From landform to process: Morphology and formation of lake-bed barchan dunes, Makgadikgadi, Botswana

Geomorphology 161–162 (2012) 1–14

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From landform to process: Morphology and formation of lake-bed barchan dunes, Makgadikgadi, Botswana Sallie L. Burrough a,⁎, David S.G. Thomas a, b, Richard M. Bailey a, Lauren Davies a a b

School of Geography and Environment, University of Oxford, South Parks Road, Oxford OX1 3QY, UK Visiting Professor, Department of Environmental and Geographical Science, University of Cape Town, Rondebosch 7701, South Africa

a r t i c l e

i n f o

Article history: Received 20 July 2011 Received in revised form 18 March 2012 Accepted 27 March 2012 Available online 4 April 2012 Keywords: Barchan dunes OSL Morphometry Landscape dynamics Makgadikgadi pans

a b s t r a c t A suite of crescentic landforms is visible from remotely sensed imagery within the Ntwetwe panPan in the Makgadikgadi basin, Botswana. We investigate the most distinct of these landforms using morphometric measurements, sedimentary data and Optically Stimulated Luminescence (OSL) signal analysis. Comparative analysis with previously published barchan morphological data sets suggest the Ntwetwe features fall within the spectrum of morphometric parameters found in a range of barchan dunefields from around the world. There is currently insufficient comparative morphometric data from sub-aqueous dunefields to be able to distinguish the particular formative environment of the dune. OSL signal analyses however, support the hypothesis of Grove (1969) [Grove, A.T., 1969. Landforms and climatic change in the Kalahari and Ngamiland. Geographical Journal, 135: 191–212] that the last deposition of the sediments within the Ntwetwe forms was most likely aeolian in origin. Luminescence signal analysis is employed to investigate potential transport and bleaching environments of the sediments forming the features, but results in this case do not shed further light on the formative conditions of these enigmatic landforms. © 2012 Elsevier B.V. All rights reserved.

1. Introduction In many semi-arid or subtropical regions, environmental changes during past millennia are often represented in the development of landform suites relating to formerly more humid or more arid conditions. Lake basin landforms and dune systems have, respectively, become major geomorphic proxies for environmental reconstructions of these hydrological states, although their interpretation in climatic terms is not necessarily straightforward (Thomas and Burrough, 2012). Indeed, dunes can be disaggregated into forms that have differing hydrological contexts, such that some duneforms are not clear indicators of arid conditions. This issue is well illustrated in the context of lunette dunes, found on the margins of small closed lake basins (pans) and subject to varying interpretations of the hydrological conditions within pans at the time of their formation (see e.g. Bowler, 1973, 1976, 1986 in Australia, and Lancaster, 1978, 1979 in the Kalahari, southern Africa). There are in fact a variety of geomorphological relationships between aeolian dunes and lake basins. In this paper we briefly outline the associations between lakes and dunes that are considered in the literature, before moving to analyse the relationships and palaeoenvironmental significance of the least-studied of these associations: lake floor dunes. There has been little systematic analysis of these forms to

⁎ Corresponding author. E-mail address: [email protected] (S.L. Burrough). 0169-555X/$ – see front matter © 2012 Elsevier B.V. All rights reserved. doi:10.1016/j.geomorph.2012.03.027

date. However, their potential significance for interpretations of lake hydrological changes has been highlighted, for example, by Grove (1969), while chronologies of development for such features can establish relationships to other changes recorded within lake basin deposits (e.g. Hesse, 2009). One explanation for the general lack of analyses for lake bed dunes is that the mode of origin of the dunes has not been clearly identified, and in some cases, such as in the Kalahari, it has been hypothesised that they are not aeolian landforms, but features developed in subaqueous conditions (Cooke, 1980). In other contexts, the variable nature of lake floor environments particularly in dryland regions, can provide increased preservation potential for landforms that are otherwise ephemeral and ultimately will lead to complex histories that must be unpicked. For example, remnants of barchans have been identified on the floor of Lake Chad (Bristow et al., 2009), where they are interpreted as being exhumed from beneath lake floor deposits by contemporary deflation. Our aim, in the context of the Kalahari where lake-floor barchan forms have been identified, is to establish their likely mode of formation, and therefore their potential use as a geoproxy (sensu Thomas and Burrough, 2012) in Quaternary palaeoenvironmental studies. 1.1. Dune and lake associations Associations between aeolian dunes and lakes are not uncommon and take various forms: interdunal lakes (e.g. Yang and Williams, 2003), basin perimeter dunes (e.g. Lancaster, 1978), downwind dunefields (Stevens, 1991), and lake-floor dunes (Grove, 1969).

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In a number of dunefields worldwide, but best represented analytically in the Badain Jaran Desert, western China (Yang et al., 2010), 200–300 m-high linear and star dunes provide the geomorphological constraints on the development of a large number of groundwater seepage lakes, which occupy interdunal depressions. On a smaller scale, a similar association exists between degraded linear dune ridges in western Zimbabwe and strings of small pan depressions in the interdune corridors, which are supplied both by groundwater seepage and seasonal rains (Goudie and Thomas, 1985). Dunes on the perimeter of lake basins commonly take the form of lunettes (Hills, 1940), found on the downwind margin of basins and first systematically analysed by Bowler (1973) in southern Australia, Lancaster (1979) in the southern Kalahari, and Huffman and Price (1979) in Texas, and subsequently analysed in terms of their potential as indicators of hydrological changes during the late Quaternary. Based on morphology, orientation and sedimentology, Bowler (1973) regarded the formation of lunettes to be associated with the deflation of sediment from basin ‘beaches’, in a manner not dissimilar to the formation of coastal dune systems. Clay-rich components of lunettes however were regarded as indicative of deflation of clay pellets from

50km

A

seasonally-dry pan floors, the explanatory mode of lunette formation preferred by Lancaster (1978, 1979) in a Kalahari context. With the application of OSL dating to both pan floor and lunette sediments, it has become further apparent that the development of lunette-pan associations is both spatially and temporally complex (Telfer and Thomas, 2006). A further association between sediment derived from a basin and downwind dune formation was mapped but not described by Grove (1969) in the Makgadikgadi basin of the Kalahari. Here, significant beach ridges have developed, through the accumulation of wavetransported sediments during multiple lake high stands, on the western margin of the palaeo-lake basin (Burrough et al., 2009). Grove (1969) also mapped an arcuate field of small longitudinal dunes radiating and extending west from the shoreline sequence, that 50 km westwards transforms into hummocky transverse features. While these features have not yet been the subject of systematic investigation, they may suggest that sediment within the shoreline sequence may have provided material for specific dune field development. Dunes were also described by Grove (1969) on the floor of the western-most part of the Makgadikgadi basin (Fig. 1d), where ‘the

50km

B

b

a

6km

50km

C

c

6km

D

50km

d

6km

6km

Fig. 1. Unidentified concentric landforms on a selection of dry lake beds within the subtropical desert belts. Inset rectangles are shown at smaller scale below images A, B, C, D as a, b, c and d. A) Lake MacKay, Northern Territory, Australia, B) Lake Amadeus, Northern Territory, Australia; C) Lake Chad, North Africa; D) Palaeolake Makgadikgadi, Botswana, Southern Africa.

S.L. Burrough et al. / Geomorphology 161–162 (2012) 1–14

3

20 27’0”S

A

C

20 45’0”S

20 40’0”S

20 36’0”S

20 31’30”S

B

24 44’0”E

24 48’0”E

24 52’0”E

24 56’0”E

25 00’0”E

25 40”E

25 6’0”E

25 12’0”E

25 16’0”E

Fig. 2. a) Location of dunes identified and mapped in the Kalahari (Ntwetwe Pan study area shown as inset rectangle). b) Dunes sampled for morphometric analysis in Ntwetwe. c) Aerial view of Ntwetwe barchan dunes (see Fig. 3 for scale).

north and north-west shore of the Ntwetwe Pan is decorated by a large number of fairly small dunes transverse to the easterly winds. Those furthest out into the pan appear to be the youngest and have typical barchan forms’ (p.205). Though ascribed an aeolian origin, this was disputed by Cooke (1980) who considered them to potentially be subaqueous rather than wind-formed features. Features similar in context and form are not confined to Makgadikgadi, and can be seen in aerial imagery to occur on the floors of other dry lake basins such as those of Lake Chad in North Africa (Fig. 1c) and in central and western Australia (Fig. 1a,b). In some cases, the features are significantly larger than those in Makgadikgadi, and appear to have linear dune ridges superimposed on them (Fig. 1a), adding further to the uncertainty over formation mechanisms.

These lake-floor features, together with other associated dunes described above, are potentially extremely important in terms of understanding hydrological changes within basins, major environmental changes associated with transitions between lake full/empty conditions, and background geomorphic environments that have occurred in the past. However, it is often very difficult to determine the geomorphological processes leading to the formation of such landforms in these complex environments, limiting our ability to use them as indicators of past environmental change. These ‘geoproxies’ are particularly important in continents such as Africa and Australia where drylands are extensive under contemporary climate regimes (~62% (UNSO, 1997) and 75% (Croke, 1997) of the landmass of Africa and Australia, respectively, is classified as hyper-arid, arid or semi-

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crest

stoss slope slipface

A

B

Horn to horn width, W

horns

Wind Direction Dune Length, L

Lh

Lb

i) Active barchan dune

C Toe to Court

ii) Degraded barchan dune

D

iii) Plan view of Ntwetwe barchan dune (CNES/ Spot imagery, 2011)

1km Fig. 3. a) Key features of typical aeolian barchan dunes; b) parameters measured in the morphometric analysis; c) hypothesised model of degradation to current form i–iii); d) High resolution 2011 CNES/SPOT image from Google Earth of a typical Ntwetwe barchan dune.

arid). In these environments, due to the paucity of alternative biological proxies, the ability to understand past Quaternary processes without recourse to landform dynamics is extremely challenging. In the context of the Makgadikgadi lake-floor features in Botswana, Grove (1969) noted that ‘these dunes would seem to provide evidence of the climate having been drier than the present day in the interval since the great Makarikari [sic] lake disappeared, with subsequent wetter episodes’ (p205). There is, however, currently no systematically collected evidence to assess the formation and significance of these features. This paper attempts to do this, using data from the Makgadikgadi basin, but with relevance to similar features found in other sub-tropical and dryland dry lake basins. 1.2. Barchan features of Makgadikgadi The Makgadikgadi depression is the largest of three discreet subbasins of a palaeo-mega-lake system in the Middle Kalahari (Fig. 2a). Major variations in the water level of the Makgadikgadi basin during the late Quaternary have been demonstrated using chronologies of

relict fossil shorelines that demarcate the limits of the lacustrine basin (Burrough et al., 2009) and via investigations of the basin inflows (Nash et al., 1994) and diatomaceous deposits within the megalake system (Shaw et al., 1997; Shaw et al., 2003). Under present-day conditions, the Makgadigkadi basin is generally dry though seasonally holds localised shallow standing water (up to a metre in depth). Local rainfall (usually falling between November and April) averages 400 mm yr − 1 (Ringrose et al., 2003) but evapotranspiration currently exceeds this rate by a factor of 3 (Bhalotra, 1987). Under contemporary conditions the surface water of the pans varies between fresh (during the wet season) and highly alkaline (during the dry season), and when dry, the surface of the pans becomes extensively salt encrusted. The basin encompasses two principal pans, Sua Pan in the east and Ntwetwe Pan in the west, together with a number of smaller pans closer to the periphery (Fig. 2a). The barchan ‘islands’ cluster on the western side of the Ntwetwe Pan (Fig. 2b). These are crescentic, unconsolidated accumulations of sand consistently oriented in-line with the prominent (east-to-west) wind direction. From east to west, their forms tend to

S.L. Burrough et al. / Geomorphology 161–162 (2012) 1–14

A

5

B

C

Fig. 4. a) Example plot of De with signal integration time (circles) and Z ratio (squares) for an aliquot (sample MAK/09/15/3) immediately following a daylight partial bleach; b) the same aliquot partially bleached by blue LEDs. In both a and b, the Z ratio plotted is the De at each integral (plotted as the mid-point in seconds) over the De calculated for the first integral. c) Mean proportional contribution of each signal component: fast (F), medium (M) and slow (S) to the bulk continuous wave OSL signal (SUM) during illumination. The integral time tz used to calculate the Z ratio is indicated by the dashed lines.

get progressively less discrete and increasingly dense, amalgamating into a series of merged landforms. Their vegetated surfaces, which can sustain a dense covering of halotrophic grasses (pre-dominantly Odyssea paucinervis and Sporobolus ioclados (J. Bradley, pers. comm.) and occasional acacia trees, make these features visibly distinct in remotely sensed imagery, in contrast to the generally de-vegetated and inhospitable conditions of the pan floor (Fig. 2c). Grove (1969) noted the presence of what appear, from aerial photographs, to be terraces or benches on the slopes of these sand accumulations (Fig. 2c) and suggested this was evidence for one or more wet phases post-dating the dune formation, when a standing body of water would have cut into the topography. Grey and Cooke (1977) and Cooke (1980) also note the large numbers of broad ‘roughly crescentic islands’, which tend to merge together to the west. Cooke (1980) suggested that the “sand spreads” must post-date the pan surface, and theorised that if the barchan-like forms of Ntwetwe were once true barchans then they must have been severely reduced in height and bulk since they became inactive.

forms (though they can merge into barchanoid ridges), with limited sediment supply being the characteristic that most likely explains the discreteness of form (Ash and Wasson, 1983). Barchans occur in many sand seas on earth (Lancaster, 2011), and the simplicity of pure barchan forms has made them a favourite of text books, field-survey, descriptive (e.g. Finkel, 1959; Long and Sharp, 1964; Hastenrath, 1987) and modelling studies (e.g. Kroy et al., 2005). Despite this, compared with other forms of transverse features and linear dunes, they are not especially common in terrestrial systems, forming only ~10% (by area) of all desert dunes on earth (Thomas, 2011) and containing less than 1% of all dune sand (Wilson, 1973). Fig. 3 shows features of a classic barchan dune form. In summary, barchans comprise: a gently sloping stoss (windward) slope up which arriving sand migrates, a brink line often but not always coincident with the crestline, and a slip face with a concave profile,

1.3. Barchan dune morphology 1.3.1. Aeolian barchans In environments that are conducive to aeolian sand transport, the type and form of any resultant dune bedform are determined by a number of factors, of which wind energy, wind directional variability and sediment availability are recognised as most important. Dunes that are not significantly constrained by topography develop, over time, into self-organised patterns (dune fields), the morphometry of which is shaped by the interactions of these wind and sediment factors (Werner, 1995). The actual form of dunefields and the individual dunes within them varies markedly across sand seas globally (Breed and Grow, 1979), as well as in the extra terrestrial systems of Mars, Venus and Titan where dunes are a spatially significant landform (Bourke et al., 2010; Clarke, 2011). Transverse dunes, which have crestlines normal to the resultant transport direction, and migrate in-line with the fluid flow, develop where sand transport directional variability is limited (sometimes described as unidirectional). Crescentic barchans dunes are discrete

Fig. 5. High precision differential GPS data for single dune form (greyscale indicates height above sea level in metres (see legend).

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Table 1 Morphometric data for barchans dunes, from selected sites worldwide where ‘Length’ is the length of the windward slope; ‘Width’ is the horn to horn width and ‘Wavelength’ is the crest to crest distance between dunes . Data from this study shown in grey. All other data taken from Breed and Grow (1979). Region

Latitude

Longitude

Mean length Mean width Mean wavelength Length/width Width/wave-length Length/wave-length (km) (km) (km) (L/W) (L/λ) (W/λ) 1.31

2.22

2.27

0.59

0.98

0.58

101°–102° W

1.65

3.14

1.73

0.53

1.82

0.95

3 4

Eastern TaklaMakan (China) Ala Shan Desert (China)

0.94 2.88

0.89 2.92

0.85 0.75

1.06 0.99

0.90 0.74

Western TaklaMakan (China) Cherchen Sand Sea Takla Makan (China) Peski Karakumy (near Mary, Russia) White Sands (USA) Algodones (California, USA)

87° 15′–88° E 101° 30′– 102° 30′ E 79° 30′–80° E 86° 45′–87° 30′E 62° 30′–64° 15′ E 106° 15′ W 114° 45′– 115° 15′W 08° 30′–09° 30′E 113° 30′– 114° W 39°–40° E

0.80 2.16

5 6

42° 30′–43° 30′ N 41° 30′–42° 30′ N 40° 15′ N 40°–40° 30′ N 39° 45 N 39° 30′–40° 30′ N 37°–37° 45′ N 32° 45′ N 32° 45′–33° 10′N 32° 20′–30° N 31° 45′–32° N 27° 30′–29° N 26° 30′–27° 30′N 26°–26°45′ N 25° 45′N

62° 30′–65° E

2

Pesky Karakumy (near Bukantau, Russia) Nebraska Sand Hills (USA)

0.84 2.20

1.66 3.24

1.10 3.00

0.51 0.68

1.51 1.08

0.76 0.73

1.16

1.96

1.76

0.59

1.11

0.66

0.07 0.88

0.11 1.61

0.11 1.07

0.65 0.55

1.00 1.50

0.65 0.82

0.65

1.43

1.24

0.45

1.15

0.52

0.66

1.04

1.38

0.63

0.75

0.48

0.80

1.41

1.84

0.57

0.77

0.43

68° 30′–69° 30′ E 69°–70°E

1.30

1.50

1.44

0.87

1.04

0.90

0.47

0.93

0.58

0.51

1.60

0.81

01° 30′–02° W 53°–54° 30′E

0.62

0.77

–

0.81

0.59

0.80

0.63

0.74

1.27

0.94

0.44

0.71

0.59

0.62

1.20

0.75

21° 23′ N

51° 30′–52° 30′ E 53° 30′–55° E

2.09

2.76

2.56

0.76

1.08

0.82

20° 30′N

55° 15′–56° E

0.67

1.43

1.76

0.47

0.81

0.38

17° 30′–19° N 20°–20° 45′ S 23° 30′–24° 30′ S

09°–13° W

1.59

4.10

1.71

0.39

2.40

0.93

24° 35′–25° 15 E 14° 30′–14° 45′ E

0.45

0.67

1.28

0.67

0.52

0.35

0.68

1.12

0.87

0.61

1.29

0.78

1

7 8 9

10 Northeast Sahara (Algeria and Tunisia) 11 Gran Desierto (Sonora, Mexico) 12 An Nafud (northern Saudi Arabia) 13 Western Thar Desert (near Sukkur, Pakistan) 14 Thar Desert (near Umarkot, Pakistan) 15 Western Sahara (Mali) 16 Al Jiwa (Saudi Arabia) 17 Persian Gulf (UAE) 18 Eastern Rub'al Khali (Saudi Arabia) 19 Southeast Rub' al Khali (Saudi Arabia) 20 Aoukar sand sea (Mauritania) 21 Ntwetwe (Botswana) 22 Namib Coast (Namibia)

23° 15′–23 30′N 22°–25° N

located between two horns that face down-wind, which develop due to flow streamlining and greater velocities along dune edges, and which are points of sand loss from the dune. Studies of terrestrial barchan dune movement report barchan migration rates of 5 m to over 60 m per year (e.g. Norris, 1966; Sarnthein and Walger, 1974). To maintain a steady form/volume of sediment, a barchan needs either to retain all of its sediment, or to receive from fluid flow the same amount of sand as is lost from the horns during the process of migration. However, form variations are in reality considerable, both in terms of situations where barchans retain an isolated presence within dune fields (Sauermann et al., 2000) or when forms merge, either by small dunes feeding into and being subsumed within larger forms due to differential migration rates (e.g. Simons, 1956), or because sediment availability leads to individual barchans merging to form barchanoid or transverse ridges (e.g. Derickson et al., 2008). ‘Mega-barchans’ may be up to 500 m high (Warren, 1976). 1.3.2. Subaqueous barchans Subaqueous barchan forms have also been recognised from various environments including fluvial channels (e.g. Nittrouer et al., 2008), tidal regions (Bolduc and Duchesne, 2009) and the ocean floor under conditions of limited sand supply, hard substrate surfaces and unidirectional water currents (Todd, 2005; Daniell and Hughes, 2007. Comparatively fewer field studies have been undertaken due largely to the difficulties of observation and measurement in subaqueous environments

but laboratory flume experiments of subaqueous forms have become commonplace (e.g. Endo et al., 2004, 2005). Such experiments have demonstrated the morphological similarity of these forms to aeolian dunes, suggesting it is sand flux and the geometry of fluid flow which is key to the morphology of barchan dunes rather than the physical mechanisms involved in particle motion (which are significantly different between the fluid media of air and water) (Hersen et al., 2002).

1.4. Barchan dune occurrence Barchans occur in many sand seas (Breed and Grow, 1979) and subaqueous localities on earth, but the requirement of unidirectional sand transport and limited sediment availability inhibits where they form and persist. Their migratory nature gives a clue to where and why they occur, as well as to why they are not more widely occurring. Being migratory forms, barchans can appear to transport sand from source areas to locations where, due to greater sand transport directional variability and/or sediment accumulation occurring, forms that contain a higher volume of sand or that are more persistent (i.e. show a lower rate of sediment turn over) develop. Thus, in a downwind direction, barchans tend to lead into areas where barchanoid/transverse ridges, as observed in the Namib Sand Sea (Lancaster, 1989; Livingstone et al., 2010), or linear dunes, as in Sinai (Tsoar, 1974) occur.

S.L. Burrough et al. / Geomorphology 161–162 (2012) 1–14

10.0

A

7

Aeolian dunefields

9.0

Ntwetwe dunefield Sub-aqueous dunefields

8.0

Mean Width (km)

7.0 6.0 5.0 4.0 3.0 2.0 1.0 0.0 0.0

1.0

2.0

3.0

4.0

5.0

Mean Length (km) 4

8

B

y = 0.9031x + 0.2597 R² = 0.278

3.5

Wavelength (km)

Width (km)

3 2.5 2 1.5 1 0.5 0

C

y = 1.9655x + 0.5459 R² = 0.312

7 6 5 4 3 2 1

0

0.5

1

1.5

Length (km)

2

0

0

0.5

1

1.5

2

Length (km)

Fig. 6. a) Ntwetwe barchans shown within the mean and range of dune width/length relationships of other barchan dunefields (see Table 1); b) dune length/width relationship and c) dune length/wavelength relationship for all measured Ntwetwe dunes.

2. Methods 2.1. Morphometric analysis Morphometry can be used to classify, characterise and distinguish landform suites in terms of formative environments and processes. Attempts to characterise barchan dunes and to explain relationships between development, movement and environmental factors, as well as form variance, have often focussed on morphometric analysis. In the context of aeolian sand dunes, this approach has been employed since the early work of Long and Sharp (1964) who measured critical parameters of barchan dunes to allow form classification; an approach that has subsequently been used to compare the characteristics of forms between dunefields on earth and in an inter-planetary context (e.g. Bourke and Goudie, 2009). Typical relationships have been determined between height, width and length parameters, though Sauermann et al. (2000) suggest that scale invariance can occur, even within a single dunefield. Long and Sharp (1964) and subsequent applications of their approach (e.g. Norris, 1966; Hastenrath, 1987), used the ratio of two key parameters, horn-to-horn dune width (a) and dune length (c, in line with migration direction) within areas of distinct barchan dunes. A more comprehensive assessment of the full suite of transverse dune

forms, including transverse ridges and merged barchans rather than solely isolated forms, was facilitated by the Landsat-based global survey of Breed and Grow (1979). Their approach focussed on a smaller suite of measurements, readily determined from imagery, comprising dune width and length, defined in the same manner as Long and Sharp (1964), and the crest-to-crest wavelengths of dunes. For transverse ridges that comprised multiple crescentic sections, each horn-to-horn element was treated as a separate form. In the present work, following Long and Sharp (1964) and Breed and Grow (1979), and using 2011 high resolution CNES/Spot imagery (~2.5 m spatial resolution) via the Google Earth platform, morphometric data were collected for all (149) distinct dune forms clearly identified on the Ntwetwe pan surface and exhibiting a characteristic crescentic form (Fig. 2b). Horn to horn widths, dune lengths and inter-dune wavelengths were quantified using Google Earth's interactive measurement tool. Comparative analyses of key morphometric parameters with previously published barchan morphological data were then undertaken. A separate 3 arc-second Digital Elevation Model (DEM) was created from Shuttle Radar Topographic Mission (SRTM) data (Farr and Kobrick, 2000) with a regional resolution of approximately 90 m by 94 m crudely capturing the relative topography of even the smallest dune features measured. Elevation data were extracted from pan floor and

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Laboratory experiments

Subaerial Dunes

R1 R2 R3 R4 R5

A1 A2 A3 A4 A5 A6 A7

Subaqueous Dunes S1 S2

Ntwetwe dunes (this study)

Lh/W

1

0.5

0

0

0.5

1

1.5

2

2.5

Lb/W Fig. 7. Ntwetwe dune measurements in the context of parameters measured by Endo et al. (2005) to distinguish between subaqueous and subaerial flows. Experiments R1-4 (Endo et al., 2005) were conducted in the laboratory under unidirectional currents of varying speeds using a recirculating flume. Those of Endo et al. (2004) (R5) were also conducted in the laboratory but used oscillatory water wave motion. Subaqueous dunes: S1 (Allen, 1968); S2 (Lonsdale and Malfait, 1974). Subaerial dunes: A1 (Finkel, 1959); A2 (Long and Sharp, 1964); A3 (Hastenrath, 1967); A4 (Khalaf and Al-Ajmi, 1993); A5 (Hesp and Hastings, 1998); A6 (Gay, 1999); A7 (Sauermann et al., 2000).

barchan sites using ArcGIS providing estimates for dune heights with approximately 1 m a.s.l. precision. The accuracy of these measurements was validated using high resolution differential GPS survey within a localised area of the pan. 2.2. Field survey Differential GPS surveys were carried out by mounting dGPS rover equipment on quad bikes. Base stations were each surveyed for a minimum of 19 h. Post-processing was carried out in the field using the Leica GeoOffice post-processing system. Measurements were tied to the Botswana Government reference beacon 588 surveyed in 1978 (912 m a.s.l.). The precision of this reference is unknown but since the overall aim of this survey was to investigate the relative elevation relationships between landforms, the absolute error on the reference point is unimportant. Repeat measurements suggest a maximum latitudinal and longitudinal variation of 0.000015 (1.7 m) and 0.0000924° (0.96 m), respectively. Average absolute deviations, however, were 0.34 m horizontally and 0.25 m vertically. Measurement during times of poor satellite coverage exceeded these values with maximum vertical differences at these times up to 1.1 m. These data were not included in the survey. The resulting uncertainty of the presented measurements is therefore generally below +/−0.5 m and adequate to address the questions being investigated in this study. 2.3. Sedimentological analyses The carbonate and organic content of the sediment was determined using sequential loss on ignition at 550 °C and 950 °C (described in Heiri et al., 2001). Sediment colour was determined both in the field and after drying in the laboratory using standard Munsell colour charts. Particle size analysis (2–2000 μm grains) was carried out using a Malvern Laser Particle Size Analyser (Hydro 2000MU) and sediment statistics calculated using the Folk and Ward formulae (Folk and Ward, 1957).

2.4. Optically Stimulated Luminescence signal analysis Optical dating provides the potential to place depositional landforms within the chronology of long term environmental change. However, without first understanding the process which leads to their formation, the significance of such features as independent Quaternary geoproxies is diminished. The objective of this paper is to attempt to establish the process leading to the last depositional event of middle Kalahari lake bed barchans and in particular to distinguish between the competing aeolian and subaqueous formation hypotheses put forward respectively by Grove (1969) and Cooke (1980). Here we explore the potential to apply Optically Stimulated Luminescence (OSL) analysis as a tool to allow these geomorphic process-based hypotheses to be addressed rather than as a chronometric technique per se. The decay rate of the quartz OSL signal under illumination is strongly dependent on the energy of incident light (Spooner, 1994). Over visible wavelengths, the higher the photon energy (shorter wavelength), the greater the rate of signal depletion. The OSL signal comprises several components, named fast, medium and slow components, due to their relative decay rates during OSL measurement (Bailey et al., 1997). The main OSL signal used for dating is dominated by the fast and medium components. The dependence of the decay rate on illumination wavelength is, however, different for each of these components. Singarayer and Bailey (2003) demonstrated that using a stimulation wavelength of 600 nm (red light), the medium component bleaching rate is only ~ 5% that of the fast but ~13% at 500 nm (blue/green), ~ 47% at 400 nm (violet) and 100% at 358 nm (invisible UV). In terms of geomorphological processes, the implication of this is that in conditions of aeolian/sub-aerial deposition, where the UV (short wavelength ~10–400 nm) component of the spectrum is significant, it is likely that the bleaching rate of the medium component will be similar to that of the fast component (although not identical, as the daylight spectrum contains wavelengths greater than 358 nm, shifting the mean effective wavelength of daylight in to the region where the fast component is bleached faster than the medium component). During subaqueous

25.120 25.120 25.120 25.120 25.211 25.211 25.165 25.165 25.165 25.165 25.165 24.879 24.879 24.879 24.879 24.866 24.866 24.866 24.866 25.213 25.216 25.216 25.216 − 20.570 − 20.570 − 20.570 − 20.570 − 20.552 − 20.552 − 20.555 − 20.555 − 20.555 − 20.555 − 20.555 20.683 20.683 20.683 20.683 − 20.671 − 20.671 − 20.671 − 20.671 − 20.542 − 20.641 − 20.641 − 20.641 MAK/08/1/1 MAK/08/1/2 MAK/08/1/3 MAK/08/1/4 MAK/08/7/1 MAK/08/7/2 MAK/08/9/1 MAK/08/9/2 MAK/08/9/3 MAK/08/9/4 MAK/08/9/5 MAK/09/15/1 MAK/09/15/2 MAK/09/15/3 MAK/09/15/4 MAK/09/17/1 MAK/09/17/2 MAK/09/17/3 MAK/09/17/4 MAK/09/29/1 MAK/09/34/1 MAK/09/34/2 MAK/09/34/3

1.0 2.0 3.0 4.0 0.6 1.2 0.7 1.2 2.0 2.5 3.0 0.5 0.9 1.2 0.6 0.6 1.3 1.8 2.3 0.5 0.5 1.2 2.0

22.2 17.3 18.0 22.5 23.3 19.1 19.8 8.5 20.7 19.0 24.8 11.5 18.7 21.7 23.2 9.0 23.5 17.7 15.9 8.6 33.3 31.6 33.7

3.1 1.3 1.0 2.4 1.3 2.3 2.6 1.8 2.1 1.6 2.0 4.2 1.8 7.2 8.3 1.7 3.9 4.9 5.1 1.2 4.1 4.4 7.7

3.0 1.8 1.9 3.0 3.8 3.0 2.8 2.1 2.0 1.9 2.2 4.6 3.7 4.5 1.9 3.2 2.5 3.2 3.4 3.5 6.8 3.5 4.1

132.2 210.9 173.0 207.4 105.6 180.5 203.9 175.9 164.6 222.7 267.3 71.6 139.0 127.2 201.9 86.9 102.9 193.2 169.7 274.7 40.9 35.1 37.7

4% 4% 3% 20% 5% 9% 12% 7% 6% 8% 22% 1% 9% 10% 1% 2% 1% 10% 2% 25% 13% 1% 1%

18% 34% 25% 23% 14% 26% 29% 25% 22% 34% 31% 11% 19% 19% 32% 11% 12% 28% 27% 29% 13% 1% 2%

53% 58% 66% 46% 48% 47% 46% 58% 66% 54% 42% 42% 45% 39% 60% 50% 60% 42% 49% 31% 15% 25% 28%

22% 3% 6% 10% 28% 16% 12% 9% 6% 4% 5% 39% 24% 27% 7% 34% 25% 18% 19% 13% 50% 64% 57%

2% 0% 1% 1% 4% 2% 2% 0% 0% 0% 0% 8% 3% 5% 1% 3% 2% 2% 3% 2% 9% 9% 11%

1.03 1.00 1.01 1.04 0.91 1.04 0.85 1.14 1.02 1.05 1.04 1.01 1.03 0.97 1.06 1.02 1.06 1.09 1.00 1.02 1.01 1.05 0.94

0.04 0.03 0.08 0.08 0.81 0.25 0.10 0.13 0.08 0.10 0.19 0.16 0.17 0.20 0.05 0.09 0.14 0.09 0.10 0.70 0.07 0.09 0.04

Longitude Depth Carbonates Organics Sorting Median (microns) Coarse sand Medium sand Fine sand Silt Clay Z ratio Equivalent dose over-dispersion (°E) (m) (% by mass) (% by mass) (% by volume) (% by volume) (% by volume) (% by volume) (% by volume) (natural) Latitude (°S) Sample ID

Table 2 Sedimentary data, Ntwetwe barchan features. Over-dispersion and Z ratios were calculated from De measurements on 20–30 multigrain aliquots per sample.

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geomorphological processes, however, the daylight spectrum is significantly altered due to attenuation of shorter wavelengths (Berger, 1990; Sanderson et al., 2007) and there is also a reduction in the intensity of light reaching the quartz grains. Berger (1990) suggests the higher energy portion of the incident solar spectrum (b400 nm) is effectively removed by more than 30 cm depth of water. Compared to subaerial transport, subaqueous transport is therefore not only less efficient at resetting OSL signals in quartz in general but is likely to reset the fast component of the signal significantly more rapidly than the medium component (as the mean effective illumination wavelength is shifted to longer wavelengths than is the case for unattenuated daylight). A consequence of this difference is that the relationship between the illumination spectrum and bleaching rate may, in theory, be used to help distinguish between aeolian and subaqueous depositional processes. If a sample is illuminated during transport to the extent that both the fast and medium OSL signal components are fully at rest (zero signal) prior to deposition, this is referred to as ‘full resetting/bleaching’. In this case, all else being equal, the age derived from each component following the burial period, would be identical. In cases where the OSL signal is only partially reset prior to sediment deposition (i.e. the duration of light exposure was insufficient to reduce the signal to zero), a residual signal remains (‘partial resetting/bleaching’). If partial bleaching occurs under aeolian conditions, where the resetting rate is similar for both components, it is expected that both components would be reset to a similar level, meaning that the De values (De is a quantity derived from the measured OSL signal magnitude, which equates laboratory-measured radiation dose response to the dose received in nature during the burial period, and is used in calculating the burial age) of each component would be similar (though not identical). By comparison under subaqueous deposition it is expected that there will be a more significant difference in the bleaching rates of the fast and medium components, such that for incompletely reset samples, the apparent De derived from the medium component will be significantly greater than that from the fast component (and this difference will be greater than in the case of bleaching under aeolian conditions). 2.4.1. Sampling and De measurement OSL sampling was carried out at seven sites within accessible discrete barchan forms using a hand operated Dormer drill with a lighttight sampling head. Twenty three samples in total were taken from these seven sites within a broad 35 km east–west transect in the Ntwetwe pan. At two sites (MAK/08/1 and MAK/08/9) the full depth of the dune profile (4 m and 3 m, respectively) was sampled. At other sites the dune was sampled where it was exposed. For each sample, chemical pre-treatment, sieving, density separation and surface etching were undertaken to isolate the 180–210 μm quartz fraction before the Equivalent Dose (De) was estimated from small (2 mm diameter mask size) multigrain aliquots using the Single-Aliquot Regenerative-dose (SAR) protocol (Wintle and Murray, 2006). Measurements were made using automated Risø TL/OSLDA-15 readers (Botter-Jensen, 1997; Botter-Jensen et al., 2003). 2.4.2. Z ratio determination Differences in the apparent De of different components were assessed using successive time intervals of the measured OSL decay curve for De calculation (Bailey, 2003a). The form of the plot of De as a function of measurement time, De(t), (e.g. Fig. 4a,b) was used to infer the contributions of the components. The fast component typically dominates the earlier part of the OSL signal, giving way to the medium and eventually the slow components as successively later parts of the signal are integrated. For each sample a Z-ratio was calculated: a ratio of De calculated from the initial (0–0.05 s) part of the signal to that from the later (2–3 s) part of the signal (Fig. 4c). The integration window for the Z ratio was chosen by selecting the time (tz) at which the medium component was large relative to the fast

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Linear dunes Beach ridges Lake bed Lunette dunes Barchan dunes 0 Coarse sand Medium sand Fine sand

0 Silt Clay

50 Mean % by volume

100 0

300 0 Median grain size (microns)

10

20

0

Mean % carbonates by mass

10 Mean % organics by mass

Fig. 8. Comparative sediment characteristics for different landform types within the Makgadikgadi region.

component whilst still maintained a usable signal to noise ratio. tz was defined by deconvolving the OSL signal into its constituent exponential components, using a standard Levenberg–Marquart minimisation algorithm, and finding the time at which the value of the function ‘medium component minus fast component’ was at a maximum. In the case of aeolian sediments partially-bleached under subaerial conditions, De should be less dependent on the illumination time and the Z-ratio should be closer to 1 compared to the case of subaqueous bleaching. In both cases we expect Z > 1, but Z (subaqueous bleaching) should be significantly greater than Z (subaerial bleaching), in cases where the signals are partially-bleached to approximately the same level prior to measurement. 2.4.3. Experimental bleaching scenarios The success of using De(t) and Z-ratio calculations to ascertain depositional conditions is dependent on the relative contribution of the fast of the medium component signals to the bulk OSL signal. This is expected to vary both at the grain and sample level. A first order check of the feasibility of using this approach was undertaken by monitoring the effects of partially bleaching multigrain aliquots using daylight and simulated subaqueous wavelength bleaches on laboratory administered doses. To avoid pre-sensitisation, a room temperature (20 °C) optical stimulation, with the blue diodes used in standard OSL measurement (λ=470 nm) for 200 s (60% power), was used to bleach the multigrain aliquots (Roberts et al., 1999; Wallinga et al., 2000), followed by a 10 ks room temperature storage (allowing the thermal redistribution of charge

photo-transferred during bleaching). This re-trapping of charge into OSL traps during storage (reported as up to 5%, (Smith et al., 1990)) was minimised by a second bleach and 10 ks pause such that the maximum total amount of charge re-trapped by recuperation becomes insignificant. Laboratory administered beta doses (5 Gy) were followed by a 260 °C, 10 s preheat. Partial bleaching was then carried out either in the laboratory illumination (470 nm at ~18 mW/cm2 for 1 s at 20 °C) or natural daylight illumination (exposure to daylight for 4 s). The ratio of pre- to post-exposure signal was monitored using measured OSL from short (0.1 s) blue LED stimulations. Partial bleaching in both experiments was chosen to deplete the signal by 50% and typically exhibited a post-/pre-exposure ratio of between 0.4 and 0.7. Doses were then recovered using a single aliquot regeneration (SAR) protocol (c.f. Murray and Wintle, 2003) using an exponential background subtraction (Bailey and Burrough, in preparation) to reduce slow component contributions to the bulk signal. 3. Results and discussion 3.1. Field survey results Field observation reveals the Ntwetwe landforms to be relatively flat features, with an average elevation above the pan surface of ~ 7 m, and typically covered with grasses and occasional acacia thorn shrubs. Also striking are ubiquitous parallel features contouring the islands, which stand out on remotely sensed images because of slight

Fig. 9. Comparison of Ntwetwe Z values to samples partially bleached under daylight and simulated subaqueous conditions. The parameter Z provides a representative measure of multigrain aliquot De variation with illumination time for each Ntwetwe barchan sample (4 aliquots). Subaqueous conditions were simulated by partially bleaching aliquots with blue LEDS (470 ± 20 nm) at 50% power and measuring resulting Des and Z values. An additional set of measurements was performed on aliquots partially bleached by daylight.

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Fig. 10. De plotted against Z-ratios for samples taken from down-core auger profiles in six barchan dune sites. Individual aliquot measurements are shown (circles) as well as the mean De/Z ratio for each sample (squares).

vegetation differences but are, however, obscure to ground level observation. The dGPS measurements suggest that these concentric features do have topographic expression and are not merely vegetation patterning (Fig. 5). The significance and cause of this topographic variation remain unknown. We hypothesise that these features relate to differential degradation of sedimentary structures associated with former barchan dune accretion, such that the underlying westwarddipping cross-stratification produced during active mobile phases of the barchan dune (and indicating dune migration from east to west), has been subsequently exposed at the surface (Fig. 3c). The last significant position of the crest during a mobile phase of barchan development is thus identified by the last concentric feature on the western side of the barchans. Geometrically, the Ntwetwe dune forms differ from many active barchan dunes by being typically more circular in shape, as opposed to being elongated along their stoss slopes. It is possible that this apparent geometry is caused by post-

depositional degradation, in particular the seasonal flooding (up to a metre of water) that would render the shallowest parts of the dunes most vulnerable to erosion. The redistribution of sediment from the steepest western side of the dunes, presumably during the heaviest rains of the wet season, is evident as ephemeral sand drapes on the pan floor within and beyond the lee of the dune. 3.2. Morphometry The eastern–central part of the pan contains morphological forms that are more isolated and widely spaced compared to the western areas where dunes appear closer and more coalesced. Mean morphometric values and value ranges are given in Table 1. The spacing differences are reflected in wavelength measurements, but not in other parameters since, following Breed and Grow (1979), each element of a transverse ridge is treated as a separate landform. The Ntwetwe features ranged in width from 0.1 to 3.5 km and in length from 0.1 to

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1.6 km. The distance between dune crests (wavelength) ranged from 0.4 to 6.9 km. The analysis of transverse dune features (including barchans) by Breed and Grow (1979), which encompassed a wide range of contexts and analyses of over 20 sand sea contexts, provides a data set to which Ntwetwe features can be compared (Fig. 6). Table 1 provides a comparison of the mean calculated dune morphometric ratios of length/width, width/wavelength and length/wavelength for these dunefields. Whilst the morphometry of analogous subaqueous dunes is frequently discussed, comparable morphometric measurements are rarely presented, making any analyses of their similarity difficult. The exception to this is a study of a field of subaqueous sand dunes on the Carnegie Ridge (equatorial Pacific) measured using side looking sonar (Lonsdale and Malfait, 1974) and barchan-shaped sand banks observed in the north-west of Torres Strait (Daniell and Hughes, 2007). Summary data from this study are also included in Fig. 6 for comparison with measurements from Ntwetwe. As shown in Fig. 6, morphometric parameters (length, L; horn to horn width, W (Fig. 3) and wavelength (λ)) for the Ntwetwe features fall within the published ranges for barchan dunefields. The range of values for each of the three defined ratios (Table 1) for the 21 dunefields is relatively large, with L/W means ranging from 0.39 to 0.87 (mean L/W ratio for Ntwetwe is 0.67) W/λ from 0.75 to 2.4 (Ntwetwe = 0.52) and L/λ from 0.38 to 0.95 (Ntwetwe = 0.35). Given the degree of scale invariance expected within dune forms however, the Ntwetwe dunes bear out the expected relationship between horn width and length for aeolian dunefields. It is likely that the basis of some of the difference between the Ntwetwe ratios and those of other dunefields relates to the degraded nature of the dune forms in comparison to active aeolian systems. It is unlikely that the wavelength of dunes would have changed since stabilisation but L/λ and W/λ ratios are relatively low compared to other aeolian barchan dunes, suggesting that the L and W values have both reduced by degradation. Thus the L/W ratio is maintained but the absolute values lie at the smaller end of the scale. There are insufficient data to make this comparison with sub-aqueous dunes. Following Endo et al. (2005), we define two additional parameters: i) body length Lb, the length of the dune from toe to horn and ii) horn length Lh, the mean projective length between the horn tip and the crest (Fig. 3b). These parameters were rendered dimensionless by normalising to the dune width: Lb/W and Lh/W. Plotted in this way, the dimension-ratios of forms can be compared to other aeolian and subaqueous barchan forms including those developed by laboratory controlled subaqueous unidirectional and oscillatory flows (Fig. 7). Endo et al. (2005) suggest that, in the context of morphometric ratios, the body of barchans formed by unidirectional water currents is longer than barchans formed under water waves and subaerial flows (Endo et al., 2005; Fig. 7). The Ntwetwe landform ratios span the divide between subaerial flows and the two categories of water-formed features. A problem is that we are comparing the morphometric-ratios of degraded forms in Ntwetwe with data for active landforms or features that were newly formed in laboratory experiments. In particular, in the case of the Ntwetwe barchans, outwash into the bayhead area and degradation at the narrower points (i.e. the horns) are likely to have decreased the Lh/W ratio and increased the Lb/W ratio. Three principal points can be noted from the comparative analysis of Ntwetwe morphometry with other published data. First, aeolian dunefields dominated by transverse forms display a wide range of dune sizes, shapes and spacings, but the relationship between mean width and mean length and between mean wavelength and mean length can be described using a linear relationship. Secondly, the Ntwetwe features fall within the spectrum of morphometric parameters found in a range of sand seas but still form a statistically significant sub-population. Thirdly, there is insufficient published comparative data from subaqueous dunefields to be able to statistically separate these populations from subaerial dunefields. Comparable morphometric data (horn width, dune-length and dune wavelength) from two sets

of subaqueous barchanoid forms suggest that subaqueous forms potentially fall within the range of parameters observed from aeolian sand seas rendering morphometric analyses along these lines a poor discriminator of formative process. 3.3. Sedimentological analysis The barchan dunes are consistently composed of medium to fine grained sand with a median grain size that ranged from 35 to 275 μm with a strong dependence on sampling depth (Table 2, Fig. 8). Sediment taken from less than 1 m depth contained a greater fine grained component than samples below this depth. Many samples contained silt/clay aggregates and silcrete fragments, the latter originating from the floor of the pan where it precipitates into a range of forms (Shaw et al., 1990). In comparison to other landforms existing within the basin, mean barchan landform grain size parameters did not exhibit any significant difference to within basin sediments (lunettes and lake bed material) but all were significantly different from mean beach ridge sediment characteristics, which delimit the perimeter of the basin, and from linear dune sediment which lies outside the basin (Fig. 8). This difference lies predominantly in the proportion of the finer fraction making up the sediment. Within-basin sediments are consistently composed of greater silt and clay sized fractions and contain a much greater proportion of carbonate-rich and organic material. We interpret this as the mobilisation of sediment deposited under highly seasonal conditions not dissimilar from those found in the basin today, where shallow waters are rich in carbonate ions that precipitate to form sodium and calcium carbonates under high salinities (Shaw et al., 1990; McCulloch et al., 2008) and sustain a seasonal salt-tolerant algal mat growth principally made up of filamental chloroflexus colonies (Shaw et al., 1990). Under freshwater subaqueous conditions, it is likely that carbonate ions would remain in solution and carbonate content would be low as in the case of beach ridge sediments. Microfossils were observed in very low abundance (a single ostracod valve and a single diatom frustule were found under scanning electron microscopy) but were not systematically investigated because of the likelihood for such material to be reworked with sedimentary particles either via aeolian or subaqueous transport and thus offer little potential to differentiate between aeolian and lacustrine processes. 3.4. Using Optically Stimulated Luminescence to test formative geomorphological processes 3.4.1. Experimental bleaching experiments The parameter Z (summing over several independent measurements from each condition/sample) provides a representative measure of De as a function of illumination time and is plotted in Fig. 9 for partial blue diode bleaches, partial daylight bleaches and fully bleached aliquots of each barchan sample. These data show a dependence on the residual De (degree of partial bleaching) as expected (Bailey, 2003a,b; Bailey et al., 2003) but no signifcant difference in Z ratio between the blue diode partial bleaches and the daylight partial bleaches, with both experiments resulting in Z > 1. The lack of expected differentiation is explored in greater detail in (Bailey and Burrough, in preparation) but suggests that in this instance, due to the balance of component sizes and the difficulties of adequate slow component (background) subraction, De(t) relationships (e.g. Fig. 4a, b) may be unsuitable as a method to distinguish the geomorphological nature of the bleaching event. Full bleaching scenarios however, carried out during dose recovery tests, indicate Z ratios much closer to 1 (ranging from 0.90 to 1.10) (Fig. 9) and extremely similar to the Z ratios from the natural signal (which ranged from 0.85 to 1.14) This suggests that these samples were very well bleached in the natural environment and,

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thus, were either deposited by aeolian processes or withheld no residual dose prior to subaqueous deposition. Both scenarios are to some degree possible given the long transport pathway of the Okavango and Boteti fluvial systems that would allow sediment from distant sources to be completely bleached before deposition. In the latter case however, we would not expect the seidment source to be entirely allochthonous but given the sediment limiting conditions of barchan formation would expect signficant mobilisation and redeposition of existing lake floor sediments that would have occurred at a depth sufficient to allow barchan dune formation. The incorporation of reworked lake bed material would most likely lead to wide De distributions and some indication of partial bleaching. Negative values of the Z ratio occur where De declines as a function of integration time and may be due in part due to expected statistical noise, but may also suggest some thermal instability in the medium component, for which there is mounting empirical evidence (e.g. Tsukamoto et al., 2003; Bailey, 2010). Note that in standard SAR measurements, for dating purposes, the erroneous effects of the medium component are minimised by using only the initial part of the decay curve or by isolating a fast-component-only signal (Bailey, 2010). 3.4.2. Interpretation of De distributions Fig. 10 shows De plotted against Z-ratios for samples taken from down-core auger profiles in six barchan dune sites. Four samples (MAK/08/7/1, MAK/08/7/2, MAK/08/15/2 and MAK/09/29/1) exhibit over-dispersion values >20%. There is no significant relationship between individual De values and their Z-ratios in any quartz samples, suggesting the wider distribution of De values in these samples is a function not of partial bleaching but of grain mixing (e.g. bioturbation) or variation in the received dose by grains due to dosimetric variability within the sediment matrix (Bailey, 2003a). The mean over-dispersion for the remaining 19 samples is just 9% and De values are consistent within errors throughout the entire depth of the dune body, suggesting a translocation mechanism for these landforms which is effectively instantaneous in geological time. 4. Conclusion Although they have received very little previous analysis, barchanlike dune forms are in fact a relatively common occurrence on dry lake floors in the low latitudes. Many of these dunes are now vegetated and not actively forming but their analysis may provide a valuable insight into environment al conditions and processes of the past. Ntwetwe Pan in northern Botswana contains a complex suite of degraded, crescentic landforms. Morphometric analysis suggests that whilst the form of these features fits squarely within the range of aeolian barchan dune morphologies observed from a wide sample of global dunefields, alternative hypotheses for their formation, including the possibility of a subaqueous origin, cannot be ruled out. There is currently insufficient morphometric data available to be able to demonstrate any significant morphological difference between the Ntwetwe landforms and those of comparable subaqueous forms. Distinguishing process domains from morphological observations presents an interesting challenge for geomorphologists engaged in establishing inaccessible and unknown environmental conditions including those in extra-terrestrial contexts. The use of Optically Stimulated Luminescence (OSL) signal analysis also now offers the potential to provide information on the processes leading to the last depositional event of sedimentary quartz grains. Typically this is achieved via moving the integration window for De calculations from the bulk signal of OSL measurements on quartz, making it feasible to compare the De values of different signal component and thus obtain additional information on the nature of the bleaching spectrum during the last depositional process being investigated. This is possible because the bleaching rates under

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aeolian conditions are more efficient than those under sub-aqueous conditions leading us to expect the De of all signal components to be equal if aeolian processes were responsible for the formative depositional event. If sediments were deposited subaqueously, we would expect the De of the slower components to be less well bleached than those of the fast component and thus the Des would not be equal (Bailey, 2003a). However, in a set of experiments carried out to verify these underlying assumptions we demonstrate that due to the balance of component sizes and the difficulties of adequate slow component (background) subraction this technique is not currently viable with the sediments that the Ntwetwe barchan features are constructed from. Despite the development of sophisticated tools during the last 40 years, the hypothesis of Grove (1969) that discrete crescentic forms of the Makgadikgadi basin are aeolian in origin can neither be accepted nor rejected conclusively on the basis of the data presented in this paper. Their relevance as geoproxy indicators, particularly useful for establishing the envelope of environmental variability in the Kalahari region, therefore remains enigmatic, providing an ongoing challenge to geomorphologists to find adequate generally applicable tools for distinguishing between records of subaerial and subaqueous processes. Acknowledgements This project was supported by a Royal Geographical Society small grant, a NERC Geophysical Equipment Facility Loan 897 and the Boise Fund, University of Oxford. We thank two anonymous reviewers and Andy Plater for very helpful comments that greatly improved this manuscript. References Allen, J.R.L., 1968. Current Ripples: Their Relation to Patterns of Water and Sediment Motion. North-Holland Publish Corporation, Amsterdam, p. 433. Ash, J.E., Wasson, R.J., 1983. Vegetation and sand mobility in the Australian desert dunefield. Zeitschrift Fur Geomorphologie Supplementband 45, 7–25. Bailey, R.M., 2003a. Paper I: the use of measurement-time dependent single-aliquot equivalent-dose estimates from quartz in the identification of incomplete signal resetting. Radiation Measurements 37, 673–683. Bailey, R.M., 2003b. Paper II: the interpretation of measurement-time-dependent single-aliquot equivalent-dose estimates using predictions from a simple empirical model. Radiation Measurements 37, 685–691. Bailey, R.M., 2010. Direct measurement of the fast component of quartz optically stimulated luminescence and implications for the accuracy of optical dating. Quaternary Geochronology 5, 559–568. Bailey, R.M., Singarayer, J.S., Ward, S., Stokes, S., 2003. Identification of partial resetting using De as a function of illumination time. Radiation Measurements 37, 511–518. Bailey, R.M., Smith, B.W., Rhodes, E.J., 1997. Partial bleaching and the decay form characteristics of quartz OSL. Radiation Measurements 27, 123–136. Bailey, R.M., Burrough, S.L., in preparation. The utility of De(t) data in optical dating. Quaternary Geochronology. Berger, G.W., 1990. Effectiveness of natural zeroing of the thermoluminescence in sediments. Journal of Geophysical Research 95, 12,375–12,397. Bhalotra, Y.P.R., 1987. Climate of Botswana, Part II: Elements of Climate. Meteorological services, MWTC, Gaborone. Bolduc, A., Duchesne, M.J., 2009. Discovery of megadunes in the middle estuary of the St. Lawrence River, Quebec, Canada (Decouverte de megadunes dans l'estuaire moyen du fleuve saint-laurent, Quebec, Canada). Journal of Water Science 22, 125–134. Botter-Jensen, L., 1997. Luminescence techniques: instrumentation and methods. Radiation Measurements 27, 749–768. Botter-Jensen, L., Andersen, C.E., Duller, G.A.T., Murray, A.S., 2003. Developments in radiation, stimulation and observation facilities in luminescence measurements. Radiation Measurements 37, 535–541. Bourke, M.C., Goudie, A.S., 2009. Varieties of barchan form in the Namib Desert and on Mars. Aeolian Research 1, 45–54. Bourke, M.C., Lancaster, N., Fenton, L.K., Parteli, E.J.R., Zimbelman, J.R., Radebaugh, J., 2010. Extraterrestrial dunes: an introduction to the special issue on planetary dune systems. Geomorphology 121, 1–14. Bowler, J.M., 1976. Aridity in Australia: age, origins and expressions in aeolian landforms and sediments. Earth-Science Reviews 12, 279–310. Bowler, J.M., 1986. Spatial variability and hydrologic evolution of Australian lake basins: analogue for Pleistocene hydrologic change and evaporate formation. Palaeogeography, Palaeoclimatology, Palaeoecology 54, 21–41.

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