacial Maximum 21,000 cal yr B.P. experiment Upper-level winds Southward displacement and strengthening of the jet stream in both January and July Sea-level pressure and surface winds Glacial anticyclone over ice sheet; prevailing easterlies in PNW, strong onshore flow in SW; strong Aleutian low in January, southerly flow into Alaska in January Large-scale vertical velocity Strong rising motions in SE in July, and in SW in January Net radiation Strong negative anomaly over ice sheet in July Atmospheric moisture Much drier than present throughout Surface climates Greatest cooling near ice sheet (T ∼ −10 ◦ C); less cooling farther away; generally dry to south of ice sheet; wet in SW Late-Glacial 14,000 cal yr B.P. experiment Upper-level winds Jan. and July jet stream at present latitude, and stronger than present Sea-level pressure and surface winds STHs in July as strong or stronger than present Large-scale vertical velocity Stronger-than-present rising motions in continental interior in January Net radiation Continued strong negative anomaly over ice sheet in July, with positive anomaly along southern edge of continent Atmospheric moisture Continued dryness Surface climates Continued cold near ice sheet; July temp. near or higher than present in SW U.S., SE U.S., and Alaska; Jan. temp. below or near present throughout Early Holocene 11,000 cal yr B.P. experiment Upper-level winds Upper-level circulation generally near present configuration; ST ridge over SW U.S. in July Sea-level pressure and surface winds Strong STHs in July; strong onshore flow into SW US Large-scale vertical velocity Rising motions in the SW and sinking in PNW and eastern North America in July Net radiation Continued strong negative anomaly over ice sheet in July, with strong positive anomaly over continental interior in July Atmospheric moisture Continued drier than present in Jan; wetter than present over much of continent in July Surface climates July temp. higher than present everywhere except along edge of ice sheet; Jan. temp. near present; dry in PNW and Alaska; wet in SW US Mid-Holocene 6000 cal yr B.P. experiment Upper-level winds Upper-level circulation generally near present configuration; ST ridge over SW U.S. in July Sea-level pressure and surface winds Strong STHs in July; strong onshore flow into SW U.S. and S US Large-scale vertical velocity Rising motions in continental interior and sinking in PNW in Jan., rising motions in the SW and sinking in PNW and eastern North America in July Net radiation Strong positive anomaly at high latitudes and negative in interior in Jan., strong positive anomaly throughout in July Atmospheric moisture Generally moister than present Surface climates Wetter than present in SW U.S. in July; warmer than present in July throughout a Abbreviations: cal (calendar years before present); PNW (Pacific Northwest); SW (Southwestern U.S.); SE (Southeastern U.S.);

STH (subtropical high-pressure system); ST (subtropical). and the glacial anticyclone as the ice sheet retreated, and strengthening of the East Pacific and Bermuda subtropical high pressure systems in summer as the (positive) insolation anomaly increased, followed by weakening as the insolation anomaly decreased. (5) increases in summer temperature earlier in the sequence in regions distant from the ice sheet. (6) development of a thermally induced low surface pressure over the continent in summer when the insolation anomaly was at its maximum, and conse-

quent enhancement of the summer monsoon in the southwestern U.S. (7) concurrent increases in effective moisture in the southwestern U.S. and decreases in the Pacific Northwest and continental interior when the monsoonal circulation was amplified during the time of summer insolation maximum (Harrison, 2003). (8) generally lower-than-present winter temperatures over the continent throughout the sequence of simulations.

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These responses are quite robust, appearing in most simulations or partial sequences of simulations, and suggest that a substantial part of the regional-scale patterns in paleoclimatic data that have been reviewed in this volume are explainable in terms of the direct and indirect effects of insolation and the direct effect of the ice sheet (see also Bartlein et al., 1998; Thompson et al., 1993; Webb et al., 1993, 1998). More recent simulations and their comparisons with paleoclimatic data have attempted to show how secondary climatic variations across different regions may be mechanistically linked. For example, simulations for 6000 yr ago show a decrease of surface low pressure and an increase in the height upper-level ridge over the western U.S. These changes induce large-scale subsidence (or sinking of air) in the interior of North America (Harrison et al., 2003), thereby linking climate anomalies that are opposite in sign across different regions.

Existing and Emerging Issues in Paleoclimatic Modeling Comparison of paleoclimatic observations with the conjectures of conceptual models or simulations from numerical models have long been part of the practice of Quaternary science. The rapid development of the different classes of models and syntheses of paleoclimatic data, presented either as maps for key times (e.g. Kohlfeld & Harrison, 2000), or time series at key locations (e.g. Alley & Clark, 1999), ensure that formal comparisons between simulations and observations in datamodel comparisons will continue to increase in frequency.

There are a number of issues that arise in such comparisons in specific and in paleoclimatic modeling in general that when addressed will enhance the effectiveness of those activities.

Model and Data Resolution One issue that frequently arises in data-model comparisons is disparity in the spatial and temporal resolutions of paleoclimatic data and model output. Spatial-resolution mismatches are probably most evident in the comparisons of EMIC and GCM output with networks or syntheses of paleoclimatic data. Most present-day GCMs (Fig. 4) have grid-cell resolutions coarser than 2 degrees of latitude or longitude (about 200 km at the equator) while EMICs are still coarser (10 degrees or more). Topography in EMICs and GCMs is highly generalized, and much smoother than the real terrain – topographically complex regions like the Cordillera may be represented as broad featureless domes in a GCM. RCMs mitigate this issue somewhat, and can represent features such as the Sierra Nevada, Cascade Range, or Columbia Basin, but resolutions in RCMs are still fairly coarse (grid cells greater than 25 km on a side). Paleoclimatic data, in contrast, is often site specific, and a given indicator may represent the environment of a watershed or of a smaller area. Consequently, some kind of “downscaling” of the simulations is necessary, even for RCMs, if the object is the direct comparison of observations and simulations at particular locations. The current approach (Harrison et al., 1998) is to apply the model’s “anomalies”

Fig. 4. Land-surface elevations for two climate models and for the western U.S. on a 5-min latitude by longitude grid for comparison. The NCEP Reanalysis Model (Kistler et al., 2001) is a GCM with a rectangular resolution of 2.5 degrees, and represents the topography of the western U.S. as a broad dome centered over western Colorado. Varying elevations over the Pacific Ocean are an artifact of the spectral representation of the atmosphere in the model. RegCM2, a regional climate model (e.g. Hostetler & Bartlein, 1999), as depicted here for the western U.S. has a resolution of 36 km. Viewed at arm’s length, the RegCM grid captures much of the physiographic detail for the region, such as the Sierra Nevada and Cascade Range, and the Snake River Plain and Central Valley, that is evident in the actual 5-min elevations.

Modeling Paleoclimates 579 (differences between a paleoclimatic and modern “control” simulation) to either a gridded higher-resolution observed data set for the present day, or to data for a specific location, both derived by spatial interpolation within a network of modern climate stations. This approach necessarily assumes that the models have sufficient resolution to describe adequately those anomalies, and that mediation of large-scale anomaly patterns does not change over time. Increases in computational capacity, which in turn will allow finer model resolutions, should ultimately minimize this issue. Simple spatial averaging of the paleoclimatic data to approximate the resolution of a model does not suffice to reduce the mismatch. A topographically complex region, for example, may vary from place to place in its paleoclimatic response to localized physiographic effects on atmospheric circulation (Whitlock & Bartlein, 1993). Paleoclimatic data may also be quite sparse in some regions, and simple interpolation among sites may not be appropriate (Broccoli & Marciniak, 1996). Mismatches in resolution also arise between time series of climate simulations and observations. Chronological and sampling issues that limit the “downcore” resolution of paleoclimatic data can make comparisons with model output of annual- (or shorter-) timescale resolution difficult. Conversely, there may be limitations in the temporal resolution of a model simulation related to an inability to specify changes in boundary conditions frequently enough. In both the spatial or temporal cases, some thought must therefore be given to placing the simulations and observations on a similar framework or timescale by appropriate filtering or aggregation. An example of such an approach involving comparisons among time series of data with differing resolutions is given by Tinner & Lotter (2001).

Variables Another potential mismatch in comparing simulations and observations arises when considering the variables that can be simulated or reconstructed. Climate models such as EMICs and GCMs can be quite specific in what they simulate, including variables (like atmospheric vertical motions) that cannot be directly observed. In contrast, a specific paleoclimatic record may require some kind of transformation in order to be interpreted in quantitative terms, and by themselves, the interpretations cannot discriminate among multiple controls. For example, the observation of a positive glacier mass balance can signify either increased winter precipitation or decreased summer temperature. This ambiguity can be removed, however, through the application of process models (e.g. Hostetler & Clark, 2000) that explicitly quantify the dependence of mass balance on a number of controls.

Experimental Design The design of a modeling experiment is an important issue that must be considered in comparing simulations and

observations. From the paleoclimatic data perspective, this includes the specific protocols that are used to synthesize data, and the scheme for describing chronological control (Kohlfeld & Harrison, 2000). It should be possible for the reader to trace the development of a particular interpretation or reconstruction (see Farrera et al., 1999, for an example; Harrison, 2003). From the perspective of a model, experimental design is for the most part synonymous with the choice of which boundary conditions are changed and by how much, though it also includes selection of the length of a simulation in the case of spatial-pattern applications of models. Recall that there are two approaches to the design of an experiment – the sensitivity test in which one or more boundary conditions are changed, and the full simulation in which all of the boundary conditions are changed (Peteet, 2001). In the sensitivity-test approach, leaving a boundary condition (say, land-surface cover, or the size and location of ice sheets) unchanged from its present value is the same as assuming that it does not vary over time, or if it does vary, that it does not have any influence on climate. The comparison of a sensitivity-test simulation with observations must therefore consider the extent to which the simulations should be expected to resemble the observations in the first place.

Paleoclimatic Diagnostics Although it is satisfying when a simulation agrees favorably with some observations, that situation may not be the most informative one – it is the mismatches that indicate something is wrong or needs improvement. There are three sources of apparent mismatches between simulations and observations: inadequacy of the climate model, misinterpretation of the data, and shortcomings in the experimental design, as discussed above. (Note that incorrect “false positive” comparisons could also arise from the same sources.) From a distance, it might be perceived that the goal in data-model comparisons is simply to discriminate among the three sources, but a better way of thinking of the general exercise is as paleoclimatic diagnostics. Analogous to its shorter-timescale cousin climate diagnostics (apart from its longer temporal focus), paleoclimate diagnostics could use almost the same description of its objectives, which are: “. . . to identify the nature and causes of climate variations on time scales ranging from a month to centuries . . . [and] to develop the ability to predict important climate variations on these time scales (NOAA Climate Diagnostics Center Web Page).” If we extend this definition to longer time scales, data syntheses and model simulations can be viewed as complementary tools that can be used to understand past climatic variations. The motivation for understanding past climatic variations is now stronger than ever, in light of the realization that humans may be producing Quaternary-size changes in climate. The full range of climate models is being used in making projections of future climate, and those models need testing, something that can be done using syntheses of paleoclimatic data and paleoclimatic “natural experiments.”

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Paleoenvironmental observations also demonstrate that no environmental system is completely insensitive to climatic variations, which raises questions about the magnitude of that sensitivity and how specific systems respond, questions that modeling approaches are well suited to answer. Together, climate models and paleoclimatic data are bringing scientists closer to the goal of understanding climate well enough to predict its future course.

Summary The synthesis of paleoclimatic data sets and the simulation of past climates using climate models are a complimentary set of activities that lead to better understanding of the climate system. The objective of paleoclimate modeling is to quantify the behavior and variations of the components that describe the climate system. These components include: (1) boundary conditions or external controls, such as solar irradiance. (2) slow-response variables that characterize the general state of the climate system, such as continental ice sheets. (3) fast-response variables that comprise what is ordinarily thought of as weather, such as the configuration of the jet stream and location and strength of surface high- and low-pressure centers. (4) variables that provide the forcing for the many environmental subsystems that depend on climate, such as lakes. Climate models can be classified according to the applications to which they are put, which include simulating the temporal evolution and spatial patterns of the climate system, and the attendant responses of environmental subsystems. They may also be classified by their comprehensiveness into several clusters, which include conceptual models, elemental models, Earth-system models of intermediate complexity, comprehensive models represented by coupled general circulation models, and ultimately, full Earth-system models. With one exception, the classes of climate models and the manner in which they are applied were evident in the publications of the INQUA Congress in 1965. Example simulations of the Quaternary climates of North America illustrate regional patterns of climate. These respond directly to continental ice sheets, and both directly and indirectly to changes in insolation.

Acknowledgments We thank Bryan Shuman, J.J. Shinker, and Debra Zahnle, and two anonymous reviewers for comments on the manuscript, and Sandy Harrison for discussion. Research was supported by the National Science Foundation (ATM-9910638) and by the U.S. Geological Survey.

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