Climate change: an integrated perspective

LINKING CLIMATE CHANGE TO LAND SURFACE CHANGE

ADVANCES IN GLOBAL CHANGE RESEARCH VOLUME 6

Editor-in-Chief Martin Beniston, Institute of Geography, University of Fribourg, Perolles, Switzerland

Editorial Advisory Board B. Allen-Diaz, Department ESPM-Ecosystem Sciences, University of California, Berkeley, CA, U.S.A. R.S. Bradley, Department of Geosciences, University of Massachusetts, Amherst, MA, U.S.A. W. Cramer, Department of Global Change and Natural Systems, Potsdam Institute for Climate Impact Research, Potsdam, Germany. H.F. Diaz, NOAA/ERL/CDC, Boulder, CO, U.S.A. S. Erkman, Institute for Communication and Analysis of Science and Technology – ICAST, Geneva, Switzerland. M. Lal, Centre for Atmospheric Sciences, Indian Institute of Technology, New Delhi, India. M.M. Verstraete, Space Applications Institute, EC Joint Research Centre, Ispra (VA)‚ Italy.

The titles in this series are listed at the end of this volume.

LINKING CLIMATE CHANGE TO LAND SURFACE CHANGE

Edited by

Sue J. McLaren and

Dominic R. Kniveton Department of Geography, University of Leicester, Leicester, England, U.K.

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

eBook ISBN: Print ISBN:

0-306-48086-7 0-7923-6638-7

©2003 Kluwer Academic Publishers New York, Boston, Dordrecht, London, Moscow Print ©2000 Kluwer Academic Publishers Dordrecht All rights reserved No part of this eBook may be reproduced or transmitted in any form or by any means, electronic, mechanical, recording, or otherwise, without written consent from the Publisher Created in the United States of America Visit Kluwer Online at: and Kluwer's eBookstore at:

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TABLE OF CONTENTS

Table of contents

v

Preface

vii

Contributing Authors

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SECTION A: SHORT-TERM CLIMATE VARIABILITY

Chapter 1 Brooks, N. and Legrand, M. Dust variability over Northern Africa and rainfall in the Sahel

1

Chapter 2 Agnew, C. T. and Chappell, A. Desiccation in the Sahel

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Chapter 3 Yair, A. and Bryan, R. B. Hydrological response of desert margins to climate change: The Effect of Changing Surface Properties

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Chapter 4 Viles, H. and Goudie, A. H. Weathering, geomorphology and climatic variability in the Central Namib Desert

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Chapter 5 Adegoke, J. O. and Carleton, A. M. Warm season land surface-climate interactions in the United States Midwest from mesoscale observations

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Chapter 6 Wilby, R. L. and Dettinger, M.D. Streamflow changes in the Sierra Nevada, California, simulated using a statistically downscaled General Circulation Model scenario of climate change

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Chapter 7 Schmidt, M. and Dehn, M. Examining links between climate change and landslide activity using GCMS: Case Studies from Italy and New Zealand

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SECTION B: LONG-TERM CLIMATE VARIABILITY

Chapter 8 Bachhuber, F. W. and Catto, N. R. Geologic evidence of rapid, multiple and high magnitude climate change during the last glacial (Wisconsinan) of North America

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Chapter 9 Catto, N. R. and Bachhuber, F. W. Aeolian geomorphie response to climate change: an example from the Estancia valley, Central New Mexico, U.S.A.

171

Chapter 10 White, K., McLaren, Black, S. and Parker, A. Evaporite minerals and organic horizons in sedimentary sequences in the Libyan Fezzan: implications for palaeoenvironmental reconstruction

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Chapter 11 Gurney, S. D. Relict cryogenic mounds in the UK as evidence of climate Change

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Chapter 12 Burgess, P. E. ,Palutikof, J. P. and Goodess, C. M. Investigations into Long-Term Future Climate Changes

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SECTION C: SUMMARY

Chapter 13 Kniveton, D. and McLaren, S. 247 Geomorphological and climatological perspectives on land surface – climate change

Index

261

PREFACE The relationships that exist between changes in climate and land surface change are topical issues, but research and collaboration between researchers from the different disciplines of climatology, geomorphology and Quaternary Sciences, is often hampered by the different approaches; the incompatibility of scales of involvement (both spatial and temporal) of the various models used; and by differences of interest in such topics as mean values for climatic parameters and the probabilities of extreme events. In terms of approaches there are those researchers who have tried to model past, present and future climatic changes, and there are people who have used proxy data (such as sediments and landforms) to reconstruct past climates. Only relatively recently have attempts been made to integrate the two distinct approaches. In order to improve our understanding of the relationships that exist between changing climates and land surfaces, a number of factors need to be considered including: - the spatial and temporal scales of climate variability and geomorphological change; the impacts of climate change on various landforms; the modification of climate by surface processes; modelling climate change on a global scale as well as downscaling of such model outputs so that they are applicable on regional scales; and prediction and management of land surface changes as a result of future climate changes. These factors will be discussed further in Chapter 13. To understand how climate is likely to change in the future, it is necessary to have an understanding of how climate has changed in the past in order to identify any underlying trends in natural climatic change. Many of the studies that use various proxies to make interpretations of past environmental conditions from landforms and other land surface features, as well as the small scale recent process-based research all need to be placed in a larger framework to aid our understanding of global climate change. Palaeoreconstructions are needed to provide evidence of past changes, to help in the comprehension of the responses of terrestrial surfaces and to help validate predictive models of climate change. Present day studies rely on the processes of observation, measurement (using both field work and analysis of remotely sensed images) as well as modelling. This book by no means attempts to be a summary of the main research on looking at the relationship between climate change and land surface change, but rather gives a selection of papers that show some of the different approaches that have been

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undertaken to address the many issues and to highlight the importance of multidisciplinary research over different timescales (from 101 to 106 years) and from the scale of local catchment studies to global processes. Recent technical advances in techniques such as absolute dating; geochemical analyses, remote sensing and climate modelling have aided these studies. The book stems from a one-day conference held at the Royal Geographical Society with the Institute of British Geographers (R.G.S with I.B.G.) Annual Conference held in Leicester on January 5th 1999. The symposium was jointly organised by the Association of British Climatologists (A.B.C.) and the British Geomorphological Research Group (B.G.R.G.), and was organised by Sue McLaren, Dominic Kniveton and John McClatchey. The selection of peer-reviewed papers included in this book address a wide range of issues ranging from looking at long-term climate changes through modelling (Burgess et al), palaeoenvironmental reconstructions (e.g. White et al, Catto & Bachhuber, Bachhuber & Catto) through to evidence of short-term climatic variability (e.g. Adegoke & Carleton, Brooks & Legrand, and Yair & Bryan) and attempts to downscale from General Circulation Models (GCM’s) to allow modelling of regional-scale patterns of climatic change and the effects on various surface and geomorphological processes (e.g. Wilby & Dettinger and Schmidt & Dehn). The chapters show just a small selection of the wide-ranging nature of research currently being undertaken in the general area of climate change and terrestrial surface processes. The main division of papers has been made in terms of the spatial and temporal scales of the studies rather than between climatology and geomorphology because the editors of the book wish to stress the importance of trying to link these two areas. The final chapter develops the main themes of the preceding chapters in the context of the wider field of scientific literature. The authors hope that this book makes an early attempt to present some recent advances in understanding the linkages between climates and land surfaces in order to further our ability to predict environmental change. The success of the conference and the production of the book were as a result of many people. We would like to thank the A.B.C., the B.G.R.G. and the R.G.S. (with the I.B.G.) for providing funds for guest speakers to attend the conference. We are grateful to John McClatchey (Nene), Dave Thomas (Sheffield), Alan Werritty (Dundee) and Norm Catto (Newfoundland), for acting as Chairpersons at the symposium. In terms of the preparation of the book we would like to thank all the contributors (especially for

PREFACE

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meeting all the deadlines); the many reviewers; Susan Draycott, Ruth Pollington and Kate Moore for help with printing and cartography; and to Mariette Ph de Jong and Astrid Zandee who approached the authors with the offer of publishing the book with Kluwer Academic Publishers.

SUE McLAREN DOMINIC KNIVETON Department of Geography, University Of Leicester, Leicester LE1 7RH

CONTRIBUTING AUTHORS: JIMMY ADEGOKE: Department of Geography and Earth System Science Center, The Pennsylvania State University, University Park PA 16802, U.S.A. CLIVE AGNEW: Department of Geography, University College London, 26 Bedford Way, London WC1H OAP U.K. FRED BACHHUBER: University of Nevada, Las Vegas, Las Vegas, NV, USA, 89154-4010. STUART BLACK: Postgraduate Research Institute for Sedimentology, The University of Reading, Whiteknights, Reading, RG6 6AB, U.K. NICK BROOKS: Climatic Research Unit, School of Environmental Sciences, University of East Anglia, Norwich, NR4 7TJ, U.K. RORKE BRYAN: Faculty of Forestry, The University of Toronto, Toronto, Canada. PAUL BURGESS: Climatic Research Unit, University of East Anglia, Norwich NR4 7TJ, U.K. ANDREW CARLETON: Department of Geography and Earth System Science Center, The Pennsylvania State University, University Park PA 16802, U.S.A. NORM CATTO: Memorial University of Newfoundland, St. John’s, Canada, A1B 3X9 ADRIAN CHAPPELL: Telford Institute of Environmental Systems, Department of Geography, University of Salford, Manchester, M5 4WT U.K. MARTIN DEHN: Dept. of Geography, University of Bonn, Meckenheimer Allee 166, D-53115 Bonn, Germany MICHAEL DETTINGER: U.S. Geological Survey, Water Resources Division, California, Scripps Institution of Oceanography, 9500 Gilman Drive, La Jolla, California, 92093-0224 CLARE GOODESS: Climatic Research Unit, University of East Anglia, Norwich NR4 7TJ, UK

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ANDREW GOUDIE: School of Geography, University of Oxford, Mansfield Road, Oxford OX1 3TB DOMINIC KNIVETON: Department of Geography, University of Leicester, University Road, Leicester LE1 7RH MICHEL LEGRAND: Laboratoire d’Optique Atmosphérique Université de Sciences et Technologies de Lille-1, F59655 Villeneuve d’Ascq cedex, France. SUE McLAREN: Department of Geography, University of Leicester, University Road, Leicester, LE1 7RH, U.K. JAN PALUTIKOF: Climatic Research Unit, University of East Anglia, Norwich NR4 7TJ, U.K. ADRIAN PARKER: Geography Department, Oxford Brookes University, Gipsy Lane Campus, Headington, Oxford, OX3 0BP, U.K. MICHAEL SCHMIDT: Dept. of Geography, University of Bonn, Meckenheimer Allee 166, D-53115 Bonn, Germany HEATHER VILES: School of Geography, University of Oxford, Mansfield Road, Oxford OX1 3TB, U.K. KEVIN WHITE: Landscape and Landform Research Group, Department of Geography, The University of Reading, Whiteknights, Reading, RG6 6AB, U.K. ROBERT L. WILBY: Division of Geography, University of Derby, Kedleston Road, Derby, DE22 1GB, UK. National Center for Atmospheric Research Boulder, Colorado, 80307-3000, USA AARON YAIR: Department of Geography, The Hebrew University, Jerusalem, Israel.

DUST VARIABILITY OVER NORTHERN AFRICA AND RAINFALL IN THE SAHEL NICK BROOKS Climatic Research Unit School of Environmental Sciences, University of East Anglia, Norwich NR4 7TJ, U.K. MICHEL LEGRAND Laboratoire d’Optique Atmosphérique Université de Sciences et Technologies de Lille-1, F59655 Villeneuve d’Ascq cedex, France.

Abstract The Infra-Red Difference Dust Index (IDDI) is a new dataset that uses reductions in atmospheric brightness temperature (derived from METEOSAT IR-channel measurements) to map the distribution of mineral aerosols over continental Africa. The IDDI dataset is described, and the IDDI data are used to identify the major African dust sources, located in the Sahel-Sahara zone. The seasonal variations in these sources are discussed. Annual, seasonal and monthly dust indices are constructed from the IDDI data for different latitudinal zones in the Sahel-Sahara zone. The temporal and spatial variability of dust production in the Sahel and Sahara is inferred from these indices and the latitudes of maximum dust production are identified. Interannual variability of dust production is described in conjunction with a consideration of variations in annual rainfall over the Sahel. Relationships between rainfall and subsequent dust production in the Sahel are investigated by correlating zonally averaged rainfall and IDDI values at lags of one and two years. The spatial and temporal patterns of dust production suggest that spring and summer deflation is associated with the passage of convective disturbances across the Sahel. There is evidence that wet-season rainfall totals have an impact on dust production in the later part of the following dry season. The results also suggest a cumulative impact of rainfall on December dust production. However, there is no evidence from this study that dust production is associated with widespread land degradation. KEY WORDS: dust, rainfall, Sahel, Sahara, variability 1

S.J. McLaren and D.R. Kniveton (eds.), Linking Climate Change to Land Surface Change, 1–25. © 2000 Kluwer Academic Publishers. Printed in the Netherlands.

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Introduction

The Sahel is the semi-arid transition zone between the Sahara desert and humid equatorial Africa. It is characterised by a steep north-south temperature gradient and high interannual rainfall variability. The timeseries of spatially aggregated rainfall anomalies for the Sahel (Figure 1) suggests that the region has experienced a desiccation since the late 1960s. Rainfall has been below the regional twentieth century mean for most years since 1968. Large rainfall deficits in 1972 and 1973 contributed to famine in the Sahel, and the largest rainfall deficit this century was associated with the Ethiopian famine of 1984. In both of these cases the impact of drought was exacerbated by other factors.

West African visibility data indicate that levels of atmospheric dust over the Sahel throughout the year have increased dramatically since the 1950s, and it has been suggested that dust loadings over the Sahel now exceed those over the Sahara (N’Tchayi et al., 1994, 1997). Middleton (1985) found an increase in dust storm activity in certain parts of the Sahel during drought years. Prospero and Nees (1986) reported elevated dust concentrations in the atmosphere over the North Atlantic after the deficient wet seasons of the early 1970s. More recently, Tegen and Fung (1995) and Tegen et al. (1996) have suggested that 30-70% of the global mineral aerosol budget is the result of deflation from soils which have been degraded by climate change and/or human activity. They invoke human activity in the Sahel, and a climatic shift in the boundary

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3

between the Sahel and Sahara, as major factors in determining the global atmospheric dust budget. These studies have resulted in the widely held opinion that dust production in northern Africa has largely shifted from the Sahara to the Sahel as a result of climatic desiccation and inappropriate land-use practices. Until now, it has been difficult to assess such assumptions using observational data as such data have been somewhat limited in spatial extent. However, a new proxy dust-loading dataset for continental Africa now exists, based on METEOSAT infra-red channel measurements. This dataset is known as the Infra-Red Difference Dust Index (IDDI). While the IDDI detects any aerosols which reduce the infra-red radiance at the top of the atmosphere, it may be interpreted in terms of dust concentrations over the arid and semi-arid regions of northern Africa, where mineral dust is the dominant atmospheric aerosol. The IDDI dataset has been used in a preliminary investigation of spatial and temporal dust variability over the Sahel-Sahara zone of northern Africa (i.e. Africa north of the Equator). This paper presents results detailing the spatial and temporal variability of atmospheric dust loadings for the period 1984-1993. Spatial variability and seasonality are addressed via a visual analysis of dust/IDDI fields. A more quantitative presentation of seasonality and meridional variation in dust production is achieved by plotting mean monthly IDDI values, spatially averaged over different latitudinal zones, against time. A qualitative interpretation of dust variability in response to rainfall is presented, followed by a discussion of correlations between wet-season rainfall and subsequent dust loadings as represented by zonally averaged IDDI values. The short length of the IDDI time series means that many of the conclusions are speculative. However, a consideration of the results within the context of existing knowledge enables a plausible conceptual model of rainfall influences on dust production to be constructed. This study concentrates on the aerosol signal in the IDDI fields over the Sahel-Sahara zone, because of the recent changes in observed dust concentrations and also because this region contains the major African dust sources. We may also be confident that signals in the IDDI fields over the arid and semi-arid regions of northern Africa are the result of the episodic transport of dust (see below). However, IDDI signals over other parts of Africa are also discussed where appropriate. Possible explanations for the presence of strong signals in the IDDI data where dust is unlikely to be a major atmospheric constituent are presented.

2.

The Infra-Red Difference Dust Index

The IDDI dataset has been developed at the Laboratoire d’Optique Atmosphérique at the Université des Sciences et Technologies de Lille, France (Legrand et al., 1994). IDDI data represent the reduction in the measured infra-red (IR) brightness temperature (BT) of the atmosphere from that which would result from an aerosol-free atmosphere. Brightness temperature values are derived from METEOSAT IR-channel radiometric count measurements taken daily at approximately 11:30 UTC. Fields of maximum brightness temperature over non-overlapping 15-day periods are constructed. Fields of differences between these composite fields and daily brightness temperature fields

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within each 15-day period are then calculated. The resulting difference fields are divided into 10x10-pixel boxes and a statistical algorithm based on the spatial coherence method (Legrand et al., 1994) is used to classify pixels as cloudy or non-cloudy. Cloudy pixels are assigned a cloud-masking code, and the remaining pixels represent the IDDI values, where brightness temperature reductions are due to the presence of aerosols alone. The IDDI signal results from the reduction in the temperature of the underlying land surface by reduced solar insolation (resulting in less emitted IR radiation), and also from the attenuation of the outgoing longwave radiation (OLR) by the aerosol layer. Attenuation of OLR will be greatest when the aerosol particles have effective diameters of the same order of magnitude as the wavelength of the radiation, i.e. of the order of 10 Sub-micron particles are transparent in the infra-red (Maley, 1982). Theoretical considerations and recent, as yet unpublished, modelling studies (Legrand, pers. comm.) indicate that, in the case of mineral aerosols, small dust particles cause the greatest reduction in daytime temperatures, while coarse dust causes the greatest daytime reduction in IR radiance at the top of the atmosphere (TOA). The strongest signals in the IDDI will therefore result from dust events with a high proportion of large particles, although events comprised of small particles in high concentrations will be detected due to the reduction in emitted IR radiation from the cooler surface. The IDDI data are converted to a 1° latitude x 1° longitude geographical grid, and exist over land regions only. The geographical coverage extends from 35° south to 38° north and 18° west to 45° east, covering all of Africa and parts of the Middle East (see Figure 2). The dataset will be updated to the present day in the near future. The IDDI data have been validated against ground-based visibility and aerosol optical depth (AOD) measurements at a number of sites throughout West Africa (Legrand et al., 1994). During these validation studies, it was found that IDDI values correlated well with near-surface visibilities. IDDI values of 5 K and above corresponded to dusty conditions, when visibility was reduced below 10 km, and values of 10 K and above corresponded to severely dusty conditions, with visibility reduced below 5 km. IDDI images have also been compared with fields of AOD over the eastern tropical Atlantic in order to verify continuity across the West African coast. Nonetheless, there are several potential pitfalls to be considered when interpreting the IDDI data. The detection of aerosols depends on the variability in their concentration. If concentrations are generally elevated over the whole of the 15-day reference period, they will be interpreted as part of the “clear-sky” background, reducing the BT values of the reference field. Long-term dust haze is therefore unlikely to be detected. A similar problem may occur over regions which are covered by cloud throughout the reference period. Long-term cloud cover will result in misleading reference values, and may also affect the efficiency of the cloud-detection algorithm, leading to the erroneous identification of cloud as IDDI (i.e. aerosol) data. Over very cloudy regions such as those near the Equator, the IDDI data may be unreliable due to this “cloud

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contamination”. Problems may also arise where low, relatively warm, clouds are present; these may be identified as aerosols, resulting in large IDDI values. Also to be considered is the presence of aerosols resulting from biomass burning, which is widespread throughout much of Africa in the dry season. Such aerosols typically have dimensions of less than (Artaxo et al, 1994); they will have some impact on the OLR, but their dominant effect will be one of cooling of the land surface. These particles should therefore have a similar effect on the measured TOA radiance to fine dust aerosols. However, because of the extent of burning, they may constitute a constant smoke haze lasting for periods of days to weeks, resulting in their not being detected in the IDDI fields, but rather being incorporated into the reference fields. The above considerations notwithstanding, the IDDI data represent a useful semiquantitative measure of dust loadings over the arid and semi-arid regions of Africa. Over the Sahara and Sahel, dust events are highly episodic and contain high proportions of aerosols large enough to strongly attenuate the OLR, resulting in strong IDDI signals. The incidence of cloud over these regions is low enough to present no significant problems of cloud contamination. The issues of biomass burning aerosols and fine dust haze are discussed in more detail below, although these features do not appear to inhibit the detection of episodic dust events over the main regions of interest in this study, which lie north of 10° N. To date, IDDI fields over the Sahel and Sahara have not been converted to AOD values, and cannot be interpreted in terms of specific volumes of dust or thicknesses of dust layers. The reduction in brightness temperature due to dust aerosols will depend on the vertical distribution of the dust, the particle density and the particle size distribution, as well as the reflective properties of the underlying surface. Nonetheless, IDDI fields reliably reflect the distribution and abundance of atmospheric mineral aerosols over northern Africa, and exhibit a sufficient degree of spatial and temporal invariance to be used in studies of large-scale dust mobilisation and transport (Legrand et al., 1994).

3.

Distribution of Saharan and Sahelian dust sources

It may be assumed that dust concentrations and particle sizes will be greatest closest to dust source regions. Fields of IDDI data may therefore be employed to identify the major source regions throughout Africa. Use of different averaging periods enables the temporal variation in the activity of dust sources to be analysed. The major dust sources in northern Africa have been identified in this fashion by Legrand et al. (1994). This section elaborates on their description, within the context of other studies of dust sources and climatological considerations of known or likely dust mobilisation processes. Discussion of the major dust sources is restricted to northern Africa, focussing on the Sahelian and Saharan zones. Monthly mean IDDI fields were created by averaging daily IDDI fields for cells where fewer than eighty per cent of days were classed as cloudy. Over most of the Sahel-

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Sahara zone, where cloud is scarce, this approach results in continuous spatial coverage in the monthly fields. Annual mean IDDI fields were created for each year by averaging the monthly mean fields over twelve-month periods. Seasonal mean fields were created by averaging the monthly fields over shorter periods for each year. Mean annual, seasonal and monthly fields were created by averaging the yearly fields over the period 1984-1993. The mean annual IDDI field for 1984-1993 is shown in Figure 2.

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Three broad regions in which IDDI values exceed 5 K are apparent. This threshold is arbitrary, but delineates distinct zones within which mean dust levels are elevated above the background. Further detail within these zones is apparent in the form of areas with IDDI values in excess of 5.5 K. These regions are interpreted as coinciding broadly with areas containing dust sources. One such region is the north-central Sahel, between about 5° E and 20° E, and 13° N and 18° N. Two maxima are apparent within this region, centred approximately at 16° E., 17° N and 9° E, 15° N. The former maximum extends over parts of the Erg of Bilma and the alluvial plain northwest of the town of Largeau in Chad. This region has been identified as an important dust source by other authors (e.g. McTainsh, 1980; Drees et al., 1993). The latter maximum lies to the south of the Aïr Mountains in Niger, in the vicinity of a region of enhanced generation of convective disturbances (Rowell and Mitford, 1992) which result in spring and summer dust mobilisation (Dubief, 1979; McTainsh, 1996). A second source region (or collection of sources), which may be labelled the West Sahara region, lies between about 7°-0° W and 20°-25° N. This area corresponds to a region that includes the Erg Iguidi and Erg Chech of northern Mali, northern Mauritania and southwestern Algeria. A nearby maximum in the IDDI field lies over a region of seasonal watercourses in the Morocco-Western Sahara border region. Dust transported large distances over the Atlantic and to Europe has been identified as originating in these regions (Reiff et al., 1986; Coudé Gaussen et al., 1987; Chiapello et al., 1997). The third major source region extends from about 13° N to 25° N, and some 1° to 3° either side of the 30° E meridian, from northern Sudan into southern Egypt. Hereafter this is referred to as the East Sahel-Sahara region. This region is characterised by the Haboob dust storms of the Nile Valley (McTainsh, 1996), and dust from the northeastern Sudan has been transported to the eastern Mediterranean (Middleton, 1986, 1997). A minor region of activity is indicated by high IDDI values over a small area centred on 14° E, 22.5° N, between the Plateau de Djado in northern Niger and the Idhan Murzuq erg in southwestern Libya. This region is hereafter referred to as the northern Niger region. All the source regions identified above are characterised by fields of sand dunes or seasonal watercourses, or both. This suggests that erodible material is supplied by dune fields or by water erosion, or a combination of the two. The source region near the Aïr Mountains extends into the zone of degraded soils as suggested by UNEP (1992), suggesting that land use and climatic desiccation of soils may be partly responsible for deflation in this area. However, the region is dominated by numerous water channels and few permanent human settlements, suggesting that water erosion is an important factor in providing erodible material.

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Also present in the annual field are strong IDDI signals over the Horn of Africa, westcentral Africa and southeastern Africa. The two former regions exhibit IDDI values as high or higher than the highest Sahelo-Saharan values. Dust transport over the Horn of Africa is associated with the Asian Monsoon circulation in summer (Husar et al., 1997). The west-central African signal is unlikely to be due to dust aerosols, while the reason for the southeastern African signal is open to debate. The high IDDI values over these three regions are discussed further in Section 4, within the context of seasonal changes in the regional environment. 3.1. DUST SOURCES AND LAND DEGRADATION The northern limit of the region characterised by land degradation is placed in the region of 17° N on soil degradation maps published by UNEP (1992). However, estimates of the extent of soil degradation in the Sahel are extremely unreliable and subjective (Warren, 1996; Williams and Balling, 1996). In the absence of reliable soil degradation data it is impossible to identify new dust sources arising from land-use practices or climatic desiccation, or to quantify the contribution of disturbed soils to the regional dust budget. However, soil degradation is likely to be minimal in regions of low rainfall and outside of the zone of rainfed agriculture, the limit of which is placed at the location of the 300 mm isohyet by WMO (1976). Fields of annual rainfall totals derived from the dataset of New et al., (1999, not shown) indicate that the 300 mm isohyet lies to the south of 17° N. These considerations suggest that the 17° N latitude represents a reasonable and liberal (if somewhat arbitrary) working limit for the zone containing degraded soils. This limit will be employed when the role of soil-state in dust production is considered in Sections 5-8. Examination of the mean annual IDDI field suggests that the major dust source regions in the Sahel and Sahara conform to the accepted, or “classical”, sources of dust, created by “natural” processes of sediment production and deflation. A possible exception is the source region in the north-central Sahel in the vicinity of the Aïr Mountains. It is possible that material from anthropogenically degraded soils does not produce a strong signal in the IDDI data, resulting in an underestimation of the extent of the major source regions. Aerosols from degraded soils are likely to be very different in nature from those deflated from arid to hyper-arid desert regions. Dust consisting of such aerosols will contain more organic material and have a higher clay content, resulting in a high proportion of small aerosol particles (McTainsh and Walker, 1982). Organic material has been detected in dust deposited in Niger (Drees et al., 1993) and northern Nigeria (McTainsh and Walker, 1982). However, it is not clear whether the organic input is due to the long-term desiccation of vegetated areas or if it is a long-term feature of the soil-dust cycle. As previously discussed, the IDDI signal from dust with a low mean particle size will be predominantly the result of surface cooling. McTainsh and Walker (1982) report a tendency for lower visibility and reduced solar radiation to be associated with finer mean particle sizes. The correlation of IDDI values with measured visibilities (Legrand et al., 1994) suggests that the IDDI are capable of

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detecting such fine material. It is possible that such fine material from degraded soils exists as a semi-permanent dust haze throughout much of the year, in which case it would not be detected by the IDDI for the reasons outlined above. However, it is reasonable to suppose that episodic dust events would originate over such degraded land in the same fashion as over other regions, and as a result of the same atmospheric processes. This would be particularly true outside of the wet season, when the Sahara and Sahel are both subject to the Harmattan circulation. The lack of a regional signal in the IDDI data over the hypothesised regions of widespread land degradation (the vicinity of the Aïr Mountains notwithstanding) therefore calls into question the assumption that aerosols from degraded soils contribute significantly to the regional dust budget, and the budget of material exported from northern Africa (Tegen and Fung, 1995).

4.

Seasonal variations in dust production and non-dust IDDI signals

Seasonally averaged fields of IDDI are presented in Figure 3. Again a threshold of 5 K delineates broad regions of dust activity, with further detail apparent in the form of IDDI values in excess of 6 K. While this analysis focuses on northern Africa, the structure of the seasonal IDDI fields in southern, eastern and central Africa is also discussed where appropriate. In JFM the most active areas are the East Sahel-Sahara, the north-central Sahel and the northern Niger regions. A broad shift in dust activity from east of 5° E in JFM to west of 15° E in AMJ is apparent. The East Sahel-Sahara sources remain active in AMJ, although the geographical extent of IDDI values greater than 6 K is reduced. AMJ IDDI values are high over southern Morocco and western Algeria, and also in the western part of the north-central Sahara. JAS represents the peak of the Sahelian wet season, when the surface discontinuity between the West Africa Monsoon airmass and the dry Saharan airmass lies at its northernmost limit, around 20° N in August (Hastenrath, 1991). IDDI values greater than 6 K are confined to the west of 5° E and between 17° N and 25° N. The southern limit of this zone is very distinct; the 4 K/5 K boundary occurs close to the 5 K/6 K boundary at approximately the same latitude from the West African coast to 7° E. This suggests that large dust loadings are prevented from occurring south of the northern limit of the monsoon rains, which extend to within several hundred kilometres south of the surface discontinuity (Hastenrath, 1991). The southern latitudes of this region coincide with an area identified by Rowell and Milford (1992) as a region of enhanced generation of convective disturbances or disturbance lines (DLs), encompassing the plains to the north of the Niger Bend.

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Another major feature of the JAS field is the region of very high IDDI values over the Horn of Africa in JAS. These values are considerably higher than the maximum values over the Sahel and Sahara. This signal over the Horn of Africa coincides with very high equivalent aerosol optical thickness (EAOT) measurements over the Arabian Sea (Husar et al., 1997 – based on data from July 1989 to June 1991). The parts of Arabia visible in the IDDI fields exhibit low IDDI values, suggesting that dust transport over the Arabian Sea is predominantly from the Horn of Africa (Sirocko and Sarnthein, 1991). Mobilisation and transport of dust is aided by the East African (or Somali) low-level jet,

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which is active at this time of year as part of the summer monsoonal circulation (Hastenrath, 1991, Husar et al., 1997). Transport of dust over large distances occurs above the monsoon inversion, in a fashion analogous to the transport of Saharan dust above the trade wind inversion in the Saharan air layer (Kalu, 1979; Sirocko and Sarnthein, 1989). The OND field exhibits low IDDI values except in a small region within the north-central Sahel zone and another over northern Niger. Examination of the mean monthly fields (not shown) illustrates that dust loadings are lowest in November, and that the high-IDDI regions in the OND field are due to the “switching on” of sources in these regions in December. The major IDDI signals outside of the regions discussed above are detailed and interpreted below. 4.1. THE GUINEA COAST In JAS a zone of relatively high IDDI values exists over the Guinea Coast region, extending in places to some 12° N and exhibiting a maximum in the east over Nigeria. The period JAS corresponds to the “Little Dry Season” (Barry and Chorley, 1995) in this region and it might therefore be expected that widespread biomass burning would be prevalent. Monthly maps of fire distribution are available for some years from the World Fire Atlas, compiled by the European Space Agency and the European Space Research Institute (ESA/ESRIN) as part of the Ionia programme (Arino and Melinotte, 1995; Arino et al., 1997). These maps have been produced from AVHRR and ATSR satellite data. A visual comparison of the monthly IDDI fields with monthly fire maps for 1993 suggests that the JAS high IDDI values over the Guinea Coast are not due to combustion products, as fires are almost entirely absent from this region in this period according to the fire maps. At this time of year detectable fires are concentrated between the Equator and 20°S, where IDDI values are low. Strong fire signals in the ESA/ESRIN data occur over and to the east of the Guinea Coast throughout the winter, with fires being most widespread in January. Again, the regions of high IDDI values do not correspond to those characterised by fires; the January 1993 IDDI field exhibits low values over the Guinea Coast. However, the relationship between the distributions of fires as detected by satellite remote sensing methods, and high concentrations of biomass burning aerosol products is not necessarily straightforward. Fires will only be detected if they exist under relatively clear-sky conditions. Both clouds and high concentrations of airborne combustion products will obscure the ground from satellite detectors operating in the visible part of the electromagnetic spectrum. Thus, fires that produce large quantities of aerosols may not be detected. It is plausible that material from such fires is responsible for some of the high IDDI signals apparent in figures 5.4 to 5.6, providing at least a partial explanation for the summer Guinea Coast signal. Another plausible explanation for the high IDDI values over the Guinea Coast in summer is that dust is transported from the Sahel-Sahara to a zone of relatively stagnant air over this region, where it remains in the atmosphere for some time. Between the Guinea Coast and the Sahel-Sahara transition zone, dust will be removed from the atmosphere by rainfall, resulting in short residence times, low aerosol concentrations

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and hence low IDDI values. A further possibility is that of cloud contamination arising from persistent cloudy conditions throughout the periods used to create the reference fields. This is most likely over Nigeria, where the highest regional IDDI values exist in the vicinity of a region of frequent cloud cover. 4.2. WEST-CENTRAL AFRICA Large quantities of combustion aerosols also provide a plausible explanation for high IDDI values over regions where dust is unlikely to be a major feature of the atmosphere. Such values are seen over west central Africa (stretching from Gabon to the Democratic Republic of Congo and southwards over Angola) in all the fields, and are greatest in OND, JFM. Some biomass burning occurs in this region in these periods, particularly in October (based on 1993 data from ESA/ESRIN). However, the frequency and density of fires during the periods in question is far greater between 0° and 15° N, where IDDI values remain low. Again, these discrepancies between the IDDI and fire data may be due to the complex relationship between fire and smoke aerosol distributions. This region is adjacent to a region of frequent cloud cover in JFM and OND (i.e. southern hemisphere spring and summer), when the IDDI values are highest. It is possible that some cloud contamination occurs in these periods. 4.3. EASTERN AFRICA High IDDI values also occur over many of the eastern coastal regions of Africa south of the Equator, particularly in AMJ and JAS. These regions contain no extensive deserts, but do include semi-arid and dry sub-humid zones. The boreal summer high IDDI signal occurs during the dry season in East Africa. It is possible that dust mobilisation occurs from disturbed soils in these regions, although a complex biomass burning aerosol signal is again highly plausible, as burning is widespread in the dry season. Cloud contamination is likely in JFM and OND, but during AMJ and JAS the elevated IDDI values exist well away from areas of frequent cloud cover. 4.4. SOUTHERN AFRICA Finally it is worth mentioning the southern hemisphere African deserts in terms of dust sources as defined by the IDDI data. These regions do not stand out in the seasonal or annual fields, although elevated IDDI values are apparent over the Kalahari in JFM. It is striking that the Namib Desert does not appear to be a significant source of dust. The cold Benguela Current to the immediate west of the desert results in a highly stable atmosphere that is not conducive to the generation of the type of large convective events that are responsible for dust mobilisation and transport in northern Africa. While dust storms do occur over the sandy desert in the Namibian interior, it appears that the spatial and time scales associated with these events are such that they do not produce a major signal in the mean IDDI fields. The coastal atmosphere is very different from that over West Africa, and it is likely that the atmospheric environment over the Namib desert is such that dust aerosols are not carried to the elevations necessary for long-range transport. Middleton (1997) states that dust mobilisation and transport from the southern

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African deserts is poorly understood, but suggests that the scale of such phenomena is not comparable with that which characterises the northern African regions. 4.5. SUMMARY The above discussion further illustrates some of the caveats to be considered when interpreting the IDDI data. The question of whether the IDDI is a reliable means of detecting combustion aerosols remains open, and will only be resolved when the relationship between detected fires and the nature and distribution in the atmosphere of their products is better understood. It also appears that the IDDI is less reliable under persistently cloudy conditions. Further work is required to decouple the effects of biomass burning products and cloud contamination from the impacts of dust on the IDDI signal. However, over the regions of interest in this study, the IDDI appears to perform well, exhibiting cumulative signals from large dust events and identifying the major sources of dust aerosols. It may therefore be used with confidence in studies of Saharan and Sahelian aerosols and their relationships with the regional climate. Seasonal and geographical variations in the IDDI data may also be used to infer information concerning the behaviour of the major aerosol sources in northern Africa.

5.

Meridional variation in dust production

In order to assess the seasonal variation in dust production in northern Africa in a more quantitative fashion, several different zones were defined. These zones are the aggregated Sahel (10° - 20 ° N), the aggregated Sahara (20° - 30° N), the South Sahel (10° - 15° N), the North Sahel (15° - 20° N), the South Sahara (20° - 25° N), the North Sahara (25° - 30° N), the zone from 15° - 17° N and the zone from 18° - 20° N. The last two zones are used to examine dust seasonality either side of the suggested limit of soil degradation (Section 3.1). Spatially aggregated, mean monthly IDDI values over the zones described above (Figure 4) illustrate a broad commonality of dust loadings over the Sahel-Sahara region. Values are generally high in the first half of the calendar year, falling to a minimum in October or November and rising again in December. However, the evolution of the North Sahara zone departs from that of the other zones outside of March-June. This is to be expected as a result of the influence of mid-latitude weather systems such as Mediterranean and Atlantic cyclones.

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From June to September, dust levels are higher over the Sahel than over the Sahara. Sahel dust loadings peak in June; Saharan dust loadings are at a maximum in March and April. The lowest dust levels occur in November over the Sahel, and in October over the Sahara. IDDI values are consistently higher over the North Sahel than over the South Sahel, and the North Sahel exhibits the highest values of all the 5°-latitude zones in December and January and from June to September. The North Sahel contains the transition zone between the Sahel and Sahara and the nominal northern geographical limit of soil degradation. The 15°-17° N band lies to the south of this limit, so variations of IDDI within this band may be interpreted as reflecting variability of dust production from potentially disturbed soils, with a component due to advection from zones to the north, particularly during the dry-season. IDDI values in the 18°-20° N band may be assumed to reflect variability of dust production from undisturbed soils. However, the uncertainties in the estimates of the extent of soil degradation (Section 3.1) should be recalled. The 15°-17° N band yields the larger IDDI signal from December to March and in May and June. (The April value is similar to that in the 18°-20° N. band.) This indicates that the meridional maximum in dust loadings lies in the 15°-17° N band in December, January and June, when the maximum values in the 5° latitude zones occur over the North Sahel. Similarly, dust loadings are highest in the 18°-20° N band from July to September. (These results are unchanged if other 2°-latitude bands within the North Sahel are considered.) During the summer the 2°-latitude bands exhibiting the highest IDDI values lie to the south of the average position of the surface discontinuity (Tetzlaff and Peters, 1988; Hastenrath, 1991), i.e. within the monsoonal air mass. It is arguable

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that these high IDDI values represent advected material from the Sahara overlying the monsoon air. However, if this is the case, still higher IDDI values should be apparent closer to the northerly source regions. Therefore, these meridional maxima in IDDI values may be interpreted as representing meridional maxima in dust levels resulting from dust mobilisation in the shallow northern part of the monsoon air layer. Thus dust mobilisation is at a maximum within the zone containing potentially degraded soils in the early to mid dry season and in the early phase of the wet season. Mobilisation may remain high in this zone in JAS, but rainfall will remove dust from the atmosphere, shifting the maximum in the IDDI signal to the northern fringes of the active rainfall zone. It is likely that the June maximum in the 15°-17° N band is due to the intensity of the deflation processes and the balance between dust mobilisation and removal, rather than the sensitivity of the soils to deflation. In June this band corresponds to the northernmost extent of the wedge of monsoonal air (Tetzlaff and Peters, 1988), where the convective disturbances that generate rainfall and mobilise dust are weak due to the small thickness of the monsoon air layer (Hastenrath, 1991). Such weak disturbances may be sufficient to cause deflation, but too weak to produce significant amounts of precipitation. Thus the June maximum may be simply a manifestation of the regional climatology. The same processes are likely to be responsible for deflation in the 18°-20° N band in JAS. In December and January both the Sahel and Sahara are subject to the regional-scale Harmattan circulation, characterised by northeasterly winds over most of northern Africa (McTainsh, 1996). Deflation processes are therefore associated with large-scale atmospheric circulation patterns, suggesting that dust mobilisation will be greatest where soils are most vulnerable. The December-January maximum in dust production between 15° and 17° N is therefore likely to represent a meridional maximum in the availability of erodible material. This may be due to the fragility of degraded soils in this region, or a maximum in water erosion arising from the action of rainfall and rainfall-runoff on semi-arid surfaces. Low vegetation cover may also play a role; it is likely that the combination of relatively high rainfall (when compared with the dry Sahara) and the lack of vegetation protection of the land surface together result in high water erosion rates. Degraded soils will be more susceptible to water erosion, but it is not necessary to invoke land degradation in order to explain this maximum in dust production. 6.

Interannual variability of dust and rainfall

Figure 5 shows rainfall anomalies for the period 1983-1994, standardised with respect to the 1983-1984 mean. This period represents the period over which IDDI data are available, and includes 1983 in order to show all years that may affect dust values at a lag of +1 year.

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Figure 6 shows yearly, spatially averaged annual IDDI anomalies calculated over four different periods, for the various latitudinal zones described in the previous section. The primary objective of such a representation is to illuminate interannual variability of atmospheric dust loadings over bands subject to different rainfall regimes. While the main zones of interest are those in the Sahel, values for Saharan zones are included so that rainfall-dominated regions may be compared with arid regions. The annual period represents the mean IDDI values over the period November-October, chosen to commence around the beginning of the dry season. The wet-season is liberally defined as the period May-October, during which deflation mechanisms are most likely to be associated with the westward travelling disturbance lines (DLs), which bring the majority of rainfall to the Sahel (Rowell and Milford, 1992). The early dry-season is defined as November-December, the part of the dry season in which the vegetation cover as represented by NDVI values is significantly greater than the dry-season minimum (Hess et al., 1996). The late dry-season is defined as January-April, during which vegetation cover is close to the dry-season minimum, and in which dust mobilisation and transport in both the Sahel and Sahara are subject to the Harmattan circulation (Adeyfa and Holmgren, 1996).

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In many cases the anomalies over one region reflect those over the other regions, suggesting a common atmospheric influence on the primary deflation mechanisms. Major differences between Sahelian and Saharan regions are likely to be due to the influence of rainfall in the Sahel. The impact of the severe 1984 drought is evident in annual, early dry-season and late dry-season anomalies in 1984/85, 1984 and 1985 respectively. Over the Sahel the anomalies for these years are large and positive. Over the Sahara these anomalies are small or negative. Rainfall influences therefore serve to decouple the Saharan and Sahelian dust signals. 6.1. ANNUAL ANOMALIES The three driest years in the Sahel in the 1984-93 period were 1984, 1987 and 1990 (Figure 5). The annual periods following these wet-seasons exhibit the largest positive IDDI anomalies in the South Sahel series (Figure 6). The largest-magnitude negative anomalies in the Sahel occur after the wet-seasons of 1985, 1989 and 1991. These IDDI anomalies occur after dry or intermediate-rainfall years. This pattern of large negative IDDI anomalies is also reflected in the South Sahara, suggesting that atmospheric influences (for example a low frequency of strong surface winds) may be partly responsible for these periods of low dust loadings. 6.2. WET SEASON ANOMALIES The largest positive wet-season IDDI anomalies in the South Sahel occur in 1988, 1989 and 1991, the wettest years in the 1984-93 period. This further supports the hypothesis that DLs (which are more frequent and intense in wet years) are largely responsible for dust mobilisation in the wet-season. For these three relatively wet years, IDDI anomaly magnitude decreases with increasing rainfall. While three years do not represent sufficient data to constitute a trend, this result suggests the possibility that summer dust loadings may be generally higher in wetter years but that, above a certain rainfall threshold, dust levels decline as rainfall increases. This is physically plausible: intense DLs will mobilise greater quantities of dust than weak DLs, but will also produce more rainfall, which will remove dust from the atmosphere. Thus spring/summer Sahel atmospheric dust loadings are likely to be controlled by two processes that act in opposition to each other. The relative strengths of these processes will depend on the frequency and intensity of the DLs in any given wet-season. This conceptual model has important implications for the identification of the mechanisms behind the observed increases in Sahelian dust production. Lamb et al. (1998) report a decrease in both the frequency and intensity of DLs over the Sahel since the onset of dry conditions in the late 1960s. Enhanced spring and summer dust loadings over the Sahel may therefore be the result of a change in the balance between processes controlling the mobilisation and removal of dust particles, rather than, or in addition to, changes in soil properties. In the North Sahel and South Sahara the dustiest wet-seasons occur in 1987, 1988 and 1991. Dust mobilisation in these regions is likely to be related to DL activity within the vicinity of the surface discontinuity, where the monsoon air layer is not thick enough to

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allow rainfall generation. Rowell and Milford (1992) have identified August DLs generated as far north as 20° N. Wet-season dust levels are lowest in 1985, 1986 and 1990 (South Sahel) or 1989 (North Sahel). All of these years are dry except 1989, which follows the wet year of 1988. Low levels of dust in dry years may be explained by weak or infrequent DLs. The low level of dust in 1989, a relatively wet year (Figure 5), suggests that there was not much material available for deflation in this year. This may be due to the removal of such material after heavy water erosion in 1988 and/or a recovery in the vegetation cover of the Sahel over 1988 and 1989. An alternative explanation is that dust levels are high in wet-seasons dominated by weak DLs (which do not produce much rainfall) and low in wet-seasons dominated by intense DLs. If the rainfall in 1989 was the result of a predominance of the latter, removal of dust by rainfall may have dominated over dust mobilisation. 6.3. EARLY DRY SEASON ANOMALIES The most striking aspect of the November-December anomalies is the switch from positive anomalies in the Sahel until 1988 to negative anomalies from 1989 onwards. This pattern is punctuated by small positive anomalies in the South Sahel in 1985 and 1986 and a small positive anomaly in the North Sahel in 1992. A single year of drought may not have a long-term impact on soil or vegetation (Bullard, 1997). However, several consecutive years of drought, as occurred in the early-mid 1980s, are likely to have a cumulative impact on vegetation and hence on the organic matrix of the soil, leading to loss of soil cohesion. It is suggested that the wetter conditions prevailing from 1988 onwards led to a recovery in soil cohesion by encouraging vegetation cover, which would result in a greater degree of protection of Sahelian soils from deflation (Bullard, 1997). The dry year of 1990 occurred in isolation, and would not have had a long-term impact on soil properties. The positive IDDI anomaly following the 1988 wet-season is probably due to water erosion caused by the action of heavy summer rainfall on soils with little vegetation cover (either because of the distribution of rainfall or due to the dying off of vegetation under the previous dry conditions). In the short-term this would lead to an increase in the amount of erodible material (Baird, 1997). The anomaly series for the Saharan regions do not closely reflect those for the Sahelian regions, further reinforcing the interpretation that dust production in this period is largely a function of earlier rainfall. However, the large negative IDDI anomaly in 1989 is apparent in all the series except that representing the South Sahel (where the anomaly is negative but not of great magnitude), suggesting that the regional-scale circulation also modulates dust production in this period. The positive IDDI anomaly of 1988 is also not confined to the Sahel, suggesting a possible atmospheric influence on dust levels throughout the region.

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6.4. LATE DRY SEASON ANOMALIES The Sahel exhibits lower interannual variability in dust loadings over the January-April period than over the other periods described here. As rainfall exhibits considerable interannual variability, these results suggest that rainfall generally has a small impact on January-April dust production. However, a very large positive IDDI anomaly is evident in all Sahelian zones in 1985, after the extremely dry years of 1983 and 1984, providing compelling evidence for a cumulative impact of multiple years of large rainfall deficits. 1987 and 1990 are years characterised by extreme rainfall deficits that follow years that are dry, but not extreme in terms of rainfall. 1987 and 1990 are not followed by large positive late dry season IDDI anomalies. It is speculated that the small 1988 IDDI anomalies in the Sahelian regions may be due to the lack of rainfall and the consequent reduction in erodible material produced by water erosion.

7.

Rainfall-dust correlations

Rainfall over the May-October period and monthly mean IDDI values were spatially averaged over the zones defined in Section 5. The resulting timeseries, representing aggregated dust and rainfall values over spatially coincident areas, were correlated. Correlations were performed between rainfall and monthly IDDI values representing twelve months commencing in the November immediately following the wet-season (lag = +1 year), and between rainfall and IDDI values representing twelve months commencing in November of the following year, i.e. thirteen months after the end of the wet-season (lag = +2 years). The results were tested for statistical significance using a simple monte-carlo style randomisation procedure. For each correlated pair, one of the timeseries was randomised 10,000 times and the two series correlated for each randomisation. If the original correlation was exceeded fewer than 500 out of 10,000 times the result was deemed to be significant at the 5 per cent level. Correlations at the 1 per cent level were also noted. Correlations not significant at the 5 per cent level were rejected. Correlations were calculated for Saharan zones for purposes of comparison: significant relationships would not be expected over Saharan regions where rainfall is low and infrequent. The rainfall averaging period is arbitrary in the case of Saharan rainfall, further reducing the likelihood of meaningful statistical dust-rainfall relationships over Saharan regions. If such relationships were found, they would suggest that statistically significant results over the both the Sahara and the Sahel were artefacts of the statistical procedure employed. Comparisons with the Sahara notwithstanding, the short length of the timeseries means that the resulting correlations should not be interpreted as demonstrating definite physical relationships between rainfall and dust loadings. Nevertheless, considerations of the probable mechanisms of dust production provide a conceptual context within which such correlations may be interpreted. Significant correlations may therefore be used to infer likely impacts of rainfall on dust production, as well as the temporal

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distribution of such lagged relationships. Such an approach is useful in reinforcing or rejecting existing hypotheses, and suggesting new hypotheses, of dust variability. 7.1. LAG 1-YEAR RELATIONSHIPS No significant correlations result from the lag 1-year analysis for the Saharan zones. This is encouraging as it suggests that the Sahel correlations described below are not merely coincidental results arising from an analysis of short time series.

Significant negative correlations at a lag of one year were found for all the Sahelian zones in March, and for the South Sahel in April (Table 1). The strongest apparent relationships occurred in March over the aggregated Sahel, the North Sahel and the 15°17° N band. These results suggest that what variability there is in dust production throughout the Sahel in March is significantly influenced by the previous year’s rainfall. March falls within the period characterised by large dust loadings and low dust variability, so the proportion of the March dust production that results from the influence of rainfall on the soil-state is likely to be relatively small. Significant positive correlations occur in October over the aggregated Sahel and the South Sahel. This result is difficult to explain. If it is physically meaningful, it may be due to rainfall-driven soil erosion in one year sensitising the soil to the particular deflation mechanisms operating in the following October. These mechanisms are likely to be related to DL activity at the end of the wet-season. Such DLs may be strong enough to mobilise dust but too weak to produce much rainfall. The same deflation mechanisms will operate throughout the wet-season, but removal of dust from the atmosphere by rainfall will result in a weak IDDI signal, masking the relationship between soil-state and dust mobilisation. This conceptual model is likely to be appropriate only in an arid regime where soils are fragile and susceptible to rainfall

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erosion. The higher correlation in the wetter South Sahel therefore suggests that advection from the more arid northern zones may be responsible for this relationship. This assumes that rainfall variability is coherent between the South Sahel and the more northerly regions, as the correlation is the result of consideration of South Sahel rainfall only. This interpretation is highly speculative. 7.2. LAG 2-YEAR RELATIONSHIPS Significant negative correlations for the 2-year lagged timeseries are observed in December over the North Sahel and over the two narrower bands lying within the North Sahel (Table 2). The North Sahel signal gives rise to a smaller significant negative correlation over the aggregated Sahel. These results suggest that rainfall variability has a cumulative impact on December soil properties and hence on dust production in the northern latitudes of the Sahel, where rainfall is low and where some regions may be characterised by soil degradation (UNEP, 1992).

A similar relationship is suggested for the North Sahel in September. This may represent the impact of desiccation-related soil degradation on dust production. However, this result is not reflected in the correlations for the 15°-17° N and 18°-20° N bands. Also of note is the fact that a significant positive correlation is observed over the North Sahara for August. This signal results in a smaller significant correlation for the aggregated Sahara. For such short time series, any physical interpretation of this isolated significant Saharan result would be wildly speculative. It is highly plausible that it is a physically meaningless artefact of the statistics. Hence the isolated September result for the North Sahel should also be treated with caution. A positive correlation for May over the North Sahel may reflect a multi-year sensitising of soils to deflation by rainfall erosion (as suggested for the lag 1-year October results), or may also be spurious.

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Discussion and Conclusions

The IDDI data enable the major dust sources to be identified, and the seasonal evolution of these sources to be described. Dust sources are identified with regions of sandy desert and regions of seasonal watercourses. The distribution of airborne dust in summer is closely associated with the position of the surface discontinuity between the monsoon and Saharan air masses, and suggests a role for convective disturbance lines (DLs) in summer dust production. The role of DLs in dust production is further supported by the northward migration of the meridional dust maximum over the Sahel in the summer months. The monthly zonally averaged IDDI values indicate that dust production in the SahelSahara zone is at a maximum between 15° and 17° N in December, January and June, and between 18° and 20° N from July to September. Dust levels over the Sahara exceed those over the Sahel in much of the dry Season, and mean zonally averaged dust values between 20° and 25° N (Sahara) exceed those between 10° and 15° N (Sahel) in all months except May and June. The zonal maximum in dust production is therefore located in Sahelian latitudes only during part of the year, and in the zone of potential land degradation for only three months of the year. In June, this maximum is likely to be the result of the balance between dust mobilisation and wet deposition as determined by the prevailing meteorology. In December and January, the strongest IDDI signals occur over the accepted natural dust sources located in the north-central Sahel. While material from disturbed soils may contribute to the dust budget in these months, the notion that such processes have created major new source regions in areas not previously associated with dust production, and extending throughout much of the Sahel, is not supported by this study. It is possible that the maximum in December/January dust production within the 15°-17° N zone may be the result of a meridional maximum in the generation of deflatable material by water erosion, arising from the balance between rainfall intensity (greater than in more northerly regions) and vegetation cover (less extensive than in more southerly regions). It should also be noted that mean dust concentrations are relatively low in December, and that dust activity in northern Africa as a whole is greater throughout much of the year than in January. This is particularly so in regions near the West African coast. It is therefore unlikely that dust production from disturbed soils in these two months makes a large contribution to the regional and global annual dust budget. Rainfall-dust correlations indicate that enhanced dust production in April and May is associated with reduced rainfall in the previous year. Limited evidence for a cumulative impact of drought on December dust production is provided by correlations between rainfall and IDDI values at a lag of two years. April falls within the period during which interannual variability in Sahelian dust concentrations is low and dust events in northern Africa are frequent. The component of the April dust budget associated with rainfall in

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preceding years is therefore likely to represent relatively small percentage changes in the quantities of dust mobilised. The results of this study suggest that rainfall does exert some influence on dust production during certain parts of the year, although rainfall does not appear to be the dominant factor in determining the amount of dust mobilised in the Sahel on interannual timescales. The research described here does not support the notion that dust-event frequency in the Sahel has increased as a result of widespread land degradation, nor that the Sahel has become a more important source of mineral aerosols than the Sahara. It appears that the role of the land surface (and particularly of human activity) in modulating atmospheric dust concentrations has been over-emphasised, while too little attention has been paid to the role of meterological processes in determining the regional dust budget. In particular, observed changes in the nature of summer rainbearing disturbances may have played a key role in decadal-scale changes in the amount of dust mobilised within, and exported from, Sahelian regions.

Acknowledgements This work was completed as part of a PhD project funded by the Climatic Research Unit in the form of the Hubert Lamb Studentship, and supervised by Dr Mike Hulme. IDDI data were obtained from the Laboratoire d’Optique Atmosphérique, Université de Sciences et Technologies de Lille. Rainfall data were provided by Mark New at the Climatic Research Unit. Thanks are also extended to an anonymous reviewer for their comments.

References Adeyfa, Z. D. and Holmgren, B. (1996) Spectral solar irradiance before and during a Harmattan dust spell, Solar Energy 57 (3), 195-203. Arino O., Melinotte, J-M. Rosaz, J. M. and Monjoux, E. (1997) ESA Fire Product, Proceedings of the 7th ISPRS conference on Physical Measurement and Signatures in Remote Sensing, 7-11 April, Courchevel. Arino O. and Melinotte, J-M. (1995) Fire Index Atlas, Earth Observation Quarterly 50, T.D. Guyenne (ed.), ESA Publications Division, ESA/ESTEC, Keplerplaan 12200 AG, Noordwijk, The Netherlands. Artaxo, P., Gerab, F., Yamasoe, M. A. and Martins, J. V. (1994) Fine mode aerosols compositions at 3 longterm atmospheric monitoring sites in the Amazon Basin, Journal of Geophysical Research - Atmospheres 99 D11, 22857-22868. Baird, A. J. (1997) Overland flow generation and sediment mobilisation by water, in Thomas, D. S. G. (Ed.) Arid zone geomorphology: Process, form and change in drylands, 2nd Edition, John Wiley and Sons Ltd, 165-184. Barrey, R. G. and Chorley, R. J. (1995) Atmosphere, weather and Climate, 6th Edition, Routledge, p. 259. Bullard, J. E. (1997) Vegetation and geomorphology, in Thomas, D. S. G. (Ed.) Arid zone geomorphology: Process, form and change in drylands, 2nd Edition, John Wiley and Sons Ltd, 109-131. Chiapello, I., Bergametti, G., Gomes, L., Chatenet, B., Dulac, F., Pimenta, J., Santos Suares, E. (1995). An additional low layer transport of Sahelian and Saharan dust over the North-Eastern Tropical Atlantic, Geophysical Research Letters 22, 3191-3194. Coudé-Gaussen, G., Rognon, P., Bergametti, G., Gomes, L., Strauss, B, Gros, J. M., Le Coustumer, M. N. (1987) Saharan dust on Fuertaventura Island: Chemical and mineralogical characteristics, air mass trajectories, and probable sources, Journal of Geophysical Research 92, 9753-9771.

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Drees, L. R., Manu, A. and Wilding, L. P. (1993) Characteristics of aeolian dusts in Niger, West Africa, Geoderma 59, 213-233. Dubief, J. (1979) Review of the North African climate with particular emphasis on the production of eolian dust in the Sahel zone and in the Sahara, in Morales, C. (Ed.) Saharan Dust, John Wiley and Sons Ltd. Hastenrath, S. (1991) Climate Dynamics of the Tropics, Kluwer Academic Publishers, Dordrecht. Hess, T., Stephens, W. and Thomas, G. (1996) Modelling NDVI from decadal rainfall data in the north east arid zone of Nigeria, Journal of Environmental Management 48, 249-261. Husar, R. B., Prospero, J. M. and Stowe, L. L. (1997) Characterisation of tropospheric aerosols over the oceans with the NOAA advanced very high resolution radiometer optical thickness operational product, Journal of Geophysical Research 102 D14, 16,889-16909. Kalu, A. E. (1979) The African dust plume: its characteristics and propagation across West Africa in winter, in Saharan Dust: mobilisation, transport, deposition: papers and recommendations from a workshop held in Gothenburg, Sweden, 25-28 April 1977, C. Morales (ed.), Wiley. Lamb, P. J., Bell, M. A. Finch J. D. (1998) Variability of Sahelian disturbance lines and rainfall during 19511987, Water Resources Variability in Africa during the XXth Century (Proceedings of the Abidjan ’98 Conference held at Abidjan, Cote d’lvoire, November 1998). IAHS Publications No. 252. Legrand, M., N'Doume, C. and Jankowiak, I. (1994) Satellite-derived climatology of the Saharan aerosol, Passive Infrared remote sensing of clouds and the atmosphere II, D. K. Lynch (ed.), Proc. SPIE 2309, 127-135. McTainsh, G. H. (1980) Harmattan dust deposition in northern Nigeria, Nature 286, 587-588. McTainsh, G. (1996) Dust concentrations and particle-size characteristics of an intense dust haze event: inland delta region, Mali, West Africa, Atmospheric Environment 30, 1081-1090. Maley, J. (1982) Dust, clouds, rain types, and climatic variations in tropical North Africa, Quaternary Research 18, 1-16. Middleton, N. J. (1985) Effect of drought on dust production in the Sahel, Nature 316, 431 -434. Middleton, N. J. (1997) Desert dust, in Thomas, D. S. G. (Ed.) Arid Zone Geomorphology, Wiley. New, M. G., Hulme, M. and Jones, P. D. (1999) Representing 20th century space-time climate variability. II: Development of a 1901-1996 monthly terrestrial climate fields. Journal of Climate, in press. N’Tchayi, G. M., Bertrand, J., Legrand, M. and Baudet, J. (1994) Temporal and spatial variations of the atmospheric dust loading throughout West Africa over the last thirty years, Annales Geophysicae 12, 265273. N’Tchayi Mbourou, G., Bertrand, J. J., Nicholson, S. (1997) The diurnal and seasonal cycles of wind-borne dust over Africa north of the equator, Journal of Applied Meteorology 36, 868-882. Prospero, J. M. and Nees, R. T. (1986) The Impact of the North African Drought and El-Niño on Mineral Dust in the Barbados Trade Winds, Nature 320, 735-738. Reiff, J. Forbes, S., Spieksma, T.Th. M., Reynders, J. J. (1986) African dust reaching northwestern Europe: A case study to verify trajectory calculations, Journal of Climate and Applied Meteorology 25, 1543-1567. Rowell, D. P. and Milford, J. R. (1993) On the generation of African squall lines, Journal of Climate 6, 11811193. Sirocko, F., Sarnthein. M., Lange, H. and Erlenkeuser, H. (1991) Atmospheric summer circulation and coastal upwelling in the Arabian Sea during the Holocene and the last Glaciation, Quaternary Research 36, 7293. Tegen, I. and Fung, I. (1995) Contribution to the atmospheric mineral aerosol load from land surface modification, Journal of Geophysical Research, 100, 18,707-18,726. Tegen, I., Lacis, A. A. and Fung, I. (1996) The influence on climate forcing of mineral aerosols from disturbed soils, Nature 380, 419-422. Tetzlaff, G. and Peters, M. (1988) A composite study of early summer squall lines and their environment over West Africa, Meteorology and Atmospheric Physics 38, 153-163. UNEP (1992) World Atlas of Desertification, N. Middleton and D. S. G. Thomas (Eds.), Edward Arnold, London. Warren, A. (1996) Desertification, in Adams, W. M., Goudie, A. S. and Orme, A. R. (Eds.) The Physical Geography of Africa, pp 343-355. Williams, M. A. J. and Balling, R. C. (1996) Interactions of Desertification and Climate, WMO, UNEP, Arnold, London, pp 25-28. WMO (1976) Special environmental report No. 9: An evaluation of climate and water resources for development of agriculture in the Sudano-Sahelian zone of West Africa, WMO – No. 459.

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DESICCATION IN THE SAHEL C. T. AGNEW Department of Geography, University College London, 26 Bedford Way, London WC1H OAP UK ([email protected]) A. CHAPPELL Telford Institute of Environmental Systems, Department of Geography, University of Salford, Manchester, M5 4WT UK ([email protected])

Abstract The Sahel region of West Africa is well known as a region of environmental degradation. The reported incidence of drought and desertification has been challenged but regional desiccation is still widely accepted. This paper investigates the evidence for regional desiccation and in particular the effect of aggregating rainfall statistics across the area. Regression analysis reveals that the recent regional downward trend in rainfall is not reproduced at all stations at the 1% level of significance but is significant when data is aggregated. Geostatistical methods were used to investigate the spatial variability of rainfall. The results suggest that changes in the raingauge network since 1945 rather than climate may be influencing regional rainfall statistics. It was found that the distribution of raingauges between 1945 and 1975 was not adequate to sample latitudinal changes in rainfall and that the annual rainfall for the region was largely a product of poor sampling east to west until sufficient stations were reporting data from 1970. These results raise questions over the use of regional statistics such as rainfall anomalies and the fitting of regional trend lines to depict climate change in the Sahel. Geostatistical methods offer a more complex but more reliable approach to the estimate of regional rainfall characteristics. 27

S.J. McLaren and D.R. Kniveton (eds.), Linking Climate Change to Land Surface Change, 27–48. © 2000 Kluwer Academic Publishers. Printed in the Netherlands.

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1.

Introduction

The West African Sahel is well known due to reports on drought, desertification and famine that span at least three decades (Figure 1). Evidence of a change in the region’s climate is usually portrayed by a standardised rainfall anomaly plot (Figure 2) which displays almost continuous negative anomalies since the 1960s where,

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is the standardised rainfall anomaly. is the station rainfall for the ith station and kth year. is the time mean of the ith station. is the standard deviation of ith precipitation station (after Jones and Hulme, 1996). This has been noted for some time with Lamb (1974), Nicholson (1979) and Winstanley (1973) writing about a downward trend in Sahelian rainfall after international attention focused on the drought of the early 1970s. A decade later Copans (1983), Druyan, (1989), Flohn (1987) and Tickell (1986), supported the view of lower than average rainfall and some even wrote about a persistent drought. The notion of a desiccating environment continues into the 1990s with some workers linking this to an advancing Sahara and claims that desertification is affecting the region (Hulme and Kelly, 1993; Nicholson and Palao, 1993;

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UNEP, 1992; Zheng and Eltahir, 1998) with statements such as ‘The desert is advancing partly because of recurring cycles of drought’ (Pritchard, 1990 p.246); ‘After a 20 year series of droughts, the Sudano-Sahelian region remained the most permanently vulnerable area’ (Odingo, 1992, p.6); ‘the prolonged drought that has struck the Sahel for 25 years now.’ (D’Amato and Lebel, 1998 p.956).

A major problem with the above reports is that they fail to distinguish between desiccation, drought or desertification as defined in Table 1. It is clear that in order to adopt the most appropriate response it is necessary to distinguish between these different types of environmental degradation. The aim of this paper is then to answer the question; is the Sahel climate desiccating?

2.

Why Investigate Desiccation?

It may, at first, seem a waste of effort to investigate desiccation in the Sahel given the enormous evidence in support of persistent drought and a downward trend in rainfalls. Yet there are several reasons why the question is pertinent: refutation of the idea of an advancing desert; challenges to the notion of persistent drought; changes to the base period over which climate change is measured; changes to the number of climate stations available for analysis; reported increase in the variability of the data. The notion of land degradation in the Sahel and an advancing Sahara has been challenged, both in scientific and popular publications (Agnew, 1995; Agnew and Warren, 1996; Binns, 1995; Mainguet, 1991; Thomas, 1993). The notion of persistent drought has also been questioned for some time (Agnew 1989, 1990; Franke and Chasin, 1980; Garcia, 1981;

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Glantz, 1987, 1994; Wijkman and Timberlake, 1985). High rainfalls have been evident in the Sahel during the period of desiccation e.g., 1988 and 1998, yet there has been little critical examination of the notion of widespread desiccation. In addition, the thirty year base period upon which climate change is determined has recently (Hulme, 1992) been changed by the WMO from 1931-1960 to 1961-1990 with uncertain impacts upon the determination of rainfall trends. Hulme (1992) reported that rainfall is becoming more variable in the Sahel and the downward trend more persistent. He also noted that the number of rainfall stations available is declining. Ba et al. (1995) noted in their analysis of satellite-determined rainfalls in the Sahel, that the number of stations available between 1983 and 1988 fell from 271 to 147 (a 46% reduction). A trend that UNEP (1992) noted starting in the late 1970s and continuing through to 1990. But perhaps of greater importance is the variability in the network of Sahelian rainfall stations ‘....that are unevenly distributed in space, sparse in critical regions, and/or reported irregularly...it is often impossible to obtain a sufficient rainfall dataset over wide areas from a conventional rain gauge network.’ (Ba et al, 1995 p. 412). Hence, there is some concern over the reliability of the rainfall network employed in the Sahel to assess climate change while the impact of changing the number (and location) of rainfall stations used to determine rainfall trends in the Sahel is uncertain. Doubt can also be directed to the strategy of focusing on spatially aggregated rainfall statistics for all, if not major parts, of the Sahel that ignore local variations. This paper then seeks to investigate critically the evidence for widespread desiccation during changes in the climatic base period and aggregation of trends from different sets of stations each year. We first look at the temporal rainfall patterns through an examination of recent trends in the Sahel before undertaking a geostatistical analysis of its spatial variation.

3.

Results and Discussion

3.1 RAINFALL TRENDS: IS THE CLIMATE IN THE SAHEL DESICCATING? The analysis is based on data provided by the Climate Research Unit of the University of East Anglia. There have been several attempts to define the Sahelian region ranging from climate (Sivakumar 1989) to ecological conditions (Davy et al. 1976). As we are primarily interested in rainfall we can use the results of previous work that shows Continental Sahel (Niger, Burkina Faso and Mali east of 5°W) can be viewed as a climatic region (Ba et al. , 1995 and Moron, 1994). A regression line fitted to the standardised rainfall anomalies displayed in figure 2 for the years 1961 to 1994 produces a strong negative relationship between rainfall and time

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0.73) such that rainfall decreases by 0.069 of a standardised anomaly per annum. This is equivalent to a reduction in rainfall of 8 mm each year or 244 mm between 1961 and 1990, averaged across Continental Sahel. Given that the mean rainfall for these stations and for this period is 511mm this is a massive reduction in annual precipitation. Hulme (1996) examined reports of desiccation in the world’s drylands and found little global evidence of long term drying except for the Sahel. The Sahel showed a significant downward trend of 96.8mm per century, equivalent to a decline of around 1mm a year in annual rainfall, (based on the whole of the Sahel with an annual average rainfall of 451mm). This is much less than we have reported, even if the desiccation is assumed to only take place over the last 30 years hence producing a reduction of 3mm/year. Nicholson and Palao (1993) note the downward trend started early in the 1950s. They separated West Africa into nineteen regions and calculated for each, the standardised departures of rainfall from a long-term mean between 1950 and 1985. A regression line was fitted to the standardised departures for each region (Table 2). They found the downward trend was most pronounced in the wetter south, less in drier and western areas. Rainfall for Continental Sahel was also found to decrease by 0.055 of a standardised anomaly per annum (Nicholson and Palao, 1993) which is close to the value of 0.069 we reported above for a more recent time period.

Thus, there is some general agreement over the amount of desiccation that has taken place in recent decades. However, figure 3 (and table 2) suggest that this general figure of 0.069 of a standardised anomaly each year may be misleading. Higher values are found for Mali and Burkina Faso (0.087) but much lower for Niger (0.037). It is also evident from figure 3 that in fitting the trend line it is being strongly influenced by the very high rainfalls in the early 1960s and very low rainfalls in recent years. In between these periods there is much variation, especially for Niger where the correlation coefficient is only 0.255 but significant at the 1% level. The variation raises concern over the analysis of aggregated rainfall anomalies rather than aggregated trends. The former assumes rainfall in Continental Sahel is spatially homogeneous. It is also uncertain whether extreme values are unduly influencing the fitted regression lines and misleading the desiccation trend. These two points are discussed below.

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3.2 RAINFALL TREND LINES The analysis here focuses upon the rainfall stations in Niger as the data collected have already been shown to behave in a peculiar fashion. The persistent downward trend evident in the data aggregated for Continental Sahel (Mean) is not present in the data aggregated for Niger in figure 2. This difference is clearer in figure 4 where low rainfall during the early 1970s and mid 1980s should be contrasted with the large rainfalls in the early 1960s, late 1970s and 1994. Although the correlation coefficient of the relationship between standardised anomaly and time is small it remains significant at the 1 % level. To predict the desiccation for Niger between 1961 and 1990 (latest WMO period) regression lines were fitted to the standardised rainfall anomaly for each station and for groups of stations over this period. The reduction in rainfall equivalent to the decrease in standardised anomaly produced by the regression lines is shown in figure 5. The stations are arranged in ascending order of annual rainfall (Bilma is lowest with 12 mm, Gaya is highest with 796mm). The stations are also grouped based on annual totals of less than 250mm; between 250 and 500mm and above 500mm (used by Agnew, 1990 to identify pastoral; agro-pastoral and rainfed agriculture regions). There is a mean desiccation during this period of 150 mm of annual rainfall but this varies considerably. Surprisingly the variation is not simply based on the mean annual rainfall; Maine-Soroa (annual rainfall of 342mm) experiences a predicted drying of 237mm, equivalent to that of Maradi (annual rainfall 493mm). Whereas, Gaya (796mm) has a predicted drying of 124mm which is the same as that predicted for Zinder (annual rainfall 369mm). Notably, when the rainfall stations are grouped according to Agnew’s (1990) classification the predicted rainfall decrease does vary according to mean annual rainfall. When rainfall stations are aggregated into broad groups the trend lines for all stations are found to be significant as shown in figure 6. We have plotted the regression F statistic as the analyses all have the same degrees of freedom so the results are directly comparable irrespective of the annual amount of rainfall or magnitude of the rainfall anomalies. In contrast the results for individual stations contain a high degree of variability. Some stations e.g., Tahoua and Maine-Sora display a highly significant pattern of desiccation while at others e.g., Gaya and N’Guigmi rainfall does not fit into a linear drying trend. This is demonstrated in figures 7 and 8. The key point here is that the aggregation of rainfall anomalies conceals those stations that do not display any rainfall trend. The suspicion is that those stations that contain extreme anomalies are masking other stations where there is no clear linear trend. This can be illustrated by plotting the predicted amount of drying from the desiccation trend line against the difference between the highest and the lowest observed rainfall value (figure 9). The result is significant at the 1% level of significance. If stations with a weak trend line, (N’Guigmii, Zinder, and Gaya) are excluded then the

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relationship between desiccation trend and rainfall maximum minus rainfall minimum is even stronger In a semi-arid region of highly variable rainfall it does seem unwise to accept a general statement of regional desiccation when this is based on a few extreme observations. Furthermore, when these extreme values occur at the start (early 1960s) and end (late 1990s) points of the period it can be no surprise that a linear trend line can often be fitted to the data. This raises the question as to whether the period 1961 to 1990 is appropriate for such an analysis of desiccation in the Sahel.

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No significant trend is observed for the rainfall stations with the longest run of data, Niamey and Zinder (from 1905) and Tahoua (from 1922) plotted in figure 10. The abnormally high rainfalls of the 1950s and early 1960s are then biasing perceptions of climate change in the Sahel. In addition, the affect on aggregation of changes in the location and number of rainfall events and changes to the location and number of stations, is largely unknown and will be dealt with next. 3.3 DOES AGGREGATION MISLEAD? We have demonstrated that the spatial aggregation of rainfall anomalies can lead to the erroneous conclusion that all parts of the Sahel have experienced similar rates of recent desiccation. There is then a need to understand the spatial variability of the Sahelian rainfall and the impact of changes to the location and number of stations used to aggregate climate statistics for the Sahel. Ba et al. (1995), using Meteosat data to predict seasonal rainfalls, found a marked difference between the position of the isohyets when comparing (a) only Meteosat derived rainfalls based upon the terrestrial raingauge network to (b) Meteosat derived rainfalls based on full regional coverage. The implication was that the number and distribution of raingauges influenced the estimation of rainfall totals. The suggestion is that areal estimates of rainfall are unreliable and simply taking the mean of all stations biases the

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wetter south where there are more stations. This suggestion is tested by geostatistical analyses of total boreal summer (June, July, August and September) rainfall (TBSR) for the West African Sahel (10-20°N and 20°W to 20°E) between 1931 and 1990. A summary of the results is presented here, further details of the analysis can be found in can be found in Chappell and Agnew (in review). Intuitively reasonable trend due to latitudinal variation in the annual TBSR data was removed by fitting quadratic polynomials on their spatial co-ordinates using least squares regression (Chappell et al., 1996). Ordinary experimental variograms were computed for the residuals from the trend of TBSR every year between 1931 and 1990 in the two principal directions (N-S and E-W) of spatial variation. These variograms were fitted with models, (see Chappell and Agnew, in review for further details), using weighted least squares in Genstat (Genstat 5 Committee, 1992), which all included a nugget variance, typical in sparsely sampled continuous data (Chappell and Oliver, 1997). Although measurement error and stochastic variation in data contribute to the nugget, the largest source of variation is commonly due to spatially dependent variation that occurs over distances much smaller than the shortest sampling interval (Webster and Oliver, 1990). The majority of the variograms were bounded and included sill variance and range parameters. The dissimilarity between TBSR for average separation distances between rainfall stations increases until it reaches a maximum known as the sill variance. The lag separation distance at which the variogram reaches its sill is the range; this is the limit of spatial dependence (Webster and Oliver, 1990). Beyond this limit the variance bears no relation to the separation distance. Some of the variograms are unbounded (only linear models were appropriate and the model parameter includes the gradient) and have a structure, which appears to increase indefinitely at this scale of investigation. This suggests that as the area of interest increases, so more sources of variation are encountered (Chappell et al., 1996, Webster and Oliver, 1990) i.e., TBSR remains spatially dependent with increasing distance. The model parameters of the fitted variograms are plotted for each year in figures 11 – 13. The parameters of the variograms were used for anisotropic punctual kriging to estimate TBSR residuals on a 1 degree grid across the West African Sahel. The quadratic polynomials for each year were added back to the kriging estimates. Finally, isohyets were threaded through the kriging estimates of TBSR (with the same isoline frequency) and plotted for every year between 1931 and 1990. Figure 14a, b, c, d, e, f and g are approximately decadal examples of the rainfall maps produced.

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The magnitude and variability of the E-W nugget variance (Figure 13a) decreases over time suggesting that the spatial dependence of rainfall is better sampled over time. Since the E-W range (figure 12) is not decreasing over this period it suggests that the structure of the rainfall during this time is not varying. Thus, the reduction in spatial dependence (E-W nugget variance) is due to improved sampling of rainfall by the station network. The station locations for selected years (figures 14a, b, c, d, e, f, g) show an extension into the easterly end of the region. An E-W gradient in rainfall exists as shown by the isohyets threaded through the kriging estimates of total summer rainfall for several years. The isohyets are more compressed in the west of the region than in the east. The importance of this eastwards extension of rainfall stations is that over time more rainfall stations are located in the area of lower rainfall. Thus, the spatial variation in rainfall throughout the region is increasingly better sampled.

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The inter-annual variation in the N-S nugget variance (figure 13b) is considerably larger than the E-W nugget variance and the general trend in the former is very different from the latter. The nugget variance before ca. 1945 and after ca. 1970 is generally smaller than that shown in the period between these dates. This suggests that the spatial dependence of rainfall is better sampled in the early and late periods of the rainfall record and that between ca. 1945 to 1975 the configuration of rainfall stations in the N-S direction has poorly sampled the spatial dependence of rainfall. The reason is not evident in the selection of maps showing the rainfall stations (Figure 14) because the inter-annual variation of the nugget variance is very large. Moreover, an examination of all maps suggests that there has been very little N-S variation in the location of rainfall stations. The most likely explanation is that the N-S configuration of the rainfall stations was inadequate to sample the generally large spatial dependence (range; figure 12) during the period between ca. 1945 to 1975. This would have been obvious if more of the N-S variograms had been fitted with bounded models. That they were not is evidence itself for increasing sources of variation with increasing separation distance. However, the inter-annual variation in the N-S nugget variance is larger than that in the E-W direction because the variation in the range of spatial dependence as a ratio of the total distance is much larger in the N-S direction. This interpretation is complicated by the difficulty of modelling the variograms in this direction where fewer pairs are available than in the E-W direction. The isohyets threaded through the kriging estimates of total summer rainfall for several years (Figure 14a, b, c, d, e, f, g) show the expected anisotropic variation, whereby the rainfall gradient is greater in the N-S direction than in the E-W direction. The N-S rainfall gradient is greater within the period 1945 to 1970 (Figure 14c, d, e) than outside this period (Figure 14a, b, f, g) as evident from the isohyet compression in the south of the region. It is no coincidence that the pattern of N-S nugget variance (Fig. 13b) bears a striking resemblance to the pattern of average annual rainfall aggregated for the region (Fig. 2). The different sources contributing to the nugget cannot individually be quantified. However, it is highly likely that measurement error is small and that randomness is not large due to climatic control. Most other geostatistical applications have shown that where measurement error can be ruled out and randomness assumed small the main source of the nugget is the scale and intensity of sampling. The distribution of rainfall stations over the period ca. 1945 and 1975 has not adequately sampled the N-S variation in rainfall. The average rainfall during this period is controlled by the predominance of rainfall stations in the more southerly locations associated with larger rainfall. The average annual rainfall for the region is also a product of the poor E-W sampling configuration early in the rainfall record and the general increase in the effectiveness of sampling more recently. It appears that the only period when the spatial dependence of rainfall has been adequately sampled by the rainfall station network (i.e., the nugget variance in the N-S and E-W directions have been at their smallest) is since 1970.

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Ironically, this is the period when many workers have reported a persistent downward trend in the average annual rainfall. However, the results here suggest that this trend is a return to a more precise estimate of the rainfall in this region. Other patterns are artefacts of the use of spatial aggregation which depends on the relationship between rainfall station location and rainfall spatial dependence.

4.

Conclusions

There is overwhelming evidence in research publications that the Sahel, as a region, is desiccating. Most reports place the start of this change in climate in the 1960s though some suggest a decade earlier. Examination of standardised rainfall anomalies for Continental Sahel supports this view with annual rainfalls declining through the 1970s and 1980s. The rainfall averages in 1961 to 1990 are significantly lower than during 1931 to 1960. A more localised analysis of rainfall trends has shown that this regional aggregation can mask local variations. A downward trend over the last 30 years is not significant at all stations and Niger in particular displays different patterns compared to Mali and Burkina Faso. A downward trend is also not evident over a twentieth century long perspective giving rise to concern that the abnormally high rainfalls of the 1950s and 1960s are leading us to believe that the 1970s and 1980s are abnormally low. The danger with all of these points is that they merely illustrate the variability of the data set. The results of a geostatistical analysis show that patterns in the mean annual rainfall are an artefact of the rainfall station location and rainfall spatial dependence. The recent downturn in mean annual rainfall appears to be a return to a more precise estimate as a consequence of improved station sampling of the rainfall distribution. The analysis supports the suggestion that areal estimates of rainfall are unreliable and the simply taking the mean of all stations biases the estimate (Ba et al., 1995). However it is too simplistic to assume that the bias is due to the location of more stations in the wetter south. The geostatistical analysis has shown that the temporal variation in location (N-S and E-W) of the rainfall station network has a considerable effect on the sampling distribution of rainfall across the region. Even if we accept that conditions in many parts of the Sahel are correctly represented by the claims of widespread drought and a persistent downward trend in rainfalls, this fails to explain the impact upon people living in this region. Despite alarm that the environment is degrading food production has increased over the last 30 years (Agnew, 1995). Livestock numbers have been affected by the droughts of the 1970s and 1980s but numbers quickly reestablished when rainfalls recovered (IUCN, 1989). Claims that the Sahara is advancing have been refuted (Thomas, 1993) and poor food supply has been explained by factors such as

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price, distribution and the role of institutions rather than climate (Norse, 1994 and Olsson, 1993). Those who state that the Sahel has recently desiccated but then ignore what this means for the inhabitants and the rest of the physical environment of the region are guilty of oversimplifying rainfall patterns and of assuming the landuse systems in the Sahel are based upon the higher rainfalls experienced in the middle of this Century. Taking only a 30 year period to establish climate change is then unwise for the Sahelian region and any statistic that claims to represent Sahelian conditions as a whole should be treated with caution.

Acknowledgements The monthly rainfall totals were generously provided by the Climate Research Unit at the University of East Anglia.

References Agnew, C. T. (1989) Sahel drought, meteorological or agricultural? International Journal of Climatology 9, 371 382. Agnew, C. T. (1990) Spatial aspects of drought in the Sahel. Journal of Arid Environments 18, 279-293. Agnew, C. T. (1995) Desertification, drought and development in the Sahel, in Binns, A. (ed.) People and environment in Africa. J. Wiley and Sons, Chichester p137-149. Agnew, C. T. and Warren, A. (1996) A framework for tackling drought and land degradation. Journal of Arid Environments 33, 309-320. Ba, M. B., Frouin, R. and Nicholson, S. E. (1995) Satellite derived interannual variability of West African Rainfall during 1983-88. Journal of Applied Meteorology 34, 411 -431. Binns. T. (1990) Is desertification a myth? Geography 75, 106-113. Chappell, A. and Oliver, M. A., (1997) Geostatistical analysis of soil redistribution in SW Niger, West Africa, in E. Y. Baafi and N. A. Schofield (eds.) Quantitative Geology and Geostatistics, Kluwer, Dordrecht. pp. 961-972 Chappell, A. Oliver, M. A. Warren, A. Agnew, C. T. and Chariton, M. (1996) Examining the factors controlling the spatial scale of variation in soil redistribution processes from south-west Niger. In M. G. Anderson and S. M. Brooks (eds.) Advances in Hillslope Processes. J. Wiley and Sons, Chichester. pp 429-449 Copans, J. (1983) The Sahelian drought, in Hewitt, K. (ed.) Interpretations of calamity. Allen and Unwin, London. pp 83-97 D'Amato, N. and Lebel, T. (1998) On the characteristics of the rainfall events in the Sahel with a view to the analysis of climatic variability. International Journal of Climatology 18, 955-974 Davy, E. G., Mattei, F and Solomon, S. I. (1976) An evaluation of the climate and water resources for development of agriculture in the Sudan-Sahelian zone of West Africa. WMO Special Environmental Report No.9. WMO, Geneva. Druyan, L. M. (1989) Advances in the study of sub-saharan drought. International Journal of Climatology 9, 7790. Flohn, H. (1987) Rainfall teleconnections in northern and eastern Africa. Theoretical & Applied Climatology 38, 191-7 Franke, F. and Chasin, B. (1980) Seeds of famine. Allanheld, Osman and Co., New Jersey. Garcia, R. V. (1981) Drought and Man: Volume 1: Nature Pleads Not Guilty. Pergamon, Oxford. Glantz, M. (ed.) (1987) Drought and hunger in Africa. Cambridge University Press. Cambridge.

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Glantz, M. (ed) (1994) Drought follows the plow. Cambridge University Press, Cambridge. Hulme, M. (1992) Rainfall changes in Africa. International Journal of Climatology 12, 685-699 Hulme, M. (1996) Recent climatic change in the world's drylands. Geophysical Research Letters, 23 (1), 61-64 Hulme, M. and Kelly, M. (1993) Desertification and climate change. Environment 35 (6), 39-45, International Union for Conservation of Nature 1989 Sahel Studies. IUCN, Nairobi. Jones, P. D. and Hulme, M. (1996) Calculating regional climatic time series for temperature and precipitation: methods and illustrations. International Journal of Climatology 16, 361-377 Lamb, H. H. (1974) The Earth's changing climate. Ecologist 4, 10-15 Mainguet, M, (1991) Desertification: natural background and human mismanagement. Springer-Verlag, Berlin. Moron, V. (1994) Guinean and Saharan rainfall anomaly indices at annual and monthly time scales (1933-1990). International Journal of Climatology 14, 325-341. Nicholson, S. E. (1979) Revised rainfall series for the West African subtropics. Monthly Weather Review 107, 62023. Nicholson, S. E. and Palao, I. M. (1993) A re-evaluation of rainfall variability in the Sahel. International Journal Climatology 13, 371-389. Norse, D. (1994) Multiple threats to regional food production, environment, economy, population. Food Policy 19 (2), 113-148 Odingo, R. S. (1992) Implementation of the plan of action to combat desertification (PACD) 1987-1991 Desertification Control Bulletin 21, 6-14. Olsson, L. (1993) On the causes of famine - drought, desertification and market failure in the Sudan. Ambio 22 (6), 395-403 Pritchard, J. M. (1990) Africa. Longman, Harlow. Sivakumar, M. V. K. (1989) Agroclimatic aspects of rainfed agriculture in the Sudano-Sahelina zone, in Proceedings of Workshop on Soil, crop and water management systems for rainfed agriculture in the SudanoSahelian zone, Niamey January 1987. ICRISAT Sahelian Center, Niamey. ICRISAT, Patancheru, AP 502 324 India, pp 17-38 Thomas, D. G. (1993) Sandstorm in a teacup I Understanding desertification. Geographical Journal 159 (3), 318331. Tickell, C. (1986) Drought in Africa: impact and response. Overseas Development 102, United Nations Environment Programme (1992) World Atlas of Desertification. Edward Arnold, London. Warren, A. and Khogali, M. (1992) Assessment of desertification and drought in the Sudan-Sahelian region 19851991 UNSO, New York. Webster, R. and Oliver, M. A. (1990). Statistical methods in soil and land resource survey. Oxford Univ. Press. Wijkman, A. and Timberlake, L. (1985) Is the African drought an act of God or of man ? Ecologist 15 (112), 9-18. Winstanley, D. W. (1973) Rainfall patterns and general atmospheric circulation. Nature 245, 190-194 Zheng, X and Eltahir, E. A. B. (1998) The role of vegetation in the dynamics of West African monsoons. Journal of Climate 11, 2078-2096.

HYDROLOGICAL RESPONSE OF DESERT MARGINS TO CLIMATIC CHANGE: THE EFFECT OF CHANGING SURFACE PROPERTIES A. YAIR Department of Geography, The Hebrew University, Jerusalem, Israel. R. B. BRYAN Faculty of Forestry, The University of Toronto, Toronto, Canada.

Abstract Arid and Semi-arid ecosystems are regarded by ecologists as highly resistant to stress due to their adaptation to the extreme variability in the climatic conditions over a time scale of decades. Under such conditions a rather extreme change in climate, mainly rainfall, would be required in order seriously to affect natural semi-arid and arid environments. The above approach disregards the fact that one of the forms of landsurface change that may result from climatic change in deserts, and especially at a desert fringe, is not limited to purely climatic variables such as precipitation and temperature. It is always accompanied by quite rapid alteration of surface properties, connected to deposition of loess or sand. In subtropical semi-arid and arid areas loess deposition, at a given site, is often attributed to wet periods; while sand deposition to dry periods. The new surface properties can be expected to exercise strong influence on infiltration, runoff and soil moisture. An aspect not yet answered is how much sand or loess deposition is required to affect the hydrological regime and related water resources. In order to check the effect of thin topsoil sand or fine-grained layers on infiltration and runoff sprinkling experiments were conducted in the laboratory, at various rain intensities and duration. Data obtained show that a slight change in surface properties has a rapid and significant hydrological effect. A sand layer 1-2 cm thick is enough to eliminate runoff generation; whereas a fine-grained layer 1-2 mm thick has an opposite effect, significantly increasing runoff generation. One can therefore conclude that arid and semi-arid environments, although highly adapted to extreme variations in rainfall, may be extremely sensitive to slight changes in their surface properties, which alter their hydrological regime quickly and efficiently. 49

S.J. McLaren and D.R. Kniveton (eds.), Linking Climate Change to Land Surface Change, 49–63. © 2000 Kluwer Academic Publishers. Printed in the Netherlands.

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Introduction

The term resilience is often used by ecologists (Holling, 1973) to describe the degree to which an ecosystem can be disturbed and yet return to its previous composition and structure. A system disturbed beyond its resilience will develop into a new ecosystem characterised by an altered composition and structure. Such drastic change can be triggered by human activity. For example, the introduction of grazing into a grassland area is often responsible for the replacement of a grass cover by a shrubland. The same result has been predicted for semiarid areas due to warmer and drier climatic conditions (Schlesinger et al., 1990). However, semi-arid and especially arid ecosystems are also regarded by ecologists as highly resistant to stress (Holling, 1973; Thiery, 1982; Wiens, 1985). These ecosystems are adapted to the extreme variability in the climatic conditions in the rainy season, from year to year and over a time scale of decades. The high resilience of arid environments is well demonstrated by the fact that some rocky mountainous areas within extreme deserts (such as the Negev and Sinai) include up to 30% of Mediterranean and Irano-Turanian species (Yair and Danin, 1980). The very existence of such species in an area with less than 100 mm average annual rainfall clearly proves that enough water is provided to such plants even during a sequence of dry years. Furthermore, Shmida (1982) reports the occurrence of endemic Mediterranean species in the Negev and Sinai deserts where present day average annual rainfall is 70-100 mm. As the development of endemic species requires a long period of isolation, such occurrence can be considered as indicative of stable conditions over hundreds or even thousands of years. To summarise, as stated by Thiery (1982) “species adapted to highly variable environmental conditions, and a high rate of mortality, are more likely to tolerate an extreme stress than are species from very constant environments”. In this situation a rather extreme change in climate, mainly in rainfall, would be required in order seriously to affect natural semi-arid environments. The above approach disregards the fact that climatic change in subtropical deserts, and especially at a desert fringe, is never limited to climatic and environmental variables such as precipitation, temperature, vegetation cover and soil properties. It is almost always accompanied by quite rapid alteration of surface properties. The new surface properties exercise strong influence on infiltration, runoff, soil moisture and thus on the vegetal cover (Yair, 1983; 1994; Yair and Danin, 1980). Scientists working in the Sahara (Coude-Gaussen, 1991) and in the Negev desert (Goldberg, 1981; Yaalon and Dan, 1974; Goring-Morris and Goldberg, 1990) tend to agree that one type of system change is related to aeolian deposition. According to these authors, loess deposition in the subtropical desert fringe, took place during relatively wet periods, while sand deposition is characteristic of dry climatic periods. Studying the environmental effects of loess penetration into the Northern Negev, Yair (1983; 1994), Yair and Shachak (1985) and Kadmon et al., (1995) showed that although loess penetration occurred during a wet period, it resulted in an increase in salt input (by rainfall and dust) coupled with a limited leaching depth. This scenario led to soil salinisation and desertification processes. An opposite trend occurred during the following dry period. The negative

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effect of rainfall decrease was counteracted by sand penetration that allowed deep rainwater infiltration, deep leaching and good water preservation. The effects described above reflect the response of the environments studied to climatic changes, during the late Quaternary, at a geological time scale. An aspect not yet considered is how much sand or loess deposition is required to affect the hydrological processes. Is a sand layer of several centimetres, or a thin loess layer of several millimetres, sufficient to pass a threshold, which irreversibly affects the water regime. The thinner the layer the shorter the time necessary for impact on the environment. It is obvious that aeolian deposition rates vary tremendously from one geographic area to another, as well as within a given area, in relation to the availability of the material, the distance from the source area and the regional and local wind regime. Sand accumulation by wind can be very rapid. Field monitoring in a sandy area in the Negev desert shows deposition, and / or erosion rates, of 10-100 cm during a single year, most of it in one extreme windstorm (Kadmon and Leshner, 1995). On the basis of air photos, taken in 1968 and 1982 along the Egyptian –Israeli border, Tsoar and Moller, (1986) report sand accretion up to 5 m at the crest of linear dunes over an 18-year period. Sand incursion into the area, based on C14, TL ages and prehistoric implements, is assumed to have begun up to 43,000 years ago (Goldberg, 1981; Magaritz and Enzel, 1990; Rendell et al., 1995). Boreholes dug in the area show that sand thickness is some 30 m at the crest of the linear dunes. This gives an average net accumulation rate of ~ 0.7 mm per year. Accumulation rate is lower within the interdune corridors where sand thickness is only 6-10 m. The time required for the deposition of 1-2 cm of sand is therefore very short being of the order of 15-30 years, and probably shorter during periods of high deposition rates. The accumulation rate of loess in the Negev desert, north of the sandy area, is lower. Loess deposition under present day dry climatic conditions is of the order of 0.01mm per year (Yaalon and Ganor, 1975; Bruins and Yaalon, 1979; Goosens and Offer, 1990). According to Yaalon and Dan (1974) the loess deposit in the Beersheva sedimentary basin is 12-15 m thick and its age is of the order of 100,000 years, which gives a net average accumulation rate of 0.12-0.15 mm per year. Accumulation rates are somewhat lower on hillslopes where erosion occurs. Under such conditions the time required to deposit 1-2 mm of loess covers only a few decades. The actual deposition rate must have been higher as part of the loess has been eroded and transported out of the area into the Mediterranean Sea.

2.

Aim of Present Study

In view of the relatively fast deposition rate of aeolian materials, both sand and loess, we decided to check the effect that thin layers of sand or loess exercise on infiltration and runoff processes. The hypothesis advanced here is that natural semiarid environments, although quite resilient to changes in precipitation and temperature, are more sensitive to slight changes in surface properties associated with climatic changes.

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Such persistent sedimentary changes can be expected to alter quickly the hydrological regime, drastically affecting water availability and thus the natural ecosystem over a long time scale.

3.

Experimental Design

As the hydraulic properties of sand and loess differ greatly, two different sets of experiments were planned. The first aimed at checking the effect on runoff of two sand layers (1-2 cm and 4-5 cm thick) overlying a relatively impervious substratum. By using two sand layers that differ in their thickness we expected to identify the threshold amount of sand required to have a significant effect on the hydrological regime. In the second set the effect of a thin (1-2 mm thick) fine-grained layer, whose composition is similar to that of loess material, overlying a permeable sand layer was checked. The study is based on sprinkling experiments, conducted in the laboratory, at various rain intensities and duration. Trays of 100 cm x 39.2 cm were set at a slope angle of 4 degrees (Figure 1). Rainfall amount and distribution were monitored with rain collectors placed at the edge of the boxes (Figure 1). Surface and subsurface flow samples were collected from the trays every 60-90 seconds and moisture content was determined prior to each run.

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For the study of the effects of sand cover two trays were first filled up with Kettle Creek silt loam soil, from Ontario, Canada. The particle size composition of this sediment (Figure 2) is similar to that of loess in the northern Negev desert. Boxes were sprinkled with rainfall intensity of 31 mm/hr. for two hours, until ponding occurred. The fall height of the drops was 5.5 m and kinetic energy reproduction was 75-80% of similar natural rainfall. The material was left to drain overnight. The following day the upper layer of the wetted soil was removed and replaced with dry sand whose particle size composition is shown on Figure 2. The sand was derived by sieving samples of Pontypool loamy sand, developed on kame deposits in Ontario. The particle size composition of the sand used (Figure 2) is quite similar to that of sand forming the longitudinal dunes in the north-western Negev. One box was covered with a sand layer 1-2 cm thick and the other with a layer 4-5 cm thick. Five sprinkling experiments were conducted with the sand cover. The protocol of these experiments is given in Table 1.

The two remaining experiments were conducted on the fine-grained layer, spread over a dry sand substratum. The first run was conducted on the dry, uncompacted, silt loam. Rainfall was applied at average intensity of 44.3 mm/hr for 17 minutes, representing a rain amount of 15.5 mm. The second run was conducted three days later with wetter surface conditions and compacted topsoil. Moisture content of the topsoil was 23.3%. The test lasted 10 minutes with a rain intensity of 41.2 mm/hr, producing 6.9 mm rainfall depth.

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Results

4.1. RUNOFF GENERATION ON THIN SANDY LAYERS 4.1.1. Runs with continuous rain Data obtained during the first three experiments, in the two boxes, are presented in Figure 3. Neither surface flow nor ponding occurred on the thick sand layer during any of the runs. Surface flow did occur on the thin sand layer during the first two runs at 3637 mm/hr rain intensity (Figure 3A, 3B), but not during the last run with the lower intensity at 24.5 mm/hr. Local ponding was observed after 14 minutes of rainfall. Ponding on the shallow sand layer started at minute 6 during the first run and runoff at minute 15 with a very sharp increase in discharge. Discharge decreased suddenly five minutes later, coincident with the start of subsurface flow. This phenomenon was not observed during any of the following experiments. During the second run (Figure 3B), under wet surface and subsurface conditions (Table 1), the time to runoff was shorter, total discharge higher and equilibrium conditions were reached within a minute after runoff started. Subsurface flow was observed, on both boxes, during all three runs. During the first run subsurface flow was delayed compared with surface flow. As could be expected, because of difference in pore volume, subsurface flow started later, and with a lower discharge on the thicker sand layer (Figure 3). During the second run (Figure 3B), subsurface flow in both boxes started at the same time as surface flow. Again subsurface flow discharge was higher on the thin than on the thick sand layer. On the third run (Figure 3C), conducted at lower rain intensity, trends recorded were similar to those at the first run, except for a shorter time lag until the beginning of subsurface flow that resulted in higher discharges at the two boxes. 4.1.2. Runs with intermittent rain Data obtained are presented in Figure 4. During the first run (Figure 4A), conducted at the lower intensities, in the range of 23.6-32 mm/hr, surface runoff did not develop on any of the boxes. However local ponding was observed during the two last rainshowers when moisture conditions and rain intensities were the highest. Subsurface flow occurred only on the thin sand layer and was limited to two rain-showers (Figure 4). The second run (Figure 4B) was conducted a day later under wet surface conditions. The very high rain intensities applied, coupled with the saturated silty-loam substrate, resulted in a quick response of the thin sand layer. Surface flow developed quickly with the highest discharges recorded. Despite the extreme conditions surface flow did not develop over the thick sand layer. Subsurface flow occurred simultaneously, at ponding time, on both boxes during each of the rain-showers applied. Due to the short duration of the rain-showers equilibrium conditions were not reached. Subsurface flow discharge was higher on the thin sand layer during the first rain-shower, but not during the following one (Figure 4 B).

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The hydrological response of the sandy layers to rainfall highlights the following points: despite the extreme rain conditions applied during the experiments surface flow never developed on the 4-5 cm layer, which was able to absorb and drain all rainwater at all rain intensities applied. This is due to its high pore volume and the rapid drainage at the interface with the saturated underlying layer. The response of the 1-2 cm sand layer was different. Despite the high moisture content of the underlying layer (Table 1), runoff did not develop during the run with 24.5 mm/hr, or during the low intensity intermittent rain-showers. In both cases the amount of rain applied was of the order of 15 mm. However, surface flow developed during all higher intensity rainshowers, even on dry sand. Runoff generation over the medium and fine-grained sand used cannot be ascribed to surface sealing processes, but rather to a return flow phenomenon as described by Dunne and Black (1970). A perched water table developed above the underlying saturated silty loam soil, filling the pore space in the thin sandy layer. Once the pore space was saturated water appeared at the surface, starting first with saturated wedge at the lower end of the box. During the following stage runoff rate was conditioned by the rate of rainfall application and drainage, through subsurface flow, at the interface between the sand and fine-grained soil. The whole process was faster during the runs conducted under wet surface and subsurface conditions.

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It is important to note that the perched water table developed quickly because the underlying soil was saturated or nearly saturated during all runs. Had this soil been significantly drier, runoff would not have developed on any of the runs because of the high absorption capacity of the underlying well-aggregated soil. Two hours of sprinkling at 31 mm/hr. (representing 61mm rain depth) were needed for ponding to appear over this soil. Surface and subsurface flow recorded during the experiments are considered to result from the experimental design. They are very unlikely to occur under natural desert conditions, characterised by scarce and intermittent small rainstorms, high temperatures and high evaporation rates, coupled with dry surface and subsurface soil, that would not allow the development of a perched water table or a saturated subsoil beneath a thin sand layer. 4.2. RUNOFF GENERATION ON A THIN FINE-GRAINED LAYER Two runs were conducted. The first run was performed over dry sand covered by a powdery, loose fine-grained layer. Within 90 seconds of sprinkling, cracks developed in the fine-grained material, probably due to hydro-compaction and subsequent sealing processes of the unconsolidated fine-grained material. The cracks were 1-2 mm wide. Some of them were already sealed or filled up with water when ponding occurred at minute 7 (Figure 5). Runoff started two minutes later when the accumulated rain amount was 6.5 mm. Discharge increased quickly and stabilised after 2.5 minutes. Runoff rate, at peak flow, was 68 % of the rain applied. Subsurface flow did not develop. The second run was conducted under wet conditions. Runoff started almost immediately. Runoff coefficient, at peak flow, was 74 %. A small trickle of subsurface flow was observed at the last minute. The effect of the thin fine-grained layer, on top of the highly permeable sand, was striking. The fast development of runoff on the first run is a clear indication of very effective sealing of the fine-grained layer, supported by the complete lack of subsurface flow. Runoff generation during the second run was much faster; due to wet surface conditions as well as to the development of a surface crust seal at the end of the first run.

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Discussion

The discussion will focus on the issue of extrapolating results obtained in small-scale laboratory experiments to field conditions. Such an analysis is important in an attempt to upscale results obtained to a landscape scale, while addressing the question of the long- term environmental effects of surface changes connected to climatic changes. An ideal situation for this investigation is found in the northern Negev desert, where different phases of loess and sand deposition have been recorded during the late Quaternary. In the central part of the northern Negev, the loess deposits cover the flat valley bottoms and extends over valley hillslopes, where the loess is in direct contact with older Eocene and Cretaceous bedrock. South of the loess-covered area, the landscape is rocky with deeply incised valleys. Such a situation allows studying the hydrological and environmental effects of loess deposition over rocky areas. A different landscape exists in the western Negev, along the northern part of the Egyptian- Israeli border. Here, the Nizzana sand field represents the eastern edge of the extensive Sinai erg. It is characterised by linear dunes separated by wide interdune corridors. Several trenches dug in an interdune corridor reveal a sequence of loess layers alternating with sandy units. The thickness of the loess layers is in the order of 10-40 cm. (Yair, 1990;

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Harrison and Yair, 1998). In several places the loess units outcrop at the surface, forming flat surfaces adjacent to dune ridges. The proximity of loess-covered surfaces to sandy covered surfaces allows a comparative study of these two units on the water regime and related environmental conditions. Average annual rainfall in the study area varies in the range of 90-200 mm, being higher in the northern loess covered area than in the southern sandy and rocky areas. The sprinkling experiments conducted lasted 9 days during which the total rain amount applied was 82.5 mm, very close to the long-term average annual rainfall for the southern areas. Rain amounts applied, during each of the runs, occur in this area one to three times a year. However rain intensities used, and especially their long duration, represent rather extreme to very extreme conditions. Rainfall data collected in this area during the last 30 years show that 85% of the rain fell at an intensity below 10 mm/hr (Kutiel, 1978). Rain intensities up to 30-35 mm/hr are recorded almost every year but for a short duration of 1-6 minutes. Higher intensities are rare and usually last no more than 1-2 minutes. 5.1. ENVIRONMENTAL EFFECTS OF LOESS DEPOSITION OVER ROCKY SURFACES As indicated earlier, there is a general agreement, among scientists working in the area, that loess deposition took place during a wet period whereas sand deposition occurred during dry periods. Several studies were devoted to the environmental effects of loess penetration into the northern Negev desert. These studies cover hydrological, pedological, botanical and zoological aspects. Long-term hydrological data collected at the Sede Boqer Instrumented Watershed, located in the rocky Negev where average annual rainfall is 93 mm, clearly indicate that rocky areas respond quickly to rainfall. (Yair, 1994; 1999). The threshold rain amount necessary to generate runoff is in the order of 2mm. Runoff occurs with any rainstorm having an intensity exceeding 5 mm/hr. Under such conditions runoff frequency and magnitude are high. Runoff generated over the rocky areas is absorbed on its way downslope by colluvial mantles, allowing local water concentration, deep water penetration and soil leaching. A different hydrological response is characteristic of loess covered areas. Stibbe, (1974) and Morin and Jarosh, (1977) show that the threshold rain amount required for runoff generation over the loessial soils of the Beersheva Basin is ~ 8 mm, very close to that obtained in the laboratory experiment (6.5 mm). Due to rainfall scarcity, the prevalence of low intensity rainfall, long time intervals between rainstorms, high evaporation rates, higher porosity and higher water absorption capacity of the loess, runoff frequency and magnitude are much lower than in rocky areas. However, the compacted loess prevents deep water infiltration. The depth of water penetration seldom exceeds 40 cm (Yair, 1994). Under such conditions, leaching is confined to a shallow depth, contributing to a gradual soil salinisation process. This process is further enhanced by the increase in salt input of airborne salts, from rainfall and dust, during a wet climatic period. The degradation of the water regime, and soil salinisation process, that followed loess penetration, had long lasting effects on the vegetation and on the biological activity. Comparative studies were conducted in the Negev desert between the northern, wetter,

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loess-covered area and the southern rocky climatically drier southern area. These studies deal with the density and composition of the vegetal cover, with the abundance and composition of burrowing animals and snails and with soil salinity. Data collected clearly show that soil salinity in the climatically wetter loess area is higher than that of soils in the rocky area. The northern loess area is also far more arid and less productive than the drier southern rocky area (Yair and Shachak, 1987; Kadmon et. al., 1989; Yair, 1994). In other words, although loess penetration had occurred during a wet climatic phase it resulted in an overall desertification effect. 5.2 LOESS DEPOSITION OVER A SANDY SUBSTRATUM Similar environmental studies were conducted in the Nizzana sand field. As could be expected, runoff generation is faster on the compacted loess covered than on the loose sand covered deposits (Yair, 1990). Depth of water infiltration is limited to 40 cm in the loess. Infiltrated water is quickly lost by evaporation. The soil is saline (Blume et al., 1995) and devoid of vegetation. The sandy areas represent a far better edaphic environment. Deep rainwater infiltration, up to 400 cm in rainy years, (Yair et. al., 1997) combined with low capillary water movement create a water reservoir available for plants. Plant cover, on the stabilised section of dune slopes, is 30-40 %, reaching almost 100 % at the base of the dunes. The pronounced positive effect of a shallow sand cover is evident where small sand mounds develop on the flat loess covered areas, allowing for the formation of local water lenses at the interface loess-sand. A mound 10 cm thick, is enough to allow germination of annual plants, whereas a mound 30-40 cm thick can support perennial shrubs.

6.

Conclusions

Data presented in this study support the hypothesis that a change in surface properties has a rapid environmental effect, in semiarid and arid areas, where climatic changes are accompanied by the rapid input of aeolian material. A sand layer 1-2 cm thick, even if deposited on top of a relatively impervious substratum, is thick enough to eliminate runoff generation under an arid rainfall regime characterised by infrequent, low intensity and intermittent rainstorms, with limited total rainfall. An opposite effect can be expected when fine-grained material, such as loess, is deposited above a highly permeable sandy substratum or over rocky surfaces. The sealing and compaction process on the fine-grained layer is so efficient that infiltration through this layer is drastically reduced, leading to a bad water regime and environmental drier conditions. Data obtained are in complete agreement with field studies conducted in the northern Negev desert. These studies show that, under the intermittent and infrequent rainstorms prevailing in the area, surface properties play the determinant role in the non-uniform runoff generation and in the redistribution of water resources in space, greatly affecting the whole ecosystem (Yair, 1983; 1994).

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To summarise, although semiarid ecosystems are highly adapted to extreme variations in rainfall, over a time scale of decades and centuries, they seem extremely sensitive to changes in their surface properties, which alter their hydrological regime quickly and efficiently. The time necessary to achieve such a drastic change seems to be very short, at a human rather than a geological time scale. Finally, it would be quite interesting to study, in a similar way, the impact that the deposition of loess and sand had on hydrological and ecological processes at the fringe of cold and glaciated deserts in the northern hemisphere (in Europe as well as in the American continent). Climatic conditions in latter areas differ significantly from those prevailing in the subtropical belt. Such a complementary study would provide a broader and deeper understanding of climatic changes on the environment for various climatic conditions.

Acknowledgements This study was conducted at the Soil Erosion Laboratory of the University of Toronto. The technical help of Mr. Niklaus Kuhn is greatly appreciated. Thanks are due to Mrs. M. Kidron, of the Department of Geography, Hebrew University, for drawing the illustrations.

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Rendell, H. M., Yair, A. and Tsoar H. (1993). Thermoluminescence dating of periods of sand movement and linear dune formation in the northern Negev. In A. C. Millington and K. Pye (eds) The Dynamics and Environmental Context of Eolian Sedimentary Systems. Geological Society Special Publication 72: 69-74. Stibbe, E. (1974). Hydrological balance of Limans in the Negev. Volcani Institute for Agricultural Research. Publication no 304, Beit Dagan, Israel. 35 pp. Schlesinger, W.H., Reynolds, J.F., Cunningham, G.L., Huenneke, L.F., Jarell, W.M., Virginia, R.A. and Shmida, A. (1982). Endemic plants of Israel, Rotem, Bulletin of the Israel Plant Information Centre, 3: 3-47 (in Hebrew). Thiery, R. G. (1982). Environmental instability and community diversity. Biological Reviews, 57: 691-710. Whitford, W.C. (1990). Biological feedbacks in global desertification. Science, 247: 1043-1048. Wiens, A.J. (1985). Vertebrate Responses to Environmental Patchiness in Arid and Semiand Ecosystems, in Pickett, S.T.A. and White, P.S. (eds), The Ecology of Natural Disturbance and Patch Dynamics. Academic Press, NY, pp.: 169-196. Yaalon, D.H. and Dan, J.(1974). Accumulation and distribution of loess-derived deposits in the semi-desert and desert fringe area of Israel. Zeitschrift fur Geomorphologie, Supplement Band 20: 91-105. Yaalon, D.H. and Ganor, E. (1975). Rates of aeolian accretion in the Mediterranean and desert fringe environments of Israel. International Congress of Sedimentology, Nice, France : 169-174. Yair, A. (1983). Hillslope hydrology, water harvesting and areal distribution of some ancient agricultural systems in the northern Negev desert. Journal of Arid Environments, 6: 283-301. Yair, A. (1987). Environmental effects of loess penetration into the northern Negev desert. J. of Arid Environmnets,13: 9-24. Yair, A. (1990). Runoff generation in a sandy area; the Nizzana sands, western Negev, Israel. Earth surface Processes and Landforms, 15: 597-609. Yair, A. (1992). The control of headwater area on channel runoff in a small arid watershed. In Parsons, T and Abrahams, a. (eds), Overland Flow, pp.53-68. Yair, A. (1994). The ambiguous impact of climate change at a desert fringe, Northern Negev, Israel, in Millington, A. C. and Pye, K. (eds). Environmental Change in Drylands: Biogeographical and Geomorphological Perspectives, Chichester, John Wiley and Sons, pp. 199-227. Yair, A. (1999). Spatial variability in the runoff generated in small arid watersheds: implications for water harvesting, in Hoekstra, T. M. and Shachak, M. (eds), Arid Lands Management toward Ecological Sustainability, pp. 212-222. Yair, A. and Danin, A. (1980). Spatial variations in vegetation as related to the soil moisture regime over an arid limestone hillside, Northern Negev, Israel. Oecologia, 47: :83-88. Yair, A. and Shachak, M. ( 1985). Studies in watershed ecology of an arid area. in Berkofsky, L.and Wurtele, G. (eds). Progress in Desert Research, Rowman and Littlefield, New Jersey, pp. 45-193. Yair, A., Lavee, H and Greitser, N. (1997). Spatial and temporal variability of percolation and water movement in a system of longitudinal dunes, western Negev, Israel. Hydrological Processes, 11: 43-58.

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WEATHERING, GEOMORPHOLOGY AND CLIMATIC VARIABILITY IN THE CENTRAL NAMIB DESERT HEATHER VILES and ANDREW GOUDIE School of Geography, University of Oxford, Mansfield Road, Oxford OX1 3TB

Abstract Weathering is an important component of geomorphological change in the Central Namib Desert. Previous studies have reported on the weathering role of salt and dissolution, allied with wind abrasion. However, many surface are covered by luxuriant lichen growths, fed by fog precipitation, whose weathering role has not been clarified. Here we present preliminary investigations of the role of lichens and other rock surface microorganisms in weathering and surface protection, using field observations from a range of sites between 2 and 80 km from the coast, coupled with Scanning Electron Microscope (SEM) observations of lichen:substrate interactions. A model of lichen weathering activity is proposed, illustrating the different roles of lichens on various rock types. Spatial segregation of lichen and other weathering processes is seen to occur at a range of scales. KEY WORDS: Biological weathering; salt weathering; rock-surface microenvironments.

1.

Introduction

The Central Namib Desert, Namibia, Southern Africa, covers some and is one of the driest deserts in the world. It is characterised by gravel plains often underlain by gypsum crusts (Eckardt and Spiro, 1999), interspersed with weathered rock outcrops and numerous dry riverbeds. The area is separated from the Namib sand sea to the south by the Kuiseb River, and from the coastal dunes in the north by the Huab River. Classified as hyper-arid, the area receives very little rainfall, although fog can dramatically increase overall precipitation amounts. Southgate et al., (1996) have analysed climate records over the period 1962-1991 from the desert research station at Gobabeb (23° 34’S, 15° 03’E) and found mean annual rainfall to be 19 mm (range 0 107 mm) and mean annual fog precipitation to be 37 mm (range 14 - 68 mm). Gobabeb is over 70 km from the coast (see figure 1), and areas nearer the sea will experience greater amounts of fog. The geology of the area is dominated by a suite of metamorphic 65

S.J. McLaren and D.R. Kniveton (eds.), Linking Climate Change to Land Surface Change, 65–82. © 2000 Kluwer Academic Publishers. Printed in the Netherlands.

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and igneous rocks from the Nosib and Swakop groups of the Damara System of late Proterozoic age. The Damaran metamorphics are highly variable and include mica schists, marble, granitic gneiss and quartzite. Into these rocks are intruded granites and black dolerite dykes (Goudie, 1972). As in other desert environments with major outcrops of rock and desert pavement, weathering plays a key role in both shaping residual rock outcrops and in producing fine-grained sediment. The nature and rate of weathering within arid environments has been the subject of much debate among geologists and geomorphologists over the years, with early views that deserts represented sterile, and highly restricted weathering environments being gradually replaced by the view that deserts can experience severe (if superficial and selective) weathering (Goudie, 1997; Goudie et al., 1997). However, there is still much debate about how desert weathering processes work, and the controls on them. As Smith (1994, p. 39) puts it: ‘Weathering research is thus not a question of what we know about desert weathering, but what we do not know…’ Much previous desert weathering research has failed to study different mechanisms in association, in an effort to improve knowledge on individual processes.

Desert weathering is often evidenced by flaking and disintegrating rock outcrops, and less commonly by distinctive small rills and pits. It is often difficult to ascribe the genesis of these features to particular weathering processes, or combinations of processes. Several major types of desert weathering processes have been studied in terms of their operation and importance. The role of insolation weathering (essentially caused by heating and cooling of rock surfaces producing steep temperature gradients

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towards the inner part of the rock) in shaping the desert landscape has been much debated (Cooke et al., 1993; Smith, 1994). Micro-environmental conditions (such as aspect) and rock type may be critical in determining how important insolation weathering is. Salt weathering has been proposed by many workers as a major, largely physical, weathering process in deserts, probably playing an important role in producing landforms such as alveole and tafoni (Mustoe, 1982), and as a source of large amounts of fine-grained debris (although chemical weathering processes have also been proposed as important to the formation of tafoni by Young, 1987, and Campbell, 1999). Three important groups of factors control the operation of salt weathering according to Cooke et al., (1993, p. 33), that is environmental and micro-environmental conditions, material properties and the characteristics of the salts themselves. Lichens, algae and other lower plants and microorganisms have been shown to play a significant, if often localised, role in desert weathering. Studies from the Negev (e.g. Danin and Garty, 1983) and cold deserts (e.g. Friedmann, 1982; Sun and Friedmann, 1991) indicate that microorganisms interact with rock surfaces in a range of ways. Firstly, some can bore into rock surfaces (the euendolithic niche, following the terminology of Golubic et al., 1981). Others inhabit cracks (the chasmoendolithic niche) or preformed cavities in the rock (cryptoendolithic niche). Each of these three types can contribute significantly to weathering through the production of small-scale pitting, and aiding the development of surface flaking and disintegration. Even those microorganisms which live purely on the surface (the epilithic niche) or under stones (the hypolithic niche) can alter chemical conditions at the rock surface, encouraging weathering. Lower plants and microorganisms also contribute to surface protection through the formation of biological crusts. Thus, mosses, liverworts and fruticose and foliose lichens growing on soils and rocks increase the wind resistance of the surface reducing erosion and trapping dustblown debris (Danin and Ganor, 1991, 1997). Again, micro-environmental conditions seem to be a very important factor in the nature of microorganic influences on weathering and surface stability. In the Negev Desert for example, Kappen et al., (1980) found that NE facing slopes have a much richer lichen cover than SW facing ones, because of differences in receipt of solar radiation and moisture retention. Chemical weathering may be more important in deserts than has been previously recognised (Smith, 1994) and there have been many reports of microscale karren features on limestones and marbles within deserts which may be produced by small amounts of precipitation given suitable lithological and micro-environmental conditions (Lowdermilk and Woodruffe, 1932; Sweeting and Lancaster, 1982; Smith, 1987). However, it is often difficult to differentiate rillenkarren from wind-hewn fluting, which may also be found on desert rocks. Such wind-eroded forms are likely to be concentrated on slopes facing into prevailing or dominant wind directions, and should occur across a range of lithologies. The foregoing discussion of current knowledge on desert weathering processes and controls indicates that much of the evidence is circumstantial and that although we can make some generalisations about what factors control the operation of weathering processes (often based on laboratory simulations which do not replicate the complex nature of real desert environments), we do not have

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adequate knowledge of how they interact one with another to produce landforms and debris. Associations between weathering processes may be synergistic (i.e. one may encourage the action of another), or one may slow down or prevent the action of others. Such coassociations may operate sequentially, or at the same time. Thus, during previous wetter phases, for example, more intense chemical weathering might have ‘pre-stressed’ the rock leaving it more vulnerable to weathering by salts. Figure 2 shows an example of sequential weathering environments from Swartbankberg, with lichens colonizing a previously exfoliated boulder. The exfoliation may have provided suitable roughened surfaces, thus encouraging lichen growth. On the other hand, lichens, although perhaps producing micro-scale chemical attack under their thalli, might protect rock surfaces from the action of wind and salts. Different weathering processes may also be spatially segregated, because of the varying climatic and micro-environmental controls which determine their operation. Previous work on weathering in the Central Namib Desert has highlighted the importance of salt weathering (Goudie et al., 1997) and dissolution of carbonate minerals (Sweeting and Lancaster, 1982). This paper focuses on biological weathering processes (which have received little previous attention here), and how they relate to other weathering processes. The Central Namib desert has a rich lichen flora (as described by Schieferstein and Loris, 1992 and shown in figure 3), largely supported by fog, along with a range of cyanobacteria and other microorganisms. Lichens and microorganisms are poikilohydric and can take up water from air with relative humidity higher than 70% and most can survive long periods of desiccation (Walter, 1986). The aims of this paper are to investigate three aspects of weathering in the Central Namib Desert, i.e.: 1. Spatial differences in biological and other weathering processes at a range of scales. 2. Co-associations between biological and other weathering processes. 3.

The likely impacts of temporal variability in climate on biological and other weathering processes.

Before considering the detailed evidence of biological weathering it is important to clarify the nature of environmental gradients within the Central Namib Desert.

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Environmental variability in space and time across the central namib desert

The Central Namib Desert stretches over 100 km inland from the coast and is characterised by clear E-W gradients in climate, lichen cover, lichen biomass, altitude and ground type. Rainfall increases markedly inland, whereas fog is highest nearer the coast with some 120 fog days per year at the coast, tailing off to around 40 days at 40 km inland, and 5 fog days at around 100 km from the coast (Olivier, 1995). Altitude increases inland towards the foot of the great escarpment around 80 to 140 km from the coast and peaks around 900m a.s.l. The occurrence of lichens seems to decrease inland with most luxuriant growths found within about 30 km of the coast. Detailed studies in the area around Swakopmund by Schieferstein and Loris (1992) show that maximum lichen coverage occurs around 5 km from the coast and biomass peaks around 1 km from the coast. Gypsum crusts are found predominantly within the coastal zone, and their eastern limit is around 50 - 70 km from the shore at an elevation of 400-500m (Eckardt and Spiro, 1999). The low-lying coastal areas are prone to the accumulation of a wide range of salts at the surface, producing a range of pan forms. Thus, it might be expected that the nature and intensity of weathering will also vary across these gradients because of the different controlling factors influencing the physical, chemical and biological processes thought to operate in this area. There are also much smaller scale variations in the nature and intensity of weathering in the Central Namib Desert, as aspect, lithology, and microclimate vary across the cm - m scale. For example, East-facing slopes are preferentially affected by seasonal dry easterly winds, known as Berg winds, whereas West-facing slopes are more protected. North-facing slopes experience higher levels of incoming radiation, greater fluctuations in temperature, and higher levels of evaporation than south-facing slopes in this southern hemisphere environment. Such microclimatic variability will have ecological and geomorphological impacts. Schieferstein and Loris (1992), in their study of near-coastal lichen fields, found that fruticose and foliose lichens were dominant on SW-facing slopes, whereas crustose (and often euendolithic) species dominated on NE and E-facing slopes. Sweeting and Lancaster (1982) suggested that solutional rillenkarren, and microorganic growths including crustose lichens are found predominantly on west-facing marble surfaces at around 50 km inland, with wind-abraded surfaces on east-facing slopes (see figure 4).

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There is also, at one site, some vertical differentiation of micro-environment and also probably therefore of weathering regime. Thus, rock surfaces at ground level in contact with salt-enriched soils are likely to be prone to salt weathering, whereas rock faces higher above the ground surface may be colonised by lichens, or affected by wind or solution. There are also patterns of temporal variability in climate and environment at a range of scales which will also have impacts upon weathering regimes. The climate of the Namib is influenced by the cold Benguela current off the west coast and by the El Niño Southern Oscillation. Although there are a few meteorological stations within the Central Namib Desert only the one at Gobabeb has a sufficiently long and continuous record to permit analysis of temporal trends. Southgate et al., (1996) found that, over the period 1962-1991, annual rainfall at Gobabeb had a coefficient of variability of 113% whereas that of fog was only 36%. Analysis of the cumulative deviation from the mean by Southgate et al., (1996) showed that in terms of rainfall the period from 1962/3 to 1974/5 was much drier than average, and between 1975/6 and 1978/9 much wetter than average, with below average conditions since then. Fog variability was found to be generally in inverse relation to that of rain, with low fog precipitation between 1974/5 and 1985/6 and then a marked increase after that.

3.

Methods

During two short field seasons in 1994 and 1996 a range of sites was visited across the Central Namib Desert (as shown in figure 1). At each site observations were made on

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the geology and geomorphology, and a range of small rock samples was removed for observation under the Scanning Electron Microscope (SEM). On return to the laboratory, SEM samples were obtained by fracturing the rock with a cold chisel to obtain a cross-sectional view from the surface into the interior of the rock. At two sites, Swartbankberg and Tomato Pan, field estimates of lichen cover were made using a 25 x 25 cm quadrat divided into 5 cm squares. Between 10 and 20 quadrats were randomly located and the % cover of lichens estimated by counting the number of 5 cm squares containing appreciable amounts of lichens.

4.

Results

The geomorphological and ecological characteristics of the sites studied are listed in Table 1. At all sites there was clear evidence of weathering, often in the form of surface flaking and small scale pitting. Lichens were found at all sites, although there was very little lichen cover at Mirabib (80 km from coast) and Gobabeb (70 km from coast). At many sites there was clear evidence of small-scale spatial patterning in weathering microenvironments. Thus, at Tomato Pan upstanding dolerite outcrops were covered with lichen-covered boulders (with crustose, foliose and fruticose types), as were lower outcrops of schist. However, just above the level of the pan surface lichens decreased in number, to be replaced by signs of harsh salt weathering (Goudie et al., 1997). At Swartbankberg, most rock outcrops were covered by a mosaic of crustose lichens, except where the surface was exfoliating rapidly. On the boulder strewn slopes at the base of Swartbankberg there was extensive lichen cover, even on boulders previously subjected to exfoliation (figure 2). At Gobabeb, on finely sculpted granite outcrops with alveoli and tafoni extensively developed, lichen cover was very rare - limited to a few, small green crustose types growing on dark, fine-grained parts of the rock, and on the visors of tafoni. At Vogelfederberg, where the Hamilton mountains marbles outcrop, and also at the Karibib marble outcrop there were clear signs of spatial segregation of weathering micro-environments into east-facing (with wind-blown flutes) and west facing (with brown lichen-pocked surfaces and occasional small rills) as shown in figure 4. On areas of lichen fields near Tomato Pan and the airport at Rooikop, a lush growth of foliose, fruticose and crustose lichens was found with lichens covering boulders and gypsum crusts (figure 3). The only bare patches were along commonly used tracks, around bushes, and at the bottom of ephemeral washes. At all sites there was a general trend for crustose lichens to dominate on outcrop surfaces, with foliose and fruticose lichens only becoming common on boulder-strewn and gravel plains. Some estimates of lichen cover made at Swartbankberg and Tomato Pan on a range of surface types are presented in Table 2. The high standard deviations between individual quadrats on each surface type illustrate the patchiness of cover at this scale. At Swartbankberg the highest mean % cover, and the highest variability, are found on the schist outcrop at the highest point sampled on the profile (c. 350 m a.s.l.), with the lowest mean value at the base of the slope (around 300 m a.s.l.). This would conform to

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the hypothesis that lichens are supported by fog here, as the higher parts of the hill will intercept more fog. At Tomato Pan, on the most stable, least salty surface on the top of the dolerite dyke, very high mean % cover is found (83.1%), with a similar declining trend down to the gravel plain just above the salty pan surface. In this case, salt is likely to be the controlling factor, as fog will be adequate throughout the site. Study of rock and lichen samples collected in the field indicates that lichen thalli of up to a few mm in thickness are commonly found (see table 3), although many lichen growths are extremely thin and in several cases at least partly euendolithic. There are cryptoendolithic growths present on several samples (evidenced by a green line around 1 mm below the surface of the rock), especially more porous and lighter coloured outcrops.

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Three major types of lichen growth are revealed in hand specimen. Firstly, individual epilithic lichen thalli forming circular or near circular patches, sometimes peeling away towards the centre. These growths provide only a patchy cover of the underlying substrate. Secondly, there are mosaics of adjoining eplithic and euendolithic lichen thalli which entirely cover the surface of some areas of rock, boulders or gypsum crust. These growths are often very thin (<< 500 microns) but extensive. Finally, there are dominantly euendolithic growths which are revealed on the surface only as narrow growths along cracks and grain boundaries, but which form an extensive sub-surface layer. Each growth style will have a corresponding impact on weathering. Scanning electron microscope (SEM) reveals more information about the nature of the substrate:lichen interactions as summarised in table 3. Foliose and fruticose lichens generally have only sparse attachments (through rhizines or rootlets) to the underlying surface and have not been investigated further here for any weathering role. Crustose lichens here generally have thalli of up to 1 mm thick growing on top of the rock, and beneath the surface fungal hyphae or whole parts of the lichen thallus may extend down by 1 mm or so.

Several lichens show clear borehole production by fungal hyphae in suitable crystal substrates (usually calcite within marbles) as shown in figure 5a and 5b. Such boreholes have been noted to occur in limestones and marbles throughout the world, with stressed or frequently wetted environments (such as coastal and arid environments and streams) having particularly well-developed features. In general, within the 25 samples studied here, boreholes extend down about into the calcite crystals (a lot less than commonly found on limestone coasts, e.g. by Le Campion-Alsumard, 1979, but more than found on terrestrial limestones on Aldabra Atoll by Viles, 1987). Fungal hyphal boreholes were also rare on other rock types here, notably on gypsum at Swartbankberg,

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where calcification of hyphae was also observed. On schist, dolerite, diorite and granite hyphal boreholes were not found, but there was abundant evidence of lichen thalli penetrating into the rock surface and leading to detachment of grains and flakes (in a similar fashion to the action of Lecidea auriculata on gabbroic boulders on moraines in North Norway observed by McCarroll and Viles, 1995). Figure 6a and 6b illustrate grain detachment under lichens on schist from Swartbankberg. However, this activity appeared spatially very patchy at the small scale, with some lichens not creating any obvious impact on their substrate. On all rock types many of the crustose lichens were peeling away from the surface. Several samples also showed the impact of a cryptoendolithic layer on weathering, as shown in figure 7a and 7b. Here, cryptoendolithic microorganisms within a diorite pebble from Tomato Pan fill pore spaces and appear to be aiding grain detachment.

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From these observations in the field and under the microscope, a model of lichen weathering and surface protecting activity in the area can be put forward (figure 8). Folios and fruticose lichens, because of the nature of their attachment to the rock (limited to point contact) do not create much weathering impact on the underlying surface. In contrast, they play a major role in trapping airborne debris, and in binding together any unconsolidated material on which they grow. Their dominant effect is thus to protect the underlying surface from erosion. Crustose lichens, on the other hand may be dominantly protective or preferentially cause weathering, depending largely on the nature of the underlying rock type and their growth and decay characteristics. Thus, on marble the biochemical action of lichens (forming fungal hyphal boreholes) tends to dominate, whereas on schist, dolerite, diorite and granite they may be important in grain and flake detachment - through what might be called biophysical action, although it probably involves chemical and physical processes of deterioration and disintegration of the underlying rock. On quartz pebbles, lichen weathering activity is generally minimal (although one sample showed some evidence of shallow fungal boreholes). In all three cases, and especially that of the quartz pebbles, the net effect of lichens is probably a protective one, at least while they are still alive. In this hyper-arid environment, lichen cover buffers the underlying rock from wind, changes its thermal response to heating and cooling stresses and will prevent water from making contact with the surface (thus limiting dissolution). However, when lichens die, decay or otherwise are removed from the surface (perhaps by grazing) net erosion will occur, as grains detached from the rock and encased within the lichen thallus will be removed. The overall importance of lichen weathering depends also on the nature of growths (e.g. predominantly euendolithic forms vs. epilithic types), the contribution of any cryptoendolithic community, and the percentage cover of lichens. In order to test this model, and apply a temporal framework, it is necessary to collect more information about the growth rates of the various lichens constituting the lichen mosaics. Furthermore, the impact of lichens on the thermal response of a range of different rock types needs testing. In some cases, lichen surface coloration may encourage insolation weathering by reducing or increasing albedo in comparison with the underlying rock.

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Discussion and conclusions

The model of lichen weathering and surface protection activity proposed above needs to be set within the context of other weathering and erosional processes within the Central Namib Desert, and against the backdrop of environmental variability over space and time at a range of scales discussed earlier. Figure 9 shows the three main processes seen to be sculpting rock surfaces in the Central Namib Desert plotted onto three axes. The bottom axis shows the major environmental gradient from the coastal zone to the inland extremes of the desert, as one moves from the salt and fog-affected coastal environment towards the inland area where wind (and insolation, and possibly rainfall) effects dominate. The side axes depict the smaller-scale gradients found at a site. On the left hand side is what is in reality often a vertical gradient from stable to unstable surfaces, or from sound upstanding rock down onto salty surfaces, as is found at Tomato Pan. On the right hand side is a gradient from benign to harsh, which in reality is a gradient of aspect from west-facing to east-facing (as found on the marble outcrops at Vogelfederberg). The dominance of salt, wind (and insolation) and lichen weathering processes can be mapped onto these axes as shown in figure 9.

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This figure illustrates the generally spatially-segregated nature of the different denudation processes, on both small and large scales reflecting the notable environmental gradients at these scales. It is clear that each of the three denudation domains (i.e. lichen-dominated, salt-dominated and wind-dominated) possess individual characteristics. Thus, the lichen-dominated domain includes a diverse series (as shown in figure 8) of protective and weathering effects, all of which are highly superficial affecting the top few mm at most, and showing clear differentiation in impacts on different rock types. Over time, the rate of weathering and removal of material in the lichen-dominated domain is controlled probably by the nature of fog inputs (which, as Southgate et al., 1996 show, contribute twice the moisture as rainfall here with a third of the variability) which control lichen growth. The salt-dominated domain, however, shows similar selectivity in terms of rock types affected (although for the case of salt weathering, it appears that porosity and easy ingress of water is the key factor predisposing rocks to salt attack, rather than mineralogy in the case of lichen-effects). The depth of weathering caused by salt is in the order of centimetres rather than millimetres. Over time, changes in the sources of salt and groundwater levels are likely to play a key role in determining salt weathering effectiveness. Finally, the winddominated domain also shows selectivity according to rock type (most previous reports of wind flutes have been on marble, e.g. Sweeting and Lancaster, 1982), and great superficiality. Changes in the strength of the Berg wind over time will affect the role played by wind in denudation here.

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The preliminary field and laboratory observations presented above have been used to create simple conceptual models of the role of lichens in weathering, in comparison with the role of other weathering processes. Clear evidence has been presented of a range of impacts of lichens on weathering and surface protection, backing up findings from other desert areas (e.g. the Negev). Lack of long-term climatic records, and the sparsity of meteorological sites within the Central Namib Desert, means that definitive statements about large-scale environmental gradients over space and time are hard to make. The almost complete lack of microenvironmental data (e.g. rock surface temperature and humidity) means that more concrete pronouncements on micro-scale environmental variability are also impossible. Until such data are available it will be difficult to be more precise about the overall role of lichen weathering in geomorphological change in the Central Namib Desert. However, lichens are clearly a neglected component of weathering here which deserve more attention in the future.

Acknowledgements We thank Amy and Alice Goudie for indefatigable field assistance, and Adrian Parker for help with preparing SEM samples. Steve Jones kindly drew the figures.

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Lowdermilk, J.D. and Woodruffe, A.O. (1932) Concerning rillensteine, American Journal of Science series 5, 23, 135-43. McCarroll, D. and Viles, H. A. (1995) Rock-weathering by the lichen Lecidea auriculata in an arctic-alpine environment, Earth Surface Processes and Landforms 20, 199-206. Mustoe, G. E. (1982) The origin of honeycomb weathering, Geological Society of America Bulletin 93, 108-115. Olivier, J. (1995) Spatial distribution of fog in the Namib, Journal of Arid Environments 29, 129-138. Schieferstein, B. and Loris, K. (1992) Ecological investigations on lichen fields on the Central Namib. 1. Distribution patterns and habitat conditions. Vegetatio 98, 113-128. Smith, B. J. (1987) An integrated approach to the weathering of limestone in an arid area and its role in landscape evolution: a case study in S E Morocco, in: Gardiner, V. (ed.) International Geomorphology 1986. Chichester: Wiley, Vol. II, 637-57. Smith, B. J. (1994) Weathering processes and forms, in: Abrahams, A.D. and Parsons, A.J. (eds) Geomorphology of desert environments. London: Chapman and Hall, 39-63. Southgate, R. I., Masters, P. and Seely, M. K. (1996) Precipitation and biomass changes in the Namib Desert dune ecosystem, Journal of Arid Environments 33, 267-80. Sweeting, M. M. and Lancaster, N. (1982) Solutional and wind erosion forms on limestone in the Central Namib Desert, Zeitschrift für Geomorphologie 26, 197-207. Sun, H.J. and Friedmann, E.I. (1991) Long-term ( to years) biogenous weathering and dynamics of microbial growth in Ross Desert sandstone. (Abstract). Transactions, American Geophysical Union 72, 102. Viles, H. A. (1987) Blue-green algae and terrestrial limestone weathering on Aldabra Atoll: an SEM and light microscope study, Earth Surface Processes and Landforms 12, 319-330. Walter, H. (1986) The Namib Desert. In: Evenari, M., Noy-Meir, I. And Goodall, D.W. Hot deserts and arid shrublands B, Ecosystems of the World 12B. Amsterdam: Elsevier, 245-282. Young, A. (1987) Salt as an agent in the development of cavernous weathering, Geology 15, 962-966.

WARM SEASON LAND SURFACE – CLIMATE INTERACTIONS IN THE UNITED STATES MIDWEST FROM MESOSCALE OBSERVATIONS J. O. ADEGOKE and A. M. CARLETON Department of Geography and Earth System Science Center, The Pennsylvania State University, University Park PA 16802, U.S.A.

Abstract The United States Midwest over the last two decades has experienced marked warm season climate anomalies, including droughts and major floods. While the development of these extreme events can usually be traced to anomalies in atmospheric circulation, and may include teleconnections, studies based on model simulations have shown that land surface forcing may be partly responsible for the persistence of these climate anomalies. This study evaluates the presence and strength of long-term land surface-climate interactions in the U.S. Midwest. We do this via an analysis of the cross-seasonal (spring and summer) associations between temperature and moisture (Palmer Drought Severity Index-PDSI, Crop Moisture-Z Index, and precipitation) anomalies. Direct and lag correlations for the 18951995 and 1948-1995 periods show that warm and dry summers tend to follow warm spring seasons. These results imply that springtime precipitation anomalies may help to determine the temperature regime of the following summer, possibly via the moisture content of the upper soil. We also show that broad land cover types tend to modulate summer climate anomalies in the U. S. Midwest.

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Over the last two decades, there has been great interest in understanding the nature and causes of interactions among the biosphere, oceans, ice and atmosphere. In particular, the coupling between the terrestrial biosphere and the atmosphere has been shown to operate on time and space scales ranging from hours to decades and from plot level to regional and even global scales (Yeh et al., 1984; Pielke et al., 1993). 83

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Anthropogenic activities that modify terrestrial vegetation, such as large scale deforestation for logging and agricultural purposes, urban expansion and industrial activities such as strip mining, are now known to affect local to global climate conditions by altering the surface energy and moisture budgets and, thereby, the characteristics of the planetary boundary layer (PBL). The primary physical processes involved in these vegetation-climate interactions are changes in surface albedo, soil heat flux, roughness length, and the partitioning of sensible to latent heat fluxes (or Bowen ratio), as they influence moisture and convection in the PBL. Early evidence of these effects came from studies of the devastating Sahel droughts of the early 1970s, which increased concerns about desertification in marginal areas (Hammer, 1970; Charney et al., 1977; Yeh et al., 1984). Our understanding of the scale interactions of land-atmosphere processes over different surfaces has been derived, in particular, from intensive field experiments in drier areas and for generally restricted time periods (Li and Avissar, 1994; Bonan et al., 1993). These include the International Land Surface Climatology Project (ISLSCP), HAPEX-MOBILHY (Hydrologic Atmospheric Pilot Experiment and Modelisation du Bilan Hydrique), HAPEX-Sahel and the southwest Australian “Bunny Fence” Experiment, or BUFEX (e.g., Andre et al., 1986; Sellers et al., 1992; Lyons et al., 1993; Goutorbe et al., 1989). Observational data from these field experiments have proved very useful for calibrating model simulations of interactions between the Earth’s land surface and the atmosphere. In particular, these calibrations have provided vegetation and soil property information for different land cover types, and have led to improvements in the fidelity with which surface biosphere models reproduce vegetation-climate interactions. However, the results of the observational studies are not directly transferable to more humid areas having different vegetation and soil moisture, or ambient atmospheric, conditions (e.g., Gibson and Vonder Haar, 1990; Raymond et al., 1994; Travis, 1997). Moreover, they are not easily "scaled up" to a region the size of a typical synoptic system (Shuttleworth, 1991; Molders and Raabe, 1996). Following the earlier field experiments, similar studies have been conducted or are planned for other regions, for example, NOPEX (Northern Hemisphere Climate-Processes Land-Surface Experiment) and LBA (The Large scale Biosphere-Atmosphere Experiment in Amazonia). These experiments are coordinated under the Biosphere Aspects of the Hydrological Cycle (BAHC) research initiative of the International Geosphere Biosphere Program (IGBP). These research initiatives are designed to generate new knowledge needed to understand the climatological, ecological, biogeochemical and hydrological functioning of the Earth system and the impact of land use change on these functions. They underscore the need for continued observational studies in humid mid-latitude regions. In this chapter, we present our observational results to date on the role of land surface conditions in the warm season (April through September) climate of the humid lowland region of the Midwest U.S. centered on the "Corn Belt". This area encompasses the five core

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Midwest States of Illinois, Iowa, Indiana, Ohio and Missouri, and the adjoining states of Michigan, Wisconsin, Minnesota and Kentucky (Figure 1). This economically vital agricultural region is susceptible to substantial interannual and interdecadal variations in summer climate. The severe drought of 1988 and devastating floods of 1993 are recent examples of extreme summertime climate anomalies that affected much of the Central U.S. These two events resulted in an estimated $52 billion in farm losses and property damage in the Midwest (Lott, 1993). Our recent and ongoing research efforts have focused on the interrelations between Midwest land cover (primarily vegetation, evapotranspiration: ET, surface and soil moisture) and climate parameters (temperature, rainfall and convective cloudiness) and their expression across a range of spatial and temporal scales. These investigations are based on the analysis of long term climate division and digital land cover data, used in conjunction with satellite data (the Advanced Very High Resolution Radiometer-AVHRR derived vegetation index, the Geosynchronous Operational Environmental Satellite-GOES). We used these datasets to determine land surface - climate interactions for sub-areas and time periods characterized by contrasting vegetation status, surface moisture and atmospheric conditions. 2.

Midwest Land Surface-Climate Associations from Historical Data

Over the last two decades, the U.S. Midwest (Fig. 1) has experienced marked warm season climate anomalies, including drought and major floods (Changnon and Kunkel, 1992; Lozano-Garcia et al., 1995). While the development of these extreme events can usually be traced to anomalies in atmospheric circulation, and may include remote forcing via teleconnections, studies have suggested that land surface forcing may be partly responsible for the persistence of these climate anomalies (McNab, 1989; Namias, 1991). Specifically, surface moisture availability has been implicated in the prolonged persistence of recent anomalous wet and dry spells in the Midwest (Betts et al., 1994; Wetzel et al., 1996). In approaching the role of land surface conditions in recent Midwest U.S. climate anomalies, we conducted an analysis of the cross-seasonal associations between temperature and moisture (precipitation, Palmer Drought Severity Index: PDSI, Crop Moisture or Z-index) anomalies for the 1895-1995 period and also sub-periods. The primary data set used is the monthly state climate division temperature, precipitation and drought indices (Palmer Drought Severity Index-PDSI and Crop Moisture, or Z Index) for the full period, available from the National Climatic Data Center (NCDC). The PDSI is a drought index based on the concept of the water balance, and was developed to use monthly temperature and precipitation as input data. The Palmer model uses parameters for evapotranspiration (calculated using the Thornthwaite water balance model), and parameters for soil-moisture recharge, runoff, and water capacity of the soil. The values given by the index range from -7 to +7, with negative values denoting dry spells and positive values indicating wet spells. The

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monthly Crop Moisture Index (Z) is used to assess short-term crop needs versus available water in the upper 5-feet (1.52 meters) of the soil profile, and is a measure of the departure from normal of the moisture climate for a given month (Karl, 1986). Of the two indices, the PDSI is slower to respond to changes in environmental conditions owing to the inclusion of a lag term from the previous month. Therefore, the PDSI is more indicative of climatological drought severity. Other data sets include gridded daily 700mb height data and anomaly charts for the U.S. (1948-1993), and U.S. crop moisture index charts for selected case study years. The tropospheric height data are used to identify the surface climate associations with synoptic scale atmospheric circulation.

The cross-seasonal (spring and summer) associations between the anomalies of temperature and moisture (precipitation, PDSI and Z-index) were determined via direct and lag correlations for the 1895-1995 and 1948-1995 periods. Spring (MAM) and Summer (JJA)

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seasonal averages for each climate variable were first calculated and then aggregated for five core states (Illinois, Iowa, Missouri, Ohio and Indiana) and also for the entire Midwest region (five core states plus Michigan, Wisconsin, Minnesota and Kentucky). This is to check on the stability of the correlations so derived. The results are similar to those identified in studies for other regions, and the U.S. as a whole (Namias, 1960, 1991; Karl, 1986; Huang et al., 1996). They show, in particular, that warm and dry summers tend to follow warm spring seasons (Table 1).

The correlations for the expanded region (Table 2) show very little sensitivity to domain size. In both cases (i.e., the core and expanded Midwest region) there are fewer significant correlations in the recent time period, although they are all of the same (negative) sign. These results imply that springtime precipitation anomalies may help to determine the temperature regime of the following summer, possibly via the moisture content of the upper soil.

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To address the extent to which the rainfall-temperature associations so revealed are due to land surface forcing (e.g., soil moisture), or to persistence of atmospheric circulation patterns across seasons (e.g., Simmonds, 1993), we classified surface and mid-tropospheric (500mb) synoptic circulation over the Midwest using manual methods (Arnold, 1994), for the spring (MAM) and summer (JJAS) seasons of the period 1948-95. We used a classification scheme that stratified the monthly-averaged 500mb height charts into six synoptic categories: strong zonal flow, weak zonal flow, strong meridional flow, weak meridional flow, cyclonic circulation and anticyclonic circulation. We also noted the presence or absence of storms in the study area for individual months during the 1948-95 period with a storm track index derived by simply tallying the number of storms recorded for each month. The results of this analysis show that during the wetter spring months, the circulation regime was mostly cyclonic with a strong zonal mean flow configuration. The circulation regime during the warmer summer months, on the other hand, was predominantly anticyclonic and meridional. This implies that, in the mean, monthly to seasonal large-scale circulation tends to be the major influence on the warm season climate of the Midwest. Notwithstanding this result, large-scale circulation does not fully account for the observed summer climate surface conditions in all situations, particularly during anomalously wet or dry years. Composite summer circulation indices for the eight driest (1953, 1957, 1960, 1967, 1976, 1983, 1988, 1991) and six wettest (1961, 1972, 1977, 1986, 1992, 1993) years show that the cross-seasonal coherence of surface climate anomalies indicated in Tables 1 and 2 above (e.g., dry spring and warm dry summer following) are maintained across different circulation patterns. For example, the 1988 severe summer drought persisted for several weeks following the flattening of a strong ridge that had stalled over the Midwest in early to mid summer (Figure 2 a, b, c). Throughout the central Midwest, abnormally dry soil moisture conditions prevailed throughout July and August as PDSI values stayed between – 4 and –7 (Figure 3 a, b, c), despite the switch in circulation around mid-July. The reduced soil moisture is likely to have helped to maintain or even to amplify the 1988 summer drought by enhancing surface sensible heating, and reducing the local evaporative contribution to the atmosphere (Kunkel, 1989).

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Surface Heterogeneity and Midwest Climate Anomalies

To further define the role of mesoscale land surface conditions in the warm season climate dynamics of the U.S. Midwest, we address the issue of whether the strength of the associations between land surface conditions and summer climate anomalies differ among cover types, and what may have happened to these associations as the boundaries between major types have evolved over the last century with increased human activities. We do this by stratifying the Midwest climate division temperature and surface moisture data according

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to dominant land cover types. Midwest climate divisions were stratified into one of five broad land cover categories and the 100-year NCDC climate data were re-analyzed on the basis of this stratification. The land cover categories are as follows: cropland, mixed cropland/urban, mixed cropland/forest, mixed forest/cropland/urban, and forest. The Midwest land use/land cover (LU/LC) map used to stratify the climate divisions was produced from the online LU/LC 1:250,000 digital data provided by the United States Geological Survey (USGS). The USGS compiled the LU/LC maps from aerial photographs acquired by the US National Aeronautic Space Administration (NASA) high-altitude missions during the 1980s. Secondary sources from earlier land use maps and field surveys were also incorporated into the LU/LC maps as needed. The maps were subsequently digitized to create a national digital LU/LC database. The USGS mapped and coded land use in the 1:250,000 quadrangles using the Anderson classification system (Anderson et. al., 1976), for levels one and two. The data, originally provided as quadrangles in the geographic information retrieval and analysis system (GIRAS) format, were converted into ARC/INFO export format and imported into the ARC/INFO software environment. The quads were then edge-matched and combined to produce a georeferenced digital LU/LC map of the Midwest (Fig. 4). A climate division map of the Midwest was overlain onto the LU/LC map and each division assigned into one of the five land use types, based on the dominant land cover type associated with particular climate divisions. Five-year running means of the summer temperature and Z-index time series (standardized anomalies) for the five broad land cover types (Figures 5 and 6) show significant differences between forest and the other land cover types. The major difference is the damping of the amplitudes of the anomalies of both climate variables (i.e., temperature and Z-index) in the forest regions. Cross correlation of the time series shows that forest correlates very poorly with all other land cover types (see the highlighted numbers on Table 3). Similar results were obtained for seasonal precipitation and PDSI anomalies (not shown). It is plausible that factors other than land cover, such as topography, may be partly responsible for this lack of correlation between climate variables for forest and other land cover types, although topography is likely to be a minor factor in the Midwest “flatlands”. This result suggests that the conversion of forest area to other land cover types, such as cropland and urban surfaces, may enhance the magnitude of interannual anomalies in summer-time precipitation and temperature.

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Conclusion

Land surface-climate interactions in the warm season have been relatively little studied for humid areas of middle latitudes, in contrast with those areas that are more marginal in terms of rainfall. The humid lowlands of the Midwest U.S. “Corn Belt” comprise a close to ideal laboratory for undertaking such an analysis, given the lack of appreciable topographic change over wide areas; the variety of natural and human-made land covers, and their heterogeneity on a range of spatial scales; the summertime peak in precipitation occurrence; the marked interannual variability of climate; and the economic significance of this dominantly agricultural region. Our preliminary studies, documented in this chapter, suggest that human modifications to the land surface of this region are detectable in the climate record of almost the last 100 years. These impacts are most evident when the prevailing synoptic situation is characterized by weak wind flow near the surface and aloft on daily time scales, and seasonally, during times of more extreme anomalies of precipitation (especially droughts). The availability of high resolution satellite data, GIS, and methodologies for multi-scale analysis such as fractal techniques, offer promise of quantifying the relationships between Midwest land surface conditions (land cover, soil moisture) and a range of climate and atmospheric indices including convective clouds. A key area of application of these techniques is in assessing the scale invariance of surface-atmosphere relationships across

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given scale ranges. These analyses, which we are currently undertaking, go beyond the simpler, yet still necessary, linear correlation studies that show a land cover-climate association over the U.S. Midwest region. Important questions, which we will soon be able to answer, include the following: Over what scale range(s) does the land surface convective cloud relationship hold? What is the relative importance of surface heterogeneity and atmospheric conditions in determining when and where convective clouds develop in the Midwest? The answers to these more immediate questions will enable the role of land surface–atmosphere interactions in the Midwest climate variations to be assessed, and should better permit the role of land surface–atmosphere interactions to be considered in short- to medium-range weather predictions for the Midwest. They should also help to inform modelers seeking to improve the parameterization of important surface processes in GCMs (e.g., the areal integration of energy flux point measurements), and also regional meteorological models for more humid locations.

Acknowledgements The authors wish to thank two anonymous reviewers for their valuable comments and suggestions on the manuscript. The support of Penn State’s Earth System Science Center (now Environment Institute) is also gratefully acknowledged. Partial funding for this research came from NSF grant ATM 98-76753. Jason Allard assisted in preparing some of the graphics included in this paper. References Anderson, J.R., Hardy, E.E., Roach, J.T. and Witmer, R. E., (1976). ‘A land use and land cover classification system for use with remote sensing data’ U.S. Geological Survey Professional Paper, 964, 28p. Andre, J.-C., Goutorbe, J.-P. and Perrier, A., (1986). ‘HAPEX-MOBILHY: a hydrological atmospheric experiment for the study of water budget and evaporation flux at the climatic scale’ Bull. Amer. Meteor. Soc., 67, 138-144. Arnold, D.L., (1994). Synoptic and Mesoscale Climatologies of Severe Local Storms for the American Midwest Ph.D. dissertation, Dept. of Geography, Indiana University, Bloomington. 257p. Betts, A. K., Hall, J.H., Beljaars, A.C.M., Miller, M.J., and Viterbo, P.A., (1994). ‘Coupling between land surface boundary layer parameterizations and rainfall on local and regional scales: Lessons from the wet summer of 1993’ Preprints, Fifth Symp. On Global Change Studies, Nashville, Amer. Meteor. Soc., 174-181 Bonan, G.B., Pollard, D. and Thompson, S.L., (1993). ‘Influence of subgrid-scale heterogeneity in leaf-area index, stomatal resistance, and soil moisture on grid-scale land-atmosphere interactions’ J. Climate, 6. 1882-1897. Changnon, S.A. and Kunkel, K.E., (1992). ‘Assessing impacts of a climatologically unique year (1990) in the Midwest’ Phys. Geog., 13, 180-190. Charney, J.G., Quirk, W.J., Chow, S.H., and Kornfield, J., (1977). ‘A comparative study of the effects of albedo change on drought in semi-arid region’ J. Atmos. Sci., 34, 1336-1385. Gibson, H.M. and Vonder Haar, T.H., (1990). ‘Cloud and convection frequencies over the southeast United States as related to small-scale geographic features’ Mon. Weath. Rev., 118, 2215-2227.

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Goutorbe, J.-P., Noilhan, J., Valancogne, C. and Cuenca, R.H., (1989). ‘Soil moisture variations during HAPEXMOBILHY’ Annals. Geophysics., 7, 415-425. Hammer, R.M., 1970. ‘Cloud development and distribution around Khartoum’ Weather, 28, 411-414. Huang, J., van den Dool, H. and Georgakakos, K.P., (1996). ‘Analysis of model-calculated soil moisture over the United States (1931-1993) and applications to long-range temperature forecasts’ Journal of. Climate, 9, 13501362. Karl, T.B., (1986). ‘The relationship of soil moisture parameterizations to subsequent seasonal and monthly mean temperature in the United States’ Mon. Weath. Rev., 114, 675-686. Kunkel, K.E., (1989). ‘A surface energy budget view of the 1988 Midwestern United States drought’ Bound.-Layer Meteorology, 48, 217-225. Li, B. and Avissar, R., (1994). ‘The impact of spatial variability of land surface characteristics on land-surface heat fluxes’ Journal of. Climate, 7, 527-537. Lott, N., (1993). ‘The summer of 1993: Flooding in the Midwest and drought in the Southeast’ National Climatic Data Center Research Customer Service Group Technical Report, 93-04. 21 p. Lozano-Garcia, D.F., Fernandez, R.N., Gallo, K.P. and Johannsen, C.J., (1995). ‘Monitoring the 1988 severe drought in Indiana, U.S.A. using AVHRR’ International. Journal of Remote Sensing, 16, 1327-1340. Lyons, T.J., Schwerdtfeger, P., Hacker, J.M., Foster, I.J., Smith, R.C.G. and Xinmei, H., (1993). ‘Land-atmosphere interaction in a semiarid region: the bunny fence experiment’ Bulletin of the American Meteorological Society, 74, 1327-1334. McNab, A. L., (1989). ‘Climate and drought’ EOS, Transactions of the American Geophysics Union, 70, 882-883. Molders, N. and Raabe, A., (1996). ‘Numerical investigations on the influence of subgrid-scale surface heterogeneity on evapotranspiration and cloud processes’ Journal of Applied Meteorology, 35, 782-795. Namias, J., (1960). ‘Factors in the initiation, perpetuation and termination of drought’ Extract of Publ. No. 51, IASH Commission of Surface Waters, 81-94. [Avail. from IASH, UNESCO, Paris]. Namias, J., (1991). ‘Spring and summer 1988 drought over the contiguous United States-causes and prediction’ Journal of Climate, 4, 54-65. Pielke, R.A., Dalu, G.A., Lee, T.J., Rodriguez, H., Eastman, J. and Kittel, T.G.F., (1993). ‘Mesoscale parameterization of heat fluxes due to landscape variability for use in general circulation models, In: Exchange Processes at the Land Surface for a Range of Space and Time Scales (Proc. of the Yokohama Symp., July 1993). IAHS Publ. No. 212. 331-342. Raymond, W.H., Rabin, R.M. and Wade, G.S., (1994) Evidence of an agricultural heat island in the lower Mississippi river floodplain’ Bulletin American Meteorological Society, 75, 1019-1025. Sellers, P.J., Hall, F.G., Asrar, G., Strebel, D.E., and Murphy, R.E., (1992) ‘An overview of the First International Satellite Land Surface Climatology Project (ISLSCP) Field Experiment (FIFE)’ Journal of Geophysical Research, 97, 18,345-18,371. Shuttleworth, W.J., (1991). ‘Insight from large-scale observational studies of land/atmosphere interactions’ In: Wood, E.F. (ed.) Land Surface-Atmosphere Interactions for Climate Modeling: Observations, Models and Analysis. Kluwer, Dordrecht. p. 3-20. Travis, D.J., (1997). ‘An investigation of Wisconsin’s anthropogenically-generated convergence boundary and possible influences on climate’ Wisconsin Geogr., 12, 34-46. Wetzel, P.J., Argentini, S., and Boone, A., (1996) ‘Role of land surface in controlling daytime cloud amount: two case studies in the GCIP-SW area’ J. Geophys. Res., 101, 7359-7370. Yeh, T-C., Wetherald, R.T. and Manabe, S., (1984) ‘The effect of soil moisture on the short-term climate and hydrology changea numerical experiment’ Mon. Weath. Rev, 112, 474-490.

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STREAMFLOW CHANGES IN THE SIERRA NEVADA, CALIFORNIA, SIMULATED USING A STATISTICALLY DOWNSCALED GENERAL CIRCULATION MODEL SCENARIO OF CLIMATE CHANGE

ROBERT L. WILBY Division of Geography University of Derby, Kedleston Road, Derby, DE22 1GB, UK @National Center for Atmospheric Research Boulder, Colorado, 80307-3000, USA MICHAEL D. DETTINGER U.S. Geological Survey Water Resources Division, California Scripps Institution of Oceanography 9500 Gilman Drive, La Jolla, California, 92093-0224

Abstract Simulations of future climate using general circulation models (GCMs) suggest that rising concentrations of greenhouse gases may have significant consequences for the global climate. Of less certainty is the extent to which regional scale (i.e., sub-GCM grid) environmental processes will be affected. In this chapter, a range of downscaling techniques are critiqued. Then a relatively simple (yet robust) statistical downscaling technique and its use in the modelling of future runoff scenarios for three river basins in the Sierra Nevada, California, is described. This region was selected because GCM experiments driven by combined greenhouse-gas and sulphate-aerosol forcings consistently show major changes in the hydro-climate of the southwest United States by the end of the 21st century. 99

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The regression-based downscaling method was used to simulate daily rainfall and temperature series for streamflow modelling in three Californian river basins under currentand future-climate conditions. The downscaling involved just three predictor variables (specific humidity, zonal velocity component of airflow, and 500 hPa geopotential heights) supplied by the U.K. Meteorological Office couple ocean-atmosphere model (HadCM2) for the grid point nearest the target basins. When evaluated using independent data, the model showed reasonable skill at reproducing observed area-average precipitation, temperature, and concomitant streamflow variations. Overall, the downscaled data resulted in slight underestimates of mean annual streamflow due to underestimates of precipitation in spring and positive temperature biases in winter. Differences in the skill of simulated streamflows amongst the three basins were attributed to the smoothing effects of snowpack on streamflow responses to climate forcing. The Merced and American River basins drain the western, windward slope of the Sierra Nevada and are snowmelt dominated, whereas the Carson River drains the eastern, leeward slope and is a mix of rainfall runoff and snowmelt runoff. Simulated streamflow in the American River responds rapidly and sensitively to daily-scale temperature and precipitation fluctuations and errors; in the Merced and Carson Rivers, the response to the same short-term influences is much less. Consequently, the skill of simulated flows was significantly lower in the American River model than in the Carson and Merced. The physiography of the three basins also accounts for differences in their sensitivities to future climate change. Increases in winter precipitation exceeding +100% coupled with mean temperature rises greater than +2°C result in increased winter streamflows in all three basins. In the Merced and Carson basins, these streamflow increases reflect large changes in winter snowpack, whereas the streamflow changes in the lower elevation American basin are driven primarily by rainfall runoff. Furthermore, reductions in winter snowpack in the American River basin, owing to less precipitation falling as snow and earlier melting of snow at middle elevations, lead to less spring and summer streamflow. Taken collectively, the downscaling results suggest significant changes to both the timing and magnitude of streamflows in the Sierra Nevada by the end of the 21 st Century. In the higher elevation basins, the HadCM2 scenario implies more annual streamflow and more streamflow during the spring and summer months that are critical for water-resources management in California. Depending on the relative significance of rainfall runoff and snowmelt, each basin responds in its own way to regional climate forcing. Generally, then, climate scenarios need to be specified – by whatever means – with sufficient temporal and spatial resolution to capture subtle orographic influences if projections of climate-change responses are to be useful and reproducible.

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Introduction

Simulations of future climate using general circulation models (GCMs) suggest that rising concentrations of greenhouse gases may have significant consequences for the global climate. Of less certainty is the extent to which regional scale (i.e., sub-GCM grid) environmental processes will be affected. This is because the length scales of GCMs (which are typically about 200 kilometres) are too coarse to resolve complex orography and important sub-grid scale processes such as convective precipitation. Furthermore, GCM output representing the surface climate under current conditions is commonly unreliable at the scale of individual grid points (see Table 1 below). Ironically, these are the scales that are likely to be of greatest interest to resource managers who have functional responsibilities that cover relatively small geographical areas. In other words, there is a scale mismatch between the scale of global change scenarios and the data requirements of the impacts analyst (Hostetler, 1994). “Downscaling” techniques have subsequently emerged as a means of bridging the gap between what climatologists currently are able to supply and what regional climate-change impact studies require. In this chapter, a range of downscaling techniques are critiqued. Then a relatively simple (yet robust) statistical downscaling technique and its use in the modelling of future runoff scenarios for three river basins in the Sierra Nevada, California, is described. This region was selected because GCM experiments driven by combined greenhouse-gas and sulphateaerosol forcings consistently show major changes in the hydroclimate of the southwest United States by the end of the 21st century. Shown in Figure 1, for example, are projected changes in winter (DJF) precipitation from the Canadian Centre for Climate Modelling and Analysis (CGCM1) and by the U.K. Meteorological Office's Hadley Centre for Climate Prediction and Research (HadCM2) transient climate-change simulations (Flato et al., 1999; Boer et al., 1999a,b; Johns et al., 1997; Mitchell and Johns, 1997). Both experiments have been central to the U.S. National Assessment of the Potential Consequences of Climate Variability and Change (see http://www.nacc.usgcrp.gov/), and both predict increases in winter precipitation over California by 2090-99. Although the scenarios in Figure 1 (along with accompanying temperature increases) imply major changes in regional snowpack, snowmelt, and runoff, confidence in HadCM2 scenarios at the basin scale is low. For example, Table 1 shows significant differences between observed and HadCM2-derived precipitation statistics for the heterogeneous landscape of the San Juan basin, Colorado. This deficiency is addressed herein by exploiting observed correlations between climate variables at the GCM grid-scale (such as geopotential height fields) and daily weather at the station scale (such as single-site precipitation). These empirical relationships are used to project future changes in atmospheric circulation and

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humidity in the HadCM2 climate-change scenarios to the station scale. A hydrological model is then used to simulate streamflows in each basin under the downscaled current- and future-climate conditions.

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Downscaling Techniques

The theory and practice of downscaling has been reviewed elsewhere (see Giorgi and Mearns, 1991; Wilby and Wigley, 1997; Wilby et al., 1998b). Therefore, we provide only a brief overview of the main downscaling approaches, namely (a) dynamical, (b) weather typing, (c) stochastic, and (d) regression-based methods. 2.1 DYNAMICAL Dynamical downscaling includes the nesting of a high-resolution regional climate model (RCM) within a GCM (Christensen et al., 1997; McGregor, 1997). The RCM uses the GCM to define time-varying atmospheric boundary conditions around a finite domain, within which the physical dynamics of the atmosphere are modelled using horizontal grid spacings of 20-50 km. The main limitation of RCMs is that they are as computationally demanding as GCMs (placing constraints on the domain size, number of experiments, and duration of simulations). However, RCMs can better resolve smaller scale atmospheric features, such as orographic precipitation, than the host GCM (Jones et al., 1995) and are able to respond in physically consistent ways to different external forcings such as land-surface or atmospheric-chemistry changes (Giorgi and Mearns, 1999). 2.2 WEATHER TYPING Weather-typing approaches involve the stratification of local meteorological variations by concomitant, synoptic-scale (1000 km) atmospheric circulation patterns (e.g., Hay et al., 1992; Matyasovsky et al., 1994). Future regional climate scenarios are then constructed by resampling observed variables from probability distributions conditioned on synthetic series of circulation patterns (e.g., Bardossy and Plate, 1992; Dettinger and Cayan, 1992; Goodess and Palutikof, 1998). The main appeal of circulation-based downscaling is that it is founded on sensible linkages between climate on the large scale, which GCMs are best suited to

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project, and weather at the local scale. The technique is also readily applicable to a wide variety of environmental variables and can preserve some of the spatial auto-correlation between multiple sites and multiple variables (e.g., precipitation and temperature). However, weather-typing schemes commonly are parochial, have difficulty simulating extreme events, and must assume stationary circulation-to-surface climate conditioning (Wilby, 1997). Precipitation scenarios produced by circulation changes alone are also relatively insensitive to future climate forcing (see Wilby et al., 1998b). 2.3 STOCHASTIC The most popular stochastic downscaling approach involves modifying parameters in conventional weather generators such as Richardson’s (1981) Weather-GENeration program (WGEN). The standard WGEN program simulates precipitation occurrence using a twostate, first-order Markov chain; precipitation amounts on wet days using a gamma distribution; and temperature and radiation components using first-order trivariate autoregression that is conditional on precipitation occurrence. Future-climate scenarios are generated stochastically using revised parameter sets that have been scaled in direct proportion to the corresponding variable changes in a GCM (Wilks, 1992). The main advantage of the technique is that it can exactly reproduce key climate statistics and has been widely used for climate-impact assessment (e.g., Mearns et al., 1996). The key disadvantages relate to the arbitrary manner in which model parameters are changed for future-climate conditions, to the unanticipated effects that these changes can have on conditional variables (Katz, 1996), and to the poor representation of interannual variability in stochastic models (Gregory et al., 1993).

2.4 REGRESSION Regression-based downscaling methods employ empirical relationships between local scale/single-site predictand(s) and synoptic-scale predictor(s). Techniques differ according to the choice of mathematical transfer function, predictor variable suite, or statistical-fitting procedure. Methods include linear and non-linear regression, artificial neural networks, canonical correlation, and principal components analyses (e.g., Conway et al., 1996; Crane and Hewitson, 1998; von Storch et al., 1993). The main strengths of regression downscaling are the relative ease of application and the parsimony of the models. However, regression models typically explain only a fraction of the observed climate variability. In common with weather-typing methods, stationarity of the empirical relationships is also assumed, and downscaled scenarios can be sensitive to the choice of predictor variables and regression method (Winkler et al., 1997). Still, regression models provide an efficient compromise

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between simpler, purely stochastic weather generators and computationally expensive, dynamical models. Regressions are inexpensive to apply but are able to reproduce physically realistic intervariable, temporal, and spatial relationships as well as sequences in predicted fields that are present in historical records. 3. Data and Modelling Methods In the remainder of this chapter, we describe the application of a regression-based downscaling model to streamflow simulation under current- and future-climate scenarios. Two sets of GCM output were used: the first to calibrate and then verify the coupled downscaling-hydrological model, and the second to downscale GCM output in order to simulate future streamflow in the Sierra Nevada. 3.1 PREDICTOR VARIABLES Table 2 lists 15 candidate variables that were originally selected by Wilby et al. (1999) for possible use as downscaling predictors. All variables were derived from combinations of daily grid-point estimates of mean sea level pressure (mslp); 500 hPa geopotential height (H); 2-metre (near-surface) temperature (T2m); and 0.995-sigma-level (near-surface) relative humidity (RH), obtained from the National Center for Environmental Prediction / National Center for Atmospheric Research (NCEP/NCAR) Reanalysis (Kalnay et al., 1996) of atmospheric observations for the period 1979 to 1995. The Reanalysis estimates were regridded from the NCEP grid (1.875° of latitude by 1.875° of longitude) to the 2.5° of latitude by 3.75° of longitude grid on which climate variations are represented in the HadCM2 simulations. The pressure data were used to calculate five daily airflow indices (U, V, F, Z, D) for both the surface (s) and the 500-mb level in the upper (u) atmosphere, according to a methodology described by Jones et al. (1993). Daily mean temperatures and relative humidities were used to estimate daily mean specific humidities (SH) by Richards' (1971) non-linear approximation.

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The GCM used to drive the downscaling model in climate-change experiments was the U.K. Meteorological Office Hadley Centre's coupled ocean/atmosphere model (HadCM2) forced by combined and albedo (as a proxy for sulphate aerosol) changes (Johns et al., 1997; Mitchell and Johns, 1997). In this sulphate-plus-greenhouse gas experiment (their "SUL" experiment), the model run begins in 1861 and is forced with an estimate of historical radiative conditions to 1993 followed by a projected future-forcing scenario with 1% increases in and sulphate per year from 1994 to 2100. HadCM2 output for the period 1980-99 was used as a proxy for the current climate (as in previous downscaling studies, such as Conway et al., 1996; Pilling et al., 1998; Wilby et al., 1998a,b). Output for 2080 to 2099 was used to downscale climate conditions arising from future anthropogenic emissions of greenhouse gases and aerosols. 3.2 STATISTICAL DOWNSCALING MODEL The statistical downscaling model (Wilby et al., 1999) was calibrated by regressions linking selected Reanalysis grid-point values as independent predictor variables with daily weather data for seven stations in or near the North Fork American, East Fork Carson, and Merced River basins (Figure 2) in the Sierra Nevada as dependent variables. The specific predictands for which regression models were fitted are the daily series of wet-day occurrence (O), wet-day amounts (R), and maximum (TMAX) and minimum (TMIN) temperatures. Regression relations were fitted on daily variables for the 10 years from 1979 to 1988 and were evaluated using the 7 years from 1989 to 1995. Separate regressions were

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undertaken for each station and each or the climatological seasons, winter (DJF), spring (MAM), summer (JJA), and autumn (SON).

All daily predictor variables were normalised using period means and standard deviations (as advocated by Karl et al., 1990) to increase transferability to GCM simulations, which may have different means and standard deviations from observed fields. The three most powerful predictor variables were selected following a step-wise multiple linear regression analysis of the 15 candidate variables listed in Table 2. The chosen predictors were gridded values of daily specific humidity (SH), the zonal velocity component of the surface geostrophic wind (Us, hereafter referred to as U), and 500 hPa geopotential heights (H). 3.2.1 Daily Precipitation Occurrence Daily values of a precipitation occurrence parameter (represented by a series of “1”s and “0”s) were regressed against three grid-box predictor variables SH, U, H, and a lag-1 autocorrelation function using the following regression equation:

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The parameters are fitted by using linear least squares regression. In simulation mode, a uniformly distributed random number is used to determine whether precipitation occurs. For a given site and day, a wet day is synthesized if 3.2.2 Daily Precipitation Amounts Wet-day precipitation amounts for a given day are downscaled using the three gridbox predictor variables SH, U, and H. Since is always non-zero, it is appropriate to formulate the following regression model (following Kilsby et al., 1998):

where the are parameters fitted by linear least squares regression and modelling error. The expected value is given by:

is a random

where is an empirically derived correction ratio that allows for the bias resulting from the re-transformation of ln(R) to R and the fact that comes from a skewed distribution. The value of is defined such that observed and downscaled precipitation totals are equal for the calibration period. Additionally, a random scaling factor Ø (with a mean of 1) is used to increase the variance of R to obtain better agreement with observations (as used by Hay et al., 1991). Note that a lag-1 autoregressive component is not used to model because its inclusion did not significantly improve the explained variance in wet-day amounts. However, it is acknowledged that this parameter may be appropriate at other locations.

3.2.3 Daily Temperatures and Daily maximum and minimum temperatures for a given day were downscaled using the three grid-box predictor variables SH, U, and H, and the preceding day’s maximum and minimum temperatures, respectively. The daily temperature series were modelled using the following regression equations: and

where and are parameters fitted by linear least squares regression, and and are model errors. Both and are assumed to be normally distributed with mean zero and standard deviation equal to the standard error of the regression equation. Both sets of residuals were modelled stochastically using conventional Monte Carlo methods (Wilby et al., 1998a).

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HYDROLOGICAL MODEL

The three river basins are simulated using parameterizations of daily heat and water budgets in the Precipitation-Runoff Modelling System (PRMS: Leavesley et al., 1983), a physically based, distributed parameter model of precipitation forms, snowpack evolution, and runoff generation. The spatial variability of land characteristics that affect snowpack and runoff is represented by hydrologic response units (HRUs), within which runoff responses to uniform precipitation or snowmelt inputs are assumed to be homogeneous. HRUs are characterized and delineated in terms of those physiographic properties that determine hydrologic responses: elevation, slope, aspect, vegetation, soils, geology, and climate (e.g., Smith and Reece, 1995). In the three models used here, HRUs were designed to incorporate all grid cells, on 100-m grids, that share nearly identical combinations of these seven physiographic properties, regardless of whether the grid cells in an HRU form a contiguous polygon (Jeton and Smith, 1993). The resulting "pixelated" model delineations represent the basins in terms of 50 HRUs in the American River model, 50 HRUs in the Carson River, and 64 HRUs in the Merced River. Within each HRU, the heat- and water-budget responses to daily inputs of precipitation and daily fluctuations of air temperature are simulated. The daily mixes of rain and snow are estimated from each day’s temperatures by interpolations between the temperatures at which precipitation historically has been either all snow or all rain (Willen et al., 1971). Interception losses, sublimation, and evapotranspiration are also parameterized and simulated in terms of precipitation and daily maximum and minimum temperatures. Runoff is partitioned between surface runoff, shallow-subsurface runoff, deep-subsurface runoff, and deep ground-water recharge on the basis of the simulated accumulations of soil moisture at each HRU and of water in deeper subsurface reservoirs that underlie multiple HRUs. The various processes acting on runoff generation from the basins are represented in sufficient detail that heat- and moisture-fluxes vary realistically with short- and long-term climatic variations. However, the particular model parameters (such as temperature thresholds for rain to fall) and various land-surface descriptors (such as plant-canopy densities) were not modified in the future-climate simulations. Thus, the details of the model’s temperaturebased parameterizations are assumed, in the present study, to be unchanged under the futureclimate scenarios. This simplification amounts to assumptions that precipitation would derive from the same heights in the atmosphere as at present and that land-surface properties, such as vegetation type, would not change under the future scenarios. Snowpack accumulation, evolution, and ultimately the heat and water balances of the snowmelt periods are critical components in the simulations and are driven by the daily inputs of precipitation and daily air temperatures (using the parameterizations of Obled and Rosse, 1977). Snowmelt in the Sierra Nevada is driven mostly by solar radiation rather than by direct inputs of heat from the surface-air temperatures (Aguado, 1985). Thus, the

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temperature-based parameterizations of daily solar-radiation inputs are an important component of the models. The method used is a simple correction of clear-sky insolation estimates – from latitude, HRU slope and aspect, and day of the year – using the occurrence of precipitation and daily maximum air temperatures as crude indicators of the presence or absence of cloud cover (see Leavesley et al., 1983). Heat deposited in the snowpack by each day's sunshine is either lost to the overlying atmosphere the following evening if air temperatures drop below freezing, or is stored to contribute to eventual snowmelt if evening temperatures remain warmer than freezing. These temperature-based snowpack and insolation parameterizations are assumed to be unchanged in the future-climate conditions. Because of the physical detail of the process representations in the models, this assumption is a reasonable simplification. In the models, however, the solar-radiation estimates are functionally tied to daily air temperatures. This means that simulated solar-radiation inputs increase along with air temperatures, and thus, in addition to being warmer and wetter, the future-climate condition is represented in the hydrological models – almost inadvertently – as also being less cloudy (on dry days) than the present condition. As more information describing future relations between daily temperatures, cloudiness, and solar radiation at the surface becomes available, the parameterization of solar-radiation inputs used in simulations of future climate conditions may need to be modified accordingly. In this study, the parameterization was the same in all simulations. The Carson and American River models are described in detail by Jeton et al. (1996). The Carson River model simulates historical streamflows from 1969 to 1998, and the American River model simulates streamflow from 1949 to 1998. These simulations are driven by precipitation and temperature records from two nearby weather stations in the Carson River model and by weather observations from four nearby stations in the American River model. Indications of the goodness-of-fit of these models are presented by Jeton et al. (1996), and overall the fits are satisfactory. For example, 97 percent of the observed fluctuations of annual flow totals in the American River during a validation (non-calibration) period from 1949 to 1968 are present in the simulations, and 80 percent of the annual flow fluctuations of the Carson River are present in simulations during its validation period from 1969 to 1979. The Merced River model was designed to simulate daily flows for the period from 1916 to present (Dettinger et al., 1999), and the model is driven by precipitation and temperature observations from two long-term weather stations in the Sierra Nevada for most of that time (prior to the mid 1930s, only one of the two stations had weather records and the model was driven with just one input station). From 1937 to 1996, the model captures 77 percent of the observed daily flow variability; during the same period, simulated annual flow totals capture 83 percent of the observed variations.

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Modelled Streamflow Under Current-Climate Conditions

The calibrated downscaling model was forced using normalized Reanalysis SH, U. and H predictor variables for the verification period 1989-95. Statistically downscaled series of daily PRCP, TMAX and TMIN at the seven stations then were used to drive the watershed models.

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The downscaling model, as shown in Figure 3, captures the timing of the July precipitation minimum but underestimates the magnitude of the March maximum. Overall, the model has a slight dry bias (<3% error), yielding an annual precipitation total of 1117 mm instead of

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the observed 1143 mm. The downscaling model has a warm bias for most of the winter months, November to January, but a cold bias in February (Figures 3b and 3c). On average, both the downscaled maximum and minimum temperatures are +0.2°C warmer than observed data (16.6 and 3.0°C instead of 16.4 and 2.8°C, respectively). The success of the combination of hydrological models with precipitation and temperature downscaling can be measured in terms of biases in the simulated streamflow. The hydrological simulations driven by downscaled meteorology (Figure 4) generally underestimated annual streamflow, whereas simulations driven by station observations of meteorology tended to overestimate observed flows. Flows simulated using downscaled meteorology averaged 90% of observed flows in the Merced, 93% in the Carson, and 92% in the American, in comparison with 106% for all station data. Despite differences in simulated gross yields, the downscaled data provide a good

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approximation of the seasonal streamflow regimes, most notably for the Merced. However, the month of maximum mean streamflow is too early and too low in the case of the Carson, and well timed but underestimated in the American. A more severe test of the combined hydrological-downscaling model performance is provided by analysis of simulated daily flows for the downscaling-validation period from 1989 to 1995. As shown in Figure 5, for example, simulations forced by the station data have greater skill on daily time scales than do simulations with the downscaled meteorology for the Carson River model. For both simulations, though, the fits are satisfactory, and a significant component of the overall bias can be attributed to the hydrological model and choice of stations used for model calibration. Similarly, the correlation skills of flows simulated using statistically downscaled (station) data in the Merced and American Rivers were r = +0.84 (0.89) and r = +0.67 (0.81), respectively.

5. Modelled Streamflow Under Future-Climate Conditions Having demonstrated the ability of the combination of the downscaled historical climate conditions with the hydrological models to reproduce realistic historical streamflow variations, we next simulated streamflow using downscaled GCM simulations. The downscaling model – as calibrated with the historical Reanalysis fields – was forced using daily SH, U, and H predictors simulated by HadCM2 for current (1980-99) and future (2080-99) climate conditions. The seasonal regimes of the surface variables downscaled from the two HadCM2 scenarios are shown in Figure 6, and the corresponding streamflow and snowpack changes simulated by PRMS are shown in Figure 7. The downscaled scenarios yield more than a 50% increase in the annual precipitation and a +3°C warming of maximum and minimum daily temperatures. However, indicated in Figure 6a, the bulk of the precipitation increase occurs in just three months; precipitation in December, January, and February increases by +104%. The increase in maximum temperature (Figure 6b) ranges from +5.6°C (September) to +1.6°C (February), in comparison with +4.5°C (September) and +1.7°C (May) for the minimum temperatures (Figure 6c).

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These changes in future precipitation and temperature regimes are reflected in simulated snowpack and streamflow changes (Figure 7). In response to increased precipitation totals, all three rivers show large increases in annual runoff, ranging from +107% in the Merced, through +103% in the Carson, to +82% in the American. The corresponding annual mean snowpack changes were +41% (Merced), +27% (Carson), and –6% (American). The Merced has the largest increase in winter snowpack with an earlier peak snow water content, and an increase that persists into the summer months. Increased snowpack is accompanied by a marked increase in Merced spring streamflows. Similar, but smaller, increases are simulated for the Carson River. Both of these rivers are mostly at high elevations and have cold winters and springs. Evidently, the warmer temperatures in the simulated future-climate conditions are not sufficiently warm to prevent significant increases in overall snowpack and springtime streamflow. Indeed, in spite of projected warmer conditions, the percentage of the Merced and Carson basins that is simulated as being snowcovered – on average – during December through March decreases by only about 5% of current snow-covered areas. Thus, although warmer temperatures would tend to reduce the snow-covered areas, much wetter

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winter conditions roughly compensate for the reduced snow-covered areas in these simulations. In contrast, simulations of the American River indicate only a slight increase in winter snowpack and a decrease in spring snowpack volumes. Snow-covered areas in the American River basin decrease by about 10% of current averages under the warmer and wetter futureclimate scenario. The American River basin is lower and warmer than the others, and it yields a mix of rainfall runoff and snowmelt runoff each year, under current conditions. Increased winter streamflows projected under future-climate conditions are due to increases in winter rainfall runoff rather than increased snowmelt runoff. This results in earlier peak streamflows for the American and Carson Rivers under future climate conditions. All three basins yield more cool-season flash flooding under the future-climate conditions, and all have less autumn snowpack owing to higher maximum and minimum temperatures in September and a slight decrease in autumn precipitation. 6. Discussion A recent synthesis of seasonal precipitation scenarios for North America predicted by 15 GCMs reported changes ranging from –4% to +8% per 1°C global-mean warming, with wintertime precipitation over California being the most sensitive season and region (Wigley, 1999). Particular models such as CGCM1 and HadCM2 (Figure 1) simulate even greater sensitivities. For example, changes in the winter precipitation of HadCM2 range from – 10%/°C over southern Texas to +20%/°C over California. Although such results may be useful for identifying regions that are potentially most vulnerable to climate variability and change, GCMs are unable to capture local climatic effects arising from topographic, coastal, and land-surface processes. Statistical downscaling offers a computationally efficient and robust method of generating the basin-scale climate-change estimates necessary for hydrological impact assessments. Accordingly, a regression-based downscaling method was used to simulate daily rainfall and temperature series for streamflow modelling in three Californian river basins under currentand future-climate conditions. The downscaling model employed just three predictor variables (specific humidity, zonal velocity component of airflow, and 500 hPa geopotential heights) supplied by HadCM2 for the grid point nearest the target basins. When evaluated using independent data, the model showed reasonable skill at reproducing observed areaaverage precipitation, temperature, and concomitant streamflow variations. Overall, the downscaled data resulted in slight underestimates of mean annual streamflow that can be attributed to underestimates of precipitation in spring and positive temperature biases in winter.

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Simulated streamflows provide a useful indication of the combined skill of the coupled downscaling-hydrological model for the current climate because streamflow is an effective integrator (in time and space) of the cumulative effects of multiple climate variables. Streamflow at a single gauging site can also be measured more reliably than areal-average precipitation or temperature in complex terrain. However, because rivers act as basin-scale and season-long integrators of climatic forcings, simple downscaling-skill measures applied to streamflow simulations can appear more skillful than would similar skill measures applied to the driving precipitation and temperature inputs. For example, in basins with significant snowpack, the gross accumulations of precipitation in the seasonal snowpack and the timing of spring melt are critical to the simulations of hydrological responses. Conversely, in rainfall-dominated basins, the timing and magnitude of individual storm events are of greater concern, and these individual storms are the most difficult for a downscaling model to reproduce given only synoptic-scale atmospheric-circulation inputs. Thus, if a downscaling method reproduces only seasonal totals of precipitation and realistic temperature fluctuations during the critical snowmelt periods, snowmelt-dominated rivers will be much better represented than will nearby rivers dominated by rainfall runoff. Even in a rainfall-driven system, though, soil-moisture and ground-water reservoirs will tend to smooth the hydrological response and contribute apparent skill to the downscaling The smoothing effects of snowpack on streamflow responses to climate forcings help to explain differences between the skill of simulated streamflows in the three basins. The Merced and American River basins drain the western, windward slope of the Sierra Nevada, whereas the Carson River drains the eastern, leeward slope. Hence, the Carson River basin is in the rain shadow of the Sierra Nevada and is drier than the others. The Merced and Carson River basins are high-elevation basins and cool overall, with the part of the Merced drainage simulated here ranging from 1,200 to over 4,000 m above sea level and the Carson River drainage ranging from 1,600 to 3,400 m. The American River basin is a lower elevation basin and warmer overall, ranging from 200 m to 2,500 m above sea level. Consequently, the Merced and Carson Rivers are snowmelt dominated whereas the American River is a mix of rainfall runoff and snowmelt runoff. Simulated streamflow in the American River responds rapidly and sensitively to daily-scale temperature and precipitation fluctuations and errors; in the Merced and Carson Rivers, the response to the same short-term influences is much less. Consequently, the skill of simulated flows was significantly lower in the American River model than in the Carson and Merced. The physiography of the three basins also accounts for differences in their sensitivities to future climate change. Increases in winter precipitation exceeding +100% coupled with mean temperature rises greater than +2°C result in increased winter streamflows in all three basins. In the Merced and Carson basins, these streamflow increases reflect large changes in winter snowpack, whereas the streamflow changes in the lower elevation American basin are driven primarily by rainfall runoff. Furthermore, reductions in winter snowpack in the

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American River basin, owing to less precipitation falling as snow and earlier melting of snow at middle elevations, lead to less spring and summer streamflow. Taken collectively, the downscaling results imply significant changes to both the timing and magnitude of streamflows in the Sierra Nevada by the end of the 21st Century. In the higher elevation basins, the HadCM2 scenario implies more annual streamflow and more streamflow during the spring and summer months that are critical for water-resources management in California. Not shown here, but of comparable concern, the future winter and spring flow simulations also include more sudden flood events in response to winter and spring storms that yield more rainy mixes of rain and snow than at present, and for that rain to fall on snopwpacks that are warmer than at present. Nearly all of the additional flow in the lower elevation, warmer American River basin occurs in winter and constitutes increased flood hazards. Thus, depending on the relative significance of rainfall runoff and snowmelt, each basin responds in its own way to regional climate forcing. Generally, then, climate scenarios need to be specified – by whatever means – with sufficient temporal and spatial resolution to capture subtle orographic influences if projections of climate-change responses are to be useful and reproducible. Acknowledgments This research was supported by ACACIA (A Consortium for the Application of Climate Impact Assessments, National Center for Atmospheric Research) and by the U.S. Geological Survey Global Change Hydrology Program. We also thank David Viner of the Climate Impacts LINK Project (UK Department of the Environment Contract EPG1/1/16) for supplying the HadCM2 simulations on behalf of the Hadley Centre and U.K. Meteorological Office. NCAR is sponsored by the National Science Foundation. References Aguado, E. (1985). Radiation balances of melting snow covers at an open site in the central Sierra Nevada, California. Water Resources Research, 21, 1649-1654. Bardossy, A. and Plate, E.J. (1992). Space-time model for daily rainfall using atmospheric circulation patterns. Water Resources Research, 28, 1247-1259. Boer, G.J., Flato, G.M. and Ramsden, D. (1999a). A transient climate change simulation with historical and projected greenhouse gas and aerosol forcing: projected climate for the 21st century. Climate Dynamics, submitted. Boer, G.J., Flato, G.M., Reader, M.C. and Ramsden, D. (1999b). A transient climate change simulation with historical and projected greenhouse gas and aerosol forcing: experimental design and comparison with the instrumental record for the 20th century. Climate Dynamics, submitted. Christensen, J.H., Machenhauer, B., Jones, R.G., Schär, C., Ruti, P.M., Castro, M., Visconti, G. (1997). Validation of present-day regional climate simulations over Europe: LAM simulations with observed boundary conditions. Climate Dynamics, 13, 489-506.

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Kilsby, C.G., Cowpertwait, P.S.P., O'Connell, P.E. and Jones, P.D. (1998). Predicting rainfall statistics in England and Wales using atmospheric circulation variables. International Journal of Climatology, 18, 523-539. Leavesley, G.H., Lichty, R.W., Troutman, B.M. and Saindon, L.G. (1983). Precipitation-runoff modelling system: User’s manual. U.S. Geological Survey Water-Resources Investigations Report 83-4238, 207 p. Matyasovszky, I., Bogardi, I. And Duckstein, L. (1994). Comparison of two GCMs to downscale local precipitation and temperature. Water Resources Research, 30, 3437-3448. McGregor, J.J. (1997). Regional climate modelling. Meteorol. Atmos. Phys., 63, 105-117. Mearns, L.O., Rosenzweig, C. and Goldberg, R. (1996). The effect of changes in daily and interannual climatic variability on CERES-wheat yields. A sensitivity study. Climatic Change, 32, 257-292. Mitchell, J.F.B. and Johns, T.C. (1997). On modification of global warming by sulphate aerosols. Journal of Climate, 10, 245-267. Obled, C. and Rosse, B.B. (1977). Mathematical models of a melting snowpack at an index plot. J. Hydrology, 32, 139-163. Pilling, C., Wilby, R.L. and Jones, J.A.A. (1998). Downscaling of catchment hydrometeorology from GCM output using airflow indices in upland Wales. In: Wheater, H. and Kirby, C. Hydrology in a Changing Environment, Vol.1, Wiley, Chichester, 191-208. Richards, J.M. (1971). Simple expression for the saturation vapour pressure of water in the range -50° to 140°. Brit. J., Appl. Phys., 4, Ll5-L18. Richardson, C.W. (1981). Stochastic simulation of daily precipitation, temperature and solar radiation. Water Resources Research, 17, 182-190. Smith, J.L. and Reece, B.D. (1995). Watershed characterization for precipitation runoff modelling system, North Fork American River and East Fork Carson River watersheds, California. U.S. Geological Survey Hydrologic Investigations Atlas HA-734, 1 sheet. von Storch, H., Zorita, E. and Cubasch, U. (1993). Downscaling of global climate change estimates to regional scales: an application to Iberian rainfall in wintertime. J. Climate, 6, 1161-1171. Wigley, T.M.L. (1999). The science of climate change: global and U.S. perspectives. Pew Center on Global Climate Change, Arlington, Virginia, 48 pp. Wilby, R.L. (1997). Nonstationarity in daily precipitation series: implications for GCM downscaling using atmospheric circulation indices. International Journal of Climatology, 17, 439-454. Wilby, R.L. and Wigley, T.M.L. (1997). Downscaling General Circulation Model output: a review of methods and limitations. Progress in Physical Geography, 21, 530-548. Wilby, R.L., Hassan, H. and Hanaki, K. (1998a). Statistical downscaling of hydrometeorological variables using general circulation model output. Journal of Hydrology, 205, 1-19. Wilby, R.L., Wigley, T.M.L., Conway, D., Jones, P.D., Hewitson, B.C., Main, J., and Wilks, D.S. (1998b). Statistical downscaling of General Circulation Model output: a comparison of methods. Water Resources Research, 34, 2995-3008. Wilby, R.L., Hay, L.E. and Leavesley, G.H. (1999). A comparison of downscaled and raw GCM output: implications for climate change scenarios in the San Juan River basin, Colorado. Journal of Hydrology, submitted. Wilks, D.S. (1992). Adapting stochastic weather generation algorithms for climate change studies. Climatic Change, 22, 67-84. Willen, D.W., Shumway, C.A., and Reid, J.E. (1971). Simulation of daily snow water equivalent and melt. Proceedings of the 1971 Western Snow Conference, Billings, Montana, v. 39, 1-8. Winkler, J.A., Palutikof, J.P., Andresen, J.A. and Goodess, C.M. (1997). The simulation of daily temperature series from GCM output. Part II: Sensitivity analysis of an empirical transfer function methodology. Journal of Climate, 10, 2514-2532.

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EXAMINING LINKS BETWEEN CLIMATE CHANGE AND LANDSLIDE ACTIVITY USING GCMS: case studies from Italy and New Zealand

MICHAEL SCHMIDT and MARTIN DEHN Dept. of Geography, University of Bonn Meckenheimer Allee 166 D-53115 Bonn Fax: +49-228-73 9099 e-mail: [email protected]

Abstract Climate is an important forcing parameter for landslide activity and hence of strong geomorphological interest, especially precipitation and temperature as inputs into the landslide system. Since climate change is widely accepted, its regional impacts are of major research interest. In fact it is not obviously clear which regions will be affected, in which manner and how the local environments will react to these changes. Climate scenarios can display probable outcomes and boundary conditions for linked environmental systems. In this study GCM (General Circulation Model) outputs are downscaled, with an empirical-statistical method applied to two locations. We present an activity scenario of a mudslide in the Dolomites and scenarios of landslide events calculated with different landslide models for the region of Wellington, New Zealand. Scenario results of the Alvera mudslide (Italy) show a significant activity decrease due to increasing winter temperature and reduced snow storage while precipitation changes do not show a clear trend. The projections for Wellington indicate fewer events of high landslide probability in the hemispherical winter due to decreased precipitation. Key words: Climate change impact, GCMs, downscaling, landslide scenarios 123

S.J. McLaren and D.R. Kniveton (eds.), Linking Climate Change to Land Surface Change, 123–141. © 2000 Kluwer Academic Publishers. Printed in the Netherlands.

124 1.

SCHMIDT AND DEHN Introduction

Water in the soil profile and slope hydrology generally are of major concern for the landslide system and therefore of great research interest in terms of modelling occurrence and activity of landslides. The other important climatic parameter for landslide studies is air temperature. The estimations of both parameters are outlined principally. Climate change is defined here as due to increasing emissions of atmospheric trace gases with climatic relevance, as simulated by the IPCC (Intergovernmental Panel on Climate Change) in the IS92 emission scenarios (Houghton et al., 1992). Coupled ocean and atmosphere GCMs are at the moment the best tools to describe future climatic conditions, even if the typical grid resolution of modern GCM is still too coarse for local scenarios (Trenberth, 1996). Several downscaling techniques have been developed to bridge this scale gap (Hewitson and Crane, 1996, Zorita and von Storch, 1998). In this study an analog downscaling technique (Cubasch et al., 1996, Zorita and von Storch, 1998) is used to derive local precipitation scenarios, which are then taken as input for landslide models. Changes of climatic variables like temperature, precipitation, sea level or soil moisture are likely to occur on a global and regional scale. The current average rate of change might be greater than any seen in the past 10,000 years. Although global averages are within historical rates of changes, regional changes could differ widely (Watson et al., 1998). These human-induced large scale phenomena will interact with sources of natural variability like the El Niño-Southern Oscillation (ENSO) and thus influence social and economic well-being. This is especially serious if the recently projected increases in frequency and magnitude of extreme events are considered (Cubasch et al., 1995).

2.

General methodology

The overall methodology is outlined in Fig. 1. Two site specific landslide models were used to establish local landslide scenarios. These are a threshold model of occurrence of landslide events for the region of Wellington, New Zealand, and a hydrological tank model coupled to a slope-stability model to describe the activity of a single mudslide in the Dolomites. Both approaches use regional/local precipitation and temperature as input parameters. These climatological scenarios are the result of the first three modelsteps in Fig. 1.

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The GCMs ECHAM4/OPYC3 (European Centre Hamburg Model, 4. Generation) without consideration of sulphate aerosols (Roeckner et al., 1996) and HadCM2SUL (Hadley Centre Coupled Model, 2. Generation) with sulphate aerosols (Johns et al., 1997) are forced with the emission scenarios IS92a, which results in a doubling of concentrations within the next century. It has to be taken into account, that GCMs do quite well in projecting general circulation and large scale variables, but due to the coarse operating scales (ca. 250x250 km) local variables such as precipitation or cloudiness are only roughly parameterised and of low confidence for site specific studies. Interrelations between Sea Level Pressure (SLP) as a large scale variable and precipitation as a local variable provide the possibility for a statistical downscaling (see below). The presented results and scenarios must be seen in terms of trends and absolutely not as event probabilities of single events for a given day. Precipitation, as a variable with high local variability, requires a downscaling to produce reliable scenarios based on GCMs. Temperature variability is also high temporally, but not spatially, so that it is a common technique to interpolate the local temperature directly from GCMs by considering and adjusting the elevation factor, which is not adequately captured by the GCM.

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2.1. THE ANALOG DOWNSCALING TECHNIQUE We applied an analog downscaling proposed by Cubasch et al. (1996) and Zorita and von Storch (1998). Two independent data sets of SLP and local precipitation are necessary for training and validating the analog model. Daily SLP of the training period is stored together with the local precipitation data in a catalog. Then SLP data of a validation period are compared with the training period by searching for similar circulation patterns (operated by SLP fields) and associated with this analog situation the amount of precipitation from the catalog. This is first done by using Empirical Orthogonal Functions (EOF), by which the data are projected in a space with orthogonal span vectors. These independent vectors describe the internal variability of the data.

Usually the 5 leading EOFs describe more than 90% of the data variance. The analog situation is then simply found by the next neighbour in the 5 dimensional EOF space

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(Fig. 2). The same procedure can be applied to observations as training period and SLP of various target periods of a GCM run. Results are precipitation scenarios of respective GCM periods. Before going to the specific case studies one major assumption should be pointed out. The empirically derived relationships between SLP and local precipitation must be constant over time and also during changing climate. If this is not the case the scenarios are not valid.

3.

Case study of the Alvera mudslide

3.1. DOWNSCALING MODEL In Fig. 3 the location of Cortina d’Ampezzo is pointed out in the land–sea mask of the ECHAM4/OPYC3 climate model with T42 resolution. It is obvious that a downscaling of precipitation values is necessary, especially because the whole alpine region is represented by a few grid-boxes only. The analog downscaling was carried out separately for the 4 meteorological seasons spring (MAM), summer (JJA), autumn (SON) and winter (DJF). The validation results in the form of time series correlations between observed and estimated seasonal precipitation are shown in Table 1.

Autumn and winter represent reliable results. In contrary summer precipitation is poorly estimated, which might be due to the high quantity of summer precipitation with convective origin. Convective showers are not so closely related to large-scale circulation than advective precipitation, which is the major component of winter precipitation. Summarising, it can be said that the analog technique based on SLP is able to successfully reconstruct the historical development of precipitation in Cortina for the seasons winter, spring and autumn, while precipitation amounts are generally underestimated as well as other precipitation parameters (for details refer to Dehn, 1999). Despite some deficiencies, the validation shows the skill of the analog technique and therefore it can be used for downscaling large-scale circulation features of GCM experiments.

128 3.1

SCHMIDT AND DEHN FEATURES OF ALVERA MUDSLIDE

The Alvera landslide is a slow moving mudflow (few centimetres per year), situated NE of Cortina d’Ampezzo. The geological structure of the region is characterized by a repeated succession of dolomitic and pellitic rocks (Angeli et al., 1992). The total length of the mudflow is about 1700 m and 100 m of width, with an average inclination of 7.3° (Fig. 3).

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The mudslide consists of clayey and silty material which is a weathering product of the San Cassiano formation. Groundwater records suggest a subdivision into a macroporous root zone underlain by a less permeable clayey layer (Angeli et al., 1998). The depth in which 75% of the movement takes place is at about 5 m below the surface which is connected by deep cracks with the upper system. The second sliding area is at 20-25 m depth (Gasparetto et al., 1996). A linear tank model is used for the simulation of this 2 component system (Angeli et al., 1998). The onset of displacement is triggered if groundwater exceeds a critical level of 0.5 m below ground. Concerning landslide activity the number of days with supercritical groundwater levels, that is > -0.5 m, was assessed in the scenarios. 3.2

FUTURE SCENARIOS

In the following activity scenarios for this mudslide are presented, based on downscaled precipitation and interpolated temperature for Cortina d‘Ampezzo. Precipitation scenarios are shown in Fig. 4 as seasonal changes between GCM control period (195079 for HadCM2 and 1960-89 for ECHAM4/OPYC3) and future period 2070-2099 for both GCMs. In the ECHAM4 scenario seasonal precipitation is generally decreasing but not significantly on the 95% confidence level. In downscaling results from the HadCM2 model there is no significant trend displayed.

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In contrary to precipitation, local temperature was derived by simply interpolating between several GCM grid points next to Cortina. The absolute values had to be fitted to local observations on a monthly basis. The temperature rise in HadCM2 with consideration of sulphate aerosols is smaller than in ECHAM4/OPYC3 without sulphate aerosols (Fig. 5).

Changes in landslide activity are depicted in Fig. 6. For both GCMs a significant decrease of future landslide activity in spring (MAM) is visible. Changes in other seasons are less dramatic and not significant. This strong change in MAM is not visible in the precipitation scenarios. This finding will be discussed later.

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Case study Wellington region

4.1. DOWNSCALING MODEL In Fig. 7 the land-sea mask of the HadCM2SUL model with the grid resolution of 2.5°x3.75°, which is similar to the ECHAM4/OYPC3 T42 resolution, is shown. For a better estimation of local precipitation values, an analog downscaling was carried out with SLP as the large scale variable, similar to the Cortina example. The validation results for the different seasons are shown in Table 2. Only the winter season (JJA) shows acceptable validation results. Therefore all further conclusions will be drawn only for the winter season.

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REGIONAL SETTINGS

The study area covers with steep and strongly dissected slopes, often oriented along fault lines, which produce aligned drainage. The relief is about 460 m with an average tectonic uplift of 1mm/a. In the area extensively faulted, tilted and folded dark grey argillite and greywacke sandstones can be found. On this bedrock colluvium soils and solifluction deposits are developed. The native forest was converted to pasture und

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scrub vegetation as a result of increased urbanisation since the 1840s. This area is prone to landslides if the high amount of 1250 mm yearly precipitation is taken into account.

4.2

The landslide models

4.2.1

THRESHOLD MODEL

The probabilities for landslide events are derived by dealing with daily regional maximum precipitation values of all the weather stations in the region. This precipitation amount is, in the case of an event, defined as the triggering value. The aim is to define triggering thresholds or the probability of a given amount of rainfall to trigger landslides. A first modelling approach is to divide the measured maximum precipitation into precipitation classes of 20 mm intervals. Minimum thresholds, below which a landslide had never occurred can be found as well as thresholds above which landslides always occurred. The classes below this daily minimum threshold of 20 mm precipitation have a probability of 0% and the precipitation classes above the maximum have 100% of landslide probability (Fig. 8).

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The probability of landslide occurrence is calculated as the ratio of triggering precipitation number by total number of occurrence for each precipitation class. These values were calculated by Glade (1998) for the period of 1903-1995 and were therefore taken as values on a larger statistical basis. The bars in Fig. 8 represent time periods of 30 years. The bars present observed data and data of GCM origin, which are downscaled and used to build precipitation scenarios for the control period (1950-1979) and for future conditions (2070-99). The datasets can be compared reasonably either by the same time period in terms of model quality or the same origin in terms of changing conditions. The model represents similar results for the time period 1950-1979, especially for the lower classes, where the number of elements in each class is higher. This result implies that the whole approach should be suitable. For the future run, a slight tendency towards less extreme events is visible. One major limitation of this model is that it works without preparatory factors. Only the actual daily amount of precipitation is included as triggering factor.

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THE ANTECEDENT DAILY RAINFALL MODEL

In this model the antecedent rainfall of 10 days is considered as preparatory factor in the form of the Antecedent Daily Rainfall Index (ADRI) (Crozier and Eyles, 1980). The principal idea that stands behind this landslide model is that preparatory factors are the reason why a stable slope system becomes marginally stable. In this modelling approach precipitation is the only preparatory factor for a slope to become marginally stable. Following the concept of critical water content (Crozier, 1998), the pre-existing water within the slope together with the daily precipitation as triggering factor, causes a landslide to occur. All other triggering factors are excluded for this model. Controlling factors are site specific and describe the behaviour of a sliding mass. These factors are not considered here, because in this study only the number of events is dealt with. Glade et al. (1999) model the antecedent rainfall influence by the following equations

where is the antecedent rainfall for day 0, the rainfall on the i-th day before day 0 and the number of considered days before day 0. is a site specific empirically derived factor of value –1.52 (Glade et al., 1999). To include the antecedent conditions of 10 days seems suitable because the use of longer antecedent conditions show no better results for this study region (Glade, 1997). Probabilities of landslide occurrence are thus calculated with the antecedent rainfall conditions as preparatory factors and the actual rainfall as triggering event, as it is shown in the following equation:

represents the daily precipitation, the probability of landslide occurrence of a given day and is the antecedent rainfall that is held in the soil as described in equation 1. The probabilities derived with this approach are outlined in the following figures.

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In Fig. 9a-c fixed probability lines for landslide events are drawn. Any point above a probability line represents the combination of the daily rainfall and ADRI with the same (or higher) event probability as the value of this line (Fig. 9a). The curved probability lines exhibit a negative relationship between antecedent conditions and the daily rainfall, which means that at high amounts of ADRI less precipitation is needed to trigger a landslide and vice versa.

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The bars in Fig. 10 show the number of days exceeding certain probability values. The control run (1950-79) and the observed values (same period) show similar results. The first two classes seem to be well represented, while the differences in the higher classes are more important. Comparing control run and the future landslide scenario (2070-99), a trend for decreasing winter landslide event probabilities, especially for the lower classes, is displayed. The numbers of the very high probabilities have to be handled with care, because these are single storm events with intense precipitation amounts for which is not yet clear, if GCMs can represent them well.

5.

Discussion

The main result of the Alvera climate change impact study is the significant decrease of landslide activity in spring without accompanying decrease of seasonal precipitation. Therefore it has to be explained how this dramatic decrease of landslide activity in spring is caused. The reason has to be searched for in the significant temperature increases simulated by both GCMs for all seasons. An explanation is suggested which is focusing on the role of snow storage of precipitation in winter (DJF). This effect contributes usually to high meltwater inputs into the slope system in spring and hence causes high landslide activity rates of the Alvera mudslide in this season. Global warming due to the greenhouse effect forces the DJF temperatures in Cortina of currently –1.1°C to exceed 0°C in the mid next century. This happens according to the GCM-experiments with or without consideration of the cooling effect of sulphate aerosols. In a warmer climate, therefore, less water stored as snow is available for the slope in spring with the consequence of lower groundwater levels and lowered rates of landslide activity. Treatment of snow storage, snow melt and potential evapotranspiration in the tank model is described in detail in Angeli et al. (1998). Problems of these parameterisations in the context of climate change applications are discussed in Dehn (1999). The results for the study area of Wellington, due to the low sample size of high landslide probabilities, can not be described with statistical parameters. It would be a future task to find a method that is suitable to construct scenarios for the remaining seasons as well as winter. In JJA the westerlies are the most important feature for advective precipitation, which is connected with synoptic scale features and are thus useful for this downscaling method The conditions for the other seasons are more variable. In the warmer months convective processes due to the surplus of energy in the atmosphere, reduces the performance of the analog downscaling.

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Even if in winter (JJA) almost 40% of the recorded landslides in the study area occur, the most severe storms with huge precipitation amount and direct triggering did not happen in JJA. Nevertheless landslides occurred, even with a calculated probability of less than 10%, which leads to the speculation of the coincidence of more than one triggering factor. In summary, it was shown that depending on landslide type and setting different climate elements develop as key parameters due to global warming. This problem is more generally discussed in Dehn and Buma (1999). A consequence of this finding is that successful assessments of future landslide activity require landslide modelling concepts which are able to represent process variability due to changes in precipitation and temperature.

6.

Conclusions

The following conclusions can be drawn from the two case studies: With the analog downscaling technique large-scale atmospheric conditions can be exploited for constructing local climate scenarios on a daily time-scale. The analog technique is performing best in the respective winter of Italy and New Zealand. Temperature increase, especially in winter, is an important causative factor for the decreasing landslide activity in spring of Alvera mudslide. The treatment of snow processes and evapotranspiration in the applied model requires improvements. The number of landslides for Wellington must be seen as minimum, because for every event only one landslide is recognised, even if there are several single slides. The tendency towards a decrease of winter landslide event probability is the major outcome of this regional study. The boundary conditions for landslides and thresholds are held constant over time, which is by no means guaranteed, as they could change by human influence like forest removal or slope cuts for urbanisation or infrastructure lines or change after a landslide. Temperature and evapotranspiration are not adequately considered by the empirical decay formula for the study area of Wellington. This shortcoming is improved in the Antecedent Soil Water Status model (Crozier, 1998, Glade, 1999), which will be used in a next step.

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Acknowledgements This paper is part of the CEC Environment Research Programme on ,,New technologies for landslide hazard assessment and management in Europe (NEWTECH)“ (ENV4CT96-0248). We would like to thank Eduardo Zorita and Hans von Storch, GKSS Research Centre, Geesthacht for assistance and cooperation with the analog technique, Jelle Buma, University of Utrecht, for the use of the hydrological model, Thomas Glade and Mike Crozier for the use of the New Zealand landslide models, and finally Erich Roeckner and Monika Esch from the Max-Planck-Institut für Meteorologie, Hamburg and David Viner from Climate Impacts LINK project, Norwich, for providing GCM data.

References Angeli, M.-G., Buma, J., Gasparetto, P., Pasuto, A. and Silvano, S. (1998) A combined hillslope hydrology / stability model for low-gradient clay slopes in the Italian Dolomites, Engineering Geology, 49, 1-13. Angeli, M.-G., Menotti, R.M., Pasuto, A. and Silvano, S. (1992) Landslide studies in the Eastern Dolomites Mountains, Italy. In: D.H. Bell (Editor), Proc. 6th International Symposium on Landslides. Balkema, Christchurch (New Zealand), pp. 275-282. Crozier, M.J. (1998) The climate landslide couple: a Southern Hemisphere perspective, Paleoclimate Research, 2, 329-350. Crozier, M.J. and Eyles, R.J. (1980) Assessing the probability of rapid mass movement. In: The New Zealand Institution of Engineers - Proceedings of Technical Groups (Editor), Proc. Third Australia - New Zealand Conference on Geomechanics, Wellington, pp. 2.47-2.51. Cubasch, U., von Storch, H., Waszkewitz, J. and Zorita, E. (1996) Estimates of climate change in Southern Europe derived from dynamical climate model output, Climate Research, 7, 129-149. Cubasch, U., Waszkewitz, J., Hegerl, G.C. and Perlwitz, J. (1995) Regional climate changes as simulated in time-slice experiments, Climatic Change, 31, 273-304. Dehn, M. (1999) Application of an analog downscaling technique to the assessment of future landslide activity - a case study from the Italian Alps, Climate Research, in press. Dehn, M. and Buma, J. (1999) Modelling future landslide activity based on general circulation models, Geomorphology, 30, in press. Gasparetto, P., Mosselman, M. and van Asch, T.W.J. (1996) The mobility of the Alvera landslide (Cortina d'Ampezzo, Italy), Geomorphology, 15, 327-335. Glade, T. (1997) The temporal and spatial occurrence of rainstorm-triggered landslide events in New Zealand. PhD, Victoria University, Wellington. Glade, T. (1998) Establishing the frequency and magnitude of landslide-triggering rainstorm events in New Zealand, Environmental Geology, 35, 160-174. Glade, T. (1999) Models of antecedent rainfall and soil water status applied to different regions in New Zealand, in preparation.

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Glade, T., Smith, P. and Crozier, M. (1999) Applying Probability Determination to refine LandslideTriggering Rainfall Thresholds using an Empirical ,Antecedent Daily Rainfall Model‘, Pure and Applied Geophysics, In press. Hewitson, B.C. and Crane, R.G. (1996) Climate downscaling: techniques and application, Climate Research, 7, 85-95. Houghton, J.T., Callander, B.A. and Varney, S.K. (eds.) (1992) Climate change 1992. The supplementary report to the IPCC scientific assessment. Cambridge University Press, Cambridge. Johns, T.C., Carnell, R.E., Crossley, J.F., Gregory, J.M., Mitchell, J.F.B., Senior, C.A., Tett, S.F.B. and Wood, R.A. (1997) The Second Hadley Centre coupled ocean-atmosphere GCM: Model description, spinup and validation, Climate Dynamics, 13, 103-134. Roeckner, E., Oberhuber, J.M., Bacher, A., Christoph, M. and Kirchner, I. (1996) ENSO variability and atmospheric response in a global coupled atmosphere-ocean GCM, Climate Dynamics, 12, 737-754. Trenberth, K.E. (1996) Coupled Climate System Modelling. In: T.W. Giambelluca and A. Henderson-Sellers (eds.), Climate Change. Developing Southern Hemisphere Perspectives, Wiley, Chichester, 63-88. Watson, R.T., Zinyowera, M.C., Moss, R.H. and Dokken, D.J. (eds.) (1998) The regional impacts of climate change. Cambridge University Press, Cambridge. Zorita, E. and von Storch, H. (1998) A survey of statistical downscaling techniques, GKSS Report 97/E/20, Geesthacht.

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GEOLOGIC EVIDENCE OF RAPID, MULTIPLE, AND HIGH-MAGNITUDE CLIMATE CHANGE DURING THE LAST GLACIAL (WISCONSINAN) OF NORTH AMERICA

F. W. BACHHUBER University of Nevada, Las Vegas, Las Vegas, NV, USA, 89154-4010 N. R. CATTO Memorial University of Newfoundland, St. John’s, Canada, A1B 3X9

Abstract Over the last few decades data derived from ice cores and dendrochronologic sequences indicate that former climate change during the Quaternary has been rapid and of high magnitude. The Estancia Valley, central New Mexico, U.S.A. contains a high-resolution sedimentologic and palaeontologic pluvial-lake record that supports these data. The known Estancia Valley record, in outcrop and subcrop, spans the Wisconsinan (last glacial episode of North America) and represents relatively continuous deposition through that time. The period from early to middle Wisconsinan (ca. 60,000 BP to 30,000 BP) is marked by a relatively stable climate, represented by dry playa and saline lake sediment and fossils. At the beginning of the late Wisconsinan, however, climatic instability becomes the norm. The first fresh-water pluvial lake, as evidenced by fossils, appears at ca. 24,000 BP, followed by additional discrete fresh-water lakes at ca. 21,000 BP, ca. 20,500, ca. 20,000 BP, ca. 17,000 BP, and ca. 14,000 BP. Each of these seemingly deeper-water stands was separated by desiccation events (the last dating after ca. 12,400 BP) indicating that each lake phase resulted from high magnitude climate change (full pluvial to full inter-pluvial climate). It 143

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also appears that each lake phase developed rapidly, stabilized for only a short period, and desiccated rapidly. Late Lake Estancia, the highest lake stand of about 52 m deep and dated at ca. 20,000 BP (coinciding with the late Wisconsinan glacial maximum) offers the most conclusive evidence of rapid lake development from dry-playa climatic conditions. The Late Lake Estancia sedimentary sequence consists of deposition on an unconformable surface, an interval of saline-lake adapted organisms, brackish-water-adapted organisms, and finally infrahaline-adapted organisms, all within a 10 cm stratigraphic interval. Radiocarbon ages and sedimentation rates suggest that change from dry-playa conditions to full pluvial-lake conditions (maximum lake depth) back to a dry playa occurred within a few decades to a few centuries. It is likely that the other fresh-water stands developed and disappeared in a similar time span. The rise and fall of major pluvial systems in the southwest U.S.A. demonstrate that the geologic response to rapid climate change can also be rapid.

1.

Introduction

Detailed palaeontologic analysis of late-Quaternary lacustrine sediments from the Estancia Valley, central New Mexico suggest that pluvial-lake systems respond quickly and dramatically to climate change. The geologic record of the Estancia Valley delineates not only multiple, high-level, fresh-water lakes, but also depicts waxing and waning phases of lake evolution, including ephemeral, low-level, lake phases of variable salinity, and desiccation events—all reflecting regional climate change. The climatic record inferred from palaeontologic and sedimentologic change in the Estancia Valley is of greater continuity and of significantly higher resolution than that of glacial, fluvial, and aeolian systems recorded in other areas of North America and compares to that of ice core records. It is believed that the degree of resolution in the Estancia stratigraphic section is on the order of a few centuries, and in some cases lake level appears to fluctuate dramatically within a few decades. Significantly, the Estancia record clearly demonstrates that geologic and biologic processes respond to rapid climate change, and in restricted environments this change is recorded in great detail. 1.1 REGIONAL SETTING The Estancia Valley is a broad, elliptically shaped, physiographic and structural basin near the geographic center of the state of New Mexico (Figure 1). Basin closure is afforded by the Manzano Mountains along the west margin (highest elevations above 3,050 m), the

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Pedernal Hills along the east margin (up to 2,312 m), Juames Mesa to the south (2,106 m), and a board, sloping saddle (~1,983 m) to the north. A topographic sill is located in the southeast portion of the basin at an elevation of 1,932 m. The elevation of the central-valley floor is approximately 1,853 m where a complex of over 60 playa-floored deflation basins is incised, up to 10 m deep, along the axis of the valley. The Holocene-age deflation basins are ringed by parabolic dunes, and other aeolian features (see Catto and Bachhuber, this volume). At present, temperature and precipitation varies significantly within the Estancia Valley watershed (see Catto and Bachhuber, this volume). The central portion of the valley, however, is arid with a large annual water-budget deficit of ~1.5 m.

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The water-budget deficit precludes the existence of perennial standing water other than at the highest elevations in the Manzano Mountains. The deflation basins distributed along the valley axis contain standing water only following heavy rainfall events. Within days or a few weeks, the deflation-basin floors return to dry playa conditions. The high-resolution geologic and palaeontologic record contained in the Estancia Valley is, in part, related to its geographic and climatic setting. Occurring between 35°- 36°N latitude, the valley is near the southern margin of the area characterized by optimal pluvial lake development. At comparable latitudes in Nevada, similar closed basins apparently did not contain large pluvial lakes (Mifflin and Wheat, 1979) during the Wisconsinan, the last North American glacial stage (compares to Weichselian in Europe). The relatively high elevation of the Estancia Valley floor (1,853 m), however, supports a climatic threshold that, with moderate precipitation or temperature change, produces a significant hydrologic response. Depending upon the magnitude of climate change the hydrologic thresholds crossed are dry playa to wet playa to saline lake to fresh-water lake, or this succession in reverse. The sedimentologic and biostratigraphic section of the Estancia Valley record these thresholds, and transitions between thresholds. At lower elevations, however, the response to similar climate change is not as rapid or as sensitive, and only extreme climate events will be recorded in the geologic record. In addition to the high valley elevation, which creates the potential for rapidly crossing critical hydrologic thresholds, the numerous late Holocene deflation basins provide vertical and lateral accessibility to the late Pleistocene sediment record. Up to 10 m of Pleistocene lacustrine sediment is exposed along the slopes of the interior deflation basins, with progressively reduced exposure thickness occurring in the more exterior (mainly eastern) deflation basins. Consequently, the great number of deflation basins and their juxtaposition to near-shore and deep-water lacustrine facies provides unique access to the geologic record. In outcrop, this record is relatively continuous from latest middle Wisconsinan through late Wisconsinan. 1.2 GEOLOGIC SETTING The Estancia Valley lies within a structural basin of the easternmost Basin and Range physiographic province. The east and west margins of the valley are bordered by a system of north-trending normal faults with uplifted areas exposing Precambrian- and Palaeozoicage rocks. Mesozoic-age sediments occur in subcrop.

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From early Miocene to early Pleistocene time, the Estancia Valley had an eastward-directed, through-flowing, fluvial system into the ancestral Pecos/Brazos river network (Kelley, 1972; Hawley, 1984). Basin closure, through faulting, probably had developed by middle Pleistocene time. The first sedimentologic record of closure consists of a 7 m-thick lacustrine section (Early Lake Estancia) (Figure 2) recorded in water wells (Titus, 1969). Little is known of Early Lake Estancia, but geomorphic evidence and possible reworked ostracods in early to middle Wisconsinan-age sediment (Bachhuber, 1992) suggest that the lake was a fresh-water body and of a size probably greater than any high-lake stand of the late Wisconsinan pluvial period. Early Lake Estancia is believed to be no younger than Illinoian in age, and probably of Illinoian age (Bachhuber, 1989). Early Lake Estancia sediment is overlain by 17 m of alluvium and dune material. Also described from water-well data (Titus, 1969), this succession, termed the Medial Sand (Figure 2), has not been tested for the presence of fossils. However, it is likely that the Medial Sand is no younger than Sangamonian in age (Bachhuber, 1992). 1.2.2. Early and Middle Wisconsinan The Medial Sand is overlain by 10.5 m of intercalated clay and gypsarenite, only the uppermost portion of which crops out along the slopes of the deepest-incised deflation basins. The unit, named the La Salina Complex (Figure 2), has a palaeontologic assemblage and sedimentology representing saline lakes of variable salt content, wet playas and dry playas (Bachhuber, 1992). For clarification, the wet playa is defined as a system supported by precipitation and ground-water flow. The floor of a wet playa contains water for much of the year, but desiccates periodically. The dry playa is exclusively supported by precipitation events, and, although flooded following rainstorms, the floor is dry for most of the year. The outcrop portion of the La Salina Complex is clearly middle Wisconsinan with ages of 30,440 ± 520 yr BP (Beta-25542) to 35,650 +3,000/-2,130 yr BP (A-4903). A age of >48,800 yr BP (AA-6330), along with amino acid racemization data (Rutter et al. 1992), from the middle of the La Salina Complex suggest that the lower portion of the unit is no younger than early Wisconsinan in age. As such, the La Salina Complex is a relatively continuous record of cold/dry climate from early through middle Wisconsinan time.

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Significantly, there is no record of high-level, fresh-water lakes in the La Salina Complex (Bachhuber, 1992).

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1.2.3. Late Wisconsinan The La Salina Complex is overlain by ~ 10 m of thick-bedded flint-gray clay and interbedded silty clay and gypsarenite. The complete late Wisconsinan section is exposed in outcrop in the more deeply incised deflation basins such as E-28 (Figure 1). Only the upper portion is exposed in the marginally located basins that elevationally step-up away from the valley axis, such as E-22 (Figure 1). Previously, the late Wisconsinan sediment had been subdivided into two main pluvial lake stands, Late Lake Estancia and Lake Willard, with an intervening inter-pluvial section, the Estancia Playa Complex (Bachhuber, 1989). In addition, based on a rich and varied palaeontologic assemblage, six distinct phases were recognized in Late Lake Estancia sediment (Bachhuber, 1989). The palaeontologic evidence depicts fresh-water conditions as compared to more saline conditions. The model developed by Bachhuber (1989) had Late Lake Estancia rise from an early late Wisconsinan saline-lake phase, waxing and waning into three recognized high-water stands and two partial draw down phases, and finally desiccating at the beginning of Estancia Playa Complex time. While this model is basically correct in recognizing high water and low water conditions, as inferred from lake salinity, we propose here that water level changes were more dramatic. Instead of partial draw down phases, desiccation occurred with these events separating the distinct high-water phases into spatially and temporally discrete pluvial lakes.

2.

Palaeolimnology of the Late Wisconsinan Pluvial System

The palaeolimnology of the late Wisconsinan pluvial and inter-pluvial sections is determined by the fossils, sedimentology, and stratigraphy exposed along the flanks of deflation basins. Owing to the distribution of the deflation basins inside the lowest strandlines, near-shore and deep-water facies are readily available for examination and sampling. The abundance and diversity of fossils, all extant species, permit palaeoecologic evaluation. From these data inferences can be drawn as to palaeolimnologic conditions especially those related to total dissolved solid (TDS) of the various lake stages. TDS, in turn, is believed to reflect water depth. Saline bodies are considered to have lesser depths with the surface extent of the lake being relatively small. Fresher water bodies are considered to be of greater depth, with the lake occupying a larger surface area. These contentions are partially supported by the occurrence of a series of strandlines that vary in elevation from approximately 1860 m to 1897 m along the Estancia Valley margins, and by vertical facies changes indicating transgressive and regressive sequences. While these

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relationships have been long recognized, recent detailed palaeontologic analyses and additional ages have required a re-evaluation of previous work (Bachhuber, 1989).

2.1. PALAEONTOLOGY The diversity and density of fossils in the late Wisconsinan sediment of the Estancia Valley are remarkable. Microfossils and macrofossils occur in varying abundance in all previously described Late Lake Estancia phases and in Lake Willard, from near-shore and deeper-water facies. In general, two basic assemblages are recognized: one characteristic of a fresh-water or near fresh-water environment; and a second that is indicative of more saline to even hypersaline conditions. The fresh-water indicators or inhabitants are cutthroat trout, gastropods, pelecypods, Pediastrum (an alga), and some ostracod and diatom species. Saline indicators are certain species of ostracods, diatoms and charophytes, foraminifers (Bachhuber and McClellan, 1977), and Ruppia (ditchgrass). In addition, the sediment contains head capsules, wings, and mandibles of insects, seeds from various plants, plant fragments and impressions, pollen, salamander bones, ephippia, and innumerable spheres and disks of unknown affinities. Even though some of the Estancia fossil material has not been identified or evaluated, based on work previously completed (Bachhuber, 1989; 1992), there is a good understanding of the various late Wisconsinan lake stands, and the conditions that prevailed during their existence. The problem with previous work, however, was that sediment sampling strategies were too coarse. Normally samples were collected at 5 – 10 cm intervals, with the vertical thickness of the sample encompassing up to 5 cm. It has now been realized that this sampling strategy resulted in completely missing a number of major salinity reversals, and it blurred the boundaries between distinct saline-lake and fresh-water events. Single processed samples often contained palaeontologic indicators of both saline and fresh-water environments, suggesting that the organisms lived together. In certain situations, these occurrences are real, reflecting mixing and rebedding of previously deposited material. In most cases, however, the mixing of what should be incompatible palaeoecologic assemblages was an artefact of the sampling technique. Consequently, stratigraphic sections were recollected and reanalyzed with a focus on stratigraphic detail at the 1 – 2 cm interval. It is this highly detailed analysis along with the discovery of a critical outcrop section (LP-IS) that provides the additional palaeolimnologic information for this paper. It is not the purpose of this paper however to detail the wealth of palaeoecological data that presently exists. Instead, the focus is on the rapidity and magnitude of palaeolimnological change, and subsequent inferred palaeoclimatological change.

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(Figure 3). Obviously L. staplini is extremely euryhaline, but high abundance of the ostracod in the Estancia sediment is invariably associated with mid-salinity associations around 35,000 ppm TDS. At higher salinity and in fresher-water bodies (as indicated by other organisms), L. staplini occurs in low abundance. Marine foraminifers also occur in Estancia Valley sediment, and where present a salinity of approximately 35,000 ppm is indicated (Figure 3).

Fresh-water Indicators The most typical and definitive fresh-water indicator in the Estancia Valley sediment appears to be cutthroat trout (Oncorhynchus clarki). However, while cutthroat trout bones, even skin fragments, are common in the fresh-water phases during the late Wisconsinan, its usefulness as a fresh-water indicator is, in fact, somewhat limited. Cutthroat trout, a native North American species, is most commonly found in modern infrahaline (<500 ppm) lakes

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and rivers, but it is also reported from lakes with TDS varying from 5,000 ppm to 12,000 ppm (Bachhuber, 1989). Its occurrence, therefore, in Estancia Valley sediment seemingly constrains salinity to <12,000 ppm TDS. Trout fossils, however, are a good initial indicator of pluvial conditions in general, even though water quality could have been brackish instead of fresh. Fortunately trout fossils are most often associated with other fresh-water indicators during early lake evolution. As these other indicators disappear from the record and trout continues, it is reasonable to assume that water levels dropped with resultant evolution from fresh water to brackish water. Certain ostracods in particular constrain water quality to truly fresh-water conditions, and these taxa are always found in association with trout bones. Although restricted in range, Cytherissa lacustris is the most diagnostic fresh-water indicator of all the organisms found in Estancia Valley sediment (Figure 3). C. lacustris is found only in modern, deep, coldwater, infrahaline (<500 ppm) water bodies of holarctic distribution (Delorme, 1969). Its appearance in two discrete stratigraphic intervals, therefore is unquestionable evidence of fresh-water conditions, and likely deep-water conditions. Candona caudata is another key indicator of fresh-water conditions. This ostracod is reported in modern water of salinity range between 20 and 2054 ppm TDS (Figure 3). Somewhat less useful is Limnocythere ceriotuberosa , which is found in modern aquatic bodies ranging in salinity from 500 to 25,000 ppm (Figure 3). In isolation, the more euryhaline nature of L. ceriotuberosa can reflect highly brackish conditions as well as fresh-water bodies. However, as L. ceriotuberosa always occurs with trout fossils, its presence in the Estancia sediment reflects salinity of <12,000 ppm. Two other ostracods occur in interpreted fresh-water conditions in the Estancia Valley. Candona rawsoni and L. staplini are always found in association with other fresh-water indicators. Conversely, they are also almost always found in saline environments, and thus the simple occurrence of either or both of these ubiquitous species poorly constrains water quality. Nonetheless, an important stratigraphic relationship is seen in much of the Estancia record, although not yet substantiated in the ecologic record. The ostracods Candona rawsoni and L. staplini dominate throughout the Estancia Valley stratigraphic sequence and occur in virtually all stratigraphic horizons. L. staplini is found in modern bodies of water varying in salinity from fresh water to almost 200,000 ppm TDS. C. rawsoni occurs in modern environments that range in salinity from fresh water to about 43,000 ppm TDS. Although both are euryhaline in modern environments, there appears to be a strong correlation between the relative abundance of these ostracods and lake salinity. When L.

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staplini occurs in numbers much greater than C. rawsoni , other organisms (such as foraminifers) indicate that salinity is around 35,000 ppm or higher. Conversely, when the C. rawsoni /L. staplini ratio (Cr/Ls) is approximately 1:3 to greater than 1:1, other indicators (such as cutthroat trout) demonstrate that fresh-water or slightly brackish-water conditions prevailed. This relationship still needs to be verified in modern environments, but it is a relatively constant relationship in Estancia Valley sediments. In an Estancia Valley pluvial lake, a relatively large Cr / Ls ratio is therefore interpreted to indicate fresh-water or near fresh-water conditions. The final fresh-water indicator discussed here is Pediastrum cf. boryanum, a common cosmopolitan green algae found in modern fresh-water environments in Europe and North America. Even though the species present within the Estancia sedimentary record has yet to be definitively identified, the genus as a whole is a good indicator of fresh-water environments (Figure 3). Unlike the other Estancia Valley fossil organisms that are extracted directly from dried sediment (Bachhuber, 1989), Pediastrum is a component of samples processed for pollen content. 2.3. STRATIGRAPHY AND BIOSTRATIGRAPHY The sediment samples used for this study were collected from deflation basin E-28 and section LP-IS (Figure 1). As palaeontology and sedimentology are consistent throughout the Estancia Valley, however, these exposures are representative. Importantly, E-28 contains the thickest outcrop section of late Wisconsinan sediment, and appears to represent the deepest-water sedimentation throughout this time interval. The location of E-28 coincides with [what is believed to have been] the topographic low point of the depositional floors of the various late Wisconsinan pluvial lakes. Even during desiccation events the sedimentary section at E-28 was least likely to be truncated or deflated. This geologic record is the most continuous and has the greatest stratigraphic resolution. The LP-IS section is a small outlier located on the east margin of Laguna del Perro (Figure 1). Here, erosion has removed any sediment overlying the upper part of the Late Lake Estancia section, with the exception of a thin cap of late Holocene aeolian material. Fortuitously, erosion of the upper portion of the section and its isolation from the deflation basin margin has produced a stratigraphic section that is thoroughly dry, and nearly vertical in development. It is here that the lateral aspect of specific stratigraphic intervals can be best demonstrated. Section LP-IS is approximately 1 km west of the E-28 section. As such,

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it contains a stratigraphic section that is slightly thinner and depositionally upslope from the section outcropping at E-28. The Estancia Valley stratigraphic section is characterized by intercalated flint-gray silty clay and white, thin-bedded gypsarenites. Palaeontologic data demonstrates that the alternating nature of clay units and gypsarenite units is reflective of oscillating water depth, with gypsarenite deposition regressing into the interior of the various lacustrine basins as water depth decreases. Likewise, an increase in clay deposition mirrors lake expansion. 2.3.1. Basal Unit At E-28, the late Wisconsinan section (Figure 4) rests on an unconformable surface developed on clay and gypsarenite of the early to middle Wisconsinan-age La Salina complex ages from >48,800 to ca. 30,000 yr BP). The basal unit of the late Wisconsinan consists of a poorly-sorted pebbly, sandy silt that contains charophyte casts (Bachhuber, 1989), highly abraded Ruppia seeds (Rutter et al. 1992), horse-teeth fragments, and saline-indicator ostracods. The basal unit, highly variable in thickness to a maximum of 90 cm, also contains an abundance of Palaeozoic-age flint, chert, and limestone clasts weathered from the Manzano Mountains on the western flank of the Estancia Valley. Sedimentologic aspects of the unit indicate that the palaeontologic assemblage, for the most part, represents reworked and rebedded La Salina Complex fossils deposited by flash floods when the valley floor was subaerially exposed. Within, and at the top of the section, cutand-fill structures are common. At one location, the overlying pluvial section is deposited as slumped infill in a 1 m deep trough cut into and through the basal unit. 2.3.2. Fresh-water Lake 1 A 1 m - thick flint-gray clay rests unconformably on the Basal Unit. This stratigraphic sequence, referred to as the “initial freshwater phase” of Late Lake Estancia (Bachhuber, 1989) clearly marks the first fresh-water and deeper-water event of the late Wisconsinan. Bachhuber (1989) originally envisioned this fresh-water phase arising and expanding from an established saline lake. The recognition, however, of a lower and upper unconformable surface of the Basal Unit indicates that the lake expanded from a subaerially exposed valley floor and rapidly evolved into a fresh-water body. Palaeontologic evidence traces early lake evolution from brackish-water to fresh-water conditions. The fresh-water phase is characterized by the occurrence of cutthroat trout, Pediastrum, tiger salamander (Ambystoma tigrinum), and the ostracod Candona caudata, a fairly good indicator of salinity less than about 2000 ppm TDS. Salamander bones have been dated at 24,300 ± 560 yr BP

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(AA-1868), and ostracods at the same stratigraphic level have been dated at 23,510 ± 240 yr BP(AA-14068).

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Owing to the uncertainties associated with bone dates, the latter age appears to be more reliable. The ages date the time of maximum development of fresh-water and inferred deepest water conditions. Stratigraphically, the fresh-water interval is about 10 cm thick with the total Lake 1 sequence (including brackish-water conditions) being 40 cm thick, suggesting that the high lake stand had a short life. Ostracod data indicate that following the fresh-water peak, the lake rapidly collapsed into a brackish-water phase, followed by a saline-lake phase, and finally desiccation.

2.3.3. Ruppia Zone 1 Following the collapse of Lake 1, the floor of the Estancia Valley was subaerially exposed on a number of occasions and covered intermittently by shallow, saline lakes at other times. The subaerial and saline lake events encompass approximately 1 m of stratigraphic section at E-28, consisting of silty clay deposited during saline lake phases and poorly- to wellsorted gypsarenites. At LP-IS the subaerial event is characterized by a 1 m thick unconformably-bound depositional sequence containing large numbers of abraded Ruppia seeds, ostracods M. ingens and H. salinas , and intermittent zones with very high abundances of L. staplini . The seeds and some of the ostracods were reworked and redeposited as epiclastic assemblages from one of the many ephemeral saline-lakes that existed following Lake 1 desiccation. The Ruppia Zone 1 sedimentologic sequence at LP-IS is highly variable in thickness, locally pinching out completely, juxtaposing the Lake 1 sequence directly in contact with younger Lake 2 sediment. Ostracod valves from the upper part of the interval are dated at 20,520 ± 200 yr BP (AA-14067). Although this value approximates the correct time frame, the age is slightly younger than that determined for the overlying Lake 2 sequence. The Ruppia Zone 1 age determination, even though out of stratigraphic context, suggests that the Ruppia seeds themselves have been reworked from late Wisconsinan sediment, rather than from the La Salina complex that dates older than ca. 30,000 yr BP. 2.3.4. Fresh-water Lake 2 As indicated by the LP-IS section, Lake 2 sediment rests unconformably on sediment from a final subaerial event of Ruppia Zone 1. At E-28, the Lake 2 section is only about 10 cm in thickness. The section contains cutthroat trout, and ostracods C. caudata and L. ceriotuberosa, neither of which occur in abundance. The occurrence of C. caudata constrains the upper limit of salinity to about 2,000 ppm TDS. The upper and lower bounding unconformable surfaces of the sedimentary sequence along with the occurrence of fish, L. ceriotuberosa, and C. caudata indicate the existence of a discrete fresher-water lake,

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albeit one of short-term duration. While water depth is unknown, Lake 2 is assumed to be relatively deep and of large surface extent. Ostracod valves from this horizon have been dated at 21,060 ± 180 yr BP (AA-14066). 2.3.5. Ruppia Zone 2 Lake 2 is unconformably overlain by Ruppia Zone 2. The sedimentologic aspects here are similar to those of Ruppia Zone 1, but the stratigraphic interval is much thinner. Ruppia seeds are not abundant, but epiclastic deposition is indicated. At LP-IS the zone is less than 10 cm thick, bounded by unconformable surfaces, and has a very low abundance of ostracods. In places, Lake 3 sediment seemingly lies conformably on Lake 2 sediment even though unconformable relationships are laterally well exposed. At E-28, the section is about 15 cm thick, contains saline-indicator ostracods such as M. ingens and H. salinas, and at one stratigraphic level, abundant L. staplini. The sedimentologic and palaeontologic data suggest that the section at LP-IS was subaerially exposed, but this may not have been true at E-28. Here the ostracod assemblage appears to be depositionally intact, indicating the existence of a saline lake throughout the time interval. If a continuous saline lake existed, based on basin morphometry, it would have covered no more than a few There is no chronology for Ruppia Zone 2, but based on stratigraphic thickness and the apparent lack of unconformable surfaces at E-28, the interval must have been of short duration. 2.3.6. Fresh-water Lake 3 At E-28 the flint-gray clay of Lake 3 is approximately 20 cm thick and contains cutthroat trout, Pediastrum, and the ostracod L. ceriotuberosa (Figure 4). The alga Pediastrum especially constrains salinity to infrahaline conditions. Trout bones were not found at the LP-IS section and the sediment was not processed for Pediastrum, although L. ceriotuberosa occurs. C. rawsoni /L. staplini ratios vary from 1:2 to greater than 1:1, a relationship that we interpret as a reflection of fresh-water to slightly brackish-water conditions. No age exists for Lake 3, but this phase must be younger than ca. 21,000 yr BP (the age of Lake 2) and older than ca. 20,000 yr BP (the age of the upper portion of the overlying Ruppia Zone 3). This is also the age bracket for the underlying Ruppia Zone 2. Obviously, Lake 3 is younger than the underlying saline lake and subaerially deposited sediment of that zone. It is estimated here that the age of Lake 3 is approximately 20,500 yr BP. 2.3.7. Ruppia Zone 3 Ruppia Zone 3 is similar in all aspects to the older Ruppia Zone 1. It unconformably overlies Lake 3 sediments, reflects multiple saline-lake sequences and subaerial processes,

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albeit one of short-term duration. While water depth is unknown, Lake 2 is assumed to be relatively deep and of large surface extent. Ostracod valves from this horizon have been dated at 21,060 ± 180 yr BP (AA-14066). 2.3.5. Ruppia Zone 2 Lake 2 is unconformably overlain by Ruppia Zone 2. The sedimentologic aspects here are similar to those of Ruppia Zone 1, but the stratigraphic interval is much thinner. Ruppia seeds are not abundant, but epiclastic deposition is indicated. At LP-IS the zone is less than 10 cm thick, bounded by unconformable surfaces, and has a very low abundance of ostracods. In places, Lake 3 sediment seemingly lies conformably on Lake 2 sediment even though unconformable relationships are laterally well exposed. At E-28, the section is about 15 cm thick, contains saline-indicator ostracods such as M. ingens and H. salinas, and at one stratigraphic level, abundant L. staplini. The sedimentologic and palaeontologic data suggest that the section at LP-IS was subaerially exposed, but this may not have been true at E-28. Here the ostracod assemblage appears to be depositionally intact, indicating the existence of a saline lake throughout the time interval. If a continuous saline lake existed, based on basin morphometry, it would have covered no more than a few There is no chronology for Ruppia Zone 2, but based on stratigraphic thickness and the apparent lack of unconformable surfaces at E-28, the interval must have been of short duration. 2.3.6. Fresh-water Lake 3 At E-28 the flint-gray clay of Lake 3 is approximately 20 cm thick and contains cutthroat trout, Pediastrum, and the ostracod L. ceriotuberosa (Figure 4). The alga Pediastrum especially constrains salinity to infrahaline conditions. Trout bones were not found at the LP-IS section and the sediment was not processed for Pediastrum, although L ceriotuberosa occurs. C. rawsoni /L. staplini ratios vary from 1:2 to greater than 1:1, a relationship that we interpret as a reflection of fresh-water to slightly brackish-water conditions. No age exists for Lake 3, but this phase must be younger than ca. 21,000 yr BP (the age of Lake 2) and older than ca. 20,000 yr BP (the age of the upper portion of the overlying Ruppia Zone 3). This is also the age bracket for the underlying Ruppia Zone 2. Obviously, Lake 3 is younger than the underlying saline lake and subaerially deposited sediment of that zone. It is estimated here that the age of Lake 3 is approximately 20,500 yr BP. 2.3.7. Ruppia Zone 3 Ruppia Zone 3 is similar in all aspects to the older Ruppia Zone 1. It unconformably overlies Lake 3 sediments, reflects multiple saline-lake sequences and subaerial processes,

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and has an upper unconformable surface. At LP-IS the zone is highly variable in thickness and laterally pinches in and out across the face of the LP-IS outcrop. Numerous unconformable surfaces and cut-and-fill structures are in evidence throughout the interval. These aspects are clearly observable at LP-IS, and are substantiated to a lesser degree at E28. At E-28, although an upper unconformable surface is difficult to demonstrate, multiple unconformable surfaces appear within the section. Ruppia seeds occur in large numbers in an epiclastic assemblage, and in certain intervals there is a paucity of other fossils, including very low numbers of ostracods. These low abundance intervals represent dry playa deposition. Ostracod fragments associated with abraded Ruppia seeds in a cut-and-fill structure in the lower portion of the zone have a age of 18,300 ± 170 yr BP (AA-14065). From the middle of the zone, in a similar association, Ruppia seeds have an age of 20,040 ± 240 yr BP (AA-1867). The younger age is significantly out of stratigraphic context, and contamination from modern roots is suspected. The ca. 20,000 yr age, however, is in stratigraphic context. 2.3.8. Fresh-water Lake 4 (Late Lake Estancia) Lake 4 unconformably overlies Ruppia Zone 3 at E-28 and LP-IS. The sedimentologic interval contains a wealth of fresh-water indicators that led to the initial recognition of freshwater conditions in the late Wisconsinan section of the Estancia Valley, and the designation of Late Lake Estancia (Bachhuber, 1971). Early work, however, did not recognize the complexity of the Wisconsinan-age section nor the palaeoclimatic implications. The designation Late Lake Estancia (Bachhuber, 1989) included all the previously described Late Wisconsinan sediment of this paper up to the Estancia Playa Complex (discussion follows later), a stratigraphic interval of about 6 m in thickness. Here, the term Late Lake Estancia refers only to the sediment section of Lake 4, an interval of approximately 40 cm in thickness. At E-28, the Late Lake Estancia section consists of flint-gray clay overlying the Ruppia bearing, iron stained, silty clay of Ruppia Zone 3. Analysis of sediment sampled at 2-cm thick intervals demonstrates a rapid transition into fresh-water conditions with a basal sediment unit comprised of both fresh- and saline-water indicators. A few centimeters higher in the section, Late Lake Estancia sediment consists of a low-abundance ostracod assemblage where the C. candona /L. staplini ratio is approximately 1:1, indicating developing fresh-water conditions. The assemblage 4 cm vertically above this level

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contains trout fossils and Pediastrum, and is dominated by C. caudata and Cytherissa lacustris. The latter ostracod is found only in modern, deep, cold-water, infrahaline (<500 ppm) water bodies. The occurrence of C. lacustris especially marks the glacio-pluvial maximum of the late Wisconsinan. High abundance of C. lacustris occurs only through a vertical stratigraphic interval of 6 cm, with very low abundance within the overlying 6 cm. Therefore, the freshest-water and deepest-water conditions seemingly prevailed over a very short time. It is at this time that the highest strandline ringing the basin may have formed, indicating maximum lake depth of about 52 m and surface extent in excess of At LP-IS, where the section represents an intermediate position between littoral zone deposition and the deep-water deposition of E-28 detailed palaeontologic analysis (Figure 5) reflects the same basic sequence as seen at E-28 but with higher resolution. The upper

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portion of Ruppia Zone 3 is characterized by a distinct saline-lake interval unconformably overlain by the Ruppia -bearing reworked sediment deposited subaerially on a dry playa surface. This sediment, in turn, was reworked as Late Lake Estancia began to develop. The initial depositional phase of Late Lake Estancia consists of a mixture of the saline-indicators from the uppermost portion of Ruppia Zone 3, and the initial fresh-water inhabitants of the growing and deepening lake. Ostracod values from this earliest lake-developmental phase are dated at 19,950± 190 yr BP (AA-14064). The overlying C. lacustris interval is dated at 19,740 ± 150 yr BP (AA-14063), and 19,760 ± 160 yr BP (AA-6329) with both ages derived from C. lacustris values from E-28. Following ca. 19,750 yr BP, C. lacustris disappears rapidly from the E-28 and LP-IS stratigraphic sections. This disappearance signals a decrease in depth of lake and volume with a resultant progressive increase in salinity to brackish-water conditions to 12,000 ppm TDS over the next 35-cm stratigraphic interval. Initially C. caudata remains in relatively high abundance, and after it disappears, C. candona abundance approximates that of L. staplini. Cutthroat trout bones are common throughout the highest lake stand and much of the stratigraphic interval represents decreasing water depth. The tight grouping of ages and the contrasting palaeoenvironmental conditions illustrated by Figure 6 is the clearest evidence of rapid and high magnitude palaeolimnologic change in the Estancia stratigraphic record. 2.3.9. Zone 4 At LP-IS, following the disappearance of C. lacustris, the Late Lake Estancia sequence is abruptly terminated (Figure 5). The longer stratigraphic fresh-water record exhibited at E28 apparently has been truncated by subaerial erosion at LP-IS indicating that Late Lake Estancia either desiccated or water level dropped to only a few metres in depth in the deepest portion of the basin, at E-28. A 15-cm thick gypsarenite capping the Late Lake Estancia sediment at E-28, however, suggests that desiccation occurred. 2.3.10. Lake 5 Following the collapse of Late Lake Estancia, Lake 5 developed rapidly in what appears to be 2 phases. An initial fresh-water phase is characterized by a Cr/Ls ratio of approximately 1:1 and a low-abundance of C. caudata, that indicates infrahaline or near infrahaline conditions. This initial phase was followed by a short-lived partial draw down

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phase, but water level recovered quickly as indicated by the appearance of ostracod L. ceriotuberosa. The section also exhibits a Cr/Ls ratio of greater than 1:1, and abundant fish fossils. Although L. ceriotuberosa has an extant salinity range of 500 – 25,000 ppm (Forester, 1986), its occurrence in the Estancia sediment, and its association with other fresher water organisms, seems to reflect salinity closer to fresh- or slightly brackish-water conditions. Therefore, Lake 5, even though it does not have infrahaline indicators other than in the initial phase, does appear to have been relatively fresh to slightly brackish through most of its existence. L. ceriotuberosa occurs over a stratigraphic interval of 30 cm and then disappears. In association with the disappearance of L. ceriotuberosa, total ostracod abundance significantly decreases, but the Cr/Ls ratio remains greater than 1:1, and cutthroat trout fossils are present. We believe that at this point the lake has evolved into a brackishwater system with TDS approximating 12,000 ppm, the upper limit tolerated by cutthroat trout. Although Lake 5 seemingly never approached the infrahaline, deep-water conditions of Late Lake Estancia, it remained as a persistent lake body over a stratigraphic interval of 1 m. Therefore, as compared to the other lake systems in the Estancia Valley, Lake 5 may have had the longest period of quasi-stability during the late Wisconsinan. Ostracod valves from the interval characterized by L. ceriotuberosa have a age of 16,890 ± 150 yr BP (AA14062). 2.3.11. Estancia Playa Complex The eventual desiccation of Lake 5 marked the beginning of the Estancia Playa Complex, a stratigraphic sequence characterized by subaerial erosion and deposition, and palaeontologic evidence supporting the occurrence of numerous discrete saline lakes. Gypsarenites and gypsiferous silts are dominant, and cut-and-fill structures and unconformable surfaces are common. Intervening clay-rich sediments have a great abundance of saline-indicator ostracods M. ingens and L. staplini . In addition, one saline lake from the upper portion of the Estancia Playa Complex has an abundance of 2 species of marine foraminifers (Bachhuber and McClellan, 1977). Foraminiferal tests from this interval have an age of 14,345 ± 105 yr BP (AA-6328). This age approximates the close of the Estancia Playa Complex sequence. 2.3.12. Fresh-water Lake 6 (Lake Willard) Lake Willard was originally identified by Bachhuber (1971) as consisting of a single freshwater phase. Detailed palaeontologic analyses of the flint-gray clay to gypsiferous silt

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section, however, demonstrates that the lake consisted of 2 distinct fresher-water phases separated by an intervening more saline phase, and possible desiccation. The first freshwater phase, here termed Lake Willard A, is characterized by a high-density ostracod assemblage of C. candona (Cr / Ls ratio greater than 1:1), L. ceriotuberosa, and low numbers of C. lacustris. Cutthroat trout bones and Pediastrum are common. Unquestionably, Lake Willard A was a fresh-water, infrahaline lake of approximately 500 ppm TDS. At E-28, Lake Willard A is represented by a total stratigraphic thickness of less than 10 cm, and a age derived from ostracod valves of 13,700 ± 105 yr BP (AA-6327). Following Lake Willard A’s freshest, deepest phase, water level dropped rapidly with a significant increase in salinity. A series of shallow saline lakes may have occupied the valley floor until Lake Willard B developed. Although it is suspected that desiccation did occur, sedimentologic evidence of this is not well demonstrated. Lake Willard B marks the final nearly fresh-water lake in the Estancia Valley late Wisconsinan history. Cutthroat trout reappears and C. rawsoni abundance is much greater than L. staplini abundance. But absence of other fresh-water indicators suggests that Lake Willard B was not as fresh or as deep as Lake Willard A. A age on ostracod valves from the interval is 12,460 ± 135 yr BP (Beta-25819). The stratigraphic thickness of Lake Willard B is approximately 10 cm. Salinity increased and water level decreased rapidly in Lake Willard B, with apparent evolution directly into a saline lake dated at 12,375 ± 95 yr BP (AA-6326). The absence of palaeontologic content suggests that this saline lake desiccated, followed by the development of a final discrete saline lake. Desiccation of this final saline lake marked the beginning of the development of the Willard Soil (Bachhuber, 1982), and the close of the late Wisconsinan in the Estancia Valley.

3.

Chronology and Lake Duration

The Estancia Valley sedimentologic and biostratigraphic record demonstrates that the late Wisconsinan was a time period of rapid and high magnitude palaeolimnologic change. Many of the palaeolimnologic events are well constrained by ages, with all fresh-water lake stages with the exception of Lake 3 having at least one definitive radiocarbon age. With the exception of two dates used in this paper, ages have been determined by Accelerator Mass Spectrometer (AMS) techniques on ostracod valves of single species of ostracods. The ostracod valves were meticulously hand picked from sediment samples and

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each AMS sample represents up to 2500 valves. Without the advent of AMS technology, few ages would be available from the Estancia Valley stratigraphic sequence. Based on chronologies, Lake 4 (Late Lake Estancia) and Lake Willard A and B reflect climate events that are well recorded in other areas of North America (Mickelson et al. 1983). Late Lake Estancia, likely the deepest and largest pluvial lake in the late Wisconsinan, should be time synchronous with the late Wisconsinan glacial maximum of ca. 18,000 yr BP. The age of Late Lake Estancia, however, is well constrained at ca. 19,800 yr BP. We believe the ~1800 year difference in chronology is due to an hard-water effect in the Estancia samples. A chronological difference of similar magnitude has affected the Lake Willard ages. Lake Willard appears to be correlative with the global Younger Dryas event, between ca. 10,000 to 11,000 yrs BP (Peteet, 1995). Lake Willard A dates at ca. 13,700 yr BP and Lake Willard B dates at ca. 12,500 yr BP. If Lake Willard A is the time equivalent to the Younger Dryas, the hard-water effect would represent a difference of about 2,700 years. On the other hand, if Lake Willard B represents the Younger Dryas, the hard-water effect is a difference of about 1500 years, an hard-water effect more comparable to that of Late Lake Estancia. The hard water effect and the timing of the Younger Dryas in the Estancia Valley have yet to be resolved, but it is proposed here that Lake Willard B is equivalent to the Younger Dryas. Lake Willard A is then a major event not yet recognized in the North American record. Each one of the 7 pluvial events recorded in the Estancia Valley late Wisconsinan stratigraphic section is believed to be a discrete fresh-, deep-water lake (Figure 6) separated by either a dry playa environment, or by a small, shallow saline lake. As such, the duration or life history of each of the individual lake stages can be calculated using sedimentation

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rates. Although prevalent bioturbation means that laminated clay is not a common depositional mode during the Estancia Valley lake histories, a number of laminated sections do occur. The laminated sediment usually consists of a couplet with an upper thin, black layer and a thicker lighter colored layer. Even though the laminations appear to be varves, this has yet to be verified. If the laminations do represent annual deposition, the average sedimentation rate derived from a number of laminated sequences from the various infrahaline and brackish-water lake stages is 3 mm/yr. This value is somewhat large for higher latitude North American lakes. occurring in well vegetated regions. However, 3 mm/yr of accumulation represents a low rate in environments where aeolian input would be high regardless of a xeric or mesic climate. At E-28, where the deepest water sedimentation has occurred through time, the various lake phases are well differentiated by lower or upper unconformable surfaces or palaeontologically distinct highly saline phases. Therefore, the stratigraphic thickness of each lake stage is easily and accurately calculated. Using an average annual sediment accumulation of 3 mm, the life-duration of each of the 7 lake stages as infrahaline (about 500 ppm) or brackish-water (<12,000 ppm) bodies is illustrated in Table 1.

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These calculations indicate that almost all of the lakes developed, stabilized, and desiccated in a time period near that of human perception. The total duration of the existence of freshwater or near fresh-water conditions during the late Wisconsinan is less than 900 years. As the upper part of the middle Wisconsinan La Salina complex (Bachhuber, 1992) is dated at ca. 30,000 yr BP, the 1000 years that the Estancia Valley was occupied by pluvial conditions represents only about 4 % of late Wisconsinan time. Conversely, 17,000 years or 96 % of late Wisconsinan time, was characterized by either dry playa or saline lake environments. The paradigm that the late Wisconsinan in the Southwest USA was a pluvial time period is grossly exaggerated. Pluvial conditions did exist, and the magnitude of pluvial events was great, but the climates, albeit repetitive, that give rise to these systems were of short duration. 4.

Palaeoclimate

The late Wisconsinan sedimentological and biostratigraphical record of the Estancia Valley is clear evidence of rapid and high magnitude palaeolimnological change. While the case has been made for fluctuating lake levels, this is inferred from the biostratigraphic evidence that actually reflects change in water quality (TDS), which is a function of water volume. In the environment of central New Mexico, a lake becomes alkaline enriched with decreased total salinity only by diluting existing high-salinity water with a large volume of enriched fresh water or by flooding a dry playa with fresh-water flow. This implies a climate change with resultant increased surface runoff and ground-water recharge that, in turn, produces high lake levels. Conversely, sulphate enrichment and high total salinity is a function of increased evaporation that produces lower lake levels, and eventually desiccation. Therefore, indirectly, dramatically fluctuating lake levels are documented in the Estancia Valley record. Most, if not all, of these fluctuations in lake level are climatically controlled. The volume of water in the valley at any given time is predominantly a function of surface runoff and ground-water recharge, and evaporation. As compared with present climatic conditions, the existence of a lake in the geologic record requires that the interrelationship between precipitation and temperature results in a positive water budget, with a large lake requiring significantly higher available water than that of a saline lake. Albeit an over simplification, the 7 high lake stages of the late Wisconsinan register times of increased precipitation and/or decreased mean annual temperature. Likewise, saline lakes and dry playas represent climatic conditions more similar to those of the present. In moving from dry playa to saline lake to a fresher- water lake important hydrologic thresholds must be

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crossed. As mentioned previously, the high elevation of the Estancia Valley floor is one of the important reasons contributing to the high-resolution biostratigraphic record. It could then be argued that valley-floor elevation places the region in a micro-climatic situation where crossing of hydrologic thresholds may involve only a slight change in precipitation and/or temperature. This is thought to be the case for crossing the dry playa/saline lake threshold, but evolution into a deep, fresh-water pluvial system requires a more significant climatic change. It is beyond the scope and purpose of this paper to dwell upon the various climatic scenarios that could give rise to alternating pluvial and inter-pluvial episodes. The data presented here simply supports such climatic change. However, the evidence supporting rapid, multiple, and high magnitude change during the Wisconsinan, along with the evidence that supports the view that individual pluvial lakes existed for only a short time, decades in some cases, brings into question the present disposition of the North American and European glacial climatic record. By extension, we question the inference that North American and European glacial climate during the late Wisconsinan / late Weichselian was marked by relatively uniform conditions until the commencement of glacial retreat, ca. 14,000 BP. Instead, during this time period, it is likely that numerous and significant climatic oscillations occurred that were not recorded in the low-resolution glacial record. 5.

Conclusions

An eight-metre-long high-resolution stratigraphic section from the Estancia Valley spans a 12,000 year period of late Wisconsinan time from ca. 24,000 yr BP to ca. 12,000 yr BP. The section has been analyzed in detail for its palaeontologic content. Fossils indicate that the time interval was characterized by a dynamic palaeoenvironments. Wide fluctuations in the range and abundance of various species/genera are believed to be a response to dramatically changing palaeolimnological conditions, particularly in terms of total salinity. Variations in salinity must reflect changes in water depth and volume. The record is interpreted as representing 7 discrete fresh-water lake stages and intervening periods of desiccation. Crossing hydrologic thresholds from dry playa to saline lake to fresh-water lake or this succession in reverse can only be a function of climate change. Based on the Estancia Valley palaeontologic record, the following conclusions are drawn. The Estancia Valley has a high-resolution record of rapid, multiple, and highmagnitude palaeolimnological change. During the glacial maximum, Late Lake Estancia attained a depth of about 52 m and had a surface area in excess of 1,200 The other pluvial lakes were likely to have been comparable in depth and size.

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Wisconsinan palaeolimnological change, only in part due to the high elevation of the valley floor, is mainly a product of rapid, multiple, and high-magnitude palaeoclimatic change. Individual fresh-water lakes developed quickly and survived only for short periods of time, between a few centuries and a few decades. The total time duration of the 7 pluvial lakes was less than 900 years. This represents only 4 % of late Wisconsinan time, inferring that glacio-pluvial climate, while dramatic in outcome, was not the norm. Instead, the late Wisconsinan of the Southwest USA is better characterized by the long total duration of inter-pluvial climatic conditions, probably similar to that of the present. The high-resolution preservation of the Estancia Valley’s palaeolimnological record and the inferred palaeoclimatological interpretations throughout the late Wisconsinan (late Weichselian) brings into question the inference that North American and European glacial climate was marked by fairly uniform conditions until the commencement of glacial retreat ca. 14,000 yr BP. The rapidity, variability, magnitude, and duration of climate change evidenced in the Estancia Valley should present a challenge to palaeoclimate modellers in their efforts to improve our understanding of climate change.

Acknowledgments Funding for many of the ages used in this paper was provided by the Office of Research, University of Nevada, Las Vegas and the NSF – Arizona AMS Facility at the University of Arizona. Their contributions and those of the anonymous reviewers are greatly appreciated.

References Bachhuber, F. W. (1971). Palaeolimnology of Lake Estancia and the Quaternary history of the Estancia Valley, central New Mexico. Unpublished Ph.D. thesis, University of New Mexico, Albuquerque, NM, USA, 238. Bachhuber, F. W. (1982). ‘Quaternary history of the Estancia Valley, central New Mexico’, in Grambling, J. A., and Wells, S. G. (eds.), Albuquerque Country II. New Mexico Geological Society, 343-346.

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Bachhuber, F. W. (1989). ‘The occurrence and palaeolimnologic significance of cutthroat trout (Oncorhynchus clarki) in pluvial lakes of the Estancia Valley, central New Mexico.’ Geological Society of America Bulletin, 101, 1543-1551. Bachhuber, F. W. (1992). ‘A pre-late Wisconsin palaeolimnologic record from the Estancia Valley, central New Mexico’, in Clark, P. U., and Lea, P. D. (eds.), The last interglaciation/glaciation transition in North America: Geological Society of America Special Paper 270, 289-307. Bachhuber, F. W., and McClellan, W. A. (1977). ‘Palaeoecology of marine Foraminifera in the pluvial Estancia Valley, central New Mexico.’ Quaternary Research 7, 254-267. Delorme, L. D. (1969). ‘Ostracodes as Quaternary palaeoecological indicators.’ Canadian Journal of Earth Sciences, 6, 1471-1476. Delorme, L. D. (1982). ‘Lake Erie oxygen; the prehistoric record.’ Canadian Journal of Fisheries and Aquatic Sciences, 39, 1021-1029. Delorme, L. D. (1990). ‘Freshwater Ostracodes’, in Warner, B.G. (ed.), Methods in Quaternary Ecology: Geological Association of Canada, Geoscience Canada, reprint series 5, 93-100. Delorme, L. D., Zoltai, S. C., and Kalas L. L. (1977). ‘Freshwater shelled invertebrate indicators of palaeoclimate in northwestern Canada during late glacial times.’ Canadian Journal of Earth Sciences, 14, 2029-2046. Forester, R. M. (1983). ‘Relationship of two lacustrine ostracode species to solute composition and salinity; Implications for palaeohydro-chemistry.’ Geology, 11, 435-438. Forester, R. M. (1986). ‘Determination of the dissolved anion composition of ancient lakes from fossil ostracodes.’ Geology 14, 796-798. Forester, R. M., Delorme, L. D., and Bradbury, J. P. (1987). ‘Mid-Holocene Climate in northern Minnesota.’ Quaternary Research 28, 263-273. Hawley, J. W. (1984). ‘The Ogallala Formation in eastern New Mexico’, in Proceedings, Ogallala Aquifer Symposium II, Texas Tech University Water Resources Research Center, 157-176. Kelley, V. C. (1972). ‘Geology of the Fort Sumner sheet, New Mexico.’ New Mexico Bureau of Mines and Mineral Resources Bulletin 98, 51. Mickelson, D. M., Clayton, L., Fullerton, D. S., and Borns, Jr., H. W. (1983).‘The Late Wisconsin glacial record of the Laurentide Ice Sheet in the United States’, in Porter, S. C. (ed.), Late-Quaternary Environments of the United States, Volume 1, The Late Pleistocene. University of Minnesota Press, 3-37. Mifflin, M. D., and Wheat, M. M. (1979). ‘Pluvial lakes and estimated pluvial climates of Nevada.’ Nevada Bureau of Mines and Geology Bulletin 94, 57. Peteet, D. (1995). ‘Global Younger Dryas?’ Quaternary International 28, 93-104. Rutter, N., Bachhuber, F. W., and Lyons, G. (1992). ‘The use of seeds in aminostratigraphy of a Wisconsin palaeolimnological record from central New Mexico, U.S.A.’ Sveriges Geologiska Undersokning 81, 307-312. Titus, F. B., (1969). Late Tertiary and Quaternary hydrogeology of the Estancia basin, central New Mexico. Unpublished Ph.D. thesis, University of New Mexico, Albuquerque, NM, USA, 179.

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AEOLIAN GEOMORPHIC RESPONSE TO CLIMATE CHANGE: AN EXAMPLE FROM THE ESTANCIA VALLEY, CENTRAL NEW MEXICO, USA N. R. CATTO Memorial University of Newfoundland, St. John's, Canada, A1B 3X9 F. W. BACHHUBER University of Nevada, Las Vegas, Las Vegas, NV, USA, 89154-4010

Abstract Three generations of Holocene aeolian activity are preserved along the axis of the Estancia Valley, central New Mexico, USA. The transition from large dome dunes to parabolic dunes to loess reflects differences in sedimentation and environments, controlled by climate change. The aeolian landscape developed on the final lacustrine plain of a complex series of freshwater pluvial lakes that occupied the valley from ca. 24,000 BP to ca. 10,500 BP, during the maximum development and subsequent deglacial phases of the Laurentide glacier. Climate change associated with the disappearance of the final pluvial lake initiated dome dune formation. The dome dunes are characterised by low-angle tabular cross-laminations and tabular ungraded finemedium beds. The internal structures indicate that limited sediment availability controlled dome dune formation. Development was initiated and sustained by prevailing westerly winds, consistent in direction, but marked by gustiness. Accumulation was episodic, punctuated by periods of stability with weak mollisol development. Charcoal from one mollisol dates to ca. 8,500 BP. After this time, the dome dunes attained their greatest relief and were capped by a thick, cemented gypsiferous soil. A series of at least five stacked, weak mollisols lapped on to the flanks of the dome dunes, ca. 4,660 BP. Overlying the stacked soils and dome dunes is a second generation of parabolic dunes with associated deflation basins. These dunes developed under westerly winds varying in intensity. Spatial limitation of dune migration contributed to sediment retention, forcing dunes to increase in height and allowing stacking of successive generations of landforms. Parabolic dunes originating along the eastern margin of Laguna del Perro climbed from the playa floor and overrode the dome dunes. Sediment movement by 171

S.J. McLaren and D.R. Kniveton (eds.), Linking Climate Change to Land Surface Change, 171–192. © 2000 Kluwer Academic Publishers. Printed in the Netherlands.

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traction and saltation was hindered, slowing transportation rates, enhancing retention, and resulting in rapid accumulation. Sediment supply was limited by the presence of ephemeral standing water in the playa. Deflation of the marginal areas was further limited by periodic rainfall events, allowing diagenetic overgrowths that stabilised the original gypsiferous lacustrine plain. Depositional episodes were transitory in the dune field, not associated with marked climate change, and were separated by periods of stability. Modern aeolian activity involves the deposition of loess. Loess is also reworked by sheetwash, as well as in playa marginal areas of the larger parabolic dunes. Under modern climate conditions, aeolian sedimentation in the Estancia Valley is marked by very limited sand supply, almost exclusively locally derived, and high rates of retention. The discontinuous vegetation cover provides localised zones for trapping of fine suspended gypsite, allowing loess accumulation. Loess accumulates in sites preferential for vegetation development, such as inactive trough blowouts and hollows between adjacent dunes. 1.

Introduction

Aeolian deposits and landforms illustrate the impact of previously existing climates on the environment. The growing interest in the assessment of climate change by geomorphologists, coupled with increased understanding of the processes and dynamics of aeolian sedimentology, highlight the necessity for detailed studies of aeolian successions in regions which no longer support active dunes. Aeolian dunes are common landforms in many basins throughout the southwestern United States. Although active dunes categorise many areas, such as the Mojave Desert of California and White Sands, New Mexico, many examples of inactive dunes are also present. The inactive dunes support sparse assemblages of vegetation, and locally have been subject to grazing by ungulates and domestic herd animals. Their relative stability and inactivity, therefore, must predominantly be attributed to climate change. These dunes serve as indicators of previous climate, as well as potential precursors of landforms that could develop if climate change proceeds in the 21st century. Study of the internal structures and morphology of these inactive dunes thus provides proxy data concerning past climate fluctuations. The Estancia Valley of central New Mexico is located 80 km southeast of Albuquerque (Figure 1; Plates 1 and 2). The basin floor, with a minimum elevation of 1,850 m a.s.l., is surrounded by topographic highs reaching over 3,000 m a.s.l.. At present, the Estancia Valley is thus completely isolated from surface drainage to the adjacent Rio Grande and Pecos River drainage basins. The western margin of the Estancia Valley is defined by Laguna del Perro, a 19 km long deflation basin oriented north-south, which supports shallow ephemeral lakes at sporadic intervals.

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During several phases of the Quaternary, playa and pluvial lakes developed in the Estancia Valley, depositing silt and clay-textured gypsite deposits in the centre of the basin, and shoreline deposits of gypsarenite and gravel along the basin margins. The most recent of the playa (lake) events involved the formation of a complex series of relatively short-lived but spatially dominant freshwater lakes, during the Late Wisconsinan from ca. 24,000 BP to 10,500 BP (see Bachhuber and Catto, this volume). Development of these freshwater (playa) lakes, and the fluctuations in water level and lacustrine chemistry recorded in their sediments and fossil assemblages, responded to climate changes linked to the Laurentide glacier of western Canada and the northern United States. When glacial recession occurred in Montana and Alberta, the changing climate resulted in the rapid desiccation of the Estancia Valley, exposing the playa gypsite and gypsarenite to aeolian reworking and transportation. Throughout the lattermost Late Wisconsinan and the Holocene, changing climate resulted in the development of a succession of aeolian landforms, dominated by "small dunes" and "great dunes" (Titus, 1969). That the "small dunes" and "great dunes" are of different ages was recognized by Titus, and later by Bachhuber (1971), but little subsequent aeolian research was undertaken. The small dunes, however, presented researchers with a serious problem -- why were the small dunes of such a rounded and lenticular configuration? The explanation advanced by Titus (1969) and later accepted by Bachhuber (1971) is that the dunes were modified by wave activity of a shallow Holocene lake following dune deposition. Bachhuber (1971) called this lake phase

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Lake Meinzer, but indicated that there was no palaeontologic data to support its existence. This paper recognizes that the small dunes are domal dunes and their form is not the result of wave modification.

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The dome dunes are developed on the floor of the Willard Soil (Bachhuber, 1971), a gypsite unit that caps the Late Wisconsinan Lake Willard plain and sediments. The initial dome dune sediments, in turn, are overlain by weakly developed mollisols. Subsequently, a second generation of aeolian activity, represented by parabolic dunes, shaped the Estancia Valley landscape. After a period of stability, the surfaces of the parabolic dunes were reworked by aeolian action in the late Holocene. The aeolian assemblages preserved in the Estancia Valley thus provide a proxy record of climate fluctuations throughout the last 10,500 years.

2.

Climate

Today the central portion of the Estancia Valley is arid (10( C mean annual temperature, 120 mm mean annual precipitation). The slopes flanking the valley are semiarid. The warmest month is July (21° C) and the coldest month is January (0° C). Temperature and precipitation conditions in the valley produce an annual water-budget deficit in excess of 1500 mm (Harbour, 1958). As a result, perennial water bodies do not exist at the lower elevations, and the floors of the deflation basins are dry playas throughout the year.

3.

Distribution and geomorphology of dunes

The Estancia dune field encompasses an area of 240 km2, extending 9 km east of Laguna del Perro (Plates 1 and 2). More than 60 individual dunes and associated deflation basins can be recognised. The dunes are confined to the central valley floor that has a general elevation of 1,835 m, and all overlie playa lacustrine sediments. The initial generation of aeolian activity was marked by the construction of dome dunes, which are present throughout the Estancia area. In the western part of the field, the dome dunes are overlain by younger parabolic dunes. Dome dunes from the initial phase of aeolian activity are exposed on the surface in the eastern part of the dune field, suggesting that sediment supply from Laguna del Perro during subsequent phases was insufficient to inundate these landforms. The dome dunes exposed on the surface in the eastern area are ovate, with long axes aligned east-west to ENE-WSW. Heights range from 1.5 m to 2.5 m, generally decreasing eastward. Principal axial lengths range from 15-20 m, and intermediate axial widths from 10-15 m, with typical length:width ratios of 1.2 : 1. Slope angles range from 0° to 10°. In the area directly east of Laguna del Perro, the dome dunes have been buried by subsequent parabolic dune development. Cross-sections exposed along playa lake margins indicate that the dome dunes had heights of 3 m -6 m, with slope angles ranging

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from 0°-12°. These dome dunes are ovate, with long axes oriented between 080°-090°. The maximum width exposed in cross-section is 20 m. The dominant features on the landscape in the western part of the Estancia Valley are the second generation parabolic dunes. The dunes are laterally coalescent adjacent to Laguna del Perro, forming a continuous transverse dune 17 km long and up to 44 m in height above the floor of Laguna del Perro, to an elevation of 1,898 m a.s.l.. Parabolic dunes extend 9 km to the east of Laguna del Perro, systematically decreasing in height, areal extent, and spatial density from west to east. Individual parabolic dunes are up to 25 m in height, 1.0 km in length parallel to the net transport direction, and 0.5 km in width. Stoss and lee slope angles reach maximum values of 8° and 32°, respectively. On the lee slopes, the maximum angle does not occur directly below the crestline, but generally is located in the basal half of the slipface. Crests are flat-topped, although subsequent deflation has greatly modified the surfaces. The ratios of arm length : distance between arms generally approximate 1:1. The northern and southern arms are equal in length, and the dunes are symmetrical about their long axes. Heights, lengths, and widths all decrease eastward. Concomitantly, the isolated playa-floored deflation basins with associated parabolic dunes increase in elevation towards the east. The elevation of the playa floors slope downward towards the axis of Laguna del Perro, as a result of the sloping water table that controlled the maximum deflation elevation. The easternmost parabolic dunes, with heights of 6 m, modal lengths of 0.5 km and widths of 250 m, are gradational to shield dunes, with minor slip faces and poor arm development. The shield dunes represent transitional forms to dome dunes. Maximum lee slope angles decrease eastward, varying from 29° on parabolic dunes in the central part of the Estancia valley to 20° on the shield dunes. The azimuth orientations of the modal long axes vary from 060°-095° throughout the dune field, with no systematic spatial variation. Local variations in axial orientation appear to be related to interference from adjacent upwind dunes. The orientations indicate formation by southwesterly to westerly winds. Both the initial and second generation dunes have undergone modification by more recent aeolian activity. Trough blowouts, with depths to 1.5 m and widths to 3 m, are present on the surfaces of the transverse dune and the larger parabolic dunes in the western part of the dune field. The trough blowouts have locally been reactivated as a result of cattle and human traffic. Azimuth orientations are variable, and many troughs are irregularly curved due to animal traffic, but the modal orientations are ENE-WSW Saucer blowouts, with maximum diameters of 2 m and maximum depths of 0.3 m, are also present on crestal surfaces of parabolic dunes, and on the highest flat summits of the transverse complex paralleling Laguna del Perro. The saucers vary in shape from oval, with long axes aligned WSW-ENE, to circular. In the eastern part of the Estancia Valley, a third generation of small dome dunes has developed perched on the larger underlying dome dunes of the first generation and second-generation parabolic and shield dunes. Typically, these dome dunes are oval in

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plan, with lengths up to 8 m, widths up to 5 m, and heights less than 1.5 m. The maximum lee slope angles are 6°-10°. Slip faces are very poorly defined, and are locally absent where dome dunes impinge on preexisting topographic rises. The azimuth orientations of the long axes of these ovate forms are aligned between 060° 100°. Small saucer blowouts are present on the surfaces of these dunes. Interspersed in the saddles between all the dune types are discontinuous, poorly sorted structureless infills dominated by gypsite with minor amounts of gypsarenite. Gypsarenite and gypsite slope failure deposits are also present along the margins of the larger, older dunes, and in areas where the bases have been undercut by playa-induced erosion. This material represents remobilised loess (brickearth). Unstratified gypsarenite/gypsite discontinuously infills trough blowouts on the larger parabolic dunes, and on the transverse dune complex. Gypsite and fine gypsarenite loess caps, tapering westward and up to 70 cm thick, occur on the crests of more parabolic dunes in the western part of the complex. 4.

Dune stratigraphy

The three generations of dunes present at Estancia differ in their internal stratigraphy, reflecting differences in the processes of sedimentation and in the depositional environments. A representative succession preserved adjacent to playa E 28, an isolated playa-floored deflation basin directly to the east of Laguna del Perro (see Bachhuber and Catto, this volume) allows direct comparison of the stratigraphy of successive dune generations, with younger parabolic dune deposits superimposed on older dome dune sediments. The dome dune sediments adjacent to E 28 in turn overlie lacustrine deposits of Late Wisconsinan Lake Willard B, which desiccated ca. 10,500 BP (see Bachhuber and Catto, this volume). 4.1

DOME DUNE SEQUENCE

The dome dune sediments are illustrated in Plates 3 and 4. The basal 80 cm of the succession consists of intercalated pedogenically altered gypsarenite and gypsite, with recrystallised selenite, interbedded with rippled sediments and discontinuous horizontal laminae. Recrystallised horizontal laminae of selenite, with maximum thicknesses of 2 mm, are present along the bases of some ripples. Wedge-form pendants of selenite with maximum widths of 3 mm and extending to 8 mm depth are also developed beneath some contacts marked by textural changes or displaying ripple forms. The ripples are laterally discontinuous, with individual ripple forms and pairs of ripples separated by horizontally laminated or structureless gypsite, suggesting that ripple migration exceeded sediment supply. The ripples have amplitudes of 0.8 to 1.1 cm, wavelengths of 20 to 32 cm, and ripple indices (c.f. Tanner 1967) of 25 to 40, indicating aeolian origin. Lee faces are commonly preserved in entirety, and are asymptotic to both the lower and upper surfaces, with a maximum angle of 11°-16°. The direction of ripple migration, varying from azimuth 080° to 107°, indicates formation by westerly winds.

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Sediment within the ripples consists of both gypsarenite and sand-sized aggregated spheres of gypsite and other clay minerals. Aggregated clasts of clay-sized material result from aeolian deflation of adjacent exposed desiccated lacustrine or salina sediments, locally forming distinctive clay-dominated lunette dunes (Hills 1940, Price and Kornicker 1961) or mixed textural assemblages of less defined form (e.g. Teller 1972). At Estancia, the aggregated gypsite clasts do not form morphologically distinct units, and are interbedded within the gypsarenite. Deposition of the gypsite aggregates proceeded penecontemporaneously with formation of the gypsarenite ripples, rather than as texturally distinctive lunette dunes. The surfaces of the horizontally laminated and structureless strata all are marked by rain pits, 0.5 to 1.5 mm in depth and 5-10 mm in diameter, characteristically with weakly convex bases. The surface contacts of structureless sediments display greater proportions of pits than do the rippled and horizontally laminated strata, with overlapping pits common. In form and dimensions, the rain pits resemble those developed on modern gypsum sediments at White Sands NM. Isolated adhesion wart structures (c.f. Kocurek and Fielder 1982), with axial lengths of 1 cm and heights of 2 mm, are present on the upper contacts of isolated horizontal laminae. The overlying sediments, which form the majority of the dome dune, are largely composed of weakly developed horizontal to gently dipping gypsarenite laminae, each 5-8 mm thick. The laminae are horizontal in the central part of the dome, and increase in dip to 7° along the western flank and 3° along the eastern flank. Contact surfaces along the laminae are marked by adhesion wart structures and fossilised rain pits. Selenite crystals, to 1 cm length, are present along some contacts and invariably are aligned along the dip of the laminae, with axes parallel to the inferred wind direction. Isolated ripples within the laminated sequences have heights of 7 -11 mm, wavelengths of 29-36 cm, and ripple indices of 31 - 46. Ripples on the eastern flank of the dome dune indicate transport directions varying between 107° and 166°, with lee side slopes of 5°-7°. Ripples preserved on the western side of the dune, and in the crestal area, show variable transport directions ranging from 065° to 233°, with lee side slopes varying from 2° 7°. Rippled horizons are commonly developed overlying gently concave-downwards erosional contacts, and are encased in structureless gypsarenite. Depressions are typically excavated 3-5 cm below the surface of the uppermost truncated lamination, and have axial dimensions of 90-150 cm. These depressions represent small saucer blowouts excavated on the stoss and crestal areas of the dune surface, and infilled with rippled and structureless gypsarenite. Individual ripple orientations indicate that transport was locally towards the centre of the depressions along their upwind and lateral margins, with sediment filling the depressions from three sides. An episode of stability is indicated by the presence of a disturbed, poorly sorted gypsarenite and gypsite stratum 25-30 cm thick, 2.8 m above the base of the dome dune succession. The unit contains dikaka structures, rippled strata which have been partially truncated, disturbed horizontal laminations, and adhesion warts. This bed is directly

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overlain along an erosional contact by rippled and horizontally laminated gypsarenite, similar to the underlying aeolian sediments. The maximum thickness of sediment in the dome dune is 5.3 m. The uppermost surface of the dome dune is capped by a laterally continuous stratum of selenite granules, conforming to the underlying domal morphology. Fragments of organic detritus up to 3 mm in length are also present. Isolated selenite crystals up to 1.5 cm in length, aligned 180°-220°, are present along the basal contact, and pendant selenite crystals are also present. This stratum developed during a period of stability, and represents the conclusion of dome dune development. The selenite crystals were moved by traction, rolled normal to the prevailing wind direction. 4.2

PARABOLIC DUNE SUCCESSION

The dome dunes are overlain by parabolic dune sediments in the western part of the Estancia Valley. Along the eastern margin of Laguna del Perro, laterally coalescent parabolic dunes have formed the transverse dune complex. The succession preserved adjacent to E 28 is typical of lee-side preservation in the interior of parabolic dunes throughout the Estancia area (Plate 5).

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The basal 3.1 m of the parabolic dune sequence at E 28 is dominated by moderately defined tabular sets of low-angle cross-laminated sand-sized clay pellet aggregates, some of which contain lacustrine ostracods, and gypsarenite. The laminae are arranged in sets 2-15 cm thick, marked by moderately defined contacts with rare adhesion wart structures dipping at 5°-15° towards 075°-095°. Individual laminae are 1-5 mm thick, and dip at angles between 8°- 22° towards 070°-115°. These beds are interpreted to represent lee-side deposition, directly downwind of the slipface (McKee 1966, 1979; Halsey et al. 1990). Wedge-form sets of cross-laminated sands, composed of clay pellet aggregates and gypsarenite, are also present, and are more prevalent in the lowermost 1 m of the dune sequence. Some sets of tabular laminae grade laterally into wedge-form sets away from the former dune crest line, in the inferred direction of transport and dune migration. The wedge-form sets have thicknesses of 1-6 cm, thinning towards the west (upwind). Within the sets, individual laminae have thicknesses of 2-5 mm, dip at 5°-20°, and indicate transport directions varying between 060° and 190°. These wedge-form laminae were produced by preferential downslope deposition by ripples migrating across the slipface, driven by winds moving parallel or oblique to the slope (Fryberger and Schenk 1981, Halsey et al. 1990). Also present in the uppermost 1.5 m of the parabolic dune succession are isolated tabular gypsarenite/sand-sized clay pellet and fine selenite beds, coarser than the underlying deposits (Plate 6). These beds are internally structureless, and range in

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thickness from 1-3 cm. They are bounded by sharp contacts, frequently rippled, with occasional adhesion wart and rain pit structures. Dips range from horizontal to 2° easterly. Ripples range in height from 5-11 mm, and in wavelength from 28-42 cm, with ripple indices between 28 and 60. In all instances, the trough width of the ripples was in excess of the crest width. Transport directions vary between 060° and 140°. The tabular beds were deposited by a combination of saltation and grainfall, where sand-sized clay pellets, gypsarenite, and fine selenite adhered to temporarily saturated surfaces following precipitation events. Ripples with trough width in excess of crestal width form over resistant surfaces, under conditions of low to moderate sediment flux (Sharp 1963). Moistened gypsarenite/selenite surfaces would provide suitable conditions for ripple development. The uppermost 20 cm of the parabolic dune sequence contains numerous dikaka structures, commonly flanked by plano-convex lenses of moderately sorted sand-sized particles. The prevalence of these features in the uppermost part of the dune indicates that vascular plants became established on the dune surface, leading to interference with saltation and eventual stabilisation (Goldsmith 1973, Halsey et al. 1990). The uppermost 10 cm of this stratum has been pedogenically modified. The sequence of deposition in the parabolic dune at E 28 indicates that the dune became progressively more stabilised over time. Migration was spatially limited, and the dune responded to continuing sedimentation by increasing in height. Similar sequences are preserved in exposures in parabolic dunes throughout the Estancia dune field. 4.3

YOUNGEST AEOLIAN SEDIMENTATION

The most recent aeolian event is represented by the development of trough blowout structures in the parabolic dunes, indicating that these features are not in equilibrium with the modern climate and sediment flux regime. In addition, small dome dunes and loess deposits have formed overlying the older dunes. The cores of the youngest dome dunes are dominated by structureless beds of gypsarenite and sand-sized clay pellet aggregates, with thicknesses of 5-20 cm. Gently dipping (2°-6°) tabular sets of gypsarenite laminae, with set thicknesses of 2-5 cm and laminae thicknesses of 1-3 mm, are preserved on the eastern flanks of some dome dunes. Pedogenetic disturbance and dikaka structures are common throughout the sedimentary successions. Rippled beds are not preserved. Desiccation cracks, forming polygonal networks and extending to 6-8 cm depth are common on the exposed surfaces of dome dunes Caps of loess, with maximum thicknesses of 70 cm, are present over most of the parabolic dunes. The maximum thicknesses are preserved in saddles between the dunes and as cliff-top successions. Differential thicknesses in successions developed at dune crests, analogous to cliff-top successions in other environments (c.f. Catto 1983) indicate deposition by westerly winds. The fine gypsarenite and gypsite loess is

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structureless, except where disturbed by dikaka and pedogenesis. Episodic accumulation of loess punctuated by stabilisation events is indicated by pendant selenite crystals, diagenetic gypsite crusts, and discontinuous thin horizontal layers of fine organic detritus. Remobilised loess (brickearth) is also present in lowlying areas between the parabolic dunes, as infills in trough blowouts, and along the margins of the playas. The brickearth is structureless and poorly sorted, containing selenite granules, gypsarenite and gypsite clasts, and fine organic detritus. Discontinuous lag deposits and monolayers of rounded selenite clasts are common, and fine interbedded strata of primary loess are also present. Local thicknesses exceed 2 m. 5.

Genesis and stabilization of dome dunes

The geomorphology and sedimentary structures of the dome dunes of the Estancia Valley indicate formation by westerly winds. Gypsite and gypsarenite that had accumulated in the nearshore areas of Lake Willard B, particularly in the basin occupied by the modern Laguna del Perro, were subject to deflation when the water level dropped ca. 10,500 BP (see Bachhuber and Catto, this volume). Repeated deflation events, interspersed with periods of high playa levels accompanied by high water tables along the littorals, have resulted in the construction of dome dunes and lunettes along the margins of gypsiferous playas and salinas in other arid and semi-arid regions (e.g. Chen and Barton 1991, Chen et al. 1993, Carignano 1996). The source of sediments for dome dune construction was primarily confined to the Lake Willard sediments exposed in the basinal and littoral areas of Laguna del Perro, although minor proportions of quartz and other non-evaporitic minerals indicate that distally transported clasts are present. Dome dunes are developed in three geomorphic settings: in arid interior areas under strong winds (McKee 1966, 1979); along marine and lacustrine coasts (Bigarella 1972); and in boreal interior regions (Halsey and Catto 1994). In all these settings, genesis of dome dunes depends upon the inhibition of slip face development (Halsey and Catto 1994). In areas where wind strength has increased while sediment supply has remained constant or increased, such as White Sands NM, dome dunes are produced by the bevelling of preexisting large parabolic or transverse dunes (McKee 1966, 1979). The internal structures in the cores of these dome dunes thus represent the original sedimentary processes of parabolic and transverse dune construction. High- to moderate-angle cross-strata aligned parallel to the modal transport direction grade upwards into horizontal laminations and trough blowout structures infilled with ungraded sediment. Development requires a high rate of sedimentation. In dune fields with several dune types, these dome dunes thus are commonly found in the most proximal locations. At White Sands, dome dunes develop directly adjacent to the gypsarenite source of Lake Lucero (McKee 1966). In contrast, dome dunes developed along marine and lacustrine coastal margins are characterised by horizontally laminated beds and low-angle cross-strata, concavedownward bounding surfaces, and thin accumulations of structureless fine sand and

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loess in saucer-shaped blowout depressions. High-angle cross-strata are not present, and trough blowout features are rare. These dunes developed under winds of varying velocity and orientation, with sediment availability limited by moisture in the nearshore environment (e.g. Bigarella 1972, McKee and Bigarella 1972, Catto 1994). The dunes may be distal or proximal to the source of sediment. A third style of dome dune forms in boreal interior regions, where restricted sand supply, seasonally influenced by snowfall and hail accumulation and surface freezing, in combination with gusting but persistently oriented winds precludes slipface development (David 1977, Halsey and Catto 1994). The cores of interior boreal dome dunes are dominated by low-angle tabular and wedge-shaped cross-strata, tabular ungraded sand beds, and corrugated fine laminae, with high-angle cross-strata confined to the basal successions of relatively few examples. Although trough blowout structures are rarely observed, saucer blowouts are generally present. A low sedimentation rate is the primary factor promoting dome dune development in boreal interior regions. In dune fields with several morphological types, the dome dunes are the most distal members of the assemblage (Halsey and Catto 1994). The dome dunes along the margin of Laguna del Perro are characterised by low-angle tabular cross-laminations and tabular ungraded fine-medium gypsarenite/clay pellet beds. Thin loess and ungraded fine sandy deposits are confined to the flanks of the dunes, and overlie erosional saucer blowout surfaces. The internal structures of these dunes thus resemble those formed in boreal interior and modern marine coastal environments, under conditions of limited or episodic sand supply. The sedimentary structures indicate that limited availability of sediment was the most important factor resulting in dome dune formation in the Laguna del Perro area, in contrast to the high sedimentation rates associated with the dunes developed in the drier White Sands area to the south (McKee 1979). Development was initiated and sustained by prevailing westerly winds, consistent in direction but marked by gustiness. The dunes developed directly adjacent to the playa source. Accumulation was episodic, punctuated by periods of stability indicated by corrugation of bed contacts and weakly developed pedogenetic modification of the gypsarenite. The presence of the charcoal horizon 14C dated at 8,585 ( 75 BP (AA-6325), mid-way in the dune stratigraphic sequence, also suggests that episodes of stability suitable for juniper growth, and consequent termination of aeolian sand influx, occurred during dome dune construction. Stabilisation of gypsarenite dunes involves development of menisci of mobilised and reprecipitated gypsum cement, reinforced by pendants of cement (Schenk and Fryberger 1988). In modern, essentially unvegetated environments, the rate of cementation (and hence stabilisation) depends largely on capillary movement of meteoric water. Periodic precipitation results in alternating dissolution and recementation, producing multiple diagenetic overgrowths on the near-surface gypsarenite and binding the upper surface (Warren 1982, Schenk and Fryberger 1988).

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In modern gypsarenite dunes, stabilisation proceeds most rapidly in the capillary zone and is slowest in the vadose zone. Evidence for stabilisation and cementation features throughout the cores of the dome dunes indicates that the sediments were periodically exposed to capillary zone conditions. On the developing gently sloping flanks and crests of the domes, conditions for diagenesis and stabilisation more resembled those of modern interdune and salina margin environments (e.g. Schenk and Fryberger 1988, Carignano 1996) than those found on the surfaces of more elevated parabolic and transverse dunes. Where vascular vegetation is not present, small amounts of fine gypsarenite and coarse gypsite may be trapped by mosses and non-vascular plants (Danin and Ganor 1991). Direct frictional adhesion also occurs on dry surfaces, but rates of retention are characteristically less than 20 % of sediment influx, involving accumulation of less than 50 g/m2/a (e.g. Goosens 1995). The 1.5 m thickness of gypsarenite accumulated between ca. 10,500 BP and ca. 8,500 BP demonstrates that retention rates were significantly higher, in excess of 1000 g/m2/a. Enhanced retention of gypsarenite would be facilitated by accumulation on moist or periodically wetted surfaces, as indicated by the presence of corrugated bounding surfaces in the cores of the dome dunes. In the absence of thermoluminescence or optical stimulation luminescence analysis, techniques which cannot at present be applied to gypsiferous deposits (e.g. Berger 1995), chronological assessment of dune successions involves determination of the periods of stabilisation, rather than the times of active dune construction. In the Estancia Valley, dome dune formation was initiated at some time following ca. 10,500 BP, when Lake Willard B disappeared. The charcoal layer at the stratigraphic mid-point of the dune indicates that dune accumulation temporarily halted ca. 8500 BP. Further episodic dune development began at some time after 8,500 BP, and continued until some time before pedogenic activity formed the multiple soils capping and off-lapping the dome dunes. Charcoal recovered from the second weakly developed mollisol horizon was 14C dated at 4,660 (170 BP (GX-13321). The initial Holocene episodes of gypsarenite deflation and dune construction thus are approximately correlative to the multiple, relatively poorly dated early to mid-Holocene dune reactivation events noted on the Colorado High Plains by Forman et al. (1992) and Madole (1995), and in the adjacent Nebraska Sand Hills by Loope et al. (1995). In Arizona, conditions in the Tusayan dunefield allowed remobilisation of surface sediment ca. 4,700 BP, while plinth deposits remained undisturbed (Stokes and Breed 1993). Enhanced aeolian transport by westerly winds and distal deposition is recorded in Early and Mid-Holocene lacustrine sediments throughout the High Plains and Midwest, particularly between ca. 8,000 and ca. 4,000 BP (Dean 1997). The delayed onset of desiccation and aeolian deflation in the Great Basin following the disappearance of the Laurentide glacier to the north has been related to poleward displacement of summer monsoonal winds, resulting in a period of enhanced precipitation between ca. 12,500 and ca. 8,000 BP (Spaulding 1991).

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Genesis of parabolic dunes

The parabolic dunes are composed of gypsarenite and sand-sized clay pellet aggregates derived from Laguna del Perro and the other playa-floored deflation basins during periods when the lacustrine sediment was subaerially exposed. The morphology of the dunes indicates that the modal transport direction was from west to east, although crossstratification developed on lee slopes around the convex dune margins were deposited by local sediment movements ranging from north to south-southwest. These variations in sediment transport directions are the products of dune surface morphology, which influences the directions of grain flow, and of localised wind patterns on the lee slopes of individual dunes. The orientations of individual cross-strata thus reflect the local transport direction, not the modal wind direction, as has been observed in numerous modern dune successions (e.g. McKee 1966, Howard 1978, Whitney 1978, Sweet 1992). In the lee of large parabolic dunes, winds responsible for sediment deposition on the lower third of the slopes are commonly oriented oblique to the axis of the dune. Parabolic dune development involves a combination of limited sand supply, efficient sediment retention and surface adhesion, and consistently oriented winds (David 1977, Halsey et al. 1990). At White Sands, parabolic dunes represent the most distal members of the dune assemblage, and limited sediment supply is the most critical factor in their development (McKee 1966, 1979). In contrast, parabolic dunes in boreal and marine coastal areas frequently occupy the most proximal positions, and sediment retention is critical in their establishment and maintenance (Halsey et al. 1990, Catto 1994). Wind direction is consistent in most interior boreal environments, but the effects of variable winds in coastal environments are frequently negated by the limitations in sand supply from exposed littoral sediments and the geomorphically limited areas inland that are available for dune expansion. Parabolic dunes developed proximal to the marginal area east of Laguna del Perro under the influence of winds varying in intensity, but the internal structures of the dunes indicate consistent westerly wind activity. Sand supply and / or rates of retention were thus critical for dune development, coupled with the spatially limited area to the east of Laguna del Perro that was available for eastward expansion and migration. Sediment supply was limited by the presence of ephemeral standing water in the playa and by maximum deflation reaching the local water table. Much of the potential gypsarenite source material not inundated directly was periodically within the capillary zone, further limiting sediment availability. Deflation of the marginal areas of Laguna del Perro was also limited by periodic rainfall events, allowing diagenetic overgrowths to stabilise the gypsum surface. Conditions for aeolian deflation were similar to those currently prevailing in interdune areas adjacent to Laguna del Perro. Retention of gypsarenite is enhanced on dune surfaces periodically moistened by precipitation. The crenulated surfaces promote adhesion of gypsarenite when moist, and act to trap traction sediment during dry intervals. Chemical weathering and diagenesis further enhances sediment retention (Tchakerian 1991).

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Spatial limitation of dune migration also contributes to sediment retention, forcing dunes to increase in height and allowing stacking of successive generations of landforms. In the Gran Desierto of Mexico, Lancaster (1992) observed that stacking of dune generations occurs where sand supply is high (unlike the Estancia Valley), or where expansion of the sand sea is topographically restricted. Stacking also is a consequence of the preexisting dune and playa topography. Parabolic dunes originating along the eastern margin of Laguna del Perro would be forced to climb from the playa floor and override the preexisting dome dunes, surmounting slopes of 10° - 15° and heights in excess of 5 m. Sediment movement by traction and saltation up these relatively steep slopes would be hindered, slowing transportation rates, enhancing retention, and resulting in rapid accumulation. Similar climbing dunes developed on up-wind facing slopes have been noted by Seppälä (1993) and Cros and Serra (1993). Climbing results in the dominance of traction and saltation processes in the development of parabolic stoss slopes and crests, and largely precludes grainfall. As the dune grows vertically and sediment is deposited and winnowed during traction, progressive deflation results in coarsening upward sequences and coarsening upwind from the crest along the stoss slope (e.g. Vincent 1996). Under these circumstances, deposition is influenced by sporadic wind gusts, and local deceleration events result in accumulations of sediment (Arens et al. 1995). The accumulation of sediment in the stoss and crestal areas at a greater rate than the spatially and gravitationally restricted climbing dune can migrate results in a growth in height and a tendency for lateral coalescence of adjacent parabolic dunes normal to the modal wind direction. In the Estancia Valley, the highest and most laterally extensive parabolic dunes, locally coalesced into transverse forms, are located along the eastern margin of Laguna del Perro. Depositional episodes were transitory in the dune field, and were separated by periods of stability. Similar episodic dune development, not associated with marked climate change, has been recorded in several areas (e.g. Lancaster 1992, Bullard et al. 1997). In addition, relatively minor climate changes may also result in a substantial change of stability regime (e.g. Stokes and Breed 1993, Madole 1995, Liu Jian et al. 1997), suggesting that the climatic fluctuations recorded in other palaeoenvironmental records throughout New Mexico would be sufficient to permit alternating dune reactivation and stabilisation. The parabolic dunes in the Estancia Valley were reactivated shortly after ca. 4660 BP. This reactivation is co-incident with episodes of dune construction and migration noted in the Mojave Desert (Tchakerian 1991), in northern Arizona (Stokes and Breed 1993), and in Colorado (Forman et al. 1992, Madole 1995). Stabilisation may be related to a gradual decrease in the evapotranspiration ratio, rather than an absolute increase in precipitation (c.f. Carignano 1996). Lateral expansion of the parabolic dunes impedes any surface drainage between playas and salinas, resulting in temporary impoundment and deposition of playa sediments. In the Estancia dune field, the tendency for vertical accumulation and restricted lateral

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migration resulted in the development of small playas in relatively fixed locations, rather than a series of ephemeral overridden interdune successions. Interdune areas are sites of deposition rather than periodic deflation and erosion resulting from dune migration. Thus, in contrast to the situation in actively migrating dune fields (e.g. Lancaster and Teller 1988), interdune and plinth deposits are commonly preserved. The marginal areas of the playa and plinth successions are marked by interbedded gypsarenite slope failure and surface wash deposits, along with poorly sorted structureless gypsite brickearth. The parabolic dunes of the Estancia dune field developed in combination with playa and ephemeral channel sedimentation. Impoundment of playas by parabolic dunes may result in increased accumulation of gypsite and gypsarenite in basins during periods favourable to dune development. Thus, subaqueous precipitation and aeolian sedimentation may be penecontemporaneous, and high playa levels may not necessarily represent periods of enhanced precipitation and dune stabilisation. Similar dune/interdune relationships were documented in the driest areas of the Nebraska Sand Hills by Loope et al. (1995).

7.

Most recent reactivation event

Modern aeolian activity involves the deposition of loess, and the reworking of loess by sheetwash and by playa activity along dune flanks to produce brickearth. Small dome dunes are ephemerally active throughout the eastern part of the dune field, and erosional trough blowouts are developing on the upper stoss and crestal areas of the larger parabolic dunes. Under modern climate conditions, aeolian sedimentation in the Estancia Valley is marked by very limited sand supply, almost exclusively locally derived, and high rates of retention. The irregular cover of shrubs limits surface traction and saltation of sand (Lee 1991), and hence precludes both rippling and slipface development. The discontinuous vegetation cover provides localised zones for trapping of fine suspended gypsite, allowing loess accumulation. Loess thus accumulates in sites preferential for vegetation development, such as inactive trough blowouts and hollows between adjacent dunes. Prolonged accumulation has resulted in smoothing of the topography, with depressions between adjacent dune hills partially infilled with accumulations of loess and remobilised brickearth (c.f. Goosens 1997).

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Much of the sand supply to the modern dome dunes is derived from erosion of trough blowouts on the larger parabolic dunes to the west. In the blowouts, erosion alternates with collapse of oversteepened slopes and moisture-induced colluviation, and the infill of the blowouts is dominated by brickearth rather than by primary aeolian sedimentation. The resulting brickearth resembles reworked loess deposits associated with ephemeral streams (Sneh 1983), and may be confused with reworked deposits resulting from fluvial activity (Jones and Blakey 1997) or climate change (e.g. Liu Jian et al. 1997). Local reactivation is attributed to a variety of causes. Local overgrazing and cattle tracks are responsible for the initiation of specific saucer and trough blowouts, respectively, but ranching is not intensive in the Estancia Valley. The timing of reactivation in Estancia also does not coincide with the late nineteenth century events noted in overgrazed areas of Colorado and Arizona. The distribution and configuration of the blowouts suggests that most were created by localised deflation from strong westerly winds, associated with the most exposed, highest, and driest dune surfaces (rather than with those preferentially grazed or lacking vegetation). The small modern dome dunes developed under conditions of extremely limited sediment supply, in contrast to the larger first and second generation features. Under the moister climate

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which currently prevails, deflation would be confined to the driest areas at the dune summits, and sediment flux from the periodically moistened areas surrounding Laguna del Perro would be severely limited. Under the current climatic regime, aeolian activity is limited to sporadic deflation, eastward migration of small quantities of gypsarenite, and construction of minor dome dunes.

8.

Summary

The sequence of latest Wisconsinan and Holocene events related to aeolian activity preserved in the Estancia Basin involved: (1) Desiccation of the latest Wisconsinan pluvial lake, Lake Willard B, ca. 10,500 BP; (2) development of the Willard Soil, a gypsite capping the youngest pluvial lake sediments; (3) initial development of the dome dunes, prior to ca. 8,500 BP; (4) a period of stability ca. 8,500 BP, indicated by the charcoal horizon; (5) renewed development of dome dunes; (6) intermittent episodes of stability, with formation of stacked weakly developed mollisols off lapping and capping the dome dunes, ca. 4,660 BP; (7) initiation of deflation, forming the parabolic dunes and the playa floored deflation basin sequence, beginning after ca. 4,660 BP and ceasing prior to the latest Holocene; and (8) ongoing development of the youngest dome dunes and loess sequences, along with present stabilization of most aeolian features.

Acknowledgements Production assistance was provided by Charles Conway, MUNCL, and by Photographic Services, Memorial University. Technical and editorial assistance was provided by Gail Catto. References Arens, S. M., Van Kaam-Peters, H. M. E., and Van Boxel, J. H. (1995). 'Air flow over foredunes and implications for sand transport.' Earth Surface Processes and Landforms 20, 315-322. Bachhuber, F. W. (1971). Paleolimnology of Lake Estancia and the Quaternary history of the Estancia valley, central New Mexico . Unpublished Ph.D. thesis, University of New Mexico, Albuquerque, NM, USA, 238. Berger, G. W. (1995). 'Progress in Luminescence Dating Methods for Quaternary Sediments', in Rutter, N.W., and Catto, N. R. (eds.), Dating Methods for Quaternary Deposits. Geological Association of Canada, 81-104. Bigarella, J. J. (1972). 'Eolian environments - Their characteristics, recognition, and importance', in Rigby, J. K., and Hamblin, W.K. (eds.), Recognition of Ancient Sedimentary Environments. Society of Economic Palaeontologists and Mineralogists, Special Publication 16, 12-62. Bullard, J. E., Thomas, D. S. G., Livingstone, I., and Wiggs, G. F. S. (1997). 'Dunefield activity and interactions with climatic variability in the southwest Kalahari Desert.' Earth Surface Processes and Landforms , 22, 165-174.

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Carignano, C. A. (1996). 'Evolucion geomorfologica de las planicies en la Provincia de Cordoba durante el Pleistocene superior.' Revista del Instituto de Geologica y Mineria, 11, 7-26. Catto, N. R. (1983). 'Loess in the Cypress Hills, Alberta, Canada.' Canadian Journal of Earth Sciences, 20, 1159-1167. Catto, N. R. (1994). 'Anthropogenic Pressures and the Dunal Coasts of Newfoundland.' Coastal Zone 1994 Conference, Halifax, Nova Scotia, September 1994, Bedford Institute of Oceanography, 2266-2286. Chen, X. Y., and Barton, C. E. (1991). 'Onset of aridity and dune-building in central Australia: sedimentological and magnetostratigraphic evidence from Lake Amadeus.' Palaeogeography, palaeoclimatology, palaeoecology, 84, 55-73. Chen, X. Y., Bowler, J. M., and Magee, J. W. (1993). 'Late Cenozoic stratigraphy and hydrologic history of Lake Amadeus, a central Australian playa.' Australian Journal of Earth Sciences, 40, 1-14. Cros, L., and Serra, J. (1993). 'A complex dune system in Baix Empord... (Catalonia, Spain)', in Pye, K. (ed.) The Dynamics and Environmental Context of Aeolian Sedimentary Systems. Geological Society of London Special Publication 72, 191-199. Danin, A., and Ganor, E. (1991). 'Trapping of airborne dust by mosses in the Negev Desert, Israel.' Earth Surface Processes and Landforms, 16, 153-162. David, P.P. (1977). Sand Dune occurrences in Canada: A theme and resource inventory study of eolian landforms in Canada. National Parks Branch, Department of Indian and Northern Affairs, Government of Canada, Contract 74-230, Report 183. Ottawa, Ontario, 183. Dean, W. E. (1997). 'Rates, timing, and cyclicity of Holocene eolian activity in north-central United States: Evidence from varved lake sediments.' Geology , 25, 331-334. Forman, S. L., Goetz, A., and Yuhas, R. H. (1992). 'Large-scale stabilized dunes on the High Plains of Colorado: Understanding the landscape response to Holocene climates with the aid of images from space.' Geology , 20, 145-148. Fryberger, S., and Schenk, C. (1981). 'Wind tunnel sedimentation experiments on the origins of aeolian strata.' Sedimentology, 28, 805-821. Goldsmith, V. (1973). 'Internal geometry and origin of vegetated coastal dunes.' Journal of Sedimentary Petrology, 43, 1128-1143. Goosens, D. (1995). 'Field experiments of aeolian dust accumulation on rock fragment substrata.' Sedimentology, 42, 391-402. Goosens, D. (1997). 'Long-term aeolian loess accumulation modelled in the wind tunnel: the Molenberg case (central loess belt, Belgium).' Zeitscrift für Geomorphologie NF 41, 115-129. Halsey, L.A., Catto, N. R., and Rutter, N.W. (1990). 'Sedimentology and development of parabolic dunes, Grande Prairie dune field, Alberta.' Canadian Journal of Earth Sciences, 27, 1762-1772. Halsey, L.A., and Catto, N. R. (1994). 'Geomorphology of Dome Dunes.' Géographie physique et Quaternaire, 48, 97-105. Harbour, J. (1958). Microstratigraphy and sedimentatinal studies of an early man site near Lucy, New Mexico. Unpublished M.S. thesis, University of New Mexico, Albuquerque, NM, USA, 111. Hills, E. S. (1940). 'The lunette, a new land form of aeolian origin.' Australian Geographer, 3, #7, 15-21. Howard, A. D. (1978. 'Sand transport model of Barchan dune equilibrium.' Sedimentology, 25, 307-338. Jian, L., Xihao, W., Shaoquan, L., and Mingshu, Z. (1997). 'The Last glacial stratigraphic sequence, depositional environment and climatic fluctuations from the Aeolian sand dune in Hongguang, Pengze, Jiangxi (China).' Quaternary Science Reviews 16, 535-546. Jones, L. S., and Blakey, R. C. (1997). 'Eolian-fluvial interaction in the Page Sandstone (Middle Jurassic) in south-central Utah, USA -- a case study of erg-margin processes.' Sedimentary Geology , 109, 181-198. Kocurek, G., and Fielder, G. (1982). 'Adhesion structures.' Journal of Sedimentary Petrology, 54, 1229-1241. Lancaster, N. (1992). 'Relations between dune generations in the Gran Desierto of Mexico.' Sedimentology 39, 631-644. Lancaster, N., and Teller, J. T. (1988). 'Interdune deposits of the Namib Sand Sea.' Sedimentary Geology , 55, 91-107. Lee, J. A. (1991). 'The Role of Desert Shrub Size and Spacing on Wind Profile Parameters.' Physical Geography, 12, 72-89. Loope, D. B., Swinehart, J. B., and Mason, J.P., (1995). 'Dune-dammed paleovalleys of the Nebraska Sand Hills: Intrinsic versus climatic controls on the accumulation of lake and marsh sediments.' Geological Society of America Bulletin , 107, 396-406. Madole, R. F. (1995). 'Spatial and temporal patterns of Late Quaternary eolian deposition, eastern Colorado, USA.' Quaternary Science Reviews , 14, 155-177.

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McKee, E. D. (1966). 'Structures of dunes at White Sands National Monument, New Mexico (and a comparison with structures of dunes from other selected areas).' Sedimentology, 7, 1-69. McKee, E. D. (1979). 'Sedimentary Structures in Dunes, in E. D. McKee, ed., A Study of Global Sand Seas'. United States Geological Survey Professional Paper 1052, Washington, 83-136. McKee, E. D., and Bigarella, J. J. (1972). 'Deformational Structures in Brazilian Coastal Dunes.' Journal of Sedimentary Petrology, 42, 670-681. Price, W. A., and Kornicker, S. (1961). 'Marine and lagoonal deposits in clay dunes. Gulf Coast, Texas.' Journal of Sedimentary Petrology, 31, 245-255. Schenk, C. J., and Fryberger, S. G. (1988). 'Early diagenesis of Eolian dune and Interdune sands at White Sands New Mexico.' Sedimentary Geology, 55, 109-120. Seppälä, M. (1993). 'Climbing and falling sand dunes in Finnish Lapland', in Pye, K. (ed.), The Dynamics and Environmental Context of Aeolian Sedimentary Systems. Geological Society of London Special Publication 72, 269-274. Sharp, R. P. (1963). 'Wind ripples.' Journal of Geology, 71, 617-636. Spaulding, W. G. (1991). 'Pluvial climatic episodes in North America and North Africa: types and correlation with global climate.' Palaeogeography, Palaeoclimatology, Palaeoecology, 84, 217-227. Stokes, S., nd Breed, C. S. (1993). 'A chronostratigraphic re-evaluation of the Tusayan Dunes, Moenkopi Plateau and southern Ward Terrace, Northeastern Arizona', in Pye, K. (ed.), The Dynamics and Environmental Context of Aeolian Sedimentary Systems. Geological Society of London Special Publication 72, 75-90. Sneh, A. (1983). 'Redeposited Loess from the Quaternary Besor Basin, Israel.' Israeli Journal of Earth Sciences, 32, 63-69. Sweet, M. L. (1992). 'Lee-face airflow, surface processes, and stratification types: their significance for refining the use of eolian cross-strata as paleocurrent indicators.' Geological Society of America Bulletin, 104, 1528-1538. Tanner, W. F. (1967). 'Ripple mark indices and their uses.' Sedimentology, 9, 89-94. Tchakerian, V.P. (1991). 'Late Quaternary Aeolian Geomorphology of the Dale Lake Sand Sheet, southern Mojave Desert, California.' Physical Geography, 12, 347-369. Teller, J. T. (1972). 'Aeolian deposits of clay sand ' Journal of Sedimentary Petrology, 42, 684-686. Titus, F. B. (1969). Late Tertiary and Quaternary hydrogeology of the Estancia basin, central New Mexico . Unpublished Ph.D. thesis, University of New Mexico, Albuquerque, NM, USA, 179. Vincent, P. (1996). 'Variation in particle size distribution on the beach and windward side of a large coastal dune, southwest France.' Sedimentary Geology , 103, 273-280. Warren, J. K. (1982). 'The hydrological setting, occurrence and significance of gypsum in late Quaternary salt lakes in South Australia.' Sedimentology, 29, 609-637. Whitney, M. (1978). The Role of Vorticity in Developing Lineation by Wind Erosion. Geological Society of America Bulletin, 89, 1-18.

EVAPORITE MINERALS AND ORGANIC HORIZONS IN SEDIMENTARY SEQUENCES IN THE LIBYAN FEZZAN: IMPLICATIONS FOR PALAEOENVIRONMENTAL RECONSTRUCTION

KEVIN WHITE Landscape and Landform Research Group, Department of Geography, The University of Reading, Whiteknights, Reading, RG6 6AB, U.K. SUE McLAREN Department of Geography, University of Leicester, University Road, Leicester, LE1 7RH, U.K. STUART BLACK Postgraduate Research Institute for Sedimentology, The University of Reading, Whiteknights, Reading, RG6 6AB, U.K. ADRIAN PARKER Geography Department, Oxford Brookes University, Gipsy Lane Campus, Headington, Oxford, OX3 0BP, U.K.

Abstract Compared to other parts of the Sahara, there is a paucity of reliable palaeoenvironmental information for the Libyan Fezzan. Ongoing archaeological work in the area provides an opportunity to link palaeoenvironmental and archaeological data to improve our understanding of human adaptations to dryland environmental change. A combination of stable isotope analysis, mineral magnetic analysis, an investigation of the particle size distributions and uranium-thorium dating, together with supporting image processing of remotely sensed data, indicates a trend of falling groundwater level since the late Pleistocene, manifested by former lakes, swamps and springlines. Archaeological evidence shows how human activity has adapted to these environmental changes. 193

S.J. McLaren and D.R. Kniveton (eds.), Linking Climate Change to Land Surface Change, 193–208. © 2000 Kluwer Academic Publishers. Printed in the Netherlands.

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KEY WORDS:

Sahara, environmental change, remote sensing, stable isotope analysis

1.

Introduction

There has been considerable research into Late Quaternary palaeoenvironments of the Sahara, especially connected with changing lake and groundwater levels (e.g. Brookes, 1993; Fontes and Gasse, 1991, Petit-Maire et al., 1994). However, there are still significant gaps in the spatial coverage of these investigations, notably in the Libyan Fezzan, despite the wealth of geomorphological and archaeological evidence of environmental change thereabouts (Gaven et al., 1981; Cremaschi and Trombino, 1998). The study area is centred on the town of Germa (ancient Garama) in the Wadi el Agial (Figure 1). The region lies in the hyperarid central core of Libya, receiving an average annual rainfall of less than 15 mm p.a. (Dubief, 1963), and relies on groundwater pumped from the Continental Intercalaire (Nubian Sandstone) aquifer. The extensive groundwater reserves in central and southern Libya have been important throughout historical times and are of great strategic significance today, as they are being increasingly exploited by the Great Man-Made River Project for use in the coastal zone (McKenzie and El Saleh, 1994; Pim and Binsariti, 1994). These aquifers have been recharged at various stages throughout the Quaternary (Edmunds and Wright, 1979). Geomorphological evidence of lake highstands in several parts of North Africa have been dated and these periods have been linked to Quaternary aquifer recharge episodes (Gasse et al., 1987). However, the regional pattern of Quaternary environmental change is still unclear, due to the paucity of information over very large areas, including the Libyan Fezzan. To clarify the palaeoclimate history of North Africa, evidence for past changes in groundwater hydrology in the Fezzan needs to be collected and interpreted. Preliminary results from an ongoing project on palaeoenvironmental reconstruction of the Wadi el Agial study site are presented here, along with some tentative conclusions regarding the relationship between palaeoenvironment and human adaptations.

2.

Methodology

Due to the absence of palaeoenvironmental data for the Wadi el Agial, the first task has involved a detailed reconnaissance to identify and map geomorphological evidence of environmental change. To assist in this, a Landsat Thematic Mapper image (Path 187, row 42 - Worldwide Reference System, acquired 12th October 1987) has been used to aid field survey. The main purpose of the remote sensing work is to map gypsum. There are high levels of Ca and S in contemporary groundwater, although salinity is variable

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(Table 1). If we assume that groundwater geochemistry has not changed significantly, groundwater at or near the surface in the past would have precipitated gypsum, which can be remotely sensed (Drake, 1995).

Image processing involved applying a spectral mixture model to the multispectral image (Settle and Drake, 1993). This approach has been demonstrated to be very effective at mapping surficial gypsum deposits (White and Drake 1993, Eckardt and White 1997). The technique yields a proportion estimate of gypsum content within each pixel of the image. The image products were georeferenced with a very low R.M.S. error (mean R.M.S. = 0.66 pixel) to a set of 20 ground control points collected during the initial field season (in March 1998) using a 12 channel hand-held GPS receiver. The ground control points were well distributed over the whole study area. The accuracy of the

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georeferencing was checked in the field the following year (January 1999) and was found to be excellent, with the location of all the points visited being well within the nominal accuracy of the GPS instrument (c. 80m). No attempt was made to calibrate the remote sensing output with field data, as the success of this technique of mapping gypsum has been demonstrated elsewhere (Bryant, 1996; White and Drake 1993), and here the aim is to map the distribution of gypsum deposits (which was checked in the field), not to quantify actual concentrations.

Other fieldwork focused on finding evidence of palaeolake deposits, denoted by fossiliferous dark organic horizons. These deposits are generally only exposed in section, not at the surface, so remote sensing was not employed here. Instead, detailed surveys were carried out along sections of the base of the escarpment (assumed to be the highest shoreline) and sections exposed by well excavation at numerous locations around the oasis. Two significant exposures of dark organic horizons were found at escarpment proximal locations, one above the town of Germa (26.5058°N 13.0784°E) and one above the town of el-Greifa (26.4965°N, 13.0030°E). Both sites exhibit thin dark organic horizons and contain a fossil assemblage dominated by Melanoides tuberculata (Figure 2). Other exposures of dark organic horizons were found in well sections in the oasis around the town of el-Greifa (Figure 1), though these were not noticed elsewhere in the study area. Wells with dark organic horizons are located at 26.5288°N, 12.9898°E; 26.5332°N, 12.9903°E; 26.5448°N, 12.9960°E; 26.5403°N, 13.0050°E; 26.5247°N, 12.9862°E. Samples of these dark organic horizons were collected from each of these sections for comparison with each other and with the deposits found along the escarpment, to determine whether the dark organic horizons exposed during well construction can be correlated with the fossiliferous palaeolake deposits found at the base of the escarpment.

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One of the major problems with identifying organic horizons in dryland situations where human activity has been significant in the past is the possibility of confusion with burned material and fire hearths. To avoid this possibility, we also sampled a known fire hearth located among archaeological remains at 26.5633°N, 12.9512°E. Stable isotope analyses of C and N were carried out on samples from all the dark organic horizons found in the field; both to correlate the samples and to see if specific isotope signatures could be identified and used to infer conditions of formation (Koch, 1998). In order to be sure that like was being compared with like in all the dark organic horizons, isotopic analyses were run on the organic carbon fraction only, inorganic carbonates were removed by treatment with 1.0 molar HCl (the percentage weight loss on HCl treatment was recorded). Analyses were performed using an automated carbon and nitrogen analyser and a continuous-flow isotope-ratio-monitoring mass-spectrometer (Europa ANCA Roboprep automated Dumas preparation system coupled to a Europa 20/20 mass spectrometer). Typical replicate measurement errors are of the order of 0.3 for

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generate magnetite and maghemite (Maher, 1986). The effect is normally so great that magnetic susceptibility alone can usually discriminate burned samples (Rummery et al., 1979). However, is a concentration-dependent magnetic parameter, and the burned signal can be lost in a sedimentary environment if it becomes diluted during diagenesis. To avoid this problem, we have used the ratio of the IRM at 20mT divided by the essentially a measure of the concentration of all ferrimagnetic minerals divided by a measure of the concentration of ferrimagnetic grains in the (stable single domain) size range (Oldfield, 1991). Full details of mineral magnetic techniques, parameters and their interpretation are given elsewhere (Walden et al., 1999). This study is only concerned with identifying any samples which may result from burning so that they can be excluded from further palaeoenvironmental interpretation. Particle size analysis of the dark organic sediments was undertaken to establish the textural characteristics of the sediments. After soaking the samples in sodium hexametaphosphate overnight to disaggregate the clasts, the sands and coarse silts were separated from the fine silts and clays by wet sieving. The sediments were then dried and the coarse fraction was dry sieved and separated into the following phi sizes: and and then each fraction was weighed. The finer fraction was calculated using the pipette sedimentation method and separated into and size fractions and their weights were determined. U-Th dating was applied to a Melanoides shell sample. The shell was cleaned in ultrapure (Milli-Q) water to remove as much detrital contamination from the exterior as possible. The sample was then crushed in an agate mortar to a fine powder and baked in an oven at 105°C for 24 hours to remove surface moisture. U-series radionuclides were measured by high resolution alpha spectrometry and ICP-MS. Uranium and thorium isotopes for alpha spectrometry were separated by ion-exchange resins and electrodeposited on stainless steel planchets (Black et al., 1997; Kuzucuoglu et al., 1998) after aliquots were taken for ICP-MS. yield monitor was used for alpha spectrometry with a decay and ingrowth correction applied for the daughter nuclide. Yields ranged from 88-100 % for uranium and 91-98 % for thorium. Blank and background determinations were carried out, averaging 20 counts in 10,000. EG&G® "Ultra" alpha detectors were used with very low background counts in the yield monitor energy range. ratios were determined on unspiked aliquots to ensure no recent alteration had taken place. and were also detected on a Varian Ultramass ICP-MS. One aliquot in 5 for ICP-MS was spiked with a yield monitor to assess U recoveries. In addition, internal standards were routinely run and checked against the NIST SRM3159 standard and agreed to better than 0.25 %. Detection limits at masses 232 and 238 were better than 8 ng after correcting for dead time, background and mass bias. Replication of multiple digestions was better than 0.6 % for both U and

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Th whereas replicate aliquot analyses (n = 36) was better than 0.3 % for U and 0.2 % for Th. Corrections were made for decay of excess and detrital on the assumption that these were present at formation of the shell. The correction for the detrital component was made from isochron plots after successive total dissolutions were performed following the preparation, firing and digestion outlined in Luo and Ku (1991) and Bischoff and Fitzpatrick (1991). In all cases the slopes of the isochrons are best determinations by a method of least squares fitting which takes account of the errors in both variables (York, 1969).

3.

3.1.

Results

EVAPORITE MINERALS

The gypsum proportions map produced by applying a spectral mixture model to the Thematic Mapper imagery (Figure 3) indicated a widespread distribution of gypsum throughout the study area. Field observations and XRD analysis of samples back in the laboratory indicate that this surficial gypsum is found in four different sedimentary environments. 1. As evaporitic crusts on small playas along the base of Wadi el Agial. 2. As crusts formed on fields, often adjacent to known foggara systems (Figure 1). Foggara, or qanats as they are called in Persia, are underground canals which tap higher water tables near mountain fronts and transport the water down to irrigated fields beyond the piedmont zone (Goblot, 1979). However, the formation of the field crusts may result from more recent irrigation from pumping rather than from ancient irrigation via foggara. 3. As outcrops along the base of the escarpment, indicating a palaeo-springline at the junction between the Murzuk Sandstone Formation and the underlying Germa Bed marls (Klitzsch and Baird, 1969). The fibrous habit of these gypsum crystals indicates that they have grown in-situ within the bedrock and surficial materials (Cody and Cody, 1988). In many places, the gypsum has been altered to anhydrite, presumably due to the high summer temperatures encountered in the surface environment. 4. As a constituent of the sediments deposited within relict fluvial channels on the south-facing dip-slope of the Murzuk plateau.

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These results provide strong evidence of the high Ca and S content of the groundwater which fed these fluvial networks when the water table was higher. The chemistry of the palaeodischarge would appear to have been similar to that of the contemporary groundwater (Table 1).

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201

DARK ORGANIC HORIZONS

The mineral magnetic data discriminate the known fire-hearth site at El Hatia (Figure 4), but one of the well-section dark organic horizons (known as foggara well, as a foggara channel was encountered during the sinking of the well shaft) has a similar high value. This may mean that the dark colour of these samples could be due to burning rather than enrichment with organic compounds, so they are excluded from further analyses. However, the lower value of the foggara well sample relative to the El Hatia sample (Table 2), though still significantly higher than the other dark horizons, indicates a lower concentration of ferrimagnetic minerals, possibly through transport and redeposition, or by dilution during diagenesis and weathering (Rummery et al., 1979).

The burned sediments (El Hatia 8 and Foggara well) also have a much higher HCl soluble content (30-37%) than the other dark organic horizons (Table 2). The same parameter can be used to discriminate between the el-Greifa palaeolake sediments (16-

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21%) and the oasis well section dark organic horizons (6-12%). Again this may result from primary sediment composition or from diagenesis and precipitation of evaporite minerals. The stable C isotope results (Table 2) show a strong depletion of relative to PDB standard (Peedee Formation, belemnite). All the samples cluster around the -23 threshold between C3 and C4 plant metabolic processes, suggesting either the organic matter in the dark organic horizons was derived from a mixed C3/C4 community (Bond et al., 1994; Kelly et al., 1998), or that some other fractionation process has operated during or after deposition of the sediment; the initial carbon isotope contents can be modified by various diagenetic effects such as precipitation/dissolution of secondary calcite and oxidation of reduced carbon (Fontes, 1994).

The stable N isotope results (Table 2) distinguish between the shoreline lake deposits identified along the base of the escarpment, and the dark organic horizons identified in the oasis well sections (Figure 4). This may be attributable to different palaeoenvironments (the values of the well section dark organic horizons are close to atmospheric N (Table 2), suggesting direct take-up by cyanobacteria, possibly in small swamps which have formed in the oasis), but they could equally well represent the difference between shoreline and deeper-water vegetation communities. Furthermore, post-depositional processes again have to be considered; ion exchange in soils can also result in fractionation effects affecting Unlike nitrite or nitrate ions, readily exchanges with clays or the so-called “humic-clay complex” of soils. Isotope fractionations between exchange resins and a liquid phase are known to exist, with an

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enrichment between 5 and 25 for the solid phase (Létolle, 1980). Two main factors can control the N isotopic characteristics of the organic matter. 1) Varying degrees of diagenesis. Decay produces isotopically light nitrogen compounds and the remaining organic matter becomes progressively enriched in 2) The migration and accumulation of decay products, such as humin, fulvin, humic and fulvic acids, in different horizons (Létolle, 1980). Therefore, the limited amount of stable isotope data available so far are unable to provide a basis for correlation of the different dark organic horizons. However, the limited extent of the dark organic horizons exposed in the oasis well sections, coupled with the lack of fossils and lower HCl soluble content, suggests that these may be very localised marsh or swamp deposits, unrelated to the fossiliferous palaeolake sediments found at the base of the escarpment. Table 3 shows the results of the particle size analyses that have been conducted. Samples from el-Greifa 1.2 and 1.3 were collected adjacent to one another close to the palaeolake shoreline. Their sediment distributions are similar with mean grain sizes of and they comprise mostly sand with low amounts of clay (Table 3). el-Greifa 3 is from a location slightly further away from the shoreline, in deeper water, which may account for the higher amounts of clay and silt in this sample. The dark organic horizons in the wells vary in terms of the mean grain sizes between 4.4 and with on average 45.9% sand, 37.6% silt and 16.6% clay. The two samples with the coarsest mean size fractions (El Hatia 8 and Foggara well) were highlighted by mineral magnetic analysis as containing burned materials. All of the samples are platykurtic, positively skewed and poorly or very poorly sorted. The Melanoides shell gave a U/Th age of 84.5 ± 4.6 ka BP. This falls within the widelypostulated 90-65ka BP Saharan lake highstand (Causse et al. 1989; Fontes and Gasse, 1989; Petit-Maire et al. 1980; Szabo et al., 1995), which can be correlated with oxygen isotope interglacial substage 5a (or 5.1). However, as only one age is so far available for the Wadi el Agial palaeolake deposits, this must be treated with caution. A large number of other Melaniodes shells are currently being dated, and a fuller analysis of these, and other radiometric ages, will be provided in a later publication.

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4.

Discussion

Despite the hyper-aridity of the present environment, the Fezzan has a long history of human occupation and a detailed archaeological record which dates back to the Palaeolithic (Mattingly et al., 1998) Extensive rock art, particularly in the Wadi Berjuj region, some 50 km to the south of Wadi el Agial, shows considerable evolution of technique and subject matter, indicating the long duration of human activity in this region. The earliest identified remains in the vicinity of Wadi el Agial are scatters of quartzite lithics and debitage along the edge of the Murzuk escarpment. These include hand axes, a cleaver, projectile points, scrapers and denticulates and a burin. The most probable date range for these materials is 105-40ka BP (Mattingly et al., 1998), and may indicate presence of a lake below the escarpment (where Palaeolithic material is very scarce) during some of this time. From this we can infer a higher regional water table, and greater activity of the relict fluvial systems along the south-facing dip slope of the Murzuk Plateau. Below the Murzuk plateau, Neolithic finds are concentrated around patches of duricrust surfaces around and within the Ubari sand sea. The duricrusts contain abundant mineralised plant roots and have been interpreted as palaeoswamp deposits. From this we can infer that, during the Holocene the area of surface water had diminished to an extent that large areas below the Murzuk Plateau were exposed and inhabited by humans. Numerous burials and settlements on the escarpment piedmont and adjacent oases are mostly attributed to the Garamantean occupation (c. 500 BC to 500 AD). During this period of occupation small lakes may have still been present in the low points of the Wadi El Agial; results from ongoing dating studies are awaited to confirm this; but extensive use of foggara technology appears to have been made at this time (Mattingly et al., 1998), in order to grow irrigated crops (van der Veen, 1992). In the vicinity of the

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village of Twesh (Figure 1), field surveys have indicated that motherwells of foggara channels (the collecting shafts dug at the top of each foggara to tap the groundwater) are associated with palaeo-springline gypsum deposits. From this we can infer that, by the time of foggara construction, these springs had ceased to be perennial features, but they may have still been zones of localised high groundwater levels. The foggaras were subsequently abandoned, possibly due to a continuing fall in the water table, although other factors, such as the introduction of new technologies or a collapse in population, may also be responsible. The time of foggara abandonment is unknown at present. The pattern of environmental change, and the close relationship with human activity, has many similarities with the record from the Western Desert, Egypt and Sudan, which has received more attention (see e.g. Peel, 1966; Wendorf et al., 1976; Wickens, 1975). Both areas demonstrate a close association between significant Palaeolithic and Neolithic populations and palaeolacustrine deposits (Wendorf and Schild, 1980), both exhibit complex evidence of fluctuating water table, with aeolian sands underlying lacustrine clays in the base of depressions (Haynes, 1982). In this respect, the Wadi el Agial seems typical of Late Quaternary environments from eastern North Africa.

5.

Conclusion

A preliminary geomorphological survey of the Libyan Fezzan has highlighted evidence of significant environmental changes in the region during the Late Quaternary. A lake highstand, tentatively dated at 85ka BP, deposited dark organic sediments and a fossil assemblage dominated in this region by Melanoides tuberculata. Mineral magnetic analysis enables discrimination of these dark organic sediments from those produced by burning. Stable isotope analysis of C suggest that either the organic matter in the dark organic horizons was derived from a mixed C3/C4 community or that some other fractionation process has operated during or after deposition of the sediment. Thin dark organic horizons exposed in well sections at the base of the Wadi el Agial differ in stable N composition from the fossiliferous palaeolake sediments exposed higher up on the escarpment slopes. This may indicate a difference in palaeoenvironments, although diagenetic modifications may also be responsible. HCl soluble content also enables discrimination of burned materials, el-Greifa palaeolake sediments and oasis well section dark organic horizons. Palaeo-springlines and minor upwelling zones in small playa basins have deposited evaporitic gypsum, which have been mapped from remotely-sensed data. The ongoing Fezzan archaeological mission is documenting changes in the distribution of human activity; Palaeolithic industry is restricted to the hamada above the Murzuk escarpment, this may be due to the presence of a large

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palaeolake covering the Wadi el Agial at certain times during the late Pleistocene, including at c. 85ka BP. Neolithic industry is found below the escarpment, but may also be restricted to the shorelines of much smaller Holocene lakes and swamps. Foggara technology was exploited during the period of Garamantean occupation of the area, but the foggaras were subsequently abandoned, possibly due to a continuing fall in the water table, although other factors, such as the introduction of new technologies or a collapse in population, may also be responsible.

References Bischoff J. L. and Fitzpatrick, J. A. (1991) U-series dating of impure carbonates: An isochron technique using total-sample dissolution. Geochim. Cosmochim. Acta, 55, pp. 543-554. Black, S., Macdonald, R. and Kelly, M. R. (1997) Crustal origin for peralkaline rhyolites from Kenya: Evidence from U-series disequilibria and Th-isotopes. Journal of Petrology, 38, pp. 277-297. Brookes, I.A. (1993) Geomorphology and Quaternary Geology of the Dakhla Oasis region, Egypt. Quaternary Science Reviews, 12, pp. 529-552. Bryant, R.G. (1996) Validated linear mixture modelling of Landsat TM data for mapping evaporite minerals on a playa surface: methods and applications. International Journal of Remote Sensing, 17, pp. 315-330. Bond, W.J., Stock, W.D. and Hoffman, M.T. (1994) Has the Karoo spread? A test for desertification using carbon isotopes from soils. South African Journal of Science, 90, pp. 391-397. Causse, C., Coque, R., Fontes, J.C., Gasse, F., Gibert, E., Ben Ouezdou, H. and Zouari, K. (1989) Two high levels of continental waters in the southern Tunisian Chotts at about 90 and 150 ka. Geology, 17, pp. 922925. Cody, R.D. and Cody A.M. (1988) Gypsum nucleation and crystal morphology in analogue saline terrestrial environments. Journal of Sedimentary Petrology, 58, pp. 247-255. Cremaschi, M. and Trombino, L. (1998) The palaeoclimatic significance of paleosols in southern Fezzan (Libyan Sahara): morphological and micromorphological aspects. Catena, 34, pp. 131-156. Drake, N.A. (1995) Reflectance spectra of evaporite minerals (400-2500 nm): applications for remote sensing, International Journal of Remote Sensing, 16, pp. 2255-2572. Dubief. J. (1963) Le climat du Sahara. Mem. Inst. Rech. Sahariennes Alger. 2, 270 pp. Eckardt, F. and White, K. (1997) Human induced disruption of stone pavement surfaces in the central Namib desert, Namibia: observations from Landsat Thematic Mapper. International Journal of Remote Sensing, 18, pp. 3305-3310. Edmunds, W.M. and Wright, F.P. (1979) Groundwater recharge and palaeoclimate in the Sirte and Kufra basins, Libya. Journal of Hydrology, 40, pp. 215-241. Fontes, J.C. (1994) Isotope palaeohydrology and the prediction of long-term repository behaviour. Terra Nova, 6, pp. 20-36. Fontes, J.C. and Gasse, F. (1989) On the ages of humid Holocene and late Pleistocene phases in north Africa - remarks on "late Quaternary climatic reconstruction for the Maghreb (north Africa)" by P. Rognon. Palaeogeography. Palaeoclimatology, Palaeoecology, 70, pp. 393-398. Fontes, J.C. and Gasse, F. (1991) PALHYDAF (Palaeohydrology in Africa) program: objectives, methods, major results. Palaeogeography, Palaeoclimatology, Palaeoecology, 84, pp. 191-215. Gasse, F., Fontes, J.C., Plaziat, J.C., Carbonel, P., Kaezmarska, I., De Dekker, P., Soulié-Märsche, I., Callot, Y. and Dupeuble, P.A. (1987) Biological remains, geochemistry and stable isotopes for the

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reconstruction of environmental and hydrological changes in the Holocene lakes from north Sahara. Palaeogeography. Palaeoclimatology, Palaeoecology, 60, pp. 1-46. Gaven, C., Hillaire-Marcel, C. and Petit-Maire, N. (1981) A Pleistocene lacustrine episode in southeastern Libya. Nature, 290, pp. 131-133. Goblot, H. (1979) Les Qanats, une technique d’acquisition de l’eau. Paris: Mouton Éditeur, 236pp. Haynes, C.V. (1982) The Darb el-Arba’in Desert: a product of Quaternary climatic change. In El Baz, F. and Maxwell, T.A. (Eds.) Desert Landforms of Southwest Egypt: A Basis for Comparison with Mars, Washington DC: NASA, pp. 91-117. Kelly, E.F., Blecker, S.W., Yonker, C.M., Olson, C.G., Wohl, E.E. and Todd, L.C. (1998) Stable isotope composition of soil organic matter and phytoliths as paleoenvironmental indicators. Geoderma, 82, pp. 59-81. Klitsch, E. and Baird, D.W. (1969) Stratigraphy and palaeohydrology of the Germa (Jarma) area, southwest Libya. In Kanes, W.H. (Ed.) Geology, Archaeology and Prehistory of the Southwestern Fezzan, Libya, Petroleum Exploration Society of Libya, Eleventh Annual Field Conference, Castelfranco-Veneto, pp. 6780. Koch, P.L. (1998) Isotopic reconstruction of past continental environments. Annual Review of Earth and Planetary Sciences, 26, pp.573-613. Kuzucuoglu, C., Pastre, J-F., Black, S., Ercan, T., Fontugne, M., Guillou, H., Hatté, C., Karabiyikoglu, M., Orth, P. and Türkecan, A. (1998) Identification and dating of tephras from Quaternary sedimentary sequences of inner Anatolia. Journal of Volcanology and Geothermal Research, 85, pp. 153-172. Létolle, R. (1980) N-15 in the natural environment. In Fritz, P. and Fontes, J.C. (Eds.) Handbook of Environmental Isotope Geochemistry Volume 1, Amsterdam: Elsevier, pp. 407-433. Luo, S. and Ku, T. L. (1991) U-series isochron method: A generalised method employing total-sample dissolution. Geochim. Cosmochim. Acta 55, pp. 555-564. Maher, B.A. (1986) Characterisation of soils by mineral magnetic measurements. Physics of the Earth and Planetary Interiors, 42, pp. 76-92. Mattingly, D.J., al-Mashai, M., Aburgheba, H., Balcombe, P., Eastaugh, E., Gillings, M., Leone, A., McLaren, S.J., Owen, P., Pelling, R., Reynolds, T., Stirling, L., Thomas, D., Watson, D., Wilson, A.I. and White, K. (1998) The Fezzan Project 1998, preliminary report on the second season of work. Libyan Studies, 29, pp. 115-144. McKenzie, H.S. and El Saleh, B.O. (1994) The Libyan Great Man-Made River Project 1; Project Overview. Proceedings of the Institution of Civil Engineers; Water Maritime and Energy, 106, pp. 103-122. Oldfield, F. (1991) Environmental magnetism, a personal perspective. Quaternary Science Reviews, 10, pp. 73-85. Peel, R.F. (1966) The landscape in aridity. Transactions of the Institute of British Geographers, 38, pp. 123. Petit-Maire, N., Delibrias, G. and Gaven, C. (1980) Pleistocene lakes in the Shati area, Fezzan. Palaeoecology of Africa, 12, pp. 289-295. Petit-Maire, N., Reyss, J.L. and Fabre, J. (1994) A last interglacial palaeolake in the hyperarid Sahara of Mali. Comptes Rendus de l’Academie des Sciences Serie II, 319, pp. 805-809. Pim, R.H. and Binsariti, A. (1994) The Libyan Great Man-Made River Project 2; The Water Resource. Proceedings of the Institution of Civil Engineers; Water Maritime and Energy, 106, pp. 123-145. Rummery, T.A., Bloemendal, J., Dearing, J., Oldfield, F. and Thompson, R. (1979) The persistence of fireinduced magnetic oxides in soils and lake sediments. Annales de Géophysique, 35, pp. 103-107. Szabo, B.J., Haynes, C.V.Jr. and Maxwell, T.A. (1995) Ages of Quaternary pluvial episodes determined by uranium series and radiocarbon dating of lacustrine deposits of eastern Sahara. Palaeogeography, Palaeoclimatology, Palaeoecology, 113, pp. 227-242.

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Settle, J. J. and Drake, N.A. (1993) Li near mixing and the estimation of ground cover proportions. International Journal of Remote Sensing, 14, pp. 1159-1177. Tite, M.S. and Mullins, C.E. (1971) Enhancement of the magnetic susceptibility of soil on archaeological sites. Archaeometry, 13, pp. 209-219. van der Veen, M. (1992) Botanical evidence for Garamantean Agriculture in Fezzan, Southern Libya. Review of Palaeobotany and Palynology, 73, pp. 315-327. Walden, J., Oldfield, F. and Smith, J.P. (Eds.) (1999) Environmental Magnetism, a Practical Guide. Technical Guide No. 6, London: Quaternary Research Association, 243pp. Wendorf, F. and Schild, R. (1980) Prehistory of the Eastern Sahara. New York: Academic Press, 414pp. Wendorf, F., Schild, R., Said, R., Haynes, C.V., Gautier, A., and Kobusienwicz, M. (1976) The Prehistory of the Egyptian Sahara. Science, 193, pp. 103-114. White, K. and Drake, N.A. (1993) Mapping the distribution and abundance of gypsum in south-central Tunisia from Landsat Thematic Mapper data. Zeitschrift für Geomorphologie N.F., 37, pp. 309-325. Wickens, G.F. (1975) Changes in the climate and vegetation of the Sudan since 20,000 B.P. Boissiera, 24, pp. 43-65. York D. (1969) Least squares fitting of a straight line with correlated errors. Earth and Planetary Science Letters, 5, pp. 320-324.

RELICT CRYOGENIC MOUNDS IN THE UK AS EVIDENCE OF CLIMATE CHANGE

STEPHEN D. GURNEY Landscape and Landform Research Group, Department of Geography, The University of Reading, PO Box 227, Whiteknights, Reading RG6 6AB, U.K.

Abstract Relict perennial cryogenic mounds, (the remains of pingos or palsas) have long been identified in the UK and have been attributed to the Last Glacial or Younger Dryas cold periods. These features take the form of circular to oval depressions surrounded completely or partially by a raised rim or rampart. The depressions very often contain a wetland with an organic soil or peat. At depth within the depressions there can be up to several metres of soft sediments (very often a silt-clay). The ramparted depressions are usually found in clusters where individual features may overlap such that they share ramparts or even where the depressions merge to produce a ‘figureof-eight’ shape in plan. Such relict periglacial features are extremely useful for palaeoclimatic reconstruction, since they can indicate the average thermal condition of the ground during the period in which they were actively forming. Many relict cryogenic mounds in the UK have been interpreted as the remains of hydraulic pingos, features which in contemporary cold climate areas indicate either discontinuous permafrost (mean annual air temperature of or continuous permafrost of Some of these features, however, have now been re-interpreted as the remains of ‘mineral-cored’ palsas, features that are indicative of only sporadic permafrost 0f or of or discontinuous permafrost of . Thus the varying interpretation and classification of these geomorphological features can lead to a wide disparity in their palaeoclimatic significance. Clearly, to obtain the correct palaeoclimatic information, such features 209

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must be accurately identified and appropriate modern analogues utilised in their interpretation. KEY WORDS: Relict cryogenic mounds, geomorphology, palaeoclimatic significance.

1.

Introduction

Relict perennial cryogenic mounds are very important geomorphological features because they indicate the former presence of permafrost. Unfortunately, their correct identification and interpretation is not straightforward and the use of many specific features for palaeoclimatic reconstruction has been questioned, either on the basis of their actual interpretation or the use of inappropriate modern analogues. Contemporary perennial cryogenic mounds are generally classified as either pingos or palsas, although the reality may be that these two types represent the extremes of a ‘landform continuum’. Pingos are the largest of the perennial cryogenic mounds and can be over 50 m high, or have diameters of over 500 m (see Gurney, 1998). Pingos are classified as either hydrostatic (formerly known as closed-system) or hydraulic (opensystem). Hydrostatic pingos are usually solitary and their type-site is the Mackenzie Delta/Tuktoyaktuk Peninsula area of the Northwest Territories of Canada (Mackay, 1998). Hydraulic pingos occur in groups or clusters, and their type-site is east Greenland (Müller, 1959). Both pingo types rely on a pressure system to deliver water to the core where it is frozen to form the ice-core responsible for local uplift of the overburden. Palsas are generally smaller than pingos ( or more high and or more diameter) and take two basic forms: those purely composed of peat and those which have a surficial layer of peat but whose core is developed in mineral sediments (mineral-cored palsas, e.g. Åhman, 1976). It must be noted, however, that an overlap exists with features such as ‘peat plateaux’, which are less discrete and can be much more areally extensive, although here the most ‘palsa-like’ features may be formed through degradation as opposed to aggradation (e.g. Hinkel, 1988). Palsas grow through the accumulation of a segregated ice core whose growth is sustained by cryosuction, which draws water to the freezing front (see Washburn, 1983a; 1983b). One of the main problems with the reconstruction of former permafrost areas from geomorphological features is that these features are discrete and are only locally preserved. Furthermore, to infer only former temperature conditions from the existence of these features is far too simplistic. Many periglacial features are to some extent polygenetic and therefore rely on a number of parameters for growth and maintenance, temperature being only one of these. Of considerable importance are the freeze-thaw regime, for example, is it a seasonal regime like that of the high latitudes, where solar insolation is concentrated in the summer and is absent in the winter, or is it more diurnal like that of the mid and low latitudes, where daily freeze-thaw cycles may operate increasing the total number of cycles per year. The thickness, timing, redistribution (by

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wind) and duration of snow cover are also extremely important since snow can reduce the penetration of frost into the ground surface (a process which may be vital to the creation of palsa type cryogenic mounds, see Seppälä, 1994). The presence/absence of vegetation and vegetation types and their albedo are also very important in the understanding of the ground thermal regime (e.g. Railton and Sparling, 1973). Clearly all of these factors must be taken into account if relict cryogenic mounds are to be used as indicators of some quantifiable climate change. This chapter will attempt a review of the literature of relict cryogenic mounds and will discuss their possible use as palaeoclimatic indicators. Although the focus is on relict cryogenic mounds in the UK (for the distribution of which see Figure 1), to ignore evidence from the rest of north west Europe and further afield would clearly be misleading and therefore an attempt will be made to synthesise and incorporate this material, where relevant.

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Previous reviews have been conducted by Wiegand (1965) who thoroughly catalogued the major relict ‘pingo’ sites of middle Europe favouring a periglacial ground-ice hypothesis for their formation (see also Maarleveld, 1965). Flemal (1976) attempted to provide a detailed discussion on relict ‘pingos’ (their characteristics and distribution) as well as their use in the reconstruction of former permafrost environments. Although Flemal does consider that palsas or another type of cryogenic mound might be the cause of some of the features described as relict ‘pingos’. Bryant and Carpenter (1987) reviewed all the ramparted ground-ice depressions in Britain and Ireland and provided a good distribution map and summary table of these features. The overlap between pingos and palsas was not, however, made clear. De Gans (1988) produced a comprehensive discussion of Pleistocene ‘pingo’ remnants, reviewing evidence from Europe and North America.

2. The basic geomorphology and sedimentology of relict cryogenic mounds The criteria used to identify relict cryogenic mounds should be clear. Generally these features take the form of circular to oval depressions that are surrounded completely or partially by a raised rim or rampart (see Table 1). They are generally found in clusters as opposed to single isolated features (see Figures 2 and 3). When the spatial density of the features is high, it is common for the ramparts and even depressions to become merged at their edges creating a mutually interfering pattern. The clusters of ramparted depressions can be found in a range of settings from valley bottoms to lower valley sides and occasionally on plateau sites. The basic sedimentology of the features is rather more problematic since many of the sites are understood and characterised in terms of the morphology alone. When considering the basic sedimentology, it is expected that the sediments comprising the rampart should be partially derived from material that has sloughed off the mound and partially derived from material that has been pushed outwards from the mound core during growth of the ice body (whether this was composed of massive-injection ice or segregated ice or a combination of the two). The material within the depression (often referred to as the ‘depression fill’ or ‘pingo fill’) should be related to the material within which the ice core formed and to the overburden material (since some of this may have collapsed into the depression on the decay of the mound, particularly if the decay was top-down). Above the pingo-fill there are very often pond or marsh deposits (e.g. peat), which formed after the collapse of the feature, and these are therefore not diagnostic of mound genesis. When relict cryogenic mounds have been investigated, however, clear relationships such as those expressed above have rarely been determined.

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Previous investigations of relict cryogenic mounds

3.1. EARLIEST WORK Relict cryogenic mounds (the remains of perennial cryogenic mounds which formed in a periglacial environment) were first identified in the UK by Pissart (1963). This work followed the investigation by Pissart (1956; 1958) of the “viviers” (literally ‘fish-ponds’) of the Hautes Fagnes plateau of north-east Belgium. They take the form of more or less circular depressions surrounded by a raised rim or rampart. Pissart ascribed the features to a periglacial genesis and termed them pingos, however, he stated that they were probably mostly composed of segregated ice (as opposed to massive or injection ice, more typical of contemporary pingo cores) when they were active. The ramparts were seen to be the key and these appeared to have been built by material soliflucting off the sides of the mound and perhaps also, by a lateral thrust from the interior of the mound.

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This paper was pioneering, not only for the discovery of relict cryogenic mounds in north west Europe indicating a colder, permafrost, environment of the late Pleistocene, but also in recognising that the genesis of these forms may not have been exactly like that of the contemporary arctic pingos that had been studied up to that time. Pissart’s work on relict cryogenic mounds in Belgium continued over many years and extensive investigations were undertaken which refined the hypotheses of their formation and ultimately led to their reclassification. As early as 1974, Pissart suggested that, in terms of genesis these features when active were more similar to palsas, than pingos. An important facet of the growth mechanism appeared to be lateral growth of the segregated ice core, which displaced sediment outwards (Bastin et al., 1974). This is far more closely related to palsa growth than to pingo growth where the diameter of the mounds is formed early on in the growth cycle and subsequent growth takes place upwards (Mackay, 1979). Work continued (e.g. Pissart et al., 1975) and the use of trenches through the features allowed detailed studies of the structure at depth and of the sedimentology (e.g. Pissart and Juvigné, 1980; Pissart, 1983a; Pissart et al., 1975; Juvigné and Pissart, 1979). The discovery of a peat layer that would have existed at ground level during the growth phase of the mounds (Pissart and Juvigné, 1980) and sedimentary structures in the base of the ramparts which were thought to be connected to the lateral growth of segregation ice lenses led to a reclassification of the features from the remains of hydraulic pingos to the remains of ‘mineralogic palsas’ (Pissart, 1983a).

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Initially the palaeoclimatic inference of these features was believed to be similar to that of contemporary hydraulic pingos (a MAAT of 7°C), however, this was subsequently modified to between 0°C and 5°C (Pissart, 1965). In terms of using these features to reconstruct permafrost type, this was important since it meant that these features were indicators of discontinuous rather than continuous permafrost.

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3.2. FEATURES IN THE UK The first study of relict cryogenic mounds in the UK compared the features of Llangurig in Wales with those of the Hautes Fagnes Plateau of Belgium (Pissart, 1963). The Welsh examples were formed in a valley bottom, which posed the possibility of an hydraulic pingo genesis, whereas the Belgian pingos were exclusively on the plateau of the Ardennes with a generally northerly aspect. The question must be posed here as to whether the genesis of the two groups of features can possibly be the same. The key point in this paper and the significant similarity between the two types of relict forms is the fact that the ramparts in both Belgium and Wales are thought to provide evidence of lateral thrust exerted by a growing ice lens. This may divide these types of mounds from arctic pingos where massive ice cores are most common. The importance of segregated ice in arctic pingos, as found in Scandinavian palsas, however, is now also known to play a role. Other early work on the relict cryogenic mounds was conducted by Trotman (1963) who used palynological data to report a vegetation history for the peat, which had developed inside the depression of one of the relict cryogenic mounds near Llangurig. The earliest peat was said to belong to an early pre-boreal time although this peat formed after the mound had collapsed and possibly a long time after and therefore is of no use in dating the active phase of the feature. The study of relict cryogenic mounds in the UK during the 1970s and 1980s was largely conducted by the Watsons. Watson (1971) reported a group of relict cryogenic mounds from Wales and the Isle of Man. Watson summarised the evidence at the Llangurig site and described another site in west Wales (Cledlyn) as well as the site on the Isle of Man at Ballaugh where the depressions are without ramparts. All of the features were classified as the remains of hydraulic pingos. The Cledlyn site was further documented by Watson (1972) and by Watson and Watson (1972). The latter paper detailed all of the cores taken from the depressions with respect to both the grain size of the included sediments and the profiles produced from the depth data. The major problem with these features is the large quantity of clay-silts, which compose the depression fill. This material was believed to have been deposited in the pingo pond during and after the collapse of the pingo, however, since the thickness of this deposit reaches over 7 m, questions must be raised whether this deposit, may possibly have existed at depth before the formation of the cryogenic mounds. Another site in west Wales at Cletwr was documented in Watson (1972) and in Watson and Watson (1974). Examples of the forms were mapped and levelled and extensive maps and sections were produced from these data and from that gained by coring. Linear relict cryogenic mounds at this site were likened to those described by Mückenhausen (1960) of the Hohen Venn/Eifel region, Germany that had also been classified as the remains of hydraulic pingos. The paper by Sparks et al. (1972), which describes “presumed ground-ice depressions” at East Walton Common in East Anglia, is often cited as evidence for relict cryogenic

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mounds in the UK. It was concluded that these features were most likely to have an origin in ground-ice accumulation, but whether this was of the pingo or palsa type was never speculated upon (except that springs were believed to have delivered the water to the core). Other sites in the UK with relict cryogenic mounds were documented during the 1980s. Hutchinson (1980) speculated upon possible Late Quaternary mound remnants in central London. The features take the form of drift-filled depressions and had been previously interpreted as Late Quaternary scour hollows (Berry, 1979). Hutchinson reinterpreted them as the remains of hydraulic pingos. Carpenter and Woodcock (1981) reported detailed investigations of a ramparted depression at Elstead in Surrey (still regarded as the most southerly relict cryogenic mound in the UK). The feature was interpreted as the remains of a Late Devensian hydraulic pingo. Bryant et al. (1985) presented a case study of relict cryogenic mounds from the Whicham Valley of Cumbria. These features were defined as the remains of hydraulic pingos and taken to indicate the former existence of discontinuous permafrost. The features were believed to be Younger Dryas in age. Miller (1990) documented relict cryogenic mounds near Brent Tor, western Dartmoor. It was concluded that they were remnants of hydraulic pingos whose water supply was derived from the passage of groundwater along a sub-surface thrust plane. Taylor (1987) reported the existence of features believed to be Younger Dryas age, relict hydraulic pingos at Abermad in west Wales. The coverage of this topic in Ballantyne and Harris (1994) is extremely useful, although the suggestion that the relict cryogenic mounds at Brent Tor are the most likely candidates for mineral palsa origin is questionable. It is likely indeed, that due to the geological conditions at the site, for example the existence of faults in the bedrock beneath the site, that these features are more closely related to hydraulic pingos. In Gurney (1995) the relict cryogenic mounds of mid and west Wales were re-evaluated on the basis of new investigations. Despite previous studies classifying these features as the remains of hydraulic pingos, a tentative reinterpretation was made that these features actually represented the remains of mineral-cored palsas on the basis of the sedimentological composition of the depression fills, which, it was felt did not concur with the previously hypothesised genesis of these features. Further work was thought to be warranted at this site. Gurney and Worsley (1996) documented the remains at Owlbury on the Shropshire/Powys border, which take the form of classic relict cryogenic mounds. These features were interpreted as the remains of former mineral-cored palsas as opposed to hydraulic pingos, on the basis of morphology and sedimentary setting.

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3.3 FEATURES IN THE REST OF NORTH WEST EUROPE

North west Europe has produced sites of relict cryogenic mounds other than those of Belgium and the UK. These include: Ireland (Mitchell, 1971; Mitchell, 1973; Coxon, 1986; Coxon and O’Callaghan, 1987), Denmark (Cailleux, 1957), Germany (Mückenhausen, 1960; Picard, 1961), Luxembourg (Slotboom, 1963), Poland (Dylik, 1965), France (Boyé, 1957; Rousset, 1965; Bout, 1986) and The Netherlands (Ploeger and Waateringe, 1964; Paris et al., 1979; De Gans and Sohl, 1981; De Gans, 1982; Bijlsma, 1983; Van der Meulen, 1988). Where an age has been assigned to the active phase of these features it was generally Last Glacial, although occasionally Younger Dryas. Where a genetic system was suggested it was generally of the hydraulic pingo type (e.g. De Gans and Sohl, 1981) although one these sites did appear to suggest an hydrostatic pingo genesis (Van der Meulen, 1988). The general characteristics of the relict cryogenic mounds in the rest of north-west Europe, takes a similar form to that of the UK (i.e. circular to oval ramparted depressions), however, there are notable instances where they are located on slopes or plateau sites (as opposed to lower valley sides or valley bottom locations).

3.4 FEATURES OUTSIDE NORTH WEST EUROPE

Outside north-west Europe the documentation of relict cryogenic mounds is rather scarce. Flemal et al. (1969 and 1973) summarised the characteristics of a group of raised, oval to circular, nearly flat-topped mounds, which cover at least of Bloomington ground moraine in north-central Illinois, USA. These forms are termed the ‘DeKalb mounds’ and were interpreted as the remains of pingos, and their existence was used to infer at least discontinuous permafrost in Illinois in Woodfordian time (Late Wisconsinan). These features are quite unlike any that have been described from Europe where relict cryogenic mounds are seen to be ramparted depressions not flattopped mounds. The hypothesis of origin states, however, that an ice-walled lake was created when the original cryogenic mounds collapsed and it was into this lake that the sediments comprising the mounds were deposited. What became of the rampart that can be seen associated with many collapsed contemporary pingos is not clear. Trimble (1978) described what he termed ‘pingo scars’ from southern North Dakota, USA. These features were circular and consisted of an outer ridge, which had a moat on the inner side. In the centre there was a dome and the whole of the feature could have relief of up to 17 m. Stone and Ashley (1992) described relict cryogenic mounds, which were some of the first features documented from North America (in the states of Connecticut and Massachsetts, USA) which were similar geomorphologically to the relict cryogenic mounds of north west Europe (see also Stone and Ashley, 1989 and Stone et al., 1991). These circular and sub-circular shallow depressions with subtly raised rims occur on a

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drained glacial-lake bottom. Marsh (1987) outlined the evidence for relict cryogenic mounds in Pennsylvania, USA that here take the form of elliptical basins surrounded by low ramparts. At one site (Halfway Run) 55 well-preserved remnants occur in a cluster. These features were interpreted as the remains of hydraulic pingos. A field of depressions similar to those at Halfway Run in Snyder County, however, were thought to be reminiscent of “mineralogical palsas” as described by Pissart (1983b) rather than relict pingos as described by Watson (1971).

4.

The use of modern analogues

The morphology and sedimentology of specific relict cryogenic mound sites in the UK has generally determined the choice of a modern analogue, except in some cases where particular analogues were chosen more because they had been previously employed for other sites. Once a modern analogue has been chosen it is often used as a source of information about possible geomorphological processes that would have operated during the life-cycle of the, now relict, features. The modern analogues are also used, of course, as a source of information concerning the palaeoclimatic significance of the relict features. Two large assumptions are made by the use of modern analogues: firstly that the growth and decay processes of the modern analogues are relevant to the relict features in question and secondly that the modern analogues are actually active in their present climatic setting. It is clear that in the past inappropriate modern analogues for relict cryogenic mounds have been employed on the basis of one or even both of these assumptions. This has led to misleading palaeoclimatic inferences being drawn. 4.1. PINGO ANALOGUES Most commonly, hydraulic (open system) pingos have been taken to be the most appropriate modern analogues for relict cryogenic mounds in the UK. This is not surprising since classically hydraulic pingos occur in lower valley side and valley bottom locations, in clusters which are mutually interfering (see Figure 4), characteristics which are shared with many of the relict cryogenic mounds sites. The term ‘open system’ was coined by Müller (1959) and the type site for these features is in east Greenland, although good examples are also found in central Alaska (e.g. Holmes et al., 1968). Importantly the key to understanding the growth of these features is often a knowledge of the local groundwater movement (see Worsley and Gurney, 1996; Gurney, 1998). The groundwater upwelling, which is sometimes called the ‘pingo spring’, is the source of the water, which forms the pingo ice-cores, and it is the varying location of upwelling, which causes the clustering of individual forms. Occasionally a pingo will develop within the depression created by the collapse of a previous pingo, which essentially creates ‘a pingo within a pingo’. There are some contemporary pingos whose specific genesis does not appear to fit neatly into the hydrostatic/hydraulic classification scheme (e.g. Pissart, 1967; Gurney and Worsley, 1997). Such features may be more suitable modern analogues than the classic or ‘type’ features that are more often employed.

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4.2. PALSA ANALOGUES Modern analogues of the palsa variety have also been cited for the relict cryogenic mounds of the UK. It should be made clear, however. that this is far more problematic than employing a pingo analogue because palsa terminology is arcane, even misleading, and the definitions employed are generally morphological as opposed to genetic in nature (see Nelson et al., 1992). It is not the intention here to attempt to clarify the terminological debate, but rather to indicate what the use of a palsa analogue means for interpretation and palaeoclimatic reconstruction.

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When a palsa analogue is invoked for relict cryogenic mounds in the UK, or indeed elsewhere, this is usually because it is believed that a segregated ice core fed by cryosuction was responsible for the growth and maintenance of the original mounds as opposed to a more massive ice core fed primarily by injection ice from water under pressure (whether hydrostatic or hydraulic). It is clearly not employed because it is thought that the mounds were composed of peat or were developed in a mire, or indeed had even the merest covering of peat. Therefore, for some, the use of a palsa type analogue is an inappropriate use of the terminology. A more pragmatic view, however, is that many relict cryogenic mounds, when active, had more in common genetically with palsas than with pingos and therefore a palsa analogue is more appropriate. Very often work conducted on relict cryogenic mounds in north west Europe has inspired the search for suitable modern analogues in present cold climate environments (e.g. Worsley et al., 1995). In the case of the re-interpretation of the relict cryogenic mounds in Belgium, a need was created for an examination of sites, which might form appropriate modern analogues, and these were forthcoming (e.g. Gangloff and Pissart, 1983; Pissart and Gangloff, 1984). More recently, investigations in cold climate areas have realised the requirement for more detailed morphological and sedimentological comparison of contemporary and relict cryogenic mounds. For example, Matthews et al. (1997) specifically relate the evidence of cyclic palsa growth and decay at altitude in south central Norway to relict (Pleistocene) features in the UK. Given the nature of analogue use for palaeoclimatic reconstruction thus far, such studies conducted in cold climate areas of a lower latitude (62°N) are, perhaps, more valuable than the arctic and high-arctic sites often cited in the past.

5. Alternative interpretations of depressions The alternative origins of any features that might be used to infer former climatic conditions should always be considered to ensure that the interpretation is as sound as possible. All too often features which are used as palaeoclimatic indicators will be taken up by other workers who are not familiar with their geomorphology and this use of secondary level information can perpetuate the use of features for palaeoclimatic reconstruction long after they have been re-interpreted or even falsified. Prince (1964) produced perhaps the most extensive piece of work on the depressions and pits of Norfolk, in which a periglacial origin was considered (see also Prince, 1961). The main body of the paper, however, was concerned with the possible anthropogenic origin of these depressions. The possibility of mineral workings and marl pits being the two most obvious modes of origin. The close juxtaposition of the pits with features with an obvious periglacial origin (e.g. patterned ground) is always a problem since it often leads one to believe that their existence might be linked. The comprehensive discussion of the alternative origins of the pits, however, forces one to consider the other possibilities more fully. This is just as it should be and this paper should remind us all

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that although in the literature of north west Europe there are numerous descriptions of depressions where a periglacial interpretation is enticing, this is not necessarily the correct interpretation. In conclusion Prince states that in Norfolk seemingly identical hollows in terms of basic morphology cannot possibly be all attributed to the same origin, anthropogenic, periglacial or otherwise. The hollows are simply too numerous and are much more complex than is immediately apparent.

6.

Palaeoclimatic inferences

The palaeoclimatic inferences of relict cryogenic mounds are based purely on the current climatic conditions observed that affect the contemporary cryogenic mounds that are believed to represent their modern analogues (see Pissart, 1987). Thus if a particular set of relict cryogenic mounds are interpreted as relict hydraulic pingos, then the climatic regime typical of hydraulic pingos (a MAAT of -3 to -7°C which can sustain discontinuous or continuous permafrost) will be applied. In this way the palaeoclimatic significance of relict cryogenic mounds is wholly reliant on correct interpretation and the appropriate use of modern analogues. If the features are incorrectly interpreted, or if indeed there are no suitable modern analogues, then the palaeoclimatic inferences will be poor or even misleading. Features originally interpreted as the remains of hydraulic pingos and now believed to be more closely related to the former existence of mineralcored palsas have suffered a major change in their palaeoclimatic significance (from an inferred MAAT of -7°C or less to perhaps only -1 °C). Compounding this problem is the fact that many if not most periglacial features rely on a number of climatic and environmental parameters for their growth and/or maintenance. These other climatic parameters are generally not referred to and only the former temperature regimes are considered. The number of freeze-thaw cycles or snow depth, for example, is not included in palaeoenvironmental reconstructions. An additional source of error is that it is usually the mean temperatures, which are inferred, as opposed to the lows or indeed the seasonal temperature profile (which would clearly have been markedly different in mid-latitude periglacial areas compared to that of high latitude periglacial areas). With cryogenic mounds the ice core is fully within the permafrost, well below the active layer, and therefore, mean temperatures are probably not such a poor way of characterising the thermal regime. Even here, however, these mean temperatures are still those of the air as opposed to the ground (mean annual air temperatures as opposed to mean annual ground temperatures, MAGT), which is, perhaps, the more significant problem with all works of palaeoclimatic reconstruction. For this reason, it has been suggested that the palaeoclimatic inferences are too low since the MAGT is usually higher than of the MAAT by 1 to 5°C because in winter snow protects the ground from the extreme cold and in summer the ground is warmed more by solar radiation than the air (Williams, 1975). Yet again the interpretation of the MAGT

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is a generalisation, which ignores other factors as described earlier (albedo, vegetation cover, snow etc).

7.

Problems with the correct identification of relict cryogenic mounds

At most of the sites it has been the morphology of the relict cryogenic mounds, which has primarily been used to interpret the genesis of these features. The sedimentological composition of the features has been used less in their interpretation either because its significance has been unclear or because the sediments have largely been inaccessible due to the lack of sections (natural or artificial). Whilst coring has been employed it has largely been inconclusive. The obvious exception to this, of course, is the work of Pissart on the relict cryogenic mounds of Belgium, where good sections of the sediments have been studied. In general, the sediments of relict cryogenic mounds in the UK can be divided into four basic groups: 1) the rampart sediments, 2) the intra-rampart sediments, 3) which should reflect the general surficial deposits of the surrounding area and therefore, are very often composed of slope deposits, which typically mantle the lower valley sides and valley bottom. 4) the upper basin sediments (where present) 5) where present are generally organic in nature and are commonly composed of peat. At various sites where this peat has been studied it has provided a younger age limit for the features (very often a very much younger age limit than that postulated for the date of the mound collapse), and 6) the lower basin sediments 7) which should compose the depression or ‘pingo fill’ and are very often composed of fine-grained material. Any interpretation of relict cryogenic mounds should be based upon the morphology and sedimentology of the features in question. The genetic hypotheses proposed should account for the location and sediment composition and should include the processes of collapse and also the evidence that might survive the collapse. It is important to remember that contemporary cryogenic mounds have a growth and decay cycle which operates independently of climatic change and therefore some of the relict cryogenic mounds may be the result of decay during the colder climatic phase rather then decay because of the end of the colder climatic phase (the climatic amelioration). Even in collapse, the variety of cryogenic mounds operate differently, for example, pingos tend to collapse from the summit down, whereas palsas generally suffer rupture of the material insulating the core at the side low down or even at the base of the mound.

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Relatively little is known about the growth mechanisms of relict cryogenic mounds, except perhaps, that once established they appeared to grow laterally as well as, presumably, vertically. This is far more similar to that of palsas than pingos as indicated previously. The growth mechanism is extremely important to understand. In pingos it allows development in a wide range of sediments from bedrock through coarse gravels to finegrained material. In palsas, however, there is a universal requirement for a frost susceptible, and thus fine-grained substrate, which cans either be of peat or mineral sediments, and it is only in such a medium that cryosuction and the growth of the segregated ice core can develop. Thus any study of the composition of relict cryogenic mounds should evaluate which sediments sustained the ice core (segregated or otherwise). If a fine-grained sediments body is identified then there is the possibility that cryosuction is involved, if there is not then other processes may have contributed to the growth of the ice core. The relict cryogenic mounds of the Welsh borders and of west Wales (Gurney, 1995; Gurney and Worsley, 1996) have a large volume of fine-grained sediments within their central depressions. It has been postulated that this sediment existed prior to the cryogenic mound growth and, in fact, facilitated that growth acting as a frost susceptible substrate within which cryosuction could operate. Previous interpretations have placed this sediment as accumulating from the deposition of material into the ‘pingo pond’, that is, the pond that often develops in the initial crater that develops in the summit of the pingo when it begins to collapse. This material, would, in part, be derived from the rampart deposits, however, very often there is no evidence for the coarse-grained slope deposits, which form the ramparts, in the basin fill.

8.

Conclusions

Clearly, the existence of relict cryogenic mounds in the UK has often been used to infer past climate (specifically for the Late Quaternary). Their almost unique status in providing information about the thermal state of the ground (i.e. type of permafrost necessary to create and /or maintain them) is responsible for this. It remains, however, that these features are not as straightforward to identify as was once thought and it should now be apparent that there are a range of analogues, which may be applied to their interpretation. Depending upon the choice of analogue, a very wide range of palaeoclimatic/palaeoenvironmental inferences may be drawn. It is important that we realise the limitations of using geomorphological features for such palaeoclimatic reconstructions. Such features are only preserved in favourable, discrete locations and the temperature inferences are not clear cut given the confusion over current air temperatures and the actual importance of ground temperatures. Clearly more work is needed to investigate the internal structure and sedimentology of these features in order

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so that the most appropriate analogue can be applied (if one is available) and much thought must be used when deciding on what exactly the climatic inferences might be.

Acknowledgements

The author would like to thank Graham Elliott, Jukka Käyhkö and Peter Worsley for field assistance and for many useful discussions concerning the evidence for relict cryogenic mounds in the UK and elsewhere. Thanks also to Heather Browning for creating the figures from the original tortured diagrams. Finally the comments of an anonymous referee helped to clarify the text in a number of places.

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Trimble, D.E. (1978) Pingo scars in southwestern North Dakota, United States Geological Survey Professional Paper 1100, 58. Trotman, D.M. (1963) Peat deposits within a pingo remnant near Llangurig, Wales, Zeitschrift für Geomorphologie 7, 168-171. Van der Meulen, S. (1988) The spatial facies of a group of pingo remnants on the southeast Frisian till plateau (the Netherlands), Geologie en Mijnbouw 67, 61-74. Washbum, A.L. (1983a) Permafrost features as evidence of climatic change, Earth Science Reviews 15, 327-402. Washbum, A.L. (1983b) Palsas and continuous permafrost, Proceedings of the 4th International Conference on Permafrost, 1372-1377. Watson, E. (1971) The remains of pingos in Wales and the Isle of Man, Geological Journal 7, 381-392. Watson, E. (1972) Pingos in Cardiganshire and the latest ice limit, Nature 236, 343-344. Watson, E. and Watson, S. (1972) Investigation of some pingo basins near Aberystwyth, Wales, Report of the 24th International Geological Congress {Montreal) Section 12, 212-233. Watson, E. and Watson, S. (1974) Remains of pingo in the Cletwr basin, south-west Wales, Geografiska Annaler 56A, 213-225. Wiegand, G. (1965) Fossile pingos in Mitteleuropa, Wurzburger Geographische Arbeiten 16, 152pp. Williams, P.J. and Smith, M.W. (1989) The frozen Earth. Fundamentals of Geocryology, Cambridge University Press, Cambridge, 306pp. Williams, R.B.G. (1975) The British climate during the last glaciation ; an interpretation based upon periglacial phenomena, in Wright, A. E. and Moseley, F. (Editors) Ice ages ancient and modern. Seel House Press, Liverpool, 95-120. Worsley, P. and Gurney, S.D. (1996) Geomorphology and hydrogeological significance of the Holocene pingos in the Karup Valley area, Traill Island, northern east Greenland, Journal of Quaternary Science 11, 249-262. Worsley, P., Gurney, S.D. and Collins, P.E.F. (1995) Late Holocene ‘mineral palsas’ and associated vegetation patterns: a case study from Lac Hendry, northern Québec, Canada and significance for European Pleistocene thermokarst, Quaternary Science Reviews 14, 179-192.

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INVESTIGATIONS INTO LONG-TERM FUTURE CLIMATE CHANGES P.E. BURGESS, J.P. PALUTIKOF* AND C.M. GOODESS Climatic Research Unit, University of East Anglia, Norwich NR4 7TJ, UK * Corresponding author

Abstract A two-dimensional climate model (LLN-2D) has been used to investigate long-term climate change (over periods of around years in the past and in the future) for Central England. The model is forced by periodic changes in the distribution of incoming radiation associated with periodic changes in the Earth’s orbit around the Sun and by changes in the atmospheric concentration of Future natural changes in concentration are estimated using a regression equation. Eight anthropogenic scenarios have been constructed. The LLN-2D simulations indicate that anthropogenic effects have the potential to disturb the climate system over very long time scales and demonstrate the non-linearities that can operate between cause and effect. 1.

Introduction

Over long time scales, defined here as those greater than years, climate change is of primary importance in determining the development of local and regional landforms and ecosystems. Thus, an understanding of climate change is essential in order to understand landform changes in the British Isles over the last two million years (the Quaternary period). It is also important for studies of long-term future development of landforms, such as those required for safety assessments of underground radioactive waste disposal in the UK (Thorne, 1995; Nirex, 1997). The research presented in this chapter has been funded by United Kingdom Nirex Limited and British Nuclear Fuels plc. A knowledge of the types and sequences of climatic states likely to be experienced in the UK over the next glacial-interglacial cycle (approximately the next 100,000 years (100 ka)) is relevant to quantitative radiological assessment studies in the UK. It is widely, though not universally, accepted that the major cause of the glacialinterglacial cycles observed over the Quaternary period are periodic changes in the seasonal and latitudinal distribution of incoming radiation associated with periodic 231

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changes in the Earth's orbit around the Sun (see Goodess et al., 1992a, b; Adcock et al., 1997; Goodess et al., 1999 and Palutikof et al., 1999 for reviews of orbital forcing, or Milankovitch, theory). The changes in insolation due to periodic variations in the Earth’s orbital parameters of obliquity, precession and eccentricity are predictable. This raises the possibility of modelling climate changes due to orbital forcing not only over long periods in the past, but also into the future. A two-dimensional climate model has been developed at the L'Institut d'Astronomie et de Géophysique de Georges Lemaître at the Université Catholique de Louvain, Belgium (Gallée et al., 1991) which is able to estimate parameters such as temperature and ice sheet volumes at 1000 year time steps and over 5° latitudinal bands for the Northern Hemisphere. The forcing inputs to the model are insolation and atmospheric concentration at each time step. The Louvain-la-Neuve model, referred to here as LLN-2D, has been used to investigate long-term climate change (over periods of around years in the past and in the future) for central England (Burgess, 1998; Goodess et al., 1999). These modelling studies are the subject of this chapter. First, we describe the characteristics of the LLN-2D model in Section 2. Second, the insolation and atmospheric forcing inputs due to natural variations are outlined in Section 3. Over the time scales of interest here, anthropogenically-induced changes in greenhouse-gas concentrations, principally are also considered important (Goodess et al, 1992a). Thus, the development of scenarios of anthropogenic changes in atmospheric concentrations over the next 150 ka is described in Section 4. The LLN-2D model generates output (temperature and ice volumes) for seven surface types (see Section 2) in 5° latitude bands. In order to obtain results applicable to a particular region, here central England, a method for downscaling model output is required. A rule-based approach to downscaling, in which the temperature time series from the model are divided into a succession of discrete climate states, is proposed in Section 5. Eight LLN-2D simulations using both natural and anthropogenic forcing, together with insolation forcing, have been completed. The results from these simulations are summarized in Section 6, and compared with those from a simulation using only natural and insolation forcing. Finally, the potential applications and limitations of these modelling studies are discussed in Section 7.

2.

The LLN-2D Model

To achieve simulations over the time scales of interest here, any model must be simplified in comparison to the fully three-dimensional models used for climate change predictions on time scales of a few hundred years. However, for the regional level of detail required, it is also essential that the model has greater resolution than the simple one-dimensional models used for energy balance calculations. LLN-2D is described as 2.5 dimensional because of the altitudinal, latitudinal and sectoral differentiation employed for the surface models (for the Northern Hemisphere only). To be more precise, the model has components for land (snow covered and snow free), sea (open and sea-ice covered), and the North European, North American and Greenland ice

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sheets. The ice sheet sub-model uses a vertically integrated equation for ice mass balance which assumes plastic flow (Gallée et al., 1991). There are parameterizations to account for lateral calving from ice sheets, snow ageing processes, albedo effects of taiga/tundra line shifts and terms for extension of sea ice into relatively warm waters. The sub-models in LLN-2D encompass many important interactions of the Earthatmosphere-cryosphere systems. The atmospheric model has a time stepping interval of three days for each month, synchronously passing the mid-monthly climatologies to the oceanic and land submodels. This process is repeated at 1000 year intervals, at which point information is fed to the ice sheet model. LLN-2D is particularly powerful in the way that it calculates explicit snow mass balances, as opposed to the snow line parameterizations employed in many other models. It is possible to use values from ocean cores to give a measure of global ice volume variation over time scales of the order of to years (Duplessy et al., 1988). LLN2D has been shown to reproduce well the low frequency variations in ice volume throughout the Late Quaternary and Holocene (125 ka Before Present to present, i.e. the last glacial-interglacial cycle) when assumptions are made about Antarctic ice volume variations over that period (Gallée et al., 1992). Furthermore, the model is known to reproduce present-day mid-latitude temperatures particularly well (Burgess, 1998). The strengths of LLN-2D derive from the acceptable computational power required for integration, representation of fast and slow feedback mechanisms and coupling of the major elements of the Earth-atmosphere-cryosphere systems. Weaknesses of the model include relatively crude spatial and temporal scales of resolution, a parameterized hydrological scheme based on observed present-day relationships, and a single basin ocean that cannot reproduce the thermohaline circulation.

3.

Long-term Natural Changes in Climate Forcing

The aim of this research has been to derive estimates of millennial scale climate change over the next 150 ka. In order to do this, scenarios of future climatic forcing are required. As discussed above, LLN-2D is forced by changes in seasonal, latitudinal and long-term variations in the insolation received at the top of the Earth’s atmosphere, and by changes in atmospheric concentrations. The first of these, insolation changes, are calculated during model integration (Berger, 1978; Berger and Loutre, 1997), leaving only the need for the specification of the atmospheric forcing. During the past 125 ka there have been significant variations in atmospheric as recorded in the gas bubbles trapped in ice sheets (Barnola et al., 1987; Jouzel et al., 1993). These gas inclusions provide snapshots of past atmospheres. A complex web of cause and effect in the carbon cycle is responsible for variations in atmospheric concentrations, as recorded in these bubbles. The causes of these variations involve

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interactions between the atmosphere, the terrestrial and oceanic biospheres, shelf and ocean floor sediments and the geosphere. They occur on all time scales from seasonal cycles to millions of years, the nature of their impact being dictated by such factors as temperature, nutrient availability, sea level, gas exchange at the ocean surface and various feedbacks (Knox and McElroy, 1985; Sundquist, 1990; Kier, 1993; Francois et al., 1997; see also reviews in Adcock et al., 1997; Goodess et al., 1999). Ideally, carbon cycle modelling would provide the necessary atmospheric concentrations required by LLN-2D as a forcing. However, although individual components of this cycle are understood in some detail, the present state of carbon cycle modelling is not able to capture fully the complexity of the interactions over the full range of the time scales involved. In the study presented here, a statistical approach to estimation of future concentrations is taken. The basis for this methodology lies in the observation that, during the Late Quaternary and Holocene, atmospheric concentrations varied from high values of around 280 ppmv during interglacial periods to only 190 ppmv during glacial periods. These high and low periods coincided with times at which Northern Hemisphere high latitude summer insolation receipts were, respectively, high (interglacial periods) and low (glacial periods). It should therefore be possible to link atmospheric concentration changes to spatio-temporal changes in insolation receipts. A purely statistical methodology such as this does not address the causal linkages between insolation and change. It assumes that sufficient information about those linkages can be captured in a regression analysis that relates concentrations to a small number of variables characterising insolation. Insolation time series at different latitudes and months throughout the past 125 ka (independent variables) were regressed using stepwise multiple regression against the time series (dependent variable) obtained from ice cores over the same period. After development and validation, the resultant equation for prediction of atmospheric concentrations over millennial time scales from insolation was found to contain strong signals from high latitude areas in winter months. This is in line with certain theories regarding the carbon cycle, which state that high latitude oceanic productivity is a strong control on oceanic draw-down and storage of atmospheric (Knox and McElroy, 1985). Validation of the predicted time series showed that the main shortcomings were during the warming since the end of the Last Glacial Maximum, with predicted values consistently below observed values. This period is interesting in that it represents a period of rapid environmental change typical of the end of a glacial period and very different from the slow englaciation over tens of thousands of years that preceded it (and which occupies most of the time frame over which the regression was developed). Factors such as rapid inundation of continental shelves and burial of carbon have previously been suggested as important factors in the dynamics of the carbon cycle during such periods of rapid change.

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Bearing in mind the known shortcomings of the regression equation, it was initialised using insolation time series over the next 150 ka to derive a time series of atmospheric change under natural conditions. Figure 1 shows predicted and observed (Jouzel et al., 1993) atmospheric concentrations over the past 125 and predicted values over the next 150 ka. The predicted future forcing was used to force the LLN-2D integration NAT (i.e. natural changes in atmospheric concentrations only).

4.

Derivation of Long-term Anthropogenic Forcing

The next step was to derive scenarios of anthropogenic changes in atmospheric over the next 150 ka. Relatively little modelling has been performed on anthropogenic effects over long time scales, two notable exceptions being the studies by Walker and Kasting (1992) and by Sundquist (1990). Results from the latter study were used here because they include terms to account for the various calcium reactions leading to precipitation or dissolution of carbonates in sea-floor sediments. First, two scenarios of the change in emissions were selected from Sundquist’s work, one representing high rates of fossil fuel burning, the other low rates. From these scenarios, time series of anthropogenic

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atmospheric concentrations up to ~10 ka AP (after present) were derived. Then, from 10 ka AP until the end of the simulation period (150 ka AP), four different scenarios of the decline in atmospheric concentrations were defined. These are bounding and intermediate cases thought to reasonably represent the rate at which the natural carbon cycle can sequester anthropogenic These scenarios specify total absorption of anthropogenic atmospheric at 30, 50, 100 and 150 ka AP. In this way, a total of eight anthropogenic time series were defined, (2 emissions scenarios x 4 absorption scenarios). These were added to the natural change time series described in Section 3 to define scenarios for the future (as shown in Figure 2). ANTH1, ANTH3, ANTH5 and ANTH7 represent low initial emissions scenarios, ANTH2, ANTH4, ANTH6 and ANTH8 represent high initial emissions scenarios, with the different rates of decay outlined above.

INVESTIGATIONS INTO LONG-TERM FUTURE CLIMATE CHANGE 5.

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Downscaling the LLN-2D Model Results

Using the output from three-dimensional global climate models, there are two main techniques for deriving detailed sub-grid scale climatologies. The first involves the use of a high resolution Regional Climate Model nested within the coarse resolution global model, and driven using the global model grid-scale climatologies as boundary conditions. The second technique uses statistically-derived relationships between grid scale and local climatologies. The zonally-averaged nature of the LLN-2D atmospheric model means that neither of these techniques are appropriate. A rule-based approach to downscaling was adopted in the studies presented here (Burgess, 1998). In order to derive rules for downscaling LLN-2D zonal output to central England, relationships between climatic-cryospheric output from LLN-2D and observed indices of climate were established over the last glacial-interglacial cycle. The climate indices were derived from various floral, faunal, limnological and pedological palaeo-evidence from around England (Burgess, 1998). By following this procedure, thresholds can be defined at which climate switches from e.g. boreal to temperate conditions. One example of the rules thus established is that if the February zonal land temperature in the latitude band 55-60°N is greater than -20.5°C then a temperate state is assigned to England. The full hierarchy of rules is outlined in Table 1. In Figure 3, we show the time series of climate states for England over the last 125 ka based on observed palaeo-evidence, and the time series derived by applying the rules.

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Once climate states have been assigned for England, analogue stations for each state can be selected from around the world (see Table 2 for the stations used). The meteorological observations from these sites were used to describe the monthly temperature and precipitation regimes for each state (Burgess, 1998). Means and standard deviations across a number of analogue stations for any climatic state give an idea of the possible range of variability that could be expected under such conditions.

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6.

Results

The climate change modelling studies presented here, based on scenarios of both natural and anthropogenic atmospheric concentrations, estimate the broad patterns of future climate change over central England up to 150 ka in the future. In our discussion of the results, we concentrate on four of the eight concentrations scenarios, as follows: NAT: natural concentrations, as predicted by the regression equation described in Section 3; ANTH1: low fossil fuel-based emissions, anthropogenic atmospheric concentrations of tailing off to zero at 30 ka AP; ANTH7: low fossil fuel-based emissions, anthropogenic atmospheric concentrations of tailing off to zero at 150 ka AP; and, ANTH8: high fossil fuel-based emissions, anthropogenic atmospheric concentrations of tailing off to zero at 150 ka AP.

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Figure 4 shows the results from the LLN-2D model (hemispheric ice volumes and January and July temperatures) generated by these four scenarios. Figure 5 shows the associated succession of climate states over central England. In all but the most extreme emissions scenario (ANTH8), there is a deterioration in climate leading to a maximum in hemispheric ice volume at between 55-60 ka AP. This glacial maximum is restricted in its severity in comparison to the Last Glacial Maximum. In England it is represented by a tundra environment, and only restricted upland ice formation.

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The exact pattern of climate change over central England associated with the general cooling toward this glacial period is strongly dependent on the scenario used. For the natural case, NAT, the present temperate climate is succeeded after 1000 years by boreal conditions that become steadily colder and dryer (i.e. changing from maritime to continental boreal conditions) towards about 45 ka AP. Thereafter tundra conditions persist for around 20 ka. The lower emissions ANTH scenarios, ANTH1 and ANTH7, show a period of enhanced warming until about 5 ka AP, then return to temperate conditions for a further 10 ka. Warmer boreal conditions are maintained well beyond 50 ka AP, followed by a period of rapid cooling coinciding with ice growth and culminating for ANTH1 in a brief glacial event at roughly 60 ka AP. The short lived nature of this glacial event implies that it is restricted to highland locations such as in the Lake District. Beyond the restricted glacial event at 65 ka AP in ANTH1, there is increasing harmony between different model integrations, regardless of the scenario chosen. A return to boreal conditions around 75 ka AP is followed by gradual and sustained decline to tundra and then glacial climate by 100 ka AP.

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ANTH8, the higher initial emissions scenario, shows a very different evolution of climate. The Northern Hemisphere remains largely ice free until nearly 100 ka AP, after which there is a short period of very rapid cooling. Over central England, this is reflected in enhanced warming conditions which persist until 20 ka AP, a complete absence of the glacial event seen in ANTH1 at 65 ka AP, and only a brief onset of glacial conditions at around 105 ka AP. Similarities with other, less severe, ANTH scenarios do not really develop until after 120 ka AP. It is clear from this research that anthropogenic effects have the potential to disturb the climate system over very long time scales. A somewhat surprising result is the presence of a short glacial event over central England in ANTH1 at 65 ka AP, which is not present in NAT. In ANTH1, complete Northern Hemisphere deglaciation is observed at around 5 ka AP. However, following this state of deglaciation, ice volumes in ANTH1 develop so that by 65 ka AP they are greater than in NAT (even though NAT never demonstrates complete deglaciation). The greater ice volume in ANTH1 is attributable to rapid growth of the Fennoscandian and Laurentide ice sheets starting at around 50 ka AP (growth of the Greenland ice sheet is limited due to its geographical setting). This result is interesting because it suggests that ice growth in a 'post-anthropogenic' world could be greater than that which would result from a purely natural evolution of climate. In order to discover the reasons for the additional ice growth in ANTH1, a detailed study of model behaviour was carried out, focusing on the relevant time periods. In particular the response of bedrock to ice loading and unloading in LLN-2D was investigated because it was considered that bedrock 'memory' effects might contribute to the anomalous behaviour of ANTH1 ice volumes. Actual model bedrock deflections in response to ice loading (expressed as metres below sea level) were compared to the equilibrium bedrock deflection depths that would be expected for the ice volume present (i.e. that depth which would occur if the bedrock had sufficient time to respond). The ratio of the two gives a measure of the continuing, time lagged, response of the bedrock to present and previous ice loading conditions. At 65 ka AP, the forcing conditions favour the decay of ice sheets. In NAT, full bedrock depression has been reached and strong rebound takes place, with associated high ice decay rates (due to lateral calving and desert altitude effects). In ANTH1, however, bedrock depression at 65 ka AP is still in disequilibrium with the ice loading and this means weaker isostatic rebound and hence slower rates of ice melt. A larger ice volume develops in ANTH1 when insolation conditions favour ice growth, because ablation is slower and hence ice build up proceeds at an accelerated rate.

7.

Discussion

LLN-2D simulations based on natural changes, including the NAT simulation described here, show the onset of the next glaciation at about 55-60 ka AP. The addition

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of anthropogenic effects (simulations ANTH1-ANTH8) can, in the more extreme scenarios, prevent this glaciation occurring, so that the date of the next glaciation is delayed until around 110 ka AP. Anthropogenic forcing also has a major impact at the start of the simulations, causing the Greenland ice sheet to disappear by 2 ka AP. It does not reappear until just before 20 ka AP at the earliest. Future ice-sheet development is sensitive to both the maximum concentration and the decay rate of anthropogenic and there are major uncertainties in attempting to construct emissions scenarios over such long future time periods. Nonetheless, the group of ANTH1-ANTH8 simulations demonstrate that anthropogenic impacts may severely delay the onset, and restrict the extent of, the next glaciation. They also provide an appropriate envelope of possible futures that can be considered for radiological assessment purposes (Goodess et al., 1999). In order to use LLN-2D output for radiological assessment studies, it is convenient to characterise the model time series as a sequence of discrete climate states (temperate, boreal, tundra, glacial, subtropical or enhanced warming). The climatic conditions likely to be experienced in the British Isles during each of these states can then be represented using instrumental data from selected analogue sites (although there are probably no good analogues at the present day for the British Isles during a glacial state). The analogue sites used for central England are listed in Table 2. These two steps form the basis of the rule-based downscaling technique which has been devised to provide estimates of long-term future climates for central England based on output from the NAT and ANTH1-ANTH8 simulations. Severe limitations are imposed by the model configuration – for example, because the model is only two-dimensional it is necessary to assume in the downscaling that the climates of the British Isles will move in the same sense, and with broadly the same intensity, as the zonal climate. However, these limitations cannot be overcome until such time as computing power and speed allows the application of three-dimensional climate models to these problems of very long-term climate change. The use of three-dimensional models might also allow a wider range of climate forcing mechanisms to be explicitly modelled. The LLN-2D simulations are forced by insolation and only, whereas climate varies on all time scales in response to a whole suite of random and periodic forcing factors (Goodess et al., 1992b). Further limitations arise from the model’s necessarily simplified representation of the Earth-atmosphere-cryosphere system (Burgess, 1998). For example, the single ocean basin cannot reproduce important effects such as the thermohaline circulation, precluding the modelling of events such as the Younger Dryas. At the present time, however, the downscaled output from the LLN-2D simulations provides a valuable source of information on possible climate change over very long time scales. The simulations demonstrate clearly the non-linearities that can operate between cause and effect, such that an anthropogenic-emissions scenario may at certain times exhibit greater ice sheet development than a natural-emissions scenario. They also show that the effects of anthropogenic emissions on climate can persist long after those emissions have reduced to zero, a feature depending on the length of the ‘decay’ period

INVESTIGATIONS INTO LONG-TERM FUTURE CLIMATE CHANGE 245 selected to return atmospheric concentrations of back to their natural levels. Until such time as three-dimensional climate models can be used to simulate change over the very long time scales under investigation here, the results from the simpler twodimensional models usefully inform our view of how future climates may develop.

Acknowledgements This work was funded by United Kingdom Nirex Limited as part of the Nirex Safety Assessment Research Programme and by British Nuclear Fuels plc. The authors wish to thank A. Berger and M-F. Loutre of the Université Catholique de Louvain, Belgium for making the LLN-2D palaeoclimate model available for use by Paul Burgess in his PhD studies.

References Adcock, S.T., Dukes, M.D.G., Goodess, C.M. and Palutikof, J.P. (1997) A Critical Review of the Climate Literature Relevant to the Deep Disposal of Radioactive Waste, Nirex Science Report S/97/009, United Kingdom Nirex Limited, Harwell. Barnola, J.M., Raynaud, D., Korotkevich, Y.S. and Lorius, L. (1987) Vostok core provides 160,000-year record of atmospheric Nature, 329, 408-414. Berger, A. (1978) Long-term variations of daily insolation and Quaternary climate change, J. Atmos. Sci., 35, 2362-2367. Berger, A. and Loutre, M-F. (1997) Long-term variations in insolation and their effects on climate, the LLN experiments, Surv. Geophys., 18, 147-161. Burgess, P.E. (1998) Future Climatic and Cryospheric Change on Millennial Timescales: An Assessment Using Two-dimensional Climate Modelling Studies, PhD Thesis, University of East Anglia, Norwich. Duplessey, J., Labeyrie, L. and Blanc, P. (1988) Norwegian Sea deep water variations over the last climatic cycle: Palaeooceanographic implications, in H. Wanner and U. Seigenthaler (eds.), Long and Short Term Variability of Climate, Earth Science Series, Springer Verlag, New York, pp. 83-116. Francois, R., Altabet, M.A., Yu, E.F., Sigman, D.M., Bacon, M.P., Frank, M., Bohrmann, G., Bareille, G. and Labeyrie, L.D. (1997) Contribution of Southern Ocean surface-water stratification to low atmospheric concentrations during the last glacial period, Nature, 389, 929-935. Gallée, H., van Ypersele, J.P., Fichefet, T., Tricot, C. and Berger, A. (1991) Simulation of the last glacial cycle by a coupled, sectorially averaged climate-ice sheet model. 1 The climate model, J. Geophys. Res., 96, 13139-13161. Gallée, H., van Ypersele, J.P., Fichefet, T., Marsiat, I., Tricot, C. and Berger, A. (1992) Simulation of the last glacial cycle by a coupled, sectorially averaged climate-ice sheet model. 2. Response to insolation and variations, J. Geophys. Res., 97, 15713-15740. Goodess, C.M., Palutikof, J.P. and Davies, T.D. (1992a) Studies of Climatic Effects Relevant to Deep Underground Disposal of Radioactive Waste, Nirex Report NSS/R267, United Kingdom Nirex Limited, Harwell. Goodess, C.M., Palutikof, J.P. and Davies, T.D. (1992b) The Nature and Causes of Climate Change. Assessing the Long Term Future, Studies in Climatology Series. Belhaven Press, London. Goodess, C.M., Watkins, S.J., Burgess, P.E. and Palutikof, J.P. (1999) Assessing the Long-term Future Climate of the British Isles in Relation to the Deep Underground Disposal of Radioactive Waste, Nirex Report, United Kingdom Nirex Limited, Harwell, in press. Jouzel, J., Barkov, N.I., Barnola, J.M., Bender, M., Chappellaz, J., Genthon, C., Kotlyakov, V.M., Lipenkov, V., Lorius, C., Petit, J.R. el al. (1993) Extending the Vostok ice-core record of palaeoclimate to the penultimate glacial period, Nature, 364, 407-412.

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Kier, R. (1993) Are atmospheric content and Pleistocene climate connected by wind speed over a polar Mediterranean sea?, Glob. Planet. Change, 8, 59-68. Knox, F. and McElroy, M. (1985) Changes in atmospheric Influence of marine biota at high latitudes, J. Geophys. Res., 89, 4629-4637. Nirex (1997) Nirex 97: An Assessment of the Post-closure Performance of a Deep Waste Repository at Sellafield, Nirex Science Report S/97/012, United Kingdom Nirex Limited, Harwell. Palutikof, J.P., Goodess, C.M., Watkins, S.J. and Burgess, P.E. (1999). Developments in long-term climate change, Prog. Env. Sci., 1, 89-96. Sundquist, E.T. (1990) Long-term aspect of future atmospheric and sea-level changes, in R. Revelle (ed.), Sea Level Change, National Research Council Studies in Geophysics, National Academy Press. Washington, D.C., pp. 193-207. Thorne, M.C. (ed.) (1995) Nirex Biosphere Research: Report on Current Status in 1994, Nirex Science Report S/95/003, United Kingdom Nirex Limited, Harwell. Walker, J.C.G. and Kasting, J.F. (1992) Effects of fuel and forest conservation on future levels of atmospheric carbon dioxide, Palaeogeog., Palaeoclim., Palaoecol, 97, 151-189.

GEOMORPHOLOGICAL AND CLIMATOLOGICAL PERSPECTIVES ON LAND SURFACE –CLIMATE CHANGE DOMINIC KNIVETON and SUE McLAREN Department of Geography, University Of Leicester, Leicester LE1 7RH

1.

Introduction

The climate system can be considered to consist of the atmosphere, oceans, biosphere, cryosphere (ice and snow) and lithosphere. In this book we have focused on the interactions between the first and last of these, the atmosphere and its long-term expression, the climate, and the lithosphere or more specifically, the land surface. There are numerous and often highly complex linkages between climate and land surfaces which are still relatively poorly understood. As well as trying to understand the complex processes that operate today, there are the added difficulties in trying to understand how these processes may have operated and changed in the past as well as trying to determine how they may change in the future. Improvements in our understanding of how climates have changed in the past using palaeoenvironmental reconstructions of both land-based and ocean-based records are still needed before it can be determined what causes the climate to alter. Only when the causes of past climates are understood will it be possible to fully anticipate or forecast climatic variations in the future (Bradley and Eddy, 1991); and to then identify whether it will be possible to find out whether climate predictions can be used to predict future geomorphic, sedimentary and surface changes? The cross disciplinary nature of the study of the interactions between the atmosphere and land surface can be exemplified by the contemporary debate about whether tectonics have been driven by or drive long term climate change (Whipple et al., 1999, Molnar and England 1990, Brozovic et al., 1997). The question of whether Quaternary climate change could have produced significantly increased topographic relief (as a result of increased weathering and associated isostatic uplift) resulting in accelerated rates of tectonism, can only be resolved by the contribution of climatologists defining what Quaternary climate change is as well as geomorphologists deciding whether these 247

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changes could significantly increase the volume of ‘missing mass’ between summits and ridges and finding evidence to this effect. In this chapter we highlight some of the different approaches that have been undertaken to address the many issues in land surface-climate change. In particular we examine how multidisciplinary approaches over different timescales (from to years) and from the scale of local catchment studies to global processes can improve our understanding of the complex interactions between the surface and atmosphere. 2.

Large scale research programmes on climate change

There are a number of reasons why scientists are interested in land surface and climate interactions as well as change. As key components of the larger climate system, the atmosphere and land surface influence and are influenced by each other. The role of the land surface in relation to the atmosphere ranges from the purely physical, including inducing aerodynamic drag on the atmosphere; biological, through the stomata response to environmental changes and the control of energy and mass fluxes (Pitman et al., 1999); to geochemical, through for example, evaporation of water and dust inputs. While the role of the atmosphere on the land surface can also be defined in terms of a physical interaction, through erosion and deposition; biological, through the control of growing conditions for vegetation; and geochemical, with the addition of precipitation onto and into sediments and soils as well as the deposition of dust. Through these interactions the land surface can also provide evidence of past climates (Huang et al., 2000) thus helping to determine the relative magnitudes of human induced climate change and natural climate variability. In the technical summary of Climate Change 1995: The Science of Climate Change the Intergovernmental Panel of Climate Change identified the following scientific problems as requiring the most urgent attention: (i)

the rate and magnitude of climate change and sea level rise: the factors controlling the distribution of clouds and their radiative characteristics; the hydrological cycle, including precipitation, evaporation and runoff; the distribution and time evolution of ozone and aerosols and their radiative characteristics; the response of terrestrial systems to climate change and their positive and negative feedbacks; the response of ice sheets and glaciers to climate; the influence of human activities on emissions; the coupling between the atmosphere and ocean, and ocean circulation;

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the factors controlling the atmospheric concentrations of carbon dioxide and other greenhouse gases;

(ii)

the detection and attribution of climate change: systematic observations of key variables, and development of model diagnostics relating to climate change; relevant proxy data to construct and test palaeoclimatic time series to describe natural variability of the climate system;

(iii)

regional patterns of climate change: land-surface processes and their link to atmospheric processes; coupling of scales between global climate models and regional and smaller scale models; simulations with higher resolution climate models.’ (IPCC WGI, 1995, p47)

As can be seen from this list the investigation of the relationship between the land surface and atmosphere is common to all three research problems. The IPCC suggests that to resolve these issues requires systematic and sustained global observations of relevant parameters and allied research by individual investigators in a variety of institutions as well as by co-ordinated international efforts. One of the key players organising large-scale terrestrial land surface experiments is the Global Energy and Water Cycle Experiment (GEWEX). Initiated by the World Climate Research Programme (WCRP) in 1988, the remit of the project is to observe and model the hydrologic cycle and energy fluxes in the atmosphere, at the land surface, and in the upper oceans. While each project or experiment of GEWEX has its own objectives, the overall aim of this co-coordinated programme of research is to improve prediction of global and regional climate change. Of particular interest with respect to land surfaceatmosphere interactions are the Continental-Scale Experiments (CSEs). These experiments aim to provide improved observations and coupled land-atmosphere models by studying specific hydrological regions. An example of one of these experiments is the Continental-Scale International Project (GCIP), based in the Mississippi River basin, which was designed to improve climate models by bridging the gap between small scales appropriate for modelling discrete processes over land and large scales practical for modelling the global climate system. In this book we focus less on internationally co-ordinated projects, already well documented elsewhere (e.g. Pitman et al., 1999, Gash and Kabat 1999), rather choosing to concentrate on the research activities of smaller groups of geomorphologists and climatologists, who are also contributing to the further understanding of interactions of the atmosphere and land surface. Like the larger projects however these studies reveal a variety of approaches and a number of overriding issues that need to be addressed in order to improve the understanding of the linkages between surface and atmosphere.

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Linking Climate Change with Land Surface Change

In the preface a number of issues were raised that have been covered to some degree by the selection of papers presented in this book. These will now be summarised. 3.1

SPATIAL AND TEMPORAL SCALES OF GEOMORPHOLOGICAL CHANGE AND CLIMATIC VARIABILITY

Both the atmosphere and land surface (as well as the processes that operate within and on) vary spatially from micro to global scales, and temporally over seconds to centuries and even millennia. Figure 1, from Pitman et al., 1999, details the spatial and time scales of observational and modelling evidence providing support for the influence of the land surface on weather and climate. Interestingly, given the intended aim of a number of the larger international land-atmosphere experiments to improve the modelling and prediction of the land-climate system there is relatively little overlap in terms of spatial and temporal scale between observational evidence and modelling evidence. In brief, Pitman et al., (1999) attribute this to measurements being generally on far too short a time scale to determine the role of the surface on the climate and modelling tending to focus on large perturbations. In terms of the studies presented here a key issue concerns linking micro, local and even catchment scale processes to larger scale global phenomena both in terms of the role of the atmosphere on the surface and vice versa. In particular the upscaling of processes is a central theme of many of the palaeoclimate studies. The issue of moving from one scale to another is also important in the other direction, when moving from large scale to smaller scales. This is particularly important when one considers the spatial scale of climate model output, which is typically at the resolution of 2-3° latitude/longitude grid squares. Examples of downscaling approaches to the atmosphere-surface interaction are illustrated by Wilby and Dettinger, Schmidt and Dehn, and Burgess et al.

In Figure 2 we show the spatial and temporal context of the studies outlined in this book. Spatially research has varied from small regional scale to global scale observations. Catto and Bachhuber, for example, look at the development and change in dune type as well as the periodic development of thin soils in the Estancia valley as a result of changes in climate. In contrast Brooks and Legrand have observed dust variability over the whole of North Africa. Modelling climate change has allowed scientists to increase the spatial area that they are studying to the size of a hemisphere or even globally (e.g. Burgess et al.,). In terms of timescales the approaches vary from short periods (e.g. Viles and Goudie) through to changes over periods spanning the last glacial and into the Holocene (e.g. Bachhuber and Catto and White et al.,).

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(1: Gash & Nobre, 1997; 2: Andre et al, 1989; 3: Pielke & Avissar, 1990; 4: Pielke et al, 1991; 5: Avissar & Chen, 1993; 6: Vidale et al, 1997; 7: Eastman et al, 1998; 8: Smith et al, 1992; 9: Bonan et al, 1992; 10: Polcher & laval, 1994; 11: Lean & Rowntree, 1997; 12: Henderson-Sellars et al, 1993; 14: Nicholson et al, 1998; 15: Chase et al 1996; 16: Zhao et al, 1999a; 17: Zhao et al, 1999b; 18:Eastman et al, 1999; 19: Lu, 1999; 20: Otterman et al, 1984; 21: Harvey, 1988; 22: Harvey, 1989; 23: Texier et al, 1997; 24: Claussen, 1997; 25: Claussen et al, 1999).

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3.2

EFFECTS OF CHANGING CLIMATE ON GEOMORPHOLOGICAL AND SEDIMENTOLOGICAL PROCESSES

Landforms and sediment accumulations represent the effects of various geomorphological processes that can be interpreted in the geological record from key characteristics such as texture, geochemistry, sedimentary structure, morphology and the combination of certain landforms within a region. Climatic factors are known to have an influence on the nature and the rate of operation of geomorphic processes and sedimentation. The rapid changes in climate resulted in the shifting/migration of climatic boundaries throughout the Quaternary. Geomorphological processes changed, reflecting the degree and type of climatic change. Preserved on many land surfaces today are the relicts landforms that formed under past different climatic conditions. Through the study of such remnant landforms and sediments researchers try to piece together the information preserved, to interpret past conditions and to attempt to date when they occurred. Bachhuber and Catto, for example, identified within the palaeo-record changes in conditions in the Estancia Valley in central New Mexico, from dry playa through to a fresh water lake at various times in the past 60,000 years. Bachhuber and Catto then interpreted the changes in the lake levels in terms of climatic change. White et al.,., have studied various geochemical and organic sediments to aid in the interpretation of

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higher water tables and existence of palaeolakes in the Fezzan area of southern Libya during the late Quaternary.

Shorter term variations in the land surface due to changes in the climate are shown by the work of Yair and Bryan in their assessment of the alteration of surface properties connected with the atmospheric deposition of loess or sand in subtropical semi-arid and arid areas. Their study concludes that arid and semi-arid environments, although highly adapted to extreme variations in rainfall, may be extremely sensitive to slight changes in their surface properties, which alter their hydrological regime quickly and efficiently. The influence of the atmosphere on the land surface through short-term hydrological and weathering processes is illustrated by the work of Wilby and Dettinger and Viles and Goudie, respectively. Geomorphological responses to environmental change are understood at a range of different spatial and temporal scales. Observations and model simulations can be scaled up e.g. palaeo reconstructions or scaled down from global to regional and local scale. Different approaches are useful as they allow different types of data to be integrated. Downscaling has been much used by climatologists but to date has largely been ignored by geomorphologists. However by simulating past climates using general circulation models and downscaling the results to smaller scales it might both be possible to understand the linkages between the land surface and atmosphere and interpret the land surface observations. 3.3

EFFECTS OF SURFACE ON CLIMATE

Examples of observational and modelling evidence that provide support for the role of the land surface on the weather and climate at varying time and space scales were shown in Figure 1. It is quite apparent from this body of literature that the land surface primarily affects the atmosphere through the portioning of water and energy, particularly at short time scales. What however is less clear is the question about the role of humans altering the surface and thus influencing the climate and weather (Pitman et al., 1999). In Chapter 5, Adegoke and Carleton look at the meso-scale effects of human modification of terrestrial vegetation and its affect on climate through albedo and the soil heat flux, using a combination of remotely sensed data and fieldwork. They conclude that the effects of modification are detectable within the climatic record of the last 100 years. Remotely sensed data and geographical information systems (GIS) allow the observation and analysis of surface-atmosphere processes over large, near globe coverages. Unlike point measurements remotely sensed data is often continuous in space and together with GIS provide ideal tools for assessing issues of heterogeneity and scaling characteristics of surface-climate interactions (Carleton 1999). The use of geostatistical methods to explore issues of spatial variability in rainfall is shown by Agnew and Chappell. Their study examines the issue of whether the Sahel has indeed dried out over the last fifty years or whether the variations in rainfall are due to changes in the raingauge network.

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While the use of remote sensing data has greatly enhanced the understanding of the Earth’s surface and atmosphere over large scales, there is a gap in information on the surface atmosphere interaction over the longer term. Satellite based data has only been available over the last twenty to thirty years. Longer term measurements are needed. Improved simulations of the climate at higher resolution will help develop the understanding of the surface-atmosphere interaction. Coupled with the improvement of downscaling and upscaling techniques future research will be able to ascertain the extent, magnitude and direction of the climate–land surface coupling. With the distribution of three dimensional general circulation models (GCM) by such groups as the UK Global Atmospheric Modelling Programme, geomorphologists and climatologists will be able to alter the characteristics of the land surface to assess changes in the climate; attempt to simulate changes in the climate which might result in geomorphological change and even help understand the sometimes contradicting evidence of palaeo records (e.g. Dong et al., 1996, Price et al., 1996, Toumi et al., 1996, Thorpe et al., 1997, Price et al., 1998). 3.4

PREDICTION AND MANAGEMENT OF LAND SURFACE CHANGES AS A RESULT OF FUTURE CLIMATIC CHANGES.

Climate models allow the simulation of future climate over extended time scales. The impact of future climate changes on the surface is the subject of the work of Schmidt and Dehn, Wilby and Dettinger and Burgess et al., Schmidt and Dehn looked at the response of landslides to varying climatic conditions in New Zealand and Italy. Their results show that in Italy there is a significant decrease in activity due to a temperature increase in winter and reduced snow storage. However, they could not identify any clear trend in rainfall. In New Zealand there were fewer events of high landslide probability in winter due to decreases in rainfall. If changes in climate can be predicted then management strategies designed to cope with landslide activity versus stability can be developed. The management of land surfaces and the features on and below them, under changing conditions has been recognised as being important by both Burgess et al.,; and Wilby and Dettinger. It is clearly important that prediction and management of landform environments used for purposes such as waste disposal is possible. Burgess et al.,., look at the long term future development of landforms which need to be understood for safety assessments of underground radioactive waste disposal in the United Kingdom. In particular they suggest that under enhanced greenhouse gas scenarios glaciations will not occur until 110 Ka AP. While control simulations indicate that the next glaciation will be 55-60 Ka AP. The model reveals that the Greenland ice sheet is sensitive to both maximum concentration and decay rate of anthropogenic Limitations to this approach are that the model is only two-dimensional and grossly simplifies the Earthatmosphere-cryosphere system. In all three studies mentioned in this section downscaling of the climate model output is required to allow studies on a regional scale to be conducted. Wilby and Dettinger use a regression-based downscaling model for

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streamflow simulation under current and future climate scenarios in south west USA. Their model demonstrated a reasonable skill at reproducing observed area-average precipitation, temperature and streamflow. The downscaled data resulted in slight underestimates of streamflow as a result of underestimates of spring rainfall and overestimates of temperature in winter. As stated at the beginning of this chapter successful prediction of future climate and land surface change requires an understanding of past ‘natural’ changes. A number of studies within the book have looked at elements of palaeoreconstructions. For instance Gurney shows that surface features such as pingos can be used to provide accurate indicators of past thermal conditions. While Bachhuber and Catto provide geologic evidence of rapid, multiple and high magnitude climate change during the last glacial (Wisconsinan) of North America. Considerable effort has been applied to piece together indirect or proxy evidence of longer term global and hemispheric climate change (Briffa et al., 1998, Mann et al., 1998, Huang et al., 2000). Inevitably the more evidence that is accumulated of palaeoclimate changes the greater degree of accuracy can be attached to these reconstructions. 4.

Conclusion

It is apparent from both the research outlined in this book and other studies on the subject that the primary difficulty faced in attempting to improve our understanding of climate and land surface change is reconciling the vast range in the spatial and temporal scales of processes operating between the atmosphere and surface with the observations and tools available to study them. Inevitably in a dynamic environment when trying to study past climate and surfaces we are limited in the availability of observations at the global scale over extended time periods. Remotely sensed data now provide near global coverage of the surface and atmosphere. While modern geostatistical methods and systems provide the tools with which to analyse these data at varying scales. Yet by definition remotely sensed measurements are removed from the object of observation thus restricting the detail of analysis. In addition they are of limited record length stretching back a mere 20-30 years, at best. Models of the climate and surface potentially offer a powerful tool to understand the strength, direction and extent of response of the surface and atmosphere to pertubations in their and other climate system components states. Yet again models can only provide part of the solution to questions of climate and land surface change. Partially this is because of the level of understanding or rather lack of it of the underlying processes involved and of the scales over which these processes operate. There are also restraints to the use of models because of the way in which processes are represented within the model. Computing power restraints and an incomplete knowledge of processes mean that climate models run at coarse resolution. The most advanced climate models operate at approximately 2 to 3°latitude/longitude resolution. This means that many processes between the surface and atmosphere are parametrised or simplified. While the output

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from global climate models are also of coarse resolution impacting on their use to interpret the effect of future or past simulated climate changes on the surface. The development of downscaling and upscaling techniques has attempted to bridge the gap between the scale over which models operate and over which observations are made. Indeed it could be said that it is the movement up and down scale and development of methods to do this that provides the interface between climatological and geomorphological disciplines. In the future it can be seen that the resolution of climate models will be improved. With a greater understanding of processes and improved model resolution it can be expected that a more detailed and realistic representation of surface and atmosphere processes will be incorporated into climate models leading eventually to a fully coupled three dimensional land surface-atmosphere model. It can also be envisaged that longer data sets will be available from remote sensing and field observations, while improved geostatistical methods will be developed to integrate spares and non-continuous data sets. References Adegoke, J.O and Carleton, A. M., (2000) Warm season land surface-climate interactions in the United States Midwest from meso-scale observations, in McLaren, S and Kniveton, D. (eds), Linking Climate Change to Land Surface Change, Kluwer Academic Press, Agnew, C. and Chappell, A., (2000) Desiccation in the Sahel, in McLaren, S and Kniveton, D. (eds), Linking Climate Change to Land Surface Change, Kluwer Academic Press, Andre, J-C., Bougeault, P., Mahfouf, J-F., Mascart, P., Noilhan, J., Pinty, J-P., (1989) Impact of forests on mesoscale meteorology. Philosophical Transactions of the Royal Society of London, B 324, 407-422. Avissar, R., and Chen, F., (1993) Development and analysis of prognostic equations for mesoscale kinetic energy and mesoscale (subgrid-scale) fluxes for large scale atmospheric models. Journal of Atmospheric Science, 50, 3751-3774. Bachhuber, F. W. and Catto, N., (2000) Geologic evidence of rapid, multiple and high magnitude climate change during the last glacial (Wisconsinan) of North America, in McLaren, S and Kniveton, D. (eds), Linking Climate Change to Land Surface Change, Kluwer Academic Press, Bonan, G.B., Pollard, D., and Thonpson, S.L., (1992) Effects of boreal forest vegetation on global climate. Nature, 359, 76-718. Bradley, R. S. and Eddy, J. A. (1991) Records of past global changes, in Bradley, R. S. (ed.), Global Changes of the Past. Boulder: University Corporation for Atmospheric Research, 5-9. Brooks, N. and Legrand, M., (2000) Dust variability over Northern Africa and rainfall in the Sahel, in McLaren, S and Kniveton, D. (eds), Linking Climate Change to Land Surface Change, Kluwer Academic Press, Brozovic, N., Burbank, D., Meigs, A., (1997) Climatic limits on landscape development in the Northwestern Himalaya. Science, 276, 572-574.

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Burgess, P. E. Palutikof, J. P. and Goodess, C. M., (2000) Approaches to modelling long-term climate change over the British Isles, in McLaren, S and Kniveton, D. (eds), Linking Climate Change to Land Surface Change, Kluwer Academic Press, Catto, N. and Bachhuber, F. W., (2000) Aeolian geomorphic response to climate change: an example from the Estancia valley, Central New Mexico, U.S.A. , in McLaren, S and Kniveton, D. (eds), Linking Climate Change to Land Surface Change, Kluwer Academic Press, Chase, T.N., Pielke, R.A., Kittel, T.G.F., Nemani, R., and Running, S.W., (1996) Sensitivity of a general circulation model to global changes in leaf area index. Journal of Geophysical Research, 101, 73937408. Chase, T.N., Pielke, R.A., Kittel, T.G.F., Nemani, R., and Running, S.W., (2000) Simulated impacts of historical land cover changes on global climate. Climate Dyamics, 16, 93-105. Claussen, M., Kubatzki, C., Brovkin, V., Ganopoloski, A., Hoelzmann, P., and Pachur, H.J., (1999) simulation of an abrupt change in Saharan vegetation at the end of the mid-Holocene. Geophysical Research Letters, 26, 2037-2040. Claussen, M., (1997) Modeling biogeophysical feedback in the African and Indian Monsoon region. Climate Dynamics, 13, 247-257. Dong B., P.J. Valdes and N.M.J. Hall, (1996) The changes of monsoonal climates due to Earth's orbital perturbations and ice age boundary conditions. Palaeoclimates: Data and Modelling, 1, 203-240. Eastman, J.L., Pielke, R.A., andMcDonald, D.J., (1998) Calibration of soil moisture for large eddy simulations over the Fifearea. Journal of Atmospheric Science, 55, 1131-1140. Eastman, J.L., Coughenour, M.B., and Pielke, R.A., (1999) The effects of and landscape change using a coupled plant meteorological model, Global Change Biology, submitted. Gash, J., and Kabat, P., (1999) Land-surface experiments. IGBP Global Change Newsletter, 39,12-14. Gash, J.H.C and Nobre, C.A., (1997) Climatic effects of Amazonian deforestation: Some results from ABRACOS. Bulletin of the American Meeorological Society, 78, 823-830. Gurney, S., (2000) Relict cryogenic mounds in the UK as evidence of climate change, in McLaren, S and Kniveton, D. (eds), Linking Climate Change to Land Surface Change, Kluwer Academic Press, Harvey, L.D.D., (1988) A semianalytic energy balance climate model with explicit sea ice and snow physics. Journal of Clim, 1, 1065-1085. Harvey, L.D.D., (1989) An energy balance climate model study of radaiative forcing and temperature response at 18Ka. Journal of Geophysical Research., 94, 12873-12884. Huang, S.P., Pollack, H.N., Shen, P.Y., (2000) Temperature trends over the past five centuries reconstructed from borehole temperatures. Nature, 403, 756-759. Henderson-Sellers, A., Dickinson, R.E., Durbridge, T.B., Kennedy, P.J., McGuffie, K., and Pitman, A.J., (1993) Tropical deforestation: Modelling local to regional-scale climate change., Journal of Geophysical Research., 98, 7289-7315. IPCC WGI, (1995) Climate Change 1995: - The Science of Climate Change: Contribution of Working Group I to the Second Assessment Report of the Intergovernmental Panel on Climate Change. Houghton, J.T., Meira Fiho, L.G., Callander, B.A., Harris, N., Kattenberg, A., and Maskell, K., (eds.). Cambridge University Press, New York, USA

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Lean, J., and Rowntree, P.R., (1997) Understanding the sensitivity of a GCM simulation of Amazonian deforestation to the specification of vegetation and soil characteristics. J. Climate, 10, 1216-1235. Lu, L., (1999) Implementation of a two way interactive atmospheric and ecological model and its application to the central united States, PhD thesis, Colorado State University, fort Collins, Colorado. Nicholson, S.E., Tucker, C.J., and Ba., M.B., (1998) Desertification, drought and surface vegetation: An example from the west African Sahel. Bulletin of the Ameican Meteorological Society., 79, 815-829. Molnar, P., and England, P., (1990) late Cenozoic uplift of mountain ranges and global climate change: Chicken or egg? Nature, 346, 29-34. Otterman, J., Chou, M-D., and Arking, A., (1984) effects of non tropical forest cover on climate. J. Appl. Met., 23, 762-767. Pielke, R.A, and Avissar, R., (1990) Influence of landscape structure on local and regional climate. Landscape Ecology., 4, 133-155. Pielke, R.A., dalu, G.A., Snook, J.S., lee, T.J., and Kittel, T.G.F., (1991) Nonlinear influence of mesoscale land use on weather and climate. J. Climate, 4, 1053-1069. Pitman, A., Piekle, Sr., R., Avissar, R., Claussen, M., Gash, J., Dolman, H., (1999) The role of the land surface in weather and climate: Does the land surface matter? IGBP Global Change Newsletter, 39, 411. Polcher, J., and Laval, K., (1994) A statistical study of the regional impact of deforestation on climate in the LMD GCM. Climate Dynamics, 10, 205-219. Price, G.D., P.J. Valdes and B.W. Sellwood, (1998) A comparison of GCM simulated Cretaceous and climates; implications for the sedimentary record. Palaeogeography, Palaeoclimatology, Palaeoecology, 142, 123-138. Price, G.P., B.W. Sellwood and P.J. Valdes, (1996) Sedimentological evaluation of general circulation model simulations for the "greenhouse" earth: Cretaceous and Jurrasic case studies. Sedimentary Geology, 100, 159-180. Schmidt, M. and Dehn, M., (2000) Examining links between climate change and landslide activity using GCMs, in McLaren, S and Kniveton, D. (eds), Linking Climate Change to Land Surface Change, Kluwer Academic Press, Smith, E.A., Hsu, A.Y., crosson, W.L., field, R.T., Fritscen, L.J., Gurney, R.J., kanemasu, E.T., Kutstas, w.P., nie., D., Shuttleworth, W.J., Stewart, J.B., Verma, S.B., Weaver, H.L., and Wesley, M.L., (1992) Areaaveraged surface fluxes and their time-space variability over the FIFE experimental domain. Journal of Geophysical Research., 97, 18599-18622. Texier, D., de Noblet, N., Harrison, S.P., Haxeltine, A., Jolly, D., Joussaume, S., Larrif, F., prentie, I.C., and Tarasov, P., (1997) Quantifying the role of biosphere-atmosphere feedbacks in climate change: Coupled model simulations for 6000 years BP and comparison with palaeodata for northern Eurasia and Northern Africa. Climate Dynamics, 13, 865-882. Thorpe R.B., K.S. Law, S. Bekki and J.A. Pyle, (1996) Is methane driven deglaciation consistent with the ice core record? Journal of Geophysical Research, 101-D22, 28627-28635.

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Toumi R., J.D. Haigh and K.S. Law, (1996) A tropospheric ozone- lightning climate feedback. Geophys. Res. Lett., 23, 1037-1040. Valdes, P.J., B.W. Sellwood and G.P. Price, 1996: The concept of Cretaceous climate equability. Palaeoclimates, 1, 139-158. Vidale, P.L., Pielke, R.A, barr, A., and steyaert, L.T., (1997) case study modelling of turbulaent and mesoscale fluxes over the BOREAS region. Journal of Geophyical Research, 102, 29167-29188. Viles, H. and Goudie, A. S., (2000) Weathering, geomorphology and climatic variability in the Central Namib Desert, in McLaren, S and Kniveton, D. (eds), Linking Climate Change to Land Surface Change, Kluwer Academic Press, Whipple, K.X., Kirby, E., Brocklehurst, S.H., (1999) Geomorphic limits to climate –induced increases in topographic relief. Nature, 40, 39-43. White, K., McLaren, S. Black, S. and Parker, A., (2000) Evaporite minerals and organic horizons in sedimentary sequences in the Libyan Fezzan: implications for palaeoenvironmental reconstruction, in McLaren, S and Kniveton, D. (eds), Linking Climate Change to Land Surface Change, Kluwer Academic Press, Wilby, R. L. and Dettinger, M. D., (2000) Streamflow changes in the Sierra Nevada, California, simulated using a statistically downscaled General Circulation Model scenario of climate change, in McLaren, S and Kniveton, D. (eds), Linking Climate Change to Land Surface Change, Kluwer Academic Press, Xue, Y., (1996) The impact of desertification in the Mongolian and the Inner Mongolian grassland on the regional climate. J. Climate, 9, 2173-2189. Yair, A. and Bryan, R. B., (2000) Hydrological response of desert margins to climate change, in McLaren, S and Kniveton, D. (eds), Linking Climate Change to Land Surface Change, Kluwer Academic Press, Zhao, M., Pitman, A.J., and Chase, T., (2000) The impact of land cover change on the atmospheric circulation, Climate Dyn., submitted. Zhao, M., Pitman, A.J., and Chase, T., (2000) Sensitivity of a general circlation model to global changes in leaf area index: A reassessment, J. Geophys. Res., submitted.

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Index

A aeolian activity · 166, 167, 170, 171, 183, 185 aerosols · 1, 3, 4, 5, 8, 11, 12, 13, 24, 103, 118, 121, 126, 134, 242 Algeria · 7, 9 Angola · 12 Antecedent Daily Rainfall Index · 131 atmosphere · 2-4, 11-13, 15, 18, 21, 24, 81, 82, 87, 93-95, 97,100, 102, 103, 106, 107, 117, 120, 134, 137,228, 229, 238, 242-248,251 atmospheric 227-231, 234, 239, 240

B biological weathering · 63 British Isles · 117, 226, 238, 240, 249 Burkina Faso · 28, 30-32, 45

C California · 96-98, 114, 116-118, 167, 187, 251 Chad · 7 climate anomalies · 81, 83, 86, 89

climate change · 2, 26, 29, 30, 37, 46, 47, 62, 97, 115-119, 121, 125, 127, 134, 136, 137, 139, 140, 142, 163167, 182, 184, 207, 226-228, 234, 236, 238-244, 247, 249-251 climate change impact · 119 climate model · 100, 117, 123, 124, 136, 226, 227, 239, 244, 246, 247, 250 climatic forcing · 228 Crop Moisture-Z Index 81 cryogenic mounds · 205-207, 209-211, 214-218, 220-225, 250

D deflation · 1, 2, 7, 8, 15, 16, 18, 19, 2123, 141-143, 145, 150, 151, 166, 167, 170-172, 174, 178, 180-185 deforestation · 82, 250 Democratic Republic of Congo ·12 desertification · 26-29, 46, 47, 49, 60, 61, 67, 82, 201, 251 desiccation · 2, 3, 7, 8, 22, 23, 26, 2931, 33, 35, 37, 66, 139, 140, 145, 150, 153, 158, 159, 163, 164, 168, 180 dome dunes · 166, 170, 171, 175, 177180, 182-185 downscaling · 96-98, 100-103, 108111, 114-123, 125, 127, 134-137, 227, 232, 236, 238, 244-248 downscaling techniques · 96, 98, 120, 137 droughts · 29, 45, 81, 82, 93 duricrusts · 199

262

dust · 1-16, 18-25, 49, 59, 61, 65, 80, 186, 242, 244 dust indices · 1 dust production · 1, 3, 6, 8, 9, 13, 14, 15, 18-23, 25 dust sources · 3, 5, 7

E ecosystems · 48, 49, 60, 226 Egypt · 7, 200, 201, 202 El Niño · 69, 120 endemic species · 49 environmental change · 188, 189, 200, 227, 230, 245 environmental degradation · 26, 29 Erg of Bilma · 7 evaporation · 57, 59, 60, 68, 95, 162, 242 evapotranspiration · 83, 95, 106, 135, 136, 182

general circulation models (GCMs) · 94-98, 100, 101, 114, 119-121, 125, 126, 134, 136, 245, 246, 251 Great Man-Made River Project · 189, 202 greenhouse gases · 96, 98, 103, 243 gypsarenite · 143, 145, 151, 158, 168, 172-174, 176, 178, 180-183, 185, 195, 196, 198, 199-204, 206, 210, 214 gypsum · 63, 70, 74, 75, 80, 173, 174, 179, 181, 187, 190, 194, 195, 200, 203 gypsum crusts · 68

H Holocene · 25, 141, 142, 151, 166, 168, 170, 180, 183, 185, 186, 199, 201, 202, 227, 228, 229, 244, 249 Horn of Africa · 8, 10

F

I

feedbacks (see climatic forcing) Fezzan ·188, 189, 199-203, 245, 251 fire · 11, 12, 192, 196, 202, 209 floods · 81, 83, 151 fog · 63, 66, 67, 69, 70, 78-80 foggara · 194, 196, 199

ice · 81, 139, 140, 206, 209-211, 214217, 220, 221, 223, 224, 226-229, 234-242, 247, 249-251 infiltration · 48-50, 59, 60 Infra-Red Difference Dust Index (IDDI) · 1, 3 insolation · 4, 65, 77, 78, 107, 206, 227-230, 237-239 Italy · 119, 124, 135, 136, 234, 247

G Gabon · 12

263

K

New Zealand · 119, 120, 128, 135, 136, 247 Niger · 7-9, 11, 24, 28, 30-34, 36, 37, 45, 46

Kalahari · 12, 185

O L landslide scenarios · 119 landslides · 119, 120, 129, 132, 135, 247 Libya · 7, 189, 201-203, 245 lichen · 63, 65-71, 73-77, 79, 80 loess · 48-51, 58-62, 166, 167, 172, 177-179, 183-186, 245

ocean cores · 228 optical stimulation luminescence · 180 ostracods · 143, 146-149, 151-155, 158, 160, 176

P Palmer Drought Severity Index-PDSI ·

M Mali · 7, 24, 28, 30-32, 45, 202 Mauritania · 7, 31 METEOSAT · 1, 3 microorganisms · 63, 65, 66, 75, 76, 80 Morocco · 7, 9, 80 mudslide · 119, 120, 123-125, 134, 135

N Namib Desert · 12, 63, 64, 66, 68, 69, 77-80, 251 Negev · 49-51, 58-62, 65, 79, 80, 186 Negev desert · 49-51, 58-62, 65, 66, 67 Neolithic · 199, 200, 201 New Mexico · 139-141, 163-168, 182, 185-187, 245, 249

81, 83 palsas · 205, 206, 209, 211, 214, 215, 217, 220, 221, 223-227 parabolic dunes · 166, 169, 171, 181, 182 periglacial · 205, 206, 209, 210, 220, 221, 224-227, 231, 233-236 permafrost · 205, 206, 209, 211, 212, 215, 216, 221, 223, 225-227 pingos · 205, 206, 209-212, 214-217, 219-221, 223-227, 247 playa · 139, 141-143, 153, 155, 156, 160, 162-164, 166, 167-173, 178183, 185, 186, 200, 201, 245 Pleistocene · 61, 142, 143, 166, 188, 201, 202, 209, 211, 220, 225, 227, 240 pluvial lake · 139, 142, 145, 150, 160, 166, 185 precipitation (see rainfall)

264

Q

soil degradation · 8, 13, 14, 22, 33, 34 soil-dust cycle · 8 stable isotope analysis · 188, 189 Sudan · 7, 46, 47, 200 203 surface properties · 48-50, 60, 106, 245

Quaternary · 25, 50, 58, 61, 139, 140, 145, 165, 166, 168, 185, 186, 187, 189, 200, 201, 202, 203, 215, 223, 224-229, 239, 241, 245

T tafoni · 65, 70, 80 thermoluminescence · 180

R radiation . 4, 8, 65, 68, 101, 106, 107, 116, 118, 219, 225 rainfall · 1-3, 8, 11, 15, 16, 18-38, 4051, 54, 55, 57, 59-63, 69, 78, 79, 83, 86, 93, 94, 97, 114-118, 129, 130, 132, 133, 136, 142, 167, 181, 189, 245, 246, 247, 249 raingauges · 26, 38 remote sensing ·11, 24, 94, 189-191, 201, 246, 248 rock-surface microenvironments · 63 runoff · 15, 48-51, 54, 55, 57, 59-62, 83, 96-98, 106, 113-116, 118, 163, 242

U upscaling techniques · 246, 248 U-Th dating · 193

V variograms · 38, 40-42 vegetation cover ·15, 16, 19, 49, 167, 184, 222

S Sahara ·1, 2, 5, 7-16, 18, 20, 22-24, 28, 29, 45, 49, 188, 189, 201-203 Sahel · 1-3, 5, 7-11,13-16,18-35, 37, 38, 40, 41, 45-47, 82, 246, 249, 250 Sahel-Sahara · 1, 3, 6, 7, 9, 11, 13, 23 salt weathering · 63, 65, 66, 69, 70, 79, 80 snowpack · 97, 98, 106, 107, 111, 113, 114, 115, 118

W Wisconsinan · 139, 142, 143, 145, 146, 148-151, 153, 155-164, 168, 170, 172, 183-185, 216, 226, 247, 249

Advances in Global Change Research 1. 2.

3. 4. 5.

6.

P. Martens and J. Rotmans (eds.): Climate Change: An Integrated Perspective. 1999 ISBN 0-7923-5996-8 A. Gillespie and W.C.G. Burns (eds.): Climate Change in the South Pacific: Impacts and Responses in Australia, New Zealand, and Small Island States. 2000 ISBN 0-7923-6077-X J.L. Innes, M. Beniston and M.M. Verstraete (eds.): Biomass Burning and Its InterRelationships with the Climate Systems. 2000 ISBN 0-7923-6107-5 M.M. Verstraete, M. Menenti and J. Peltoniemi (eds.): Observing Land from Space: Science, Customers and Technology. 2000 ISBN 0-7923-6503-8 T. Skodvin: Structure and Agent in the Scientific Diplomacy of Climate Change. An Empirical Case Study of Science-Policy Interaction in the Intergovernmental Panel on Climate Change. 2000 ISBN 0-7923-6637-9 S.J. McLaren and D.R. Kniveton (eds.): Linking Climate Change to Land Surface Change. 2000 ISBN 0-7923-6638-7

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