Linking Climate Change to Land Surface Change (Advances In Global Change Research, Volume 6) [1 ed.] 0792366387, 9780792366386, 9780306480867

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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

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

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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

49

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

83

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

193

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

209

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

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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

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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

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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 (