Aeolian Landforms

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IV The work of the wind

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17 Aeolian landscapes and bedforms

David S.G. Thomas

17.1 Introduction nologies now available from space; e.g. Washington et al., 2003) and (b) the internal structures of aeolian deposits Arid zone aeolian deposits have long attracted the at- to be understood (i.e. via ground penetrating radar, e.g. tention of geographers, geomorphologists and allied sci- Bristow et al., 2000); and (4) chronometric advances, such entists. They represent for many people the seminal as optically stimulated luminescence dating (e.g. Aitken, landscape of deserts, though the actual geographical dis- 1994), which are allowing the timing and frequency of tribution and extent of, for example, dune field is no more aeolian deposit accumulation to be established, as well as than 20 % of the total area of drylands today, and in some the relationships to drivers of system dynamics. deserts, such as the southwestern USA, they cover less Given these advances, we are now better placed than than 1 % of the total area. Scientificall , Bagnold’s (1941) ever to understand the spatial and the temporal behaviour seminal work on sand transport processes and dune forma- of aeolian systems, the drivers of dynamic behaviour tion established a depth of understanding and rigour that and the relationships in behaviour across scales from the set the framework for research in the ensuing half-century. movement of individual sand and silt grains to the de- Subsequently, at markedly different scales, the availabil- velopment of dunefield and sand seas. To develop an ity of remotely sensed data coving otherwise difficul to understanding of how aeolian systems behave and accu- access desert areas (e.g. Breed and Grow, 1979; Paisley mulate, it is nonetheless necessary to make separate con- et al., 1991) and the growing number of reductionist stud- siderations of processes, landforms and landscapes. In this ies involving the detailed monitoring of wind and sand chapter the larger spatial and temporal scales are exam- fl w over dunes (e.g. Weng et al., 1991; Wiggs, 1993) led ined, by looking at the distribution of aeolian deposits at to advances in aeolian research at very contrasting spatial the global scale and what determines that distribution and, scales (Thomas and Wiggs, 2008) (Figure 17.1). by examining the long-term Quaternary timescale, the de- More recently, four scientifi developments have pro- velopment of aeolian systems and their relationships to vided the opportunity for across-scale enhancements in climatic changes. The chapter considers both sandy de- understanding the nature and behaviour of aeolian systems posits – sand seas and their constituent parts – and the (Thomas and Wiggs, 2008): (1) further advances, largely landscape features comprised of wind-lain silt, termed driven by technological developments (e.g. Arens, 1996; loess. van Boxel, Sterk and Arens, 2004), in empirical measurement that allow the relationships between sediment movement and wind power to be elucidated; (2) mathematical modelling of dune (e.g. Baas and Nield, 2007) 17.2 Aeolian bedforms: scales and and atmospheric dust (e.g. Zender, Bian and Newman, relationships 2003) dynamics and their interactions with key controls; (3) technological developments in remote monitoring sys- Because of their fine-graine nature, loessic deposits (pri- tems that allow (a) high-resolution observations of sedi- marily silt-sized material with 3.9-63 µm particle diam- ment movement (via the plethora of remote sensing tech- eters) tend to be rather amorphous, landscape-blanketing

Arid Zone Geomorphology: Process, Form and Change in Drylands, Third Edition. Edited by David S. G. Thomas © 2011 John Wiley & Sons, Ltd. Published 2011 by John Wiley & Sons, Ltd. ISBN: 978-0-470-51908-0 P1: TIX/XYZ P2: ABC c17 JWST032-Thomas December 17, 2010 13:24 Printer Name: Yet to Come

428 CH17 AEOLIAN LANDSCAPES AND BEDFORMS

REDUCTIONIST INTEGRATED TEMPORAL LANDSCAPE Grain Dune SYSTEM DYNAMICS

BAGNOLD 1941 1940 Physics of windblown sand and desert dunes

1950

GROVE 1958 Sahara 1960 FROM AIR PHOTOS MAPPING SAND SEAS

JENNINGS 1968 Australia LETTAU & LETTAU GROVE 1969 1969 Kalahari 1970 Bulk sand transport WILSON 1971, 1973 Ideas across spatial scales SATELLITE IMAGERY

FRYBERGER 1979 MCKEE 1979 Dune forms & A study of global SINGHVI 1980 wind regime sand seas et al 1982 First TL dating of aeolian Process sediments studies Geological time 1990 Role of vegetation Complex WIGGS KOCUREK & HAVHOLM remote sensing OSL Parameter- et al 1995 1993 products Dating ization Long-term Sequence stratigraphy e.g. TOMS of thresholds empirical studies Modelling 2000 KOHFELD & tools Sedimentary THOMAS HARRISON 2001 systems/dunes Dirtmap monitored et al 2005 MAHOWALD Dune data survey et al 2006 responses Global dust to global dynamics warming 2010 ? ESS?

Figure 17.1 The routes and scales of aeolian research since Bagnold (1941). The themes focus on reductionist process research, research that treats aeolian systems in an integrated (time/space or scale) manner, research focused at the landscape scale and research with a strong focus on understanding the temporal history of aeolian systems. Since 2000 there has perhaps been a growing emphasis on research that attempts to integrate all these elements and which might be termed ‘earth system science’ (after Thomas and Wiggs, 2008). P1: TIX/XYZ P2: ABC c17 JWST032-Thomas December 17, 2010 13:24 Printer Name: Yet to Come

AEOLIAN BEDFORMS: SCALES AND RELATIONSHIPS 429

features (Pye and Sherwin, 1999). Loess is largely devoid of landforms that are a direct result of shaping by the wind; being comprised of fine-graine sediment, it is generally unable to hold a morphology once deposition from suspension transport by the wind has occurred, other than that determined by the underlying topography over which it lies. Thus the great loess plateau of China and the loessic deposits of eastern Europe and America are rather fla and featureless. There are exceptions to this, and low and high amplitude linear structures (‘loess hills’ and ‘loess dunes’) have been reported from the loess deposits of eastern Germany and Bulgaria. These are parallel to the wind direction that was responsible for deposition (Leger, 1990). In Bulgaria, where the features are locally known as ‘breda’, they are reported as having a wavelength of 300–1200 m and a trough–ridge amplitude of up to 40 m (Rozsycki, 1967). However, loess is very susceptible to erosion by running water, especially, paradoxically, in drier areas where vegetation cover is sparse, and most landforms associated with the material are erosional in Figure 17.2 Bedform patterns revealed in a sand sea by satel- nature, such as gullies, piping and features resulting from lite imagery. Unravelling the complex evolution of such a sys- landsliding and slumping (Derbyshire et al., 1993). tem requires scientiﬁc data at a range of temporal and spatial The wider grain-size range of sand particles transported scales. by the wind (particle diameters from 2000 to 63 µm) means that constructional aeolian forms – dunes and ripples – can develop due to the packing that can occur scale depositonal landforms – predominantly dunes – oc- between particles of different sizes. Over 99 % of arid cur. Independent, isolated, dunes can occur outside sand zone aeolian sand deposits are found in extensive sand seas, e.g. where aeolian sand supply is limited and a to- seas or ergs (Wilson, 1973), the largest landform unit of pographic obstacle causes sand accumulation, but usu- aeolian deposition (Table 17.1), within which smaller- ally dunes occur in association with other dunes. This is significan when dune-forming processes are considered because it is increasingly being recognised that the in- Table 17.1 Scales of aeolian depositional landforms. teractions that can occur between adjacent dunes, due to self-organisation (Werner, 1995), are a key component of explaining how dunes form. As such, the beautiful patterns Scale Description Namea that exist within dunefield (Figure 17.2 and see Chapter I Deposits of regional Sand seas or ergs 19), and which remote sensing from space has made so extent visible, are not merely aesthetically pleasing but result II Deposits commonly Sand sheets, streaks from the processes of organisation that occur in complex devoid of scale III and stringers environmental systems (Ewing, Kocurek and Lake, 2006; bedforms Baas, 2007). IIIa Bedforms Complex, compound The form and size of dunes can vary within and be- superimposed on and megadunes tween sand seas, through the interaction of and varia- scale I deposits IIIb As scale IIIa or Simple dunes tions in sand supply, vegetation cover and wind regime. independently While there may be spatial variability through a dunefiel IV Bedforms Normal, f uid drag and in dune size, statistical analyses tend to show that pat- 1 2 2 superimposed on megaripples terns are consistent over quite large areas of 10 –10 km scale I, II or III (e.g. Thomas, 1986; Bullard et al., 1996), resulting from deposits self-organisational principles (Ewing, Kocurek and Lake, 2006; Zheng, Bo and Zhu, 2009). Over time, patterns may aMany deposits have alternative names. These are given in the table of Breed and Grow (1979). mature and multiple patterns may become superimposed (Ewing, Kocurek and Lake, 2006), possibly resulting P1: TIX/XYZ P2: ABC c17 JWST032-Thomas December 17, 2010 13:24 Printer Name: Yet to Come

430 CH17 AEOLIAN LANDSCAPES AND BEDFORMS

from dunefiel development during different arid phases repetitive periods of accumulation, as in the case of the (Kocurek, 1998). Sand seas or parts of sand seas that are megadunes of the Badain Jaran Desert in China (Yang, Liu largely devoid of dunes are called sand sheets, stringers and Xiao, 2003; Dong, Wang and Wang, 2004). On the ba- or streaks. Ripples are the smallest bedform unit and can sis of height–wavelength relationships, Dong et al. (2009) occur on any larger-scale aeolian bedform. confir that megadunes are simply a higher hierachical Scale III and IV features (Table 17.1) are bedforms, aeolian form controlled by wind conditions. More con- a term that applies to any regularly repeated erosional troversially, by using fiel data and numerical modelling, or depositional surface feature, whether aeolian or sub- Andreotti et al. (2009) argue that megadune formation aqueous (Wilson, 1972a). This section considers general is controlled by the depth of the atmospheric boundary aspects of aeolian bedform development, while the forma- layer, while Bubenzer and Bolten (2008) suggest distinct tion of ripples and the controls on development of different atmospheric conditions have been responsible for the for- dune types are discussed in greater detail in Chapters 20 mation of draa in the eastern Sahara and the linear dunes and 21. that overlie them. These debates demonstrate that there is still much that is unclear regarding the nature of aeolian system development. 17.2.1 Scale effects in aeolian bedform Aeolian deposition and the initiation of landforms takes development place when mobile particles come to rest in a sheltered position on the ground surface, through accretion Contrary to early opinion (Cornish, 1914), ripples do not (Bagnold, 1941) or tractional deposition (Kocurek and develop over time into dunes. Ripples and dunes develop Dott, 1981). It also occurs when wind velocities decrease, independently, at different scales, in response to differ- causing a reduction in shear stress and a fall in transporta- ent factors. Wilson (1972a, 1972b) developed the concept tional potential, termed grain-fall deposition, and when of bedform scale into one of a bedform hierarchy with particles come to rest after rolling down a slope under four components: ripples (two orders), dunes and draas the effects of gravity, called avalanching or grainﬂow de- (Table 17.2), draas being large whaleback sand bodies, or position, which is important for particle movement down megadunes. topographically induced or bedform slopes (see Tsoar, The spacing of the three highest-order bedforms was 1982). In many regards, the processes of deposition and attributed to secondary atmospheric fl ws of differing accumulation are scale-independent; i.e. the same general sizes. As bigger sand grains require higher shear veloc- conditions leading to sand deposition must apply whether ities for their mobilisation, particle size was deemed to a sand sea, dune or ripple is being initiated or a nuclei set the scale of lower-atmospheric eddy currents (Wilson, or focal point for accumulation is required (Table 17.3). 1972a, 1972b). Grain size would therefore be coarser in The processes of dune and ripple formation and initiation the higher-order, more widely spaced, bedforms. are discussed in subsequent chapters; those for sand seas The universal applicability of this ‘granulometric con- are discussed below in the context of understanding their trol’ hypothesis has not survived the scrutiny of sub- global distribution. sequent investigations. Draas cannot be separated from dunes in terms of grain size (Wasson and Hyde, 1983), probably being compound or complex dune forms that 17.3 The global distribution of sand seas have developed in areas of high net sand accumulation, complex wind regimes and long-term aridity (see, Sand seas are neither clearly nor consistently define in for example, McKee, 1979; Lancaster, 1985) or through the literature. Following Wilson (1971), Fryberger and

Table 17.2 Wilson’s hierarchy of aeolian bedforms.

Order Name Wavelength (m) Height (m) Origin

1 Draas or mega dunes 300–5500 20–450 Aerodynamic instability 2 Dunes 3–600 0.1–100 Aerodynamic instability 3 Aerodynamic ripples 0.015–0.25 0.002–0.05 Aerodynamic instability 4 Impact ripples 0.05–2.0 0.0005–0.1 Impact mechanism

Source: Wilson (1972a). P1: TIX/XYZ P2: ABC c17 JWST032-Thomas December 17, 2010 13:24 Printer Name: Yet to Come

THE GLOBAL DISTRIBUTION OF SAND SEAS 431

Table 17.3 Possible nuclei for sand dune development. dominates over sand accumulation, sand sheets occur and dune forms are normally absent. 1. Topographical obstacles Figure 17.3 shows the global distribution of major ac- 2. Smaller obstacles Pebble or boulder tive and fi ed sand seas. The term ‘fi ed’ is applied here Plant or bush to sand seas where dune forms are identifiable but are Animal carcass, e.g. dead presently inactive or have limited surface activity, because camel current environmental factors do not favour the significan Human artefact, e.g. oil operation of aeolian processes. This is largely evidenced drum by the presence of a significan vegetation cover on dunes, 3. Remnant of previous e.g. in large parts of Australia and the northern Kalahari bedform sand sea in southern Africa, but also through other evi- 4. Surface irregularities dence such as the occurrence of soils on dune surfaces, and hollows and degradation of dune forms. Although it is the inter- Source: After Wilson (1972b) with modifications action between atmospheric (erosivity, by wind energy) and surface (erodibility, linked in particular to vegetation cover) conditions that determines the potential for sand transport (Lancaster, 1988; Thomas, Knight and Wiggs, Ahlbrandt (1979) and others, an active sand sea may be 2005), most of the major sand seas that are obviously define as an area of at least 125 km2 where no less than active today lie within the 150 mm mean annual isohyet 20 % of the ground surface is covered by windblown (Wilson, 1973) (Figure 17.4), where given a supply of sed- sand. Lancaster (1995) has define sand seas as ‘dynamic iment and suitable wind conditions, aeolian transport is sedimentary bodies that form part of local- and regional- uninhibited by plants. The implication of f xed sand seas scale sand transport systems in which sand is moved is that climatic and environmental conditions in the past from source zones to depositional sinks’. Dune forms are have been more conducive to aeolian sand transport than commonly present, but in areas where sand throughfl w today, particularly during the Late Quaternary Period.

SAND SEAS Active Probably limited activity Fixed LOESS Major deposits Thin or uncertain deposits

Figure 17.3 The global distribution of aeolian deposits (after Snead, 1972, modiﬁed and supplemented with data from numerous sources). P1: TIX/XYZ P2: ABC c17 JWST032-Thomas December 17, 2010 13:24 Printer Name: Yet to Come

432 CH17 AEOLIAN LANDSCAPES AND BEDFORMS

20¡W 10¡ 0¡ 10¡ 20¡ 30¡E

a) 10 Hoggar 6 2 7 20¡ Plateau Red Sea ile 1 N r 10 e iv 4 R 20¡ 11 3 14 S 5 8 en Tombouctou 9 12 eg a 13 15 l

R Khartoum . R 17 MEGA iv 16 e Daffur Blue Nile Nile r CHAD N 18 10¡ ig Lake e 18 N r Chad 19 10¡

W Ethiopian er h iv i Highlands e R te Benu N i l 0 1000 km e

RELIEF: ACTIVE DUNE SYSTEMS FIXED DUNE SYSTEMS 1. North Mauritanian Ergs 11. South Mauritanean Ergs 1000m Active dunes 2. Erg Chech - Adrar 12. Trarza and Cayor Ergs 3. Erg El Mréyé 13. Aoukâr and Aklé 400m Very limited activity and fixed dunes 4. Erg Timbouctou 14. Erg Azouad 5. Erg Bilma - Ténéré 15. Erg Gourma 6. Edeyen Murzuq 16. West Azouak 150mm annual isohyet 7. Erg Rebiana 17. East Azouak 8. Erg Foch 18. Erg Kanem Shoreline of Palaeolake Mega Chad 9. Erg Djourab 19. Sudan Qoz 10. Great Sand Sea

Figure 17.4 The distribution of active and ﬁxed sand seas in the Sahelian–Sahel belt of Africa. Note that the degree of activity/inactivity is not considered (after Grove, 1958, and Grove and Warren, 1968).

17.3.1 Sand sea development of bedrock may contribute material for sand sea development (e.g. in the eastern Namib Sand Sea; see Besler, The formation of a sand sea is contingent on sedi- 1980), but fluvial lacustrine and coastal sources are the ment availability and aeolian transport capacity. Kocurek principal locations from which sediment is derived by (1999) breaks these conditions down further such that the the wind and moved towards areas where sand seas can sediment issue is considered both in terms of its sourcing accumulate (Lancaster, 1999). (which he confusingly terms ‘supply’) and its availability Aeolian transportation sorts sediment effectively, so to the wind – i.e. just because suitable sediment exists, that usually over 90 % of particles in active sand seas are it may not be available to the wind until, for example, sand-sized. Finer silt, possibly generated by the grain-on- vegetation cover declines or sediment in a lake is exposed grain impact of sand in transport is commonly transported by declining water levels. Overtime, it is the interaction out of the desert aeolian system in suspension, where it between supply (senso Kucurek), sediment availability to may contribute to loess deposits (Crouvi et al., 2008). Clay the wind and the transport capacity of the wind that deter- is normally only transported by the wind in pellets, which mines whether a sand sea is stable, accumulating or de- then saltate and creep like the more familiar quartzose grading. Simply put, sand sea construction occurs when a sands (Bowler and Wasson, 1984; Al Janabi et al., 1988). large supply of sediment is available to wind of sufficien Dare-Edwards (1984) has called such deposits sand loess. energy to transport it to a place where transport energy Sand may move between source areas and sand seas is reduced and deposition occurs. These conditions may along preferred transport pathways. In some areas, such vary over time (Figure 17.5) at a range of scales, such as the Mojave Desert in California, these pathways may that climate cycles may affect both wind energy and the cross drainage basins and divides prior to accumulation in availability of sediment on the ground. sand seas (Zimbelman, Williams and Tchskerian, 1994). The highest rates of sand transport occur across the dune- less sand sheets and stringers, which often lead into areas 17.3.2 Sediment supply in sand seas of sand seas with dunes (Sarnthein and Walger, 1974, give a range of 62.5–162.5 m3/m width per year for Maurita- The initial sourcing of sediment for sand sea develop- nia). Rates of sand throughfl w in the Selima Sand Sheet, ment is normally non-aeolian. Locally, aeolian abrasion Egypt, for example, are as high as 1000 m/yr (Maxwell P1: TIX/XYZ P2: ABC c17 JWST032-Thomas December 17, 2010 13:24 Printer Name: Yet to Come

THE GLOBAL DISTRIBUTION OF SAND SEAS 433

Aeolian Aeolian Cumulative Aeolian Sediment sediment transport system system Climate production availability capacity response response

Destructional

Sand-starved conditions Arid

Lagged transport- limited influx Lagged availability- Constructional (dune growth & accumulation) limited influx

Weath- ering Stored Stabilization sediment Humid influx availability-limited Contemporaneous Time

0 --> + 0 --> + 0 --> +

Figure 17.5 Theoretical model of aeolian system response to climatic changes (after Kocurek, 1998). The model aims to show how sediment production, availability and aeolian transport interact to leave (or not) a sedimentary record of past conditions in sand seas.

and Haynes, 2001) and occurs through the movement valleys or seasonally inundated channel systems. Wadi of extremely low-amplitude giant ripples (500–1000 m sediments may be particularly important local aeolian wavelength; see Breed, McCauley and Davis, 1987) or forms as they dry out through drainage rather than evapo- even higher amplitude ‘chevrons’ (Maxwell and Haynes, ration and are therefore uncemented (Glennie, 1970). It is 1989). This may appear counterintuitive to earlier state- in fact likely that alluvial sources have made direct or indi- ments that sand sheets do not possess bedforms; however, rect contributions to the sediments in parts of many sand these transport features are not perceptible on the ground seas, especially in low relief situations (Cooke, Warren and are only visible via remote sensing (Maxwell and and Goudie, 1993). Haynes, 2001). When vegetation is present, it is not so Namib Desert dune sands have arrived at their des- dense as to significantl inhibit aeolian sand transport. tination through the combined effects of river, sea and Mobile barchan dunes represent bulk transport rates of wind transport (Lancaster and Ollier, 1983). In the Wahiba up to 3.49 m3/m width per year. Substantial quantities Sands of Oman, hyperaridity has precluded fluvia con- of sand may be transported huge distances across sand tributions to Holocene sand sea accumulation (Warren, sheets before accumulating into dunes, as in the case of 1988). Instead, sediment supply is attributed to coastal the Makteir erg in Mauritania (Fryberger and Ahlbrandt, erosion of soft sandy materials (Figure 17.6). The ex- 1979). Sand transport pathways may be sinuous, showing tensive dune systems of the Kalahari represent periods sensitivity to local topography and being analogous to of aeolian reworking of pre-existing continental basin ‘rivers of sand’ (Zimbelman, Williams and Tchakerian, sediments, which have accumulated by various means 1994). since the Mid-Jurassic (Thomas, 1987). These have been Dune sediments can reach sand seas along dominant supplemented in the southwestern Kalahari by aeolian non-aeolian pathways. In the Taklamakan Desert, China inputs operating along thin sand sheets (Fryberger and (Zhu, Zou and Yang, 1987), central Australia (Twidale, Ahlbrandt, 1979) (Figure 17.7). Such intraergal and 1972; Nanson et al., 2008) and other areas, so-called extraergal sands can sometimes be distinguished by ‘source bordering dunes’ have been deflate from dry river grain size and mineralogical differences (see Bowler and P1: TIX/XYZ P2: ABC c17 JWST032-Thomas December 17, 2010 13:24 Printer Name: Yet to Come

Phase 2

Phase 1 Phase 3

Gulf of Shelf Oman

n o o s n o m . . W 005500 kkmm S

Figure 17.6 Phases of deposition in the Wahiba Sands, Oman. Sands have been derived by the inland transport of coastal carbonate sands in at least three phases; in the third, Holocene, phase, cemented cliff sands have been deﬂated and transported inland (redrawn after Cooke, Warren and Goudie, 1993).

Sand stringers or sheets (extra ergal source)

Koppies Ð kraalpan

Lunette dune

Soutpan l l

v i

p o Lunette dune l o (intra ergal source) M

0 10 km Pan

Figure 17.7 Space shuttle hand-held camera photograph of part of the southwestern Kalahari duneﬁeld, southern Africa. Linear dunes dominate this sand sea, but larger, more amorphous sand sheets or stringers, also trending west–east, can also be identiﬁed. These may represent paths of extraergal sand inputs. Linear dunes also extend downwind from local sand sources, notably pan- margin lunette dunes and dry river valleys.

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THE GLOBAL DISTRIBUTION OF SAND SEAS 435

Wasson, 1984). Finally, it can be noted that for many sand seas experience less than 27 m3/m width per year poten- seas there is evidence for multiple phases of accumulation tial sandfl w (Table 17.4); however, estimating sandfl w (e.g. Figure 17.5). One reason for hiatuses in accumula- rates over large areas using limited meteorological data tion is climate change. It has been suggested that during can be problematic. Many low-energy sand seas occur the last 30 000 years active accumulation of sediment in near the centre of semi-permanent high- and low-pressure sand seas may have been restricted to less than a third of cells, while the high-energy sand seas fall in the Trade the time (Lancaster, 1990), with episodicity of accumula- Wind zones near the margins of these pressure systems. tion even affecting sand seas in the driest regions (Haney The transference of sand from high- to low-energy loca- and Grolier, 1991), because even with aridity present, tions within sand seas has been demonstrated by stud- sufficien wind energy is required for sand transport and ies in the Jafurah erg, Saudi Arabia (Fryberger et al., deposition. Other reasons for disruptions in accumulation 1984), the Namib Desert (Lancaster, 1985) and the Gran are tectonic activity and, in coastal situations, base level Desierto, Mexico (Lancaster, Greeley and Christensen, changes. The overall duration of sand sea accumulation 1987). Clearly, therefore, the general classificatio of may possibly be established by examining the mineralogi- wind environments by Fryberger and Ahlbrandt (1979) cal maturity of the materials they comprise (Muhs, 2004), masks considerable intraregional variations in sandfl w since a dominance of quartz suggests long-establishment potential. This may account for the difference between or sand seas that they have experienced long periods of their classificatio of the Simpson Desert (Table 17.3) stability, as other weaker minerals have been removed by and Ash and Wasson’s (1983) assertion that the major weathering (Muhs, 2004). limitation on sand movement there today is the low wind speeds.

17.3.3 Sandﬂow conditions and sand sea development 17.3.4 Sand sheets

It has been noted that sand sea development requires both Sand sheets can develop in aeolian environments where favourable sediment budgets and wind conditions. Where conditions do not favour dune development, though they net sand transport rates are high, throughfl w exceeds de- may exhibit low-relief aeolian features such as ripples position and bedform development is limited to sheets or and zibar (see Chapter 18). They can be small, local, fea- stringers with ripples or highly mobile dune types such as tures of a few square kilometres, often on the margins of barchans (Wilson, 1971). In such throughfl w-dominated dunefield (Kocurek, 1986) and associated with the previ- locations, sandfl w is usually unsaturated and ground sand ously referred-to movement of sand from high- to lower- cover incomplete, because unsaturated throughfl w has energy locations. Sand sheets can, however, also represent the potential to erode pre-existing deposits (Lancaster, major regional landscape components (Breed, McCauley, 1999). If surface conditions are favourable, however, iso- and Davis, 1987). For example, the Selima Sand Sheet of lated bedforms can develop so long as the sand received southern Egypt is a ‘single utterly fla sheet of fir sand, by the bedform is equal to that which is lost downwind. known to cover an area of nearly 100 000 square miles’ Wilson (1971) termed this situation metasaturated ﬂow. (Maxwell and Haynes, 2001, p.1623) (Figure 17.9). Sand seas proper accrue where sandfl w is saturated and Kocurek and Nielson (1986) recognised f ve major con- accumulation exceeds net transport, with consequential trols on aeolian sand sheet development. Vegetation,es- bedform development and the vertical or lateral accumu- pecially grasses, may encourage the accretion of low- lation of sand. angle laminae while limiting the construction of dunes. Wilson (1971) provided theoretical models of erg Coarse sand, which is not readily formed into dunes, can development in terms of wind regime and sandfl w characterise sand sheets, sometimes upwind of a dune- variations, demonstrating how sand seas should grow field where it remains as a surface lag deposit, as at the in locations of convergent windfl w and wind speed Algodone dunefield California (Kocurek and Nielson, deceleration, e.g. in intermontane basins (Figure 17.8). 1986). A coarse-sand sheet can therefore sometimes be An empirical study (Fryberger and Ahlbrandt, 1979) has regarded as a deflationa remnant, rather like some desert supported Wilson’s major assertions and identifie the pavement or regs.Anear-surface groundwater table synoptic and topographical conditions favouring sand sea (Stokes, 1968) or ‘Stokes surface’ (Fryberger, Schenk and development (Figure 17.8(b)). Krystinik, 1988) can be an important control on sand sheet Fryberger (1979) classifie sand seas into low-, development, acting as a base level to the action of wind intermediate- and high-energy environments. Most sand scour. This can occur in locations deprived of upwind P1: TIX/XYZ P2: ABC c17 JWST032-Thomas December 17, 2010 13:24 Printer Name: Yet to Come

436 CH17 AEOLIAN LANDSCAPES AND BEDFORMS

(a) NE SW Erg Erg TIBESTI Erg Erg Erg Calanscio Rebiana MASSIF Bilma Ténéré Azouak TAMGAUT MOUNTAINS

Benghazi, Libya 0 500 km

(b) Reduction of total Reduction of resultant sandflow energy sandflow energy

1 4

Figure 17.8 (a) The relationships of ergs in part of North Africa to topography (after Wilson, 1973). (b) Models of topographic influences on sand sea development (after Fryberger and Ahlbrandt, 1979). (1) In the shadow of a topographic barrier, (2) in shallow desert depressions and (3) by the direct blocking of wind, all leading to a total reduction in sand transporting energy. Resultant energy may be reduced (4) when surface winds are deflected, leading to sites of favourable accumulation, and (5) when katabatic winds off a highland oppose a dominant wind. Sand seas can also develop where water bodies interrupt regional sand flow patterns.

aeolian sand supply, or playa, sebkha or coastal sand sea probably inapplicable to the development of the hyper- situations (Fryberger, Schenk and Krystinik, 1988), where arid Selima Sand Sheet, because their effectiveness re- periodic or seasonal floodin can also prevent dune devel- quires either the presence of some moisture or proximity opment. Langford (1989) identifie a number of ways in to a dunefield Rather, Breed, McCauley and Davis (1987) which interactions between aeolian and fluvia processes ascribe its development to the blanketing of an ancient flu could lead to different sedimentary deposits in areas with vial landscape by aeolian redistribution of the sand frac- fluctuatin groundwater tables. The formation of surface tion contained within widespread alluvium deposits. The crusts, e.g. salcretes (see Chapter 6), or the growth of al- absence of major topographic barriers able to inhibit long- gal mats (Fryberger, Schenk and Krystinik,1988) can also distance wind transport accounts for the great size of this inhibit deflatio and contribute to the formation of sand sand sheet, while the coarseness of the sand has prob- sheets, as in the White Sands, New Mexico. ably caused the general lack of dune development and With the exception of the presence of coarse sand, the the formation of gently undulating subhorizontal tabular controls identifie by Kocurek and Nielson (1986) are deposits 1–10 cm thick. P1: TIX/XYZ P2: ABC c17 JWST032-Thomas December 17, 2010 13:24 Printer Name: Yet to Come

THE GLOBAL DISTRIBUTION OF LOESS 437

Table 17.4 Sand sea wind environments. that even today tend towards arid–semi-arid conditions or at least seasonal dryness – parameters that in many cases Sand sea (m3/m width per year) Rate of sand drift would have been enhanced at glacial times by enhanced continentality. 1. High-energy environments 25–40 The extensive blanket loess deposits of Central Asia Northern Arabian sand seas (particularly in the Karakum, in Tajikistan and in Uzbek- Northwestern Libya istan) extend to over 200 m and in the Huangtu Gaoyuan, Intermediate-energy environments 2. 15–25 or Loess Plateau, of China deposits attain a thickness of Simpson Desert, Australia up to 500 m (Huang, Pang and Zhao, 2000), although 150 Western Mauritania Peski Karakumy, Kazakhstan m is more representative of consistent maximum depths Peski Kyzylkum, Kazakhstan (Derbyshire and Goudie, 1998; Goudie, 2002). In both Erg Oriental, Algeria regions these deposits date back to the Pliocene, and the Erg Occidental, Algeria loess sequences in fact comprise alternating loess and Namib Sand Sea, Namibia palaeosol units, demonstrating that periods of deposi- Rub’al Khali, Saudi Arabia tion (during cold phases) alternated with warmer periods 3. Low-energy environments <15 of greater landscape stability. While these mid-latiutude SW Kalahari, southern Africa loess deposits are commonly regarded as ‘cold’ loessic de- Sahelian zone sand seas posits, their distribution actually shows close affinit to the Gobi Desert, China aeolian processes that have affected the desert areas that Thar Desert, India they border. In both regions, the loess deposits occur at the Taklimakan Desert, China downwind end of a sequence of aeolian depositional units. Thus in China (Figure 17.10) the sandy desert of the Mu Us grades southeastwards to sandy loess and then to the extensive loess plateau. This in effect represents chang- 17.4 The global distribution of loess ing aeolian transport efficien y southeastwards away from sediment source areas. Figure 17.2 shows the global distribution of loess deposits, which cover up to 10 % of the Earth’s land surface (Pecsi, 1968). Despite a complex definitiona history, 17.4.2 Peridesert loess where the role of water in loess deposition and the role of post-depositional processes has been debated, loess can In lower latitudes, dust deposits are less widespread but principally be define as a terrestrial accumulation of ae- are clearly associated with desert conditions. It is to these olian silt (dust) (e.g. Pye, 1987) or as a deposit, primarily deposits that the term ‘peridesert loess’ is more usually ap- aeolian in origin, that is principally comprised of silt-size plied. These are the loess deposits comprised of material sediment that is dominated by quartz. Fine sand and clay attributed to generation by abrasion during aeolian sand can also be present. Because of its fine-graine nature, silt saltation (e.g. Smith, Wright and Whalley, 2002). Inter- is readily transported in suspension and is prone to long- estingly, most of the major desert dust sources today that distance transport (see Chapter 18). Much dust therefore contribute material to long-distance atmospheric transport ends up being deposited in the oceans (e.g. Hesse and are associated with deflatio from dry lake basins and not McTainsh, 1999) and not on the land surface. directly from aeolian abrasion. However, the distinction between glacial and peridesert loess is not as clear-cut as sometimes presented, as implied above. Derbyshire 17.4.1 Loess production and distribution (1983) demonstrated, both mineralogically and through mapping the limits of Quaternary glaciations in China, In the mid-latitudes, loess deposits are strongly associ- that the extensive Chinese loess deposits were not solely ated with the generation, transport and deposition of silt comprised of glacial erosion material, but of sediments de- from glacial erosion systems during cold phases in the rived from desert contexts. Indeed, several authors have Quaternary (Smalley and Vita-Finzi, 1968; Smalley and now identifie a range of regularly occurring dryland pro- Krinsley, 1978). This is not to say that conditions dur- cesses that can produce ample silt that is available for ing deposition were not rather arid, since the continental dust generation and ultimately the deposition of loess on location of many of these deposits, in central Europe, in- desert margins. These include a range of weathering pro- terior North and South America and Asia is in contexts cesses, aeolian abrasion, and abrasion and fracture during P1: TIX/XYZ P2: ABC c17 JWST032-Thomas December 17, 2010 13:24 Printer Name: Yet to Come

++ + + 105¡ + + 110¡E + + + + + + + + + + + + + + + + + + + HE (Yellow + + V G R V N iv + + A e + + r HOHHOT + + U ) + H + + V V V V V V + V V V V V + V V V V V V + + VV V V V V + + V V V + V V + + V V 40¡N + + V V + + + + V V + + V + + + + V + + + + + + + + V V V + + + V + + + + + V V + + + V + + V V V + + + + + + + + V V YINCHUAN V V V V V V V V V V V V V V V V V V V V V V V V V V V V V V V V V V V V V V V V V V V + + + + + + + V V + + ++ V + V + + + + + + + + LANZHOU + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + 35¡+ + + + + + +

+ + + + + 0 200 + km XIAN

Sand dunes, semi-fixed Loess Alluvium

V V Sand dunes, slightly mobile V Pluvial Bedrock + + + Pluvial - Fluvial Faults, active and Sand dunes, moderately mobile + + hidden

Figure 17.9 Quaternary sedimentary facies of the Ordos Platform and the Loess Plateau, China, showing the spatial relationship between sand and silt units (simpliﬁed from Geomorphological Map of China and Its Adjacent Area (scale 1 : 4M), attached to Chen, 1993).

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DYNAMIC AEOLIAN LANDSCAPES IN THE QUATERNARY PERIOD 439

Increasing precipitation and vegetation cover Prevailing wind Increasing leaching and soil development

Remnant Primary Colluvially Desert source primary loess 1 Colluvially 2 loess 3 reworked of sand Active Bare reworked loess loess and dust sand dunes rock

Zone of no accumulation Zone of largely colluvially reworked loess Zone of primary loess with some colluvial reworking

Hyper-arid Arid to semi-arid Sub-humid Not to scale Climate fluctuations

1 2 3 Modern soil Modern soil Carbonate horizon Clayey Massive weathered loess loess

Massive Palaeosol coarse loess

Thick carbonate horizon

Massive loess

Figure 17.10 Schematic model of the distribution of peridesert loess in the landscape (modiﬁed after Pye and Sherwin, 1999).

episodic fluvia transport (McTainsh, 1987; Assallay et al., to trap dust: primary dust accumulation settings in vege- 1998; Smith, Wright and Whalley, 2002). tated zones of higher relief and, due to its susceptibility to The limited extent of low-latitude peridesert loess is reworking, areas of colluvially reworked loess on downs- now sometimes attributed to the lack of suitably vegetated lope locations. dust-trapping surfaces in these areas; thus most of the dust escapes into the atmosphere and is available for wider global dispersal, including into oceans (Tsoar and Pye, 17.5 Dynamic aeolian landscapes in the 1987). It may also be due to a more prosaic reason: it Quaternary period has not been properly identifie in the fiel (McTainsh, 1987; Hesse and McTainsh, 2003), partially because it is Environmental change during the Quaternary period has incorporated into soils as dust production and weathering contributed greatly to the nature of today’s deserts and tends to occur at similar rates, limiting opportunities for drylands (Chapter 3). Sarnthein’s (1978) hypothesis of blanket deposits to develop (Vine, 1987). expanded aridity during glacial times (Figure 17.11) was Pye and Sherwin (1999) have produced a schematic based both on an analysis of aeolian sediments in marine model of the characteristics of peridesert loess (Fig- cores and on identifying the extent of currently ‘fi ed’ ure 17.11). They distinguish nondepositional zones on sand dunes on the margins of today’s active sand seas. This desert margins where there is insufficien vegetation cover built very much on identifying an association between P1: TIX/XYZ P2: ABC c17 JWST032-Thomas December 17, 2010 13:24 Printer Name: Yet to Come

440 CH17 AEOLIAN LANDSCAPES AND BEDFORMS

Today

North Tropic

Equator

South Tropic

Active sand seas

At Last Glacial Maximum

North Tropic

Equator

South Tropic

Active sand seas

Figure 17.11 The distribution of active sand seas today and as suggested for the Last Glacial Maximum (after Sarnthein, 1978).

dune activity, rainfall and vegetation (e.g. Grove, 1958), sible using imagery because quartz-rich sand has a high with the role of wind energy now recognised as critically reflect vity (Muhs, 1985). important too (e.g. Lancaster, 1987). The extent of ancient dunefield is well established in The identificatio of inactive or degraded sand dunes Africa India, Australia, North America and to a lesser ex- and palaeo-sand seas has been facilitated by the extent in South America. Reviews of the fi ed sand seas of amination initially of aerial photography (Grove, 1958, the Kalahari have been provided by Thomas and Shaw 1969) and subsequently satellite imagery. In linear dune- (2002), for Australia by Wasson (1986) and for North field distinct vegetation zonation aids identificatio (Fig- America by Wells (1983). Goudie’s (2002) review of ure 17.13). Compared to interdune areas, the dune crests desert environments includes summaries of sand sea his- may be relatively sparsely vegetated, as in much of Aus- tories from around the globe. tralia (Madigan, 1936), or densely covered by trees be- A range of criteria has been used to designate f xed or cause the soft sands favour the development of deep root relict dune status (Table 17.5). In the Thar Desert (India), systems, as in the northern Kalahari (Thomas, 1984). In Makgadikgadi basin (Botswana) and Chad basin (Nige- areas where smaller dune types occur and their morpho- ria), dunes have been floode and overlain by lacustrine logical patterns are less obvious, identificatio is still fea- deposits (Singh, 1971; Cooke, 1984; Grove, 1958). Other P1: TIX/XYZ P2: ABC c17 JWST032-Thomas December 17, 2010 13:24 Printer Name: Yet to Come

DYNAMIC AEOLIAN LANDSCAPES IN THE QUATERNARY PERIOD 441

(a) 20¡E 50 100 200

Southwestern Kalahari dunes Namib 25¡S sand sea

200 100 50

(b) 50 200 100

Figure 17.12 Landsat image (5 August 1973) of part of southwestern Zambia, southeastern Angola and the Caprivi 50 Strip, Namibia. Vegetation contrasts between the crests of ﬁxed dune ridges and interdune areas allow the identiﬁcation of an- 100 cient dune systems in the area covered by the left half of the 200 image.

Figure 17.13 Suggested spatial variations in the distribu- dunes have been subjected to eluviation (Smith, 1963), tion of dune mobility index M values between today and the pedogenesis and calcification resulting from an increase last Glacial Maximum, in the southwestern Kalahari, according in available moisture and reduction in aeolian activity to Lancaster (1988). M = W/(P/PE) where W is the % time (Allchin, Goudie, and Hegde, 1978). The nature of post- wind exceeds the threshold velocity for sand transport and depositional dune degradation may, however, complicate P/PE is the ratio of annual precipitation to annual potential the subsequent interpretation of the resultant forms, as evapotranspiration. Index values are interpreted in terms of M well as the dating of phases of accumulation (Munyikwa, > 200 = duneﬁeld fully active; 100–200 = dune plinths and 2005; McFarlane et al., 2005). interdunes vegetated; 50–100 = dune crests only active and < = The presence of dune surface vegetation is perhaps the 50 dunes inactive. most widely used indicator of fi ed dune status, but needs careful interpretation as an indicator of inactivity. This is especially so for sand seas composed of linear dunes, et al., 2008) and a gradual decline in sand movement in which are in any case relatively immobile forms compared the direction of increasing vegetation density (Ash and to transverse forms (Thomas and Shaw, 1991a) and which Wasson, 1983; Wiggs et al., 1995). Depending on wind may widely support an element of plant cover even under velocities, significan sand transport may occur where up arid conditions (Tsoar and Moller, 1985). While there can to 35 % of the ground surface is plant-covered – a situation be little dispute that dunefield covered by dense wood- that occurs in several dunefields which have been desig- land are inactive, it is not so clear in the case of sparsely or nated fi ed or inactive, e.g. the southwestern Kalahari and partially vegetated dunes (Hesse and Simpson, 2006). In parts of the Australian and Indian sand seas. Some au- present arid environments, there is not a distinct threshold thors have concluded that present-day dune inactivity is between aeolian activity and inactivity, but rather different not necessarily a function of increased effective precipi- levels of activity (Livingstone and Thomas, 1993; Mason tation but may be due to a decline in windiness (Ash and P1: TIX/XYZ P2: ABC c17 JWST032-Thomas December 17, 2010 13:24 Printer Name: Yet to Come

442 CH17 AEOLIAN LANDSCAPES AND BEDFORMS

Table 17.5 Criteria used in the literature to designate dune since dunes were active, Talbot (1994), Wasson (1984), ‘relict’ status. Lancaster (1988) and Chase and Brewer (2009) have all proposed the use of dune mobility indices (Figure 17.14). 1. A cover of vegetation sufficien to inhibit sand transport Similar indices may also be used with modelled climate 2. Dune form degraded (slope angles) compared to active data to predict sand sea responses to climatic changes in dunes the future (Thomas, Knight and Wiggs, 2005; Wang et al., 3. Subject to pedogenesis 2009). 4. Presence of surface crusts/lithificatio Relict dune systems have also been used to reconstruct 5. Dune slopes gullied palaeocirculation patterns as well as the former extent 6. Internal structures destroyed by burrowing animals of aridity (Lancaster, 1981; Wells, 1983; Wasson, 1984; 7. Sediment size distribution bimodal due to post-formation Thomas and Shaw, 1991b). The orientation of the relict dust inputs dunes is compared to the resultant direction of modern 8. Dunes overlain by other depositional landforms sand-moving winds, calculated from wind data. Results (including second generation dunes) 9. Drowned in valleys/lakes/off-shore have frequently suggested that circulation patterns dif- fered from those of today when dune development occurred. Because dunes form in a wide variety of wind regimes today, often with considerable seasonal or diur- nal directional variability (see Chapter 19), the conclu- Wasson, 1983; Wasson, 1984; Bullard et al., 1996). For sions that have been drawn are frequently speculative. example, as Williams (1985, p. 226) has noted:

In the case of the source bordering dunes of Australia, 17.5.1 Dating aeolian landscape change Sudan and, presumably, of Zaire and Amazonia, the three prerequisities for dune formation seem to be Major advances have occurred in understanding Quater- a seasonally abundant supply of river sands, strong nary period sand sea dynamics through the advent and ap- unidirectional seasonal winds, and a less dense tree plication of thermoluminescence (TL) (Singhvi, Sharma cover than today. Such palaeo dunes do not connote and Agrawal, 1982) and subsequently optically stimulated extreme aridity. luminecscence (OSL) (e.g. Stokes, 1992; Stokes, Thomas and Washington, 1997) dating to aeolian sediments (see After analysis of a range of variables, Hesse and Simp- Box 17.1). Prior to the advance of these dating meth- son (2006) concluded that over millennial timescales, it is ods, quartz-rich aeolian sands and silts were datable only vegetation cover that is the primary control on linear dune- if bracketed by other sediments to which other dating fiel activity and on dune accumulation, rather than wind methods could be applied. For example, the alternating strength. A recent study by Mason et al. (2008) also sug- loess–palaeosol sequences in the Loess plateau of China gested that wind strength is not a primary control of activ- have been dated by radiocarbon or uranium-series dat- ity, though some authors attach primary importance to this ing of the palaeosol units, of which over 30 occur, by parameter (e.g. Tsoar, 2005). In an attempt to resolve the palaeomagnetic studies (Liu and Ding, 1998) or by ‘curve issue of the relative roles of erodibility (vegetation cover) fitting of sedimentological properties of the loess with and erosivity (wind energy) in explaining dunefiel inac- isotope records from polar ice cores (Balsam, Elwood tivity and the level of climatic changes that have occurred and Ji, 2005).

Box 17.1 Luminescence dating

The origins of luminescence dating lie more than four decades ago within the discipline of archaeological science and a need to determine the age of pottery. Since then it has undergone a rapid series of methodological and technological developments and has emerged in the last decade as an important and effective Quaternary environmental dating method. It is particular useful in arid zones where it is used to date directly the depositional age of sediments P1: TIX/XYZ P2: ABC c17 JWST032-Thomas December 17, 2010 13:24 Printer Name: Yet to Come

DYNAMIC AEOLIAN LANDSCAPES IN THE QUATERNARY PERIOD 443

within landforms such as sand dunes, loess deposits and lake shorelines. The application of luminescence dating has enabled a much greater understanding of environmental change in situations where the absence of organic material (e.g. in former desert contexts) prevents the use of radiocarbon dating and also where sediments are older than the age limits of radiocarbon dating.

What is luminescence dating?

Within all sediments the presence of naturally occurring radioisotopes of uranium-238, thorium-232 and potassium- 40 undergoing radioactive decay create a low-level background radiation fiel (alpha, beta and gamma radiation). Cosmic rays from the Sun also provide a significan contribution of radiation to near-surface sediments. The specifi nature of quartz and feldspar minerals allows them to behave as dosimeters: recording the total sum of ionizing radiation that they have been exposed to since they were buried (see Preusser et al., 2009, for a review). This is possible because structural defects and impurities within the minerals behave as traps in which unbound electrons (freed from their parent atom by energy derived from the radiation fiel of the sediment) will accumulate over time. The rate at which electrons are displaced depends on the strength of the level of radiation in the sediment during its burial history and is typically known as the ‘dose rate’. These displaced ‘trapped’ electrons act as a store of potential energy that can be naturally, or artificiall , dissipated in the form of photons (luminescence). Each time the minerals are exposed to sunlight during transport (e.g. the sediment is blown by the wind or eroded and washed downstream), the trapped electrons are able to obtain enough energy to be freed from this state and the dosimetric clock is reset (‘bleached’). The grains may at some point be deposited within an indicative landform such as a river terrace or sand dune and again become shielded from sunlight. The accumulation of electrons in defect sites is time-dependent and it is this property that can be utilised to obtain burial ages. A simple analogy that is frequently used to explain this process is to imagine the quartz or feldspar grains as a bucket underneath a tap into which water is dripping at a constant rate: If we know the rate of dripping (analogous to the dose rate) and we can measure the total amount of water in the bucket (analogous to the total amount of trapped electrons), then it is possible to calculate how long the bucket has been located under the dripping tap. Figure 17.14 summarises the main principles in the accumulation of a measurable luminescence signal in a sediment grain.

How are luminescence ages produced?

To avoid mineral grains being exposed to light prior to measurement, sediment samples collected in the fiel must be extracted and stored in lightproof containers. In the laboratory, under controlled wavelength low-light conditions, the sediment undergoes a series of chemical treatments to isolate the quartz or feldspar component. The total population of trapped charge within the minerals can then be measured by stimulating the material with external energy in the form of light (optically stimulated luminescence: OSL) or heat (thermoluminescence: TL). As the trapped electrons are released, energy is dissipated in the form of luminescence. Specifi measurement conditions are used that include pre-heating the sample and monitoring any change of sensitivity to laboratory stimulation during measurement in order to isolate the luminescence signal from only those traps that are thermally stable over the burial period. The natural luminescence signal of the quartz or feldspar can be measured using a photomultiplier tube (cf. Bøtter-Jensen et al., 2003) and calibrated against the luminescence emitted when the same minerals are exposed to known laboratory doses of radiation. From this it is possible to establish the equivalent laboratory radiation dose (De) to that received during the burial period (the Palaeodose) (see Wintle and Murray, 2006, for a review of dating protocols). To complete the age equation the rate of electron displacement, determined by the ionizing radiation of the burial environment (D ), must also be known. This can be estimated by either directly measuring the gamma radiation (in the fiel or laboratory) or by laboratory determination of the radioisotope concentration within the sediment (see Aitken, 1998). P1: TIX/XYZ P2: ABC c17 JWST032-Thomas December 17, 2010 13:24 Printer Name: Yet to Come

444 CH17 AEOLIAN LANDSCAPES AND BEDFORMS

Geological dose Bleaching event Burial Dose Laboratory measurement (101-6 yrs) 1 (102-4 secs) 5

4 5

2 Equivalent

Luminescence Signal Dose (incompletely Equivalent bleached 3 Dose sediment)

1 : Erosion of sediment Dose (Gy) 2 : Transportation of sediment 3 : Re-deposition of sediment 4 : Laboratory signal measurement 5 : Interpolation on to laboratory administered dose response curve

Figure 17.14 Principles of the accumulation of a luminescence signal in sediment and the effect of incomplete bleaching on determining an accurate estimate of the burial age (adapted from Bailey and Arnold, 2006). Quartz or feldspar grains deposited by the wind or another geomorphic process, once buried, begin to accumulate electrons at imperfections within the crystal structure. Over time, an accumulated dose (trapped charge population) builds up. If exposed to sunlight, these electrons are removed and the dosimetric clock is reset. Laboratory measurements of the trapped electron population allow an estimate of the radiation dose received during burial to be made. If the resetting process is incomplete (dashed line), a fraction of the geological dose will remain and the measured trapped charge population will equate to part of the geological dose plus the burial dose. Without recognition, this results in an overestimate of the sediment burial age (refer to text and Table 17.6 for details).

An age is therefore determined by

D (Gy) (total absorbed radiation dose) Age (ka) = e D (Gyka) (radiation dose rate)

Luminescence dating procedures are mineral- and site-specifi and may be affected by a number of potential difficulties some of which are detailed in Table 17.6.

Sallie L. Burrough

Luminescence dating of sand seas has shown that many are not linked to specifi morphological expressions; i.e have histories of multiple phases of accumulation. Where down-profil age changes occur within dunes, sometimes multiple generations of dune forms overlap within a sand separated by weak palaeosols (e.g. Fitzsimmons et al., sea, e.g. in Mauritania (Lancaster et al., 2002, see Chap- 2007), which identify periods of dune accumulation at ter 19, Box 19.1, for a detailed consideration of this dune 73–66, 35–32, 22–18 and 14–10 ka in the Strzlecki Desert, system) and the eastern Sahara (Bubenzer and Bolten, Central Australia. 2008), ages can appear to suggest that each dune pat- In other sand seas, as the number of direct age deter- tern represents a single phase of accumulation, when sub- minations from dune sand has increased, so the interpre- tle differences in atmospheric circulation led to different tation of sand sea development has become more com- pattern orientations. In other contexts, more complex re- plex. Initial accumulation phase histories for the Kalahari lationships are emerging, where phases of accumulation were based on a relatively limited number of OSL ages P1: TIX/XYZ P2: ABC c17 JWST032-Thomas December 17, 2010 13:24 Printer Name: Yet to Come

DYNAMIC AEOLIAN LANDSCAPES IN THE QUATERNARY PERIOD 445

Table 17.6 Causes and consequences of luminescence dating challenges.

Potential difﬁculty Physical explanation Environmental context Consequence Possible solution

Signal Geologically stable Sample is buried for a Possible to obtain a Application of new saturation electron traps f ll up very long period of minimum age techniques such as timeorthedoserateis isothermal TL may particularly high extend the datable burial age limit of some sediments Incomplete Some electrons remain Sediment is Residual dose remains Signal distribution bleaching trappedevenaftera insuff ciently exposed and De is analysis allows bleaching event to light during overestimated. identificatio of the transport Differential exposure problem. In most of grains results in a cases single grain wide distribution of dating and the use of De values age models enables the correct burial age to be determined Post- Biotic activity within the Wide (and often skewed) The use of single grain depositional sediment mixes grains De distributions dating, mixture mixing from chronologically modelling and distinct depositional Bayesian statistics events may enable the identificatio of depositional events, even when mixed Signal Electrons ‘leak’ from Dependent on specifi Incorrect age A stable part of the instability traps. Some traps properties of the signal can be isolated (responsible for a part quartz or feldspar and measured using a of the luminescence grains number of techniques signal) are more prone to this happening than others Dose rate The level of ionising Disequilibrium may Incorrect age Disequilibrium can be change over radiation is not occur between the measured and time constant over time parent and daughter corrected for in young isotopes where one or sediments. In older the other is removed sediments, from the system. equilibrium may have Ingrowth of duricrete been reestablished material during burial and past will also change the disequilibrium will dose rate remain undetected Beta The level of ionising Hot spots and buffer Wide (and often skewed) Using particular age heterogeneity radiation is not zones may occur De distributions. models, a good spatially constant within the sediment Diff cult to estimate of burial age within the sample matrix, causing some differentiate from can be obtained. grains to receive bioturbation Errors will be larger disproportionately more radiative dose during the burial history than others P1: TIX/XYZ P2: ABC c17 JWST032-Thomas December 17, 2010 13:24 Printer Name: Yet to Come

446 CH17 AEOLIAN LANDSCAPES AND BEDFORMS

C Charcoal (b) 952 896 R Root penetration 050

Ca Massive calcrete km 21 ± 8 m 77 ± 12 Dune ridges 0 (axes) 52 ± 8 27¡E 1 International Z A M B I A 895 1005 boundary 950 R R Victoria Zambezi C R Falls R 9 ± 2 15 ± 4 2 Dated sample ± 164 ± 32 28 6 (kyr ± 1 s.d.) 37 ± 8 20 ± 5 R Gwaai 3 Kazuma 42 ± 7 Depression Hwange 30 ± 4 Town 4 N 40 ± 6 951 Panda Matenga 5 R B O T S W A N A Main Camp 894 R 19¡S 1004 32 ± 6 R 11 ± 3 18 ± 5 23 ± 5 1003 28 ± 5 C C R 43 ± 9 43 ± 8 15 ± 2 Dzivanni 46 ± 8 Mud Flats 0 150 20¡E C 39 ± 5 (a) 23 ± 3 53 ± 8 km C 945 892 46 ± 7 890 103 ± 17 947 891 R R R 14 ± 4 112 ± 18 ± 22 ± 5 102 18 1.4 ± 0.5 25 ± 3 ± 23 5 C 25 ± 4 1.7 ± 0.7 Ca 22 ± 3

No (c) ss ob BOTSWANA 0 R 25¡S Mariental 943 1 23.0 ± 2.9

Tshabong Twee Rivieren 10 ± 4 NAMIBIA 2 Molopo R. 946 Keetmanshoop R 946B R 3 6 ± 2 43.9 ± 2.9 Upington

949 27 ± 7 Depth (m) 4 Orange R. 1.1 ± 0.2 948 1.1 ± 0.2 944 942 1.4 ± 0.2 (5) 1.4 ± 0.3 (4) (3) 5 (2) R S A 12 ± 2 105.6 ± 10.2 R 6 15 ± 3 28 ± 7

17 ± 4 Ca 14 ± 3 04080120 Age (kyr)

Figure 17.15 Chronology of linear dune deposition in the Kalahari Desert, based on OSL dating of sediment samples. Ages were interpreted as demonstrating discrete periods of dune deposition during the Late Quaternary (after Stokes, Thomas and Washington, 1997). P1: TIX/XYZ P2: ABC c17 JWST032-Thomas December 17, 2010 13:24 Printer Name: Yet to Come

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(a) KKalkrandalkrand

Sekretarispan Auob R. NNAM06/1AM06/1

D1268 NAM06/6NAM06/6 C20

WitvleispruWitvleispr it R.

NNAM06/4AM06/4 SStampriettampriet uit R. Auob R. NNAM06/2AM06/2 Kalahari R. Anib Lodge

25.2 km MarientalMariental

NAM06/1 NAM06/2 NAM06/4 NAM06/6 m m m m

1.6 ± 0.2 2.4 ± 0.2 1.0 3.7 ± 0.2 1.0 1.0 4.6 ± 0.9 1.0 5.6 ± 0.5 3.1 ± 0.4 5.6 ± 0.9 2.0 8.3 ± 1.0 2.0 6.6 ± 0.7 2.0 8.6± 0.9 2.0 8.6 ± 0.6 10.2 ± 1.1 12.1 ± 1.1 Structureless, 3.0 9.0 ± 1.0 3.0 11.7 ± 0. 3.0 17.9 ± 1.8 3.0 55.3 ± 4.0 homogenous, 11.2 ± 0.8 9.8 ± 0.8 23.3 ± 2.1 medium sand 4.0 10.8 ± 0.8 4.0 14.8 ± 1.3 4.0 27.7 ± 3.7 4.0 75.1 ± 5.2 23.1 ± 1.7 11.7 ± 0.9 73.2 ± 8.2 Nodular calcrete 5.0 12.3 ± 0.8 5.0 33.6 ± 3.4 5.0 64.8 ± 7.3 5.0 103.7 ± 9.5 11.2 ± 0.8 37.8 ± 3.1 74.4 ± 12.0 103.2 ± 11.7 6.0 11.3 ± 1.0 6.0 38.8± 3.3 6.0 122.9 ± 29.5 6.0 84.3 ± 5.5 11.4 ± 0.9 37.5 ± 2.8 186.2 ± 15.6 7.0 10.7 ± 0.9 7.0 45.5 ± 5.6 7.0 7.0 14.2 ± 1.0 45.6 ± 9.7 8.0 8.0 8.0 8.0 49.0± 4.5 183.3 ± 18.3

Age (ka) (b) 200 180 160 140 120 100 80 60 40 20 0

Mariental-Stampriet dunes (all ages)

? ?

Mariental-Stampriet dunes: 1 m sampling intervals, starting at 0.5 m depth

Mariental-Stampriet dunes: 1 m sampling intervals, starting at 1 m depth

MIS 6 MIS 5 MIS 4 MIS 3 MIS2IS1 M 200 180 160 140 120 100 80 60 40 20 0 Age (ka)

Figure 17.16 (a) High vertical intensity OSL dating of samples from sediment cores from linear dunes in the western Kalahari suggest a record of almost continuous deposition in the Late Quaternary, in contrast to Figure 17.15, which was based on fewer samples. (b) Demonstrates how the number of samples dated can potentially impact on the interpretation of the periodicity of dune depositional history (after Stone and Thomas, 2008). P1: TIX/XYZ P2: ABC c17 JWST032-Thomas December 17, 2010 13:24 Printer Name: Yet to Come

448 CH17 AEOLIAN LANDSCAPES AND BEDFORMS

from samples collected from shallow pits in dunes, which and the processes responsible for sediment mobilisation, gave a relatively simple picture of sand sea history, dune transport and settlement. Gaining this knowledge requires accumulation and inferred aridity (Stokes, Thomas and investigations as a range of spatial and temporal scales Washington, 1997) (Figure 17.16). (Thomas and Wiggs, 2008), and in this chapter we have As more age determinations have been achieved, pic- primarily considered what controls the nature and char- tures that are complex to interpret emerge. These include acter of the landscape features of aeolian systems: sand spatially variable accumulation histories along short dis- seas and loess deposits. These spatially extensive features tances of individual dunes (Telfer and Thomas, 2006), but provide the context in which the operation of aeolian pro- also that, as more ages are added to data sets and sam- cesses and the generation of bedforms play out, and these pling strategies become more intense, apparent discrete process- and landform-focused dimensions of aeolian sys- phases of accumulation can disappear and also appar- tems are examined in detail in the next chapters. ent continuous deposition can be represented by full sets of age determinations (Stone and Thomas, 2008) (Fig- ure 17.16). While a number of factors could contribute References to such a situation (e.g. the error range on ages, possible post-depositional sediment mixing, sampling issues, the Qua- complex suite of factors that affect dune activity and sedi- Aitken, M.J. (1994) Optical dating: a non-specialist review. ternary Geochronology, 13, 503–508. ment preservation), this demonstrates a number of impor- Aitken, M.J. (1998) An Introduction to Optical Dating: The Dat- tant factors in elucidating sand sea histories. First, simple ing of Quaternary Sediments by the Use of Photo-Stimulated relationships between dunefiel accumulation and broad Luminescence, Oxford University Press, Oxford, 267 pp. global climate trends senso Sarnthein (1978) have not oc- Al Janabi, K.Z., Jaward, A. Al-Taie, F.H. and Jack, F.J. (1988) curred. Second, variability in depositional ages can be Origin and nature of sand dunes in the alluvial plain of south- marked between dunes within an individual sand sea, be- ern Iraq. Journal of Arid Environments, 14, 27–34. cause net accumulation and preservation does not appear Allchin, B., Goudie, A.S. and Hegde, K.T.M. (1978) The Pre- to occur simultaneously at all points within a dunefield history and Palaeogeography of the Great Indian Desert, Third, care needs to be exerted in terms of how chronolo- Academic Press, London. gies are produced and how many ages are used to produce Andreotti, B., Fourriere, A., Ould-Kaddopur, F. et al. (2009) conclusions for individual sand seas: the dunefield not Giant aeolian dune size determined by the average depth of the atmospheric boundary layer. Nature, 457, 1120– the individual dune, is the natural unit from which to 1123. gain an appropriate picture of a system history (Stone and Arens, S.M. (1996) Rates of aeolian transport on a beach in a Thomas, 2008). temperate humid climate. Geomorphology, 17, 3–18. Luminescence dating is also shedding new light on Ash, J.E. and Wasson, R.J. (1983) Vegetation and sand mobility the records of past climate change preserved within loess in the Australian desert dunefield Zeitschrift fur¬ Geomor- bodies. Independent direct dating of loess has challenged phologie, Supplementband 45, 7–25. previous chronologies based on magnetic susceptibly or Assallay, A.M., Roigers, C.D.F., Smalley, I.J. and Jefferson, I. record correlation (e.g. Singhvi et al., 2001; Stevens et al., (1998) Silt: 2–64µm, 9–4ø. Earth Science Review, 4, 61–88. 2007). Contrary to views that loess deposits provide a Baas, A.C.W. (2007) Complex systems in aeolian geomorphol- spatially uniform record of deposition over large areas, ogy. Geomorphology, 91, 311–331. these studies show that accumulation histories can be Baas, A.C.W. and Nield, J.M. (2007) Modelling vegetated dune landscapes. Geophysical Research Letters, 34, L06405. as complex and variable as those derived from within DOI:10.1029/2006GL029152. sand seas. Bagnold, R.A. (1941) The Physics of Blown Sand and Desert Dunes, Methuen, London. Bailey, R.M. and Arnold, L.J. (2006) Statistical modelling of 17.6 Conclusions single grain quartz of distributions and an assessment of procedures for estimating burial dose. Quaternary Science Reviews, 25, 2475–2502. Aeolian deposits are a major – though not as extensive as Balsam, W., Ellwood, B. and Ji, J. (2005) Direct correla- widely assumed – component of dryland environments. tion of the marine oxygen isotope record with the Chinese Determining where they are, and why they are there, re- Loess Plateau iron oxide and magnetic susceptibility records. quires the application of scientifi approaches that eluci- Palaeogeography, Palaeoclimatology, Palaeoecology, 221, date the timing of deposition, the controls on deposition 141–152. P1: TIX/XYZ P2: ABC c17 JWST032-Thomas December 17, 2010 13:24 Printer Name: Yet to Come

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(1995) Eolian dunes: computer simulations and forms, 33, 1396–1418. attractor interpretation. Geology, 23, 1107–1110. Tsoar, H. (1982) Internal structure and surface geometry of lon- Wiggs, G.F.S. (1993) Desert dune dynamics and the evaluation gitudinal (seif) dunes. Journal of Sedimentary Petrology, 52, of shear velocity: an integrated approach, in The Dynamics 823–831. and Environmental Context of Aeolian Sedimentary Systems Tsoar, H. (2005) Sand dunes mobility and stability in relation (ed. K. Pye), Geological Society Special Publication 72, Lon- to climate. Physica A: Statistical Mechanics and Its Appli- don, pp. 32–46. cations, 357, 50–56. Wiggs, G.F.S., Livingstone, I., Thomas, D.S.G. and Bullard, J.E. Tsoar, H. and Møller, J.T. (1986) The role of vegetation in (1995) Dune mobility and vegetation cover in the southwest the formation of linear dunes, in Aeolian Geomorphology Kalahari Desert. Earth Surface Processes and Landforms, (ed. W.G. Nickling), Binghampton Symposia in Geomor- 20, 515–529. phology, International Series 17, Allen & Unwin, Boston, Williams, M.A.J. (1985) Pleistocene aridity in tropical Africa, pp. 75–95. Australia and Asia, in Environmental Change and Tropical Tsoar, H. and Pye, K. (1987) Dust transport and the ques- Geomorphology (eds I. Douglas and T. Spencer), George tion of desert loess formation. Sedimentology, 34, 139– Allen and Unwin, London, pp. 219–233. 153. Wilson, I.G. (1971) Desert sand f ow basins and a model Twidale, C.R. (1972) Evolution of sand dunes in the Simpson for the development of ergs. Geographical Journal, 137, Desert, central Australia. Transactions, Institute of British 180–197. Geographers, 56, 77–109. Wilson, I.G. (1972a) Aeolian bedforms – their development and van Boxel, J.H., Sterk, G. and Arens, S.M. (2004) Sonic origins. Sedimentology, 19, 173–210. anemometers in aeolian sediment transport research. Geo- Wilson, I.G. (1972b) Universal discontinuities in bedforms pro- morphology, 59, 131–147. duced by the wind. Journal of Sedimentary Petrology, 42, Vine, H. (1987) Wind-blown material and West African soils: an 667–669. explanation of the ‘ferallitic soil over loose sandy sediments’ Wilson, I.G. (1973) Ergs. Sedimentary Geology, 10, 77–106. profile in Desert Sediments Ancient and Modern (eds L. Wintle, A.G. and Murray, A.S. (2006) A review of quartz opti- Frostick and I. Reid), Geological Society Special Publication cally stimulated luminescence characteristics and their rele- 35, London, pp. 171–186. vance in single-aliquot regeneration dating protocols. Radi- Wang, X., Yang, Y.,Dong, Z. and Zhang, C. (2009) Responses of ation Measurements, 41, 369–391. dune activity and desertificatio in China to global warming Yang, X., Liu, T. and Xiao, H. (2003) Evolution of megadunes in the twenty-firs century. Global and Planetary Change, and lakes in the Badain Jaran Desert, Inner Mongolia, China, 67, 167–185. during the last 31,000 years. Quaternary International, 104, Warren, A. (1988) The dunes of the Wahiba Sands. Journal of 99–112. Oman Studies, Special Report 3, 131–160. Zender, C.S., Bian, H. and Newman, D. (2003) Mineral dust Washington, R., Todd, M., Middleton, N.J. and Goudie, A.S. entrainment and deposition (DEAD) model: description and (2003) Dust-storm source areas determined by the total 1990s dust climatology. Journal of Geophysical Research, ozone monitoring spectrometer and surface observations. 108 (D14), 4416, DOI: 10.1029/2002JD002775. P1: TIX/XYZ P2: ABC c17 JWST032-Thomas December 17, 2010 13:24 Printer Name: Yet to Come

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Zheng, X.-J., Bo, T.-L. and Zhu, W. (2009) A scale-coupled International Geomorphology (ed. V. Gardiner), John Wiley method for simulation of the formation and evolution of aeo- & Sons, Ltd, Chichester, pp. 1211–1215. lian dune field International Journal of Nonlinear Sciences Zimbelman, J.R., Williams, S.H. and Tchakerian, V.P. (1994) and Numerical Simulation, 10, 387–395. Sand transport paths in the Mojave Desert, southwestern Zhu, Z., Zou, B. and Yang, Y. (1987) The characteristics of aeo- United States, in Desert Aeolian Processes (ed. V.P. Tchakar- lian landforms and the control of mobile dunes in China, in ian), Chapman & Hall, New York. P1: TIX/XYZ P2: ABC c18 JWST032-Thomas December 15, 2010 18:2 Printer Name: Yet to Come

18 Sediment mobilisation by the wind

Giles F. S. Wiggs

18.1 Introduction Improved understanding and measurement of aeolian processes has allowed the development and calibration Wind can be a particularly effective medium for sediment of models of contemporary dune dynamics and distri- movement in drylands where a relatively sparse vegeta- bution (Andreotti, Claudin and Douady, 2002; Baddock, tion cover and thin soils combine to create highly erosive Livingstone and Wiggs, 2007; Hersen, 2005; Living- conditions on highly erodible surfaces. There is little evi- stone, Wiggs and Weaver, 2007; Parsons et al., 2004; dence to suggest that dryland winds are any stronger than Tsoar, Blumberg and Stoler, 2004; Walker and Nickling, their counterparts in humid regions, but the sparse veg- 2002; Weaver and Wiggs, 2010; Wiggs, Livingstone and etation cover in deserts allows the winds to contact the Warren, 1996), which has proved essential to our inter- surface more effectively, and the lack of root systems and pretation of palaeodune sequences (Mason et al., 2004; moisture to bind sediment together renders it far more Stone and Thomas, 2008; Thomas et al., 2000; Telfer and susceptible to erosion (Pye and Tsoar, 1990). The effi Thomas, 2006), while our improved understanding of the ciency of dryland winds is underlined by the continuous controls on aeolian processes has also allowed us to con- advance and development of desert dunes in the Earth’s sider the potential impact of global warming scenarios on sand seas (e.g. Bristow, Duller and Lancaster, 2007), the future dune activity (Knight, Thomas and Wiggs, 2004; long-range transport of Saharan dust across the Atlantic Thomas, Knight and Wiggs, 2005; Wang et al., 2009). to the Amazon basin (Swap et al., 1992), the deposition Our understanding of aeolian processes has also been of >400 m thick loess deposits in the Lanzhou region of enhanced through investigations of wind erosion from China (Porter, 2001) and the human and environmental susceptible soils and agricultural field where models of distress caused by agricultural wind erosion in the 1930s potential erosion, including the effects of implementing Dust Bowl in mid-west USA (Worster, 1979). management strategies, have been developed and cali- To interpret the evolution of landscapes such as those brated (Buschiazzo and Zobeck, 2008; Webb et al., 2009). mentioned above, the processes of aeolian erosion, trans- Previously, such models provided a fundamental frame- portation and deposition of sediment in drylands have work from which dust emission schemes have been de- to be understood. Understanding of aeolian processes has veloped for application to regional and global scale dust been inspired by the work of Bagnold (1941, 1953, 1956). dynamics models (Zender, Bian and Newman, 2003; Mar- Indeed, no discussion of aeolian processes is complete ticorena and Bergametti, 1995; Raupach, Gillette and without reference to his north African research and much Leys, 1993). Finally, our application of aeolian process can still be learned from these comprehensive works. understanding has progressed to extra-terrestrial surfaces However, there has been a great deal of progression in with a particular focus on the dunes and wind-streaks vis- our comprehension in the intervening 60–70 years and an ible on Mars (Laity and Bridges, 2009; Kok and Renno, understanding of aeolian processes has been used to clar- 2009a; Bourke, 2010; Gillies et al., 2010). ify a wide range of natural and environmental questions Sediment movement is essentially a function of both and problems. the power of the wind (erosivity) and grain characteristics

Arid Zone Geomorphology: Process, Form and Change in Drylands, Third Edition. Edited by David S. G. Thomas © 2011 John Wiley & Sons, Ltd. Published 2011 by John Wiley & Sons, Ltd. ISBN: 978-0-470-51908-0 P1: TIX/XYZ P2: ABC c18 JWST032-Thomas December 15, 2010 18:2 Printer Name: Yet to Come

456 CH18 SEDIMENT MOBILISATION BY THE WIND

holding particles in place (erodibility). In general terms, (a) (b) where erosivity exceeds the resistance due to erodibility parameters, particle dislodgement and erosion will take 40 40 place. This simple equation is complicated by a variety of surface characteristics (such as roughness, slope 30 30 and moisture content), which can affect both erosivity 20 20 and erodibility. This chapter investigates the key fac- Height (cm) tors influencin erosivity and erodibility and examines 10 10 some of the methods with which we can measure and 0 0 characterise them. In particular, there is a focus on our 0123 0123 -1 burgeoning understanding of how both the presence of Wind velocity (m s ) vegetation and wind turbulence may affect sediment en- Figure 18.1 Near-surface vertical velocity profiles showing trainment – both of which have seen considerable ad- (a) the smaller surface shear stresses in laminar flow (b) when vances in understanding since the previous edition of Arid compared to turbulent flow (after Bagnold, 1941). Zone Geomorphology. where 18.2 The nature of windflow in deserts ρ = flui density As air moves over the surface of the Earth it is retarded h = fl w depth by friction at its base and a velocity profile develops. The zone of fl w where air is affected by surface friction is V = fl w velocity called the boundary layer and within this layer there is a gradation from zero velocity in a very thin layer at ν = viscosity the surface to free stream velocity at a height beyond the effects of surface friction. The atmospheric boundary An increasing Reynolds number signifie an increasing layer (ABL) is approximately 1–2 kilometres thick. turbulence intensity and a shift towards inertial forces The structure of the velocity profil within the boundary dominating over viscous forces, such that where Re < layer is highly dependent on the type of airfl w: laminar or 500–2000 laminar fl w exists and where Re ≥ 2000 tur- turbulent. In laminar fl w, there is little mixing between bulent fl w predominates. the different layers of flui and faster layers slip over In the atmosphere, airfl w is nearly always turbulent slower layers, while at the surface itself the air is station- because air has a low viscosity and boundary layer depths ary. With this type of fl w, momentum transfer between are normally quite high. Only in very viscous, slow or thin layers of flui is accomplished by means of molecular fl ws do laminar characteristics develop. transfer. Slower moving molecules of air drift into faster moving layers, thus causing a drag on the faster overlying airfl w. Such a transfer of momentum produces shearing forces between layers of air. 18.2.1 The turbulent velocity profile Turbulent fl w also comprises zero f ow velocity at the surface, but the exchange of momentum in the boundary Under normal atmospheric conditions on flat unvegetated layer is achieved through the action of gusts and turbulent surfaces and in the absence of intense solar heating, the eddies mixing between layers. Such momentum exchange turbulent velocity profil plots as a straight line on a semi- is far more efficien than the molecular exchange seen in logarithmic chart (Figure 18.2). The gradient of the semi- laminar fl w and this is represented by a differing velocity logarithmic profil is a result of the surface roughness profile The greater mixing in turbulent fl w results in a producing a drag on the overlying airfl w. Hence, if the steeper velocity gradient at the surface and hence higher gradient of the velocity profil is known, the shear stress shearing stresses (Figure 18.1). (drag) at the surface can be determined. In sedimentolog- Laminar and turbulent fl ws can be distinguished from ical research, a common method for describing the gradi- each other by the Reynolds number (Re): ent of the velocity profil is in terms of the shear velocity (u∗). It is the value of u∗ that has been traditionally used ρ in calculations for determining thresholds of erosion and = hV Re ν (18.1) flu es of sediment transport. The shear velocity (u∗)is P1: TIX/XYZ P2: ABC c18 JWST032-Thomas December 15, 2010 18:2 Printer Name: Yet to Come

THE NATURE OF WINDFLOW IN DESERTS 457

(a) (b) 16 16

12 12 ) -1

8 8

Increasing

Velocity (m s Velocity u* 4 4

0 0 z0 z0 Height above surface Log z Height above surface Log z

Figure 18.2 Semi-logarithmic velocity proﬁles showing a focus at height z0, the aerodynamic roughness. The aerodynamic roughness in (a) is less than that in (b); hence shear velocities (u∗) in (b) are greater. See text for details.

proportional to the velocity profil gradient (Figure 18.2) otherwise known as the ‘law-of-the-wall’: and is related to the actual surface shear stress (τ )bythe 0 u 1 (z − d) following expression: = κ ln (18.3) u∗ z0 = τ /ρ u∗ 0 (18.2) where

κ = von Karman’s constant (≈0.4) In Figure 18.2(a) the semi-logarithmic turbulent velocity profil does not reach the surface. This is because very d = zero-plane displacement close to the surface the wind velocity is zero. The height of this zero-velocity region is termed the aerodynamic The zero-plane displacement (d) is a measure of the ver- roughness length (z0) and it is an important parameter for tical displacement of z0 on surfaces with large surface it is a function of the surface roughness, it has an im- roughnesses (e.g. long grass or bushes). It can largely be pact on surface erodibility and it partly controls the gra- ignored on smooth and unvegetated desert surfaces. dient of the velocity profil and hence u∗. For example, over a rougher surface (Figure 18.2(b)), the aerodynamic roughnesss (z0) is larger, and with all other parameters 18.2.2 Measuring shear velocity (u∗) remaining constant, the velocity profil gradient becomes and wind stress steeper. Hence, shear velocities (u∗) increase and, consequent upon this, sediment transport would be likely to rise. Shear velocity (u∗) is commonly determined from least Another way in which shear velocity (u∗) might increase squares regression analysis of time-averaged veloc- is by a rise in the overall environmental wind velocity, as ity measurements at several known heights. Wilkinson shown in Figure 18.2(a). Note in Figure 18.2 that height (1983/1984) reports that at least f ve velocity data points is the independent variable and is therefore plotted on the are required to determine u∗ and z0 with any statistical x axis of the graph. This is important when calculating significanc and so equipment arrays such as that shown shear velocities using a least squares regression technique in Figure 18.3 are typically erected. When plotted in the (as described below). manner of Figure 18.2, u∗ can be determined from the The relationship between aerodynamic roughness (z0), slope component of the regression equation by shear velocity (u∗) and wind velocity (u) at a height (z)are described by the Karman–Prandtl velocity distribution, u∗ = κm (18.4) P1: TIX/XYZ P2: ABC c18 JWST032-Thomas December 15, 2010 18:2 Printer Name: Yet to Come

458 CH18 SEDIMENT MOBILISATION BY THE WIND

Figure 18.3 Photograph of a 6 m high anemometer tower incorporating eight logarithmically spaced cup anemometers to measure changing shear stress (u∗) and aerodynamic roughness (z0) on the Skeleton Coast in Namibia. The array also consists of four wedge-shaped sand traps, two saltation impact responders and a vertical array of three sonic anemometers (photo: author).

where Livingstone and Warren, 1996; also see Chapter 19). In such cases, calculated values of u∗ measured above a κ = von Karman’s constant (≈0.4) dune’s surface become meaningless. Corrections to measured values of u∗ also have to be m = gradient statistic of the regression equation made for the effect of thermal stability or instability, which can significantl alter the shape of velocity profiles If The averaging time required to measure values of u∗ strong thermal heating or atmospheric stratificatio oc- meaningful for sediment transport calculations using this curs, both of which are common in arid environments least squares approach is a matter of debate. While for re- when wind velocities are low, then wind turbulence may gional scale wind erosion modelling a u∗ calculated over be driven by buoyant forces rather than mechanical forces several hours may be sufficient many geomorphologists and the velocity profil may again become non-log-linear investigating relationships between changing u∗ and sed- (Frank and Kocurek, 1994). Rasmussen, Sørensen and iment transport flu have calculated u∗ on the order of Willetts (1985) found that omitting stability corrections seconds or minutes. However, at high measurement fre- could lead to errors in the measurement of u∗ of the order quencies (∼1 second) it is clear that the log-linear nature of 15–25 %, with subsequent estimates of sand transport of the turbulent wind velocity profil collapses and it is not in error by as much as 15 times (Frank and Kocurek, possible to calculate a time-averaged u∗ (Bauer, Sherman 1994). Correction factors involve the calculation of the and Wolcott, 1992; Wiggs, Atherton and Baird, 2004). Richardson number (RI), a measure of the relative strength The highest frequencies for successful u∗ calculation are of buoyancy to shear generated turbulence (Oke, 1987), of the order of 30–40 seconds (Wiggs, Atherton and Baird, and require knowledge of the local atmospheric tempera- 2004; Weaver, 2008), although Namikas, Bauer and Sher- ture gradient. However, Rasmussen, Sørensen and Willetts man (2003) report success with measurement frequencies (1985) suggested that errors in measured velocity due to as high as 10–20 seconds. such instability was a function of height and that at ele- Fluctuating topography also hinders the successful cal- vations below ∼0.5 m the correction required was unim- culation of u∗. Process research on sand dunes has been portant. From their measurements in White Sands, New particularly hampered by the recognition that acceleration Mexico, Frank and Kocurek (1994) also found that cor- of wind fl w up the stoss slopes of dunes results in the rection factors were not necessary at wind speeds higher disruption of the log-linear velocity profil (thus negating than ∼10 m/s. the calculation of u∗), with maximum acceleration oc- An additional measure of stress within the airfl w is curring in an inner-layer depth close to the dune surface the time-averaged Reynolds shear stress (−uw). This in- (Frank and Kocurek, 1996; Lancaster et al., 1996; Wiggs, volves the calculation of the average of the instantaneous P1: TIX/XYZ P2: ABC c18 JWST032-Thomas December 15, 2010 18:2 Printer Name: Yet to Come

THE NATURE OF WINDFLOW IN DESERTS 459

fluctuatin components of the wind (in comparison to the the Reynolds shear stress are becoming more common mean value) in the horizontal (u) and vertical (w) direc- in aeolian process research (van Boxel, Sterk and Arens, tions. Such data can be measured at-a-point in the airfl w 2004; Weaver and Wiggs, 2010; Weaver, 2008), with fluc by a sonic anemometer (see Box 18.1) and so the require- tuating components measured at 10 Hz and shear stress ment for a log-linear velocity profil over a significan averaged over 10 minutes showing a good correlation with depth of boundary layer is dismissed. Measurements of sediment transport rates (Weaver, 2008).

Box 18.1 From bagnold to sonics: measuring boundary layer airﬂow

Over the last 60 years one of the greatest contributions to improving our understanding of the processes of sediment mobilisation by the wind has been the extraordinary technological development in the measurement of airfl w. From the 1940s to 1960s aeolian researchers often coped with bulky mechanical weather stations where data were recorded on paper rolls controlled by clockwork and hourly average windspeeds were considered high resolution. The bulk of these instruments meant that velocity could only reasonably be measured at a single height and often several metres away from the surface. It is a sensation then that so much of our understanding of the physics of blowing sand comes from this time period and that reference to the classic works of the time is still so prevalent. In the 1970s to 1990s small, fast response cup anemometers were developed for use with electronic logging systems. These allowed much higher resolution measurements in both space and time. The small diameter of the anemometer bodies allowed several to be placed in vertical arrays without obstructing the airfl w, thus effectively measuring the vertical velocity profil and allowing a detailed investigation of shear velocity (u∗). The small plastic cups on these instruments responded quickly to accelerations and decelerations in airfl w and these short response and lag times, combined with electronic logging, enabled data to be recorded at minute intervals or better. However, these instruments were prone to failure and required heavy maintenance if used in sandy environments for extended periods. This was particularly the case as investigators attempted to measure the enigmatic relationships between sand flu and wind velocity right at the top of the saltation layer. At this time it also became clear from wind tunnel studies, where near-surface wind speeds could be measured over stabilised (nonsaltating) sand beds with fast-response hot-wire anemometers, that measurement of turbulent frequencies in the airfl w was going to be required in the field As the twenty-firs century began, aeolian geomorphologists had started using sonic anemometers in fiel situations (Figure 18.4). These instruments measure wind velocity by calculating the delay in travel time of an ultrasonic acoustic signal over a set path between paired sonic transducer heads. The difference in time delay between both directions along the same path is proportional to the wind speed. Sonic anemometers offer huge advantages in measurement techniques to desert geomorphologists. First, they calculate wind velocity within an empty volume of air between the transducer heads, and hence there is no obstruction to the fl w. Second, they

Figure 18.4 Modern equipment: a sonic anemometer in the ﬁeld at Skeleton Coast, Namibia. P1: TIX/XYZ P2: ABC c18 JWST032-Thomas December 15, 2010 18:2 Printer Name: Yet to Come

460 CH18 SEDIMENT MOBILISATION BY THE WIND

have no moving parts prone to attack from saltating particles. Third, and most importantly, they can simultaneously measure the three high-frequency (>10 Hz) components of wind velocity in the horizontal streamwise component (u), the lateral or spanwise component (v) and the vertical component (w). From these instantaneous velocity measurements it is possible to determine turbulence statistics including the uw covariance, the Reynolds stress and coherent fl w structures, all of which are now thought to be relevant in saltation dynamics. It has taken some time for investigators to devise protocols for using the anemometers in complex fiel situations and also to develop analytical schemes to make best use of the huge volumes of data that can be collected. However, the fiel application of sonic anemometers now offers a new paradigm in aeolian geomorphology.

Given some of the difficultie in assessing the erosivity tionship assumes a homogeneous surface and well-sorted of the wind by measuring various windfl w characteristics sediment. Furthermore, it ignores the effect of roughness (as decribed above) an alternative approach is to measure element spacing and scale where the aerodynamic rough- τ the surface shear stress ( 0) directly. This can be achieved ness might be determined by larger particles on a stone by using calibrated Irwin sensors or rugged drag balances pavement. For example, Greeley and Iversen (1985) noted placed directly on the sediment surface. Irwin sensors are that z0 may reach a maximum value of d/8 for widely small, omnidirectional skin friction meters that measure spaced elements, before returning to a d/30 ratio as ele- the near-surface vertical pressure gradient (Irwin, 1980), ment spacing increased further (see Figure 18.5). Further, which, once calibrated, can be used to measure high- it is clear that z0 also varies in relation to changes in mi- frequency surface shear stresses (Wu and Stathopoulos, crotopography and this suggests that mean particle size 1994). A drag balance, buried with the measuring ele- may not be a good determinant of z0 for surfaces with a ment flus to the surface or attached to a roughness ele- large range in particle sizes (Lancaster, Greeley and Ras- ment, directly measures the force applied by the wind and mussen, 1991) or changing surface patterning (e.g. sand records it on a sensitive load cell. Both technologies have ripples). been successfully tested in wind tunnels and dryland fiel For this reason aerodynamic roughness has often been locations to measure wind forces on bare sediment and determined from the intercept of velocity profile with the vegetated surfaces in a variety of fl w conditions and sur- height axis, as shown in Figure 18.2. This intercept clearly face roughness configuration (Gillies et al., 2000; King, describes z0 as define as the depth of air at the surface Nickling and Gillies, 2005; Gillies, Nickling and King, with an effective zero velocity. When plotted in the manner 2007; Walker and Nickling, 2003). of Figure 18.2, z0 can be determined from the slope and intercept components of the regression equation by

= − / z0 exp( n m) (18.5) 18.2.3 Measuring aerodynamic roughness (z0)

The aerodynamic roughness (z0) of desert surfaces varies widely both temporally and spatially. Typical values z ~ d/30 d 0 ~ for z0 may be 0.0007 m for stationary sand surfaces (Stull, 1988), 0.003 m for surfaces with moving sand (Rasmussen, Sørensen and Willetts, 1985) to 0.2 m and greater for vegetated or semi-vegetated desert surfaces z0 ~ d/8 (Wiggs et al., 1994). Owing to its importance in determin- d ing wind shear velocity, it is a vital parameter to calculate correctly in studies of aeolian processes (Levin et al., 2008). However, on many desert surfaces it is a parameter that is difficul to resolve effectively because of its sen- z0 ~ d/30 sitivity to the frequently changing scales of nonerodible d elements and surface roughness.

Bagnold (1941) found a relationship between z0 and Figure 18.5 Aerodynamic roughness height (z0) as a func- = surface roughness. He noted that z0 d/30, where d is tion of roughness element spacing (after Greeley and Iversen, the mean surface particle diameter. However, this rela- 1985). P1: TIX/XYZ P2: ABC c18 JWST032-Thomas December 15, 2010 18:2 Printer Name: Yet to Come

THE NATURE OF WINDFLOW IN DESERTS 461

where Rasmussen, Sørensen and Willetts, 1985; Sherman, 1992; Sherman and Farrell, 2008). n = intercept statistic of the regression equation The Owen (1964) equation has been used as the basis for constructing self-regulatory models of saltation (An- m = gradient statistic of the regression equation derson and Haff, 1988, 1991; Werner, 1990; McEwan and Willetts, 1991, 1993). These steady-state models work on Such an approach is common practice and works well for the premise that as u∗ (and hence saltation) increases, so fla and noncomplex surfaces (Bauer et al., 2009; Gillies, too does the effective aerodynamic roughness (z0 ). This Nickling and King, 2007; Gillies et al., 2000; King, Nick- increased aerodynamic roughness exerts an extra drag on ling and Gillies, 2005; Sherman et al., 1998; Wiggs et al., the airfl w and this effect propagates upwards through the 1994; Weaver, 2008). However, definin a clear intercept velocity profil as an internal boundary layer. This results on complex or patterned (e.g. ploughed) surfaces is prob- in near-surface winds being reduced, eventually to reach lematic where small variations in wind direction during a steady-state value where as many grains are leaving the measurement can have a dramatic effect on calculated surface as are falling onto it (Anderson, Sørensen and values of z0 as the wind profil responds to the chang- Willetts, 1991). An equilibrium is therefore established ing effective configuratio of surface properties that pro- between u∗, z0 and saltation load (Sherman, 1992). vide the surface drag (Gillette, Herrick and Herbert, 2006; The difficultie in determining a meaningful z0, as de- Wiggs and Holmes, 2010). Furthermore, on sloping sur- scribed above, are a function of the complexity of the faces the regression equation approach to calculating z0 parameter and its response to changing surface charac- suffers from the same problems as when calculating u∗, teristics in time and space (King, Nickling and Gillies, the disruption of the log-linear velocity profil due to fl w 2006, 2008). Blumberg and Greeley (1993) regard one of acceleration. the most difficul hurdles as being able to take account of An additional issue arises when a sand bed responds the effects of a surface roughness change into the develop- to increasing wind shear by allowing grain entrainment ment of a velocity profile for as there is a transition from a and saltation to take place. In such cases the saltating smooth to rough surface (or rough to smooth), a boundary sand extracts momentum from the wind, carried in the layer grows downwind in response to that transition (Fig- form of a grain-borne shear stress, and this acts as an ure 18.6). Hence, different parts of the velocity profil are additional drag on the airfl w and so increases the value of likely to be responding to different surface roughnesses z0 (McKenna-Neuman and Nickling, 1994; Sherman and with many subsidiary boundary layers responding to Farrell, 2008; Bauer, Houser and Nickling, 2004, 2009). roughnesses provided at different scales by such elements In order to satisfy the concept of an increasing z0 with as individual grains, surface ripples or isolated vegetation. increasing u∗ when saltation is occurring, several workers In the example shown in Figure 18.6, the wind passes (Rasmussen, Sørensen and Willetts 1985; Anderson and from a smooth surface (perhaps a fla sand sheet) to a Haff, 1991; McEwan, 1993) have turned their attention to the relationship applied by Owen (1964): ) 0 = 2/ * z0 C0u∗ 2g (18.6)

u*0 where Velocity = x z0 aerodynamic roughness height during saltation x

Internal = ≈ u 0 C0 constant ( 0.02) * boundary

Surface shear velocity (u y layer In this case, the depth of saltation (associated with the aerodynamic roughness height, z0 ) is related to the lift- Smooth surface Rough surface off velocity of the individual sand grains, which in turn is Distance governed by the shear velocity (u∗). While this relationship has been widely used to determine u∗ in wind tunnel Figure 18.6 The growth of an internal boundary in response studies, it has yet to be fully utilised in fiel situations to a changing surface roughness (after Greeley and Iversen, where the constant C0 is difficul to defin (although see 1985). See text for details. P1: TIX/XYZ P2: ABC c18 JWST032-Thomas December 15, 2010 18:2 Printer Name: Yet to Come

462 CH18 SEDIMENT MOBILISATION BY THE WIND

rougher surface (such as coarse-grained ripples or stone (a) pavement). At the smooth–rough transition, airfl w close Isolated-Roughness flow to the surface is quickly decelerated by the additional drag, and the aerodynamic roughness (z0) and surface shear velocity (u∗0) increase rapidly in response. As the internal boundary layer responding to the rougher surface grows downwind, shear velocity steadily decreases to reach an (b) Wake Interference equilibrium level. It should be noted that such a rise in u∗ may not necessarily lead to an increase in sediment transport or erosion because the increase in surface roughness may result in more grains lying below the height of z0 and hence in a zone of zero wind velocity, although velocity (c) may be enhanced around individual, large, nonerodible Skimming flow roughness elements and result in additional local erosion (Ash and Wasson, 1983). From Figure 18.6, it can be seen that at points marked X the wind is still in equilibrium with the smooth upstream surface, while at point Y the wind is responding to the Figure 18.7 Flow regimes associated with different rough- z u∗ rougher downstream surface. Hence, 0 and measured ness element spacings and geometries. The shaded areas rep- at any particular point in a velocity profil are functions resent zones of reduced wind speed (after Wolfe and Nickling, not only of the size and spacing of surface roughness 1993). elements in the immediate locality but also of changes in surface roughness within the fetch of the wind (Blumberg and Greeley, 1993). trolling wind erosion (Ash and Wasson, 1983; Wasson Clearly, although z0 is a critical value to determine in and Nanninga, 1986; Wolfe and Nickling, 1993; Wiggs terms of assessing aeolian sediment transport, there are et al., 1994, 1995). However, the varying geometry, spa- many difficultie in its successful and effective calcula- tial organisation and density of such roughness elements tion. There has been some success in interpreting values are difficul to account for in models of aeolian sediment and spatial variations in z0 from remote sensing. Greeley transport (King, Nickling and Gillies, 2005). et al. (1997) found good correlations between subregional The way in which nonerodible roughness elements variations in z0 and radar backscatter in the Mojave and might interact with the wind f ow was demonstrated the- Namib Deserts, demonstrating the potential to map z0 for oretically by Wolfe and Nickling (1993) and is shown large vegetation-free areas from orbit using radar systems. diagrammatically in Figure 18.7. With an increasing den- Such an approach may make it easier to determine the sity of roughness elements the fl w regime changes from temporal variation in large-scale z0 values with changing one where the elements act individually on the f ow (iso- climatic and environmental conditions, but often the scale lated roughness f ow) to a regime where the interaction of interest in aeolian processses is small enough that fiel between the wakes downwind of the elements is such that and wind tunnel studies are the only means of providing the f ow skims across the top of the elements (skimming the relevant data. fl w). In semi-arid areas the distribution of plants is often such that the airfl w is responding to the isolated roughness or wake interference regimes (Wolfe and Nickling, 18.2.4 The effect of nonerodible roughness 1993). elements on velocity proﬁles It is clear from Figure 18.7 that an increasing roughness element density reduces near-surface fl w velocities The distribution of cobbles, rocks and vegetation on other- and this is reflecte in the corresponding velocity pro- wise homogeneous desert surfaces act as a surface rough- files Using vegetation as an example, Figure 18.8 demon- ness, providing significan drag on overlying airfl w and strates that the additional drag on the airfl w provided by considerably altering velocity profil parameters (King, a vegetated surface not only increases the aerodynamic Nickling and Gillies, 2008; Gillies et al., 2000; Wiggs roughness (z0) but also displaces it upwards by a value d, et al., 1994, 1996). By altering values of u∗ and z0 such resulting in a greater depth of f ow with zero velocity at nonerodible elements can have a considerable role in con- the surface. The greater z0 of the vegetated surface also P1: TIX/XYZ P2: ABC c18 JWST032-Thomas December 15, 2010 18:2 Printer Name: Yet to Come

THE NATURE OF WINDFLOW IN DESERTS 463

10 u 1 u 2 5 * *

1 0.5

0.1 0.05

Height (m) 0.01 0.005 z 0 h 0.001 0.0005 d z0 0.0001 0 100 200 0 100 200 Velocity (m min-1)

Figure 18.8 Vegetation effect on velocity proﬁle structure: d = zero-plane displacement, h = mean vegetation canopy height, 2 1 u∗ > u∗ .

results in an increased u∗ above the vegetation canopy. surface, a vegetated surface induces a lower near-surface From measurements on vegetated and bare dunes in the wind velocity (which would tend to diminish erosion) but Kalahari Desert, Wiggs et al. (1994) found threefold in- also a higher u∗ in the airfl w (which would tend to en- creases in above-canopy u∗ and 200 % reductions in near- hance erosion). A question therefore arises as to the height surface wind velocity on vegetated dune surfaces. above the surface at which the additional energy from an The value d in Figure 18.8 (and Equation (18.3)) is enhanced u∗ is effective. In the case of skimming fl w the zero-plane displacement height and is the level of (Figure 18.7) with a large value of d, that energy is func- the mean momentum sink and the elevation at which the tional near the top of the vegetation canopy and therefore mean drag appears to act (Thom, 1971; Jackson, 1981). has little influenc on surface erosion. However, with more A practical explanation of the physical meaning of d can isolated roughness elements the vertical displacement of be given within the context of Figure 18.7. In an isolated- the velocity profil may not be sufficien to counteract the fl w roughness regime where roughness elements (e.g. increased erosional effects of an enhanced u∗, and the ex- vegetation) are widely spaced, each individual roughness tra wind stresses may be applied on the surface sediment, element provides a drag on the airfl w and so z0 is de- thus increasing erosion potential. termined by the geometry of the roughness elements and A real difficult from a sediment transport perspective d would be small enough to ignore. At the opposite ex- wth regard to predicting both dust emission and sand en- treme, with dense roughness elements, the wakes of each trainment is in determining the proportion of u∗ that acts element interact, resulting in wake interference or skim- on the surface, compared to the proportion that acts on ming fl w (Figure 18.7). In this case it is clear that the any distributed nonerodible roughness elements. Much aerodynamic roughness (z0) would be controlled only by research has been undertaken to defin a parameter that the very tops of the vegetation canopy and so the effective would successfully take account of the impact on the air- drag provided by each individual roughness element is fl w of the geometry and spacing of nonerodible rough- considerably reduced. The value of z0 therefore decreases ness elements. Methods for partitioning the shear stress markedly (in comparison to isolated roughness f ow) but between that acting on a vegetation canopy and that af- the height at which that aerodynamic roughness oper- fecting the surface often make reference to the rough- ates is displaced upwards by the zero-plane displacement ness density (λ) (Gillette and Stockton, 1989; Musick and height (d), a value that, in this example, may approach the Gillette, 1990; Stockton and Gillette, 1990): scale of the roughness elements themselves. Vegetated surfaces are therefore seen to have conflictin influence on aeolian erosion. When compared to a bare λ = nbh/s (18.7) P1: TIX/XYZ P2: ABC c18 JWST032-Thomas December 15, 2010 18:2 Printer Name: Yet to Come

464 CH18 SEDIMENT MOBILISATION BY THE WIND

where applications of the theory have been undertaken and those that have been carried out involve a very limited range of n = number of roughness elements vegetation types (e.g. Lancaster and Baas, 1998; Craw- ley and Nickling, 2003); far more experimental data are b = width of roughness elements required to develop the model further. Furthermore, application of the Raupach, Gillette and Leys (1993) model at a h = height of roughness elements regional or global scale may be hampered by the relatively complex input parameters required. s = surface area Okin (2008) has recently presented a new model for determining the effect of vegetation on wind erosion that However, such approaches have been seen to be unsatis- partly resolves this problem. In recognition of the ob- factory when describing complex roughness distributions servation that vegetation in drylands is rarely regularly and three-dimensional objects (Minvielle et al., 2003; patterned, Okin has devised a model that focuses on the Musick, Trujillo and Truman, 1996); problems may arise size distribution of gaps between nonerodible roughness in the case of vegetation where effective height may vary elements as being the driving force controlling the amount in response to changing wind speeds and plant pliability, of potential erosion (rather than the structural parameters and the drag provided by vegetation may vary with stem of the roughness elements themselves). Initial testing of porosity. the model has shown good results but a significan ad- However, there are two models for determining the vantage is that it requires relatively straightforward input impact of shear stress partitioning in the presence of data that could be acquired from field ork or large-scale nonerodible roughness elements that have become widely image analysis. recognised in the research literature. While both models compare the ratio of stress required for erosion on a bare surface to the stress required for erosion on a surface including roughness elements, they each have a distinct 18.3 Sediment in air approach to the problem (see King, Nickling and Gillies, 2005, for a full comparison). The firs is the model of 18.3.1 Grain entrainment Marticorena and Bergametti (1995), which define the ratio between the aerodynamic roughness of the surface be- Sediment is entrained into the airfl w when forces acting tween roughness elements (z0s) and the total aerodynamic to move a stationary particle overcome the forces resisting roughness (z0). This model often forms the basis for the sediment movement. The relevant forces for the entrain- emission schemes in global-scale dust models (Zender, ment of dry, bare sand are shown diagrammatically in Bian and Newman, 2003; see Chapter 20). The other ap- Figure 18.10. proach is that of Raupach (1992) and Raupach, Gillette Particles are subjected to three forces of movement: and Leys (1993), in which the wake development behind lift, surface drag and form drag. Lift is a result of the individual roughness elements is modelled to generate air f owing directly over the particle forming a region a ratio between the erosion threshold on a bare surface of low pressure (in contrast to relatively high pressure and that incorporating nonerodible roughness elements. beneath the particle); hence there is a tendency for the The input parameters to the Raupach, Gillette and Leys particle to be ‘sucked’ into the airfl w. Surface drag is (1993) model are more complex than for the Marticorena the shear stress on the particle provided by the velocity and Bergametti (1995) model and include the vegetation profil and the form drag is also related to upwind and roughness density (λ) combined with terms describing the downwind pressure differences around the particle. When vegetation aspect ratio and erosion threshold delineated as these forces overcome the forces of particle cohesion, a function of the maximum shear stress. packing and weight, the particle tends to shake in place The Raupach, Gillette and Leys (1993) model shows and then lift off, spinning into the airstream. good general agreement with fiel and wind tunnel data It has been shown that aerodynamic entrainment is pri- (Figure 18.9) and experimental testing of the procedure marily a function of the mean grain size of the particles has confirme its robustness (Brown, Nickling and Gillies, involved combined with the erosivity of the wind, with u∗ 2008; King, Nickling and Gillies, 2005, 2006; Gillette, the preferred measure (Williams, Butterfiel and Clark, Herrick and Herbert, 2006; Gillies et al., 2000, 2010; 1990). Bagnold (1941) studied these relationships and Gillies, Nickling and King, 2007). However, very few derived values of critical threshold shear velocity (u∗ct) P1: TIX/XYZ P2: ABC c18 JWST032-Thomas December 15, 2010 18:2 Printer Name: Yet to Come

SEDIMENT IN AIR 465

1.2 Laboratory Data Marshall (1971) Crawley & Nickling (2003) 1 Guelph wind tunnel

0.8

0.6

Calculated shear stress ratio 0.4

Field Data 0.2 Wolfe & Nickling (1996) Wyatt & Nickling (1997) Lancaster & Baas (1998) Luttmer (2002) 0 0 0.2 0.4 0.6 0.8 1 1.2 Measured shear stress ratio

Figure 18.9 The relationship between measured shear stress ratios in the presence of nonerodible roughness elements and calculations using the model of Raupach, Gillette and Leys (1993) (from King, Nickling and Gillies, 2005).

for a wide range of particle sizes, using the square root of the grain diameter as the principal determinant: L σ − ρ u∗ = A gd ct ρ (18.8) Wind

where D ρ = particle density C C ρ = air density

g = acceleration due to gravity

W d = grain diameter Figure 18.10 Forces exerted on a grain by the wind: l = lift, d = surface and form drag, D = form drag, w = weight, c = A = constant dependent upon the grain Reynolds num- interparticle cohesion. ber (≈ 0.1) P1: TIX/XYZ P2: ABC c18 JWST032-Thomas December 15, 2010 18:2 Printer Name: Yet to Come

466 CH18 SEDIMENT MOBILISATION BY THE WIND

entrained only by the drag and lift forces of the moving 80 air. However, one of the most significan influence on the entrainment of grains is the existence of grains already in motion. In particular, the impact of saltating grains on the bed surface imparts momentum to previously stationary 60 grains, hence ‘splashing’ them into saltation (see the later discussion). A second threshold of grain entrainment can )

-1 therefore be identified the impact (or dynamic) threshold, which refers to the threshold of entrainment where other

(cm s 40 grains are already in motion. Anderson and Haff (1988) *ct

u Fluid thresholdImpact threshold note that the impact threshold is about 0.8 of the static threshold (see Figure 18.11) because saltating grains bring extra momentum gained from higher in the velocity profil 20 towards the surface. It is the action of impacting grains that drives the saltation process in this case, rather than the lift and drag forces of the wind. This impact mechanism of entrainment is seen to be 0 particularly important in the emission of dust-sized parti- 0 0.04 0.08 0.40 0.80 1.20 1.60 cles (<62.5 µm; see Goudie and Middleton, 2006). Many Grain diameter (mm) highly emissive sources of dust are often found where the entrainment of surface sand at relatively low wind ve- Figure 18.11 The idealised relationship between grain diam- locities results in the emission of dust via the impact of eter (d) and threshold shear velocity (u∗ct) showing the ﬂuid saltating sand grains on the surface, rather than by direct and impact thresholds (after Chepil, 1945). See text for details. flui entrainment (Rice, Willetts and McEwan, 1996; Lu and Shao, 1999; Goudie and Middleton, 2006). Figure 18.11 shows the idealised relationship between Once entrainment of grains has been accomplished, flui threshold shear velocity and grain size. From this however, the relative importance of the two processes of figure it can be seen that, in general, larger particles have grain entrainment is still not entirely clear (Ungar and a higher threshold of entrainment. However, smaller parti- Haff, 1987). Rice (1991) has shown that when impact dis- cles (with diameters less than about 0.07 mm) also require lodgement becomes established, direct flui entrainment higher shear velocities to entrain them. This is because still accounts for a significan proportion of erosion. How- particles in this size range and smaller tend to have ad- ever, the consensus from empirical observations (Bagnold, ditional molecular and electrostatic forces of cohesion. 1941; Willetts and Rice, 1985a) and theoretical examina- They also have an affinit for the retention of moisture, tions (notably by Anderson and Haff, 1988; Werner, 1990; often become protected from erosion by larger particles Haff and Anderson, 1993; McEwan and Willetts, 1994) and frequently rest in the zero-velocity layer (beneath the is that grain impact is the principal mechanism by which height of z0). The most susceptible grain size for entrain- saltation is maintained. Hence, once grains are entrained ment is seen to be between about 0.07 and 0.08 mm, i.e. at the flui threshold, a reduction in wind velocity will fin textured sand. The heightened mobility of sediment in not necessarily result in a reduction in sand transport so this size range in drylands enables large volumes of sand long as the velocity remains above the impact threshold, to be moved and accumulate, resulting in the extensive as shown in Figure 18.11. dunefield present in many drylands (see Chapter 17). The relationships shown in Figure 18.11 tend to result in a downwind finin of aeolian sediment both at the scale 18.4 Determining the threshold of the individual ripple and across whole sand seas. Such of grain entrainment a downwind sorting can be used to identify process–form relationships; sand surfaces subject to aeolian erosion tend The determination of the wind velocity (ut) or shear veloc- to consist of a coarse lag material (Nickling and McKenna- ity (u∗t ) threshold for grain entrainment is a fundamental Neuman, 1995), while those subject to aeolian deposition necessity in the understanding and prediction of the ae- are commonly composed of fin material, such as loess. olian sediment transport system. There are many means The discussion above refers to the value of the static or by which threshold values can be ascertained, although ﬂuid critical threshold of entrainment where the grains are the majority of techniques have been developed in the P1: TIX/XYZ P2: ABC c18 JWST032-Thomas December 15, 2010 18:2 Printer Name: Yet to Come

DETERMINING THE THRESHOLD OF GRAIN ENTRAINMENT 467

controlled conditions of the wind tunnel and involve the providing a range in the value of possible entrainment observation of the onset of sediment motion with in- thresholds. creasing wind velocity (see, for example, Bagnold, 1941; Stout and Zobeck (1996, 1997) and Stout (2004) tackled Kawamura, 1951; Zingg, 1953). While such studies have the problem by developing the time fraction equivalence helped in the development of relationships such as that method, which has begun to be applied in a variety of shown in Figure 18.11 they have not been able to greatly fiel situations (Wiggs, Atherton and Baird, 2004; Wiggs, improve our knowledge of thresholds in natural fiel en- Baird and Atherton, 2004; Davidson-Arnott, MacQuar- vironments where the identificatio of the onset of sedi- rie and Aagaard, 2005). The technique is based on the ment ransport is complicated by a constantly varying wind observation that sand transport is intermittent (when mea- velocity. sured at a frequency of ∼1 Hz) and the assumption that However, the development of electronic grain impact the fraction of time in which sand transport events oc- sensors such as the Sensit (Stockton and Gillette, 1990) cur (described by the intermittency factor, γ ) should be and the Safir (Baas, 2004) have enabled us to inves- equivalent to the fraction of time that the wind velocity is tigate entrainment thresholds in the fiel more fully. A equal to or greater than the threshold value for a particular typical fiel data series of wind velocity and sediment surface and at a particular time. By converting measure- transport occurrence is shown in Figure 18.12 and demon- ments of sand transport and wind speed into two binary strates the difficultie inherent in identifying one single series, the threshold value that gives rise to this equiva- entrainment threshold. While a threshold value could lence can be determined by iteration (Stout and Zobeck, be considered as the wind velocity at which the firs 1996). A modificatio of the procedure employing a cu- grains are seen to be entrained into transport (A in Fig- mulative frequency plot of wind velocity has also been ure 18.12), it is also clear that there are many instances presented by Wiggs et al. (2004a) and is demonstrated in where no sediment transport is detected despite winds Figure 18.13. being far greater than this. In order to provide a quan- An entrainment threshold calculated using the time titative means by which a single threshold can be de- fraction equivalence method is shown as B in Figure termined from such data series, Davidson-Arnott, Mac- 18.11. While the technique allows a good objective Quarrie and Aagaard (2005) recommend calculating the method by which to calculate a threshold, it is clear mean of the f ve minimum windspeeds at which trans- that the inherent variability in sand transport data again port is observed and the mean of the f ve maximum results in a threshold whereby there are instances of windspeeds at which no sand transport is observed, thus both sand transport events below threshold and no sand

40 B 30

20 A Sand transport (impacts) 10

0 012 345678 u m s-1

Figure 18.12 Time series data of wind speed and detected sand transport (at 1 Hz sampling frequency). Entrainment thresholds deﬁned by (A) the onset of sediment transport and (B) the time fraction equivalence method are shown (from Wiggs, Atherton and Baird, 2004). P1: TIX/XYZ P2: ABC c18 JWST032-Thomas December 15, 2010 18:2 Printer Name: Yet to Come

468 CH18 SEDIMENT MOBILISATION BY THE WIND

100 An alternative to comparing field-measure time series of wind and sand transport data involves the direct 80 determination of entrainment thresholds from small-scale fiel wind tunnels. Such experiments reduce wind and surface variability and so tend to give better correlations between sand transport and wind velocity. While tradi- No STEs STEs tional portable wind tunnels have provided useful data on entrainment thresholds (Belnap and Gillette, 1998) they are bulky to manoeuvre and so there are limitations on

Cumulative percentage Cumulative the surfaces that can be sampled. A recent development has been that of the Portable In-Situ Wind Erosion Lab- oratory (PI-SWERL) (Etyemezian et al., 2007), which is 0 a relatively small circular wind tunnel. The portability Low ut High of the PI-SWERL offers the ability to obtain numerous Wind speed replicates of the critical wind speed required for erosion Figure 18.13 An idealised example of the modified time frac- over a broad range of surfaces that have previously been tion equivalence method for determining the threshold of sed- impractical to measure in the field iment entrainment using a cumulative frequency plot of wind velocity and a calculated saltation intermittency value (γ )of 20 %. The threshold corresponds to the 80th percentile value 18.5 Surface modifications to of wind velocity (from Wiggs, Atherton and Baird, 2004). entrainment thresholds and transport flux

While the relationships shown in Figure 18.11 have been transport events above threshold. Wiggs, Atherton and found to be satisfactory for loose, dry, fla and homo- Baird (2004) noted that in their data set such problems re- geneous surfaces (Williams, Butterfiel and Clark, 1994), sulted in <60 % of their sand transport events accounted the critical thresholds of motion and grain transport flu es for by the calculated threshold. on natural sediment beds are also influence by variations Improvements to the explanatory power of thresholds in factors such as sediment mixtures, surface crusting, calculated in this way can be achieved if a lag time is surface slope, moisture and vegetation. The relationships assumed between the changing wind velocity and the re- between these parameters are frequently complex and are sponse of the sand transport mechanism. Many investiga- not yet fully understood. The effects of vegetation are tors have found that the response time of the sand transport dealt with above. Three other especially important mod- system to wind variability is of the order of ∼1 second ification are provided by surface crusting, surface slope (Wiggs, Atherton and Baird, 2004; Davidson-Arnott and and moisture content. Bauer, 2009; Weaver, 2008) and so lagging sand transport behind velocity by this amount results in a much closer association between these two data series. 18.5.1 Surface crusting While the identificatio of a single entrainment threshold remains an appealing prospect, it is clear that variabil- A surface crust lying above loose, erodible sediment sig- ity inherent in the natural system (in terms of paraemeters nificantl inhibits the entrainment of sediment into the such as grain size, microtopography, packing, moisture) wind. In wind tunnel studies Zobeck (1991) found that makes it an unachievable goal. Such complexity in the soils without surface crusts were 40–70 times more erodi- system means that entrainment thresholds for a particular ble than crusted soils and similar impacts on erodibility surface are perhaps more appropriately represented by a have been found by investigations in wind tunnels (Rice, range of possible thresholds (Wiggs, Atherton and Baird, Mullins and McEwan, 1997; Rice and McEwan, 2001; 2004; Davidson-Arnott, MacQuarrie and Aagaard, 2005; McKenna-Neuman and Maxwell, 2002) and the fiel Davidson-Arnott et al., 2008) whereby the probability of (Rajot et al., 2003; Houser and Nickling, 2001a, 2001b; erosion may be modelled by investigation of the cumu- Hupy, 2004). However, the effect of crust characteristics lative probability density functions of time series data of on aeolian processes is still poorly understood and so its wind velocity and sand transport (Davidson-Arnott and impact is often disregarded in aeolian transport models Bauer, 2009). (Zobeck et al., 2003). P1: TIX/XYZ P2: ABC c18 JWST032-Thomas December 15, 2010 18:2 Printer Name: Yet to Come

SURFACE MODIFICATIONS TO ENTRAINMENT THRESHOLDS AND TRANSPORT FLUX 469

Physical crusts form where limited rainfall impacts on 2.0 a soil with at least 5 % clay content (Rajot et al., 2003), 1.8 although Rice and McEwan (2001) state that the greater ucrb / ucro the proportion of fin (cementing) material, the greater 1.6 the strength of the crust with surfaces containing <12 % fine being easily eroded. Hupy’s (2004) fiel investiga- 1.4 tion found that thick, silty, physical crusts could be as protective of a surface as a cover of gravel or grass. In 1.2 contrast, McKenna-Neuman and Maxwell (2002) under- took a wind tunnel test of the breakdown of microphytic (moss) crusts and found that they were weak compared 0.8 to physical crusts and could deteriorate under the impact of saltating grains even at low wind velocities. However, 0.6 they also found that in isolated cases the elasticity of such microphytic crusts could protect the soil beneath under 0.4 heavy saltation bombardment if the integrity of the f la- 0.2 ment net in the crusts was strong. On silt/clay surfaces a crust may significantl protect the underlying fine dust-sized particulates from erosion. 30 20 10 0 10 20 30 However, where such a sediment bed is of mixed size and Upslope Downslope includes sand-sized particles that are prone to saltation, the impact of saltating particles at relatively low wind Figure 18.14 The effect of bedslope on threshold of grain = velocities can break down the crust and induce intense entrainment: solid line theoretical models of Allen (1982) and Dyer (1986), circles = experimental data (from Hardisty dust entrainment. Such bombardment is thought to be the and Whitehouse, 1988). primary process in the entrainment of dust-sized material (Houser and Nickling, 2001a), with the most severe ◦ of climbing a 20 slope under normal wind conditions, sources of dust emission from playa surfaces being those explaining the accumulation of large particles at the base associated with a surface cover of erodible sand (Cahill of dune windward slopes. et al., 1996). Surface slope may not only have an important effect on the threshold of sediment movement but also on the rate 18.5.2 Bedslope (or flux of sediment transport. Despite much discussion of the potential effect (Hunt and Nalpanis, 1985; Nalpa- nis, 1985; Hardisty and Whitehouse, 1988; Whitehouse Despite the potential importance of surface slope on the and Hardisty, 1988; Iversen and Rasmussen, 1994), con- threshold of entrainment and saltation trajectories of par- clusions are far from complete. Bagnold (1941) derived ticles (some windward dune slopes may reach angles ◦ a simple geometric relationship to describe the effect of aproaching 30 ), there has been relatively little empirical bedslope on the sand transport rate, although Howard et al. research in this area. Theoretical analyses have been pre- (1977) found that the Bagnold formula had only a minor sented by Allen (1982) and Dyer (1986) while Hardisty influenc on transport rate predictions and characterised and Whitehouse (1988), Iversen and Rasmussen (1994) the actual sand transport rate no better than if bedslope was and Rasmussen, Iversen and Rautahemio (1996) have not taken into account. In contrast, Hardisty and White- used portable fiel and tilting laboratory wind tunnels. house (1988) found that the sand transport rate was much Hardisty and Whitehouse (1988) found good agreement more dependent upon surface slope than predicted by the between theory and practice, with results indicating a Bagnold relationship (Figure 18.15). relative increase in critical threshold on positive slopes (upslope) and a decrease on negative slopes (downslope) (Figure 18.14). Computational flui dynamics (CFD) 18.5.3 Moisture content modelling by Huang, Shi and van Pelt (2008) confirme this influenc of dune slope angle on particle entrainment It is clear that a high moisture content in surface sed- while the fiel work, modelling and wind tunnel testing iment can substantially increase entrainment thresholds of Tsoar, White and Berman (1996) and White and Tsoar and reduce transport potential, except where the erosive (1998) suggest that only particles < 230 µm are capable power provided by wind speed is particularly forceful P1: TIX/XYZ P2: ABC c18 JWST032-Thomas December 15, 2010 18:2 Printer Name: Yet to Come

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24 content. Furthermore, Wiggs, Baird and Atherton (2004) noted that moisture that was adhered to saltating particles 20 could reach 2 % (gravimetric) before any influenc on the In k / k b o transport rate could be detected. Saltation of dry sediment 16 is therefore possible across wet surfaces and Sarre (1988, 12 1990) found that surface moisture contents of up to 14 % had little effect on sand already in transport. Such natural 8 inhomogeneity in surface moisture is not taken account of in the static conditions of wind tunnel tests and such data 4 tend to show a shutdown of the saltation system at lower average surface moisture values. Complexities in the influenc of soil moisture on the -4 saltation system are now being investigated on beaches. Studies by Bauer et al. (2009), Davidson-Arnott and -8 Bauer (2009), Jackson and Nordstrom (1997, 1998), Wiggs, Atherton and Baird (2004), Wiggs, Baird and et al. 30 20 10 0 10 20 30 Atherton (2004) and Davidson-Arnott (2008) have Upslope Downslope all noted the spatial and temporal variability in surface moisture content on beaches and have investigated the Figure 18.15 The effect of bedslope on sand transport rate: impacts of such variability on saltation dynamics. In solid line = geometric relationship (Bagnold, 1941), circles = particular, the sensitivity of sediment entrainment to the experimental data (from Hardisty and Whitehouse, 1988). moisture status of a very thin layer of surface sediment has been explored. Once these surface grains have dried to a sufficien extent to allow entrainment the damper (Wiggs, Baird and Atherton, 2004; Bauer et al., 2009). grains below are revealed and critical entrainment However, its influenc is complex, variable and not well thresholds rise once again. This high temporal dynamism understood. The majority of our knowledge comes from in entrainment thresholds, coupled with variation in wind tunnel investigations (e.g. Cornelis and Gabriels, wind speeds, produces a highly intermittent saltation 2003; Han et al., 2009; McKenna-Neuman and Nickling, system where erosion of surface sediment at a specifi 1989), although recent fiel studies on beaches have de- location can occur at a range of wind speeds within a very veloped our comprehension of the complexities involved short space of time. Davidson-Arnott and Bauer (2009) (Bauer et al., 2009; Davidson-Arnott and Bauer, 2009; therefore suggest that there exists a range of entrainment Wiggs, Baird and Atherton, 2004). The theoretical basis thresholds, rather than a single f xed value based on an for available predictive models is that the critical shear ve- average moisture content and grain size. Such sensitivity locity required for entrainment increases as a function of of the saltation system to the moisture status of the top the increased surface tension associated with pore mois- few grains results in the dynamics of the system being ture. However, as shown in Figure 18.16, there is much readily influence by air humidity. McKenna-Neuman disagreement as to the precise mechanism by which this and Sanderson (2008) and McKenna-Neuman (2003, should be applied. In this figur the Hotta et al. (1984) 2004) note that colder airfl ws support much higher sand and Belly (1964) data are based on empirical wind tun- transport rates than warmer air. They partly explain this nel measurements while those of Kawata and Tsuchiya by the decreased adsorption of moisture from the air to (1976) are theoretical. surface grains and hence reduced interparticle cohesion The data shown in Figure 18.16 suggest that the entrain- and lower entrainment thresholds in the cold air case. ment system is perhaps very sensitive to changes in mois- With the requirement for higher spatial and temporal ture status in the range 0–4 %. However, in fiel exper- resolution measurements of the moisture status of surface iments conducted by Wiggs, Baird and Atherton (2004) sediment new methods of measurement are being explored significan limitation on sediment entrainment was found that do not rely on gravimetric analysis of grab samples. only at higher moisture values (4–6 %). They argued that Such grab samples may include substantial amounts of this was due to spatial inhomogeneity in the wetness of subsurface (and wetter) sediment. One new method re- surface sediments such that drier sediment on topographic ceiving attention is that of measurement of surface bright- highs (such as crests of ripples) could be entrained and ness using calibrated digital photography, which has been transported at high values of average surface moisture shown to be a good indicator of the amount of pore water P1: TIX/XYZ P2: ABC c18 JWST032-Thomas December 15, 2010 18:2 Printer Name: Yet to Come

MODES OF SEDIMENT TRANSPORT 471

1.2

1.0 ) -1 0.8 (m s

ct Kawata & Tsuchiya (1976) *

0.6

Hotta (1984)

0.4 Belly (1964) Critical shear velocity u

0.2

0.0 0481216 20 24 28 Moisture content (%)

Figure 18.16 Critical shear velocity for sediment entrainment as a function of sand moisture content (from Sherman, 1990).

present (McKenna-Neuman and Langston, 2006; Darke loess or dust (see Chapter 20). At the coarse end of this and McKenna-Neuman, 2008; Darke, Robin and Oller- spectrum, some material may be transported in modifie head, 2009). saltation (Hunt and Nalpanis, 1985), where saltation particle trajectories are affected by wind turbulence. 18.6 Modes of sediment transport 18.6.2 Creep Once entrained into the airfl w, sediment may be transported by any one of four mechanisms, principally depen- Larger particles tend to be transported by one of the three dent upon the sediment grain size (Bagnold, 1941, and modes of bedload (or contact) transport: creep, reptation Figure 18.17). The following modes of transport are not or saltation. Surface creep describes the rolling action of discreet classes and the transition from one to another may coarse particles (0.5–2.0 mm) as a result both of wind not be well defined drag on the grain surfaces and the impact of high-velocity saltating grains. In the case of the fine particles in this size range, a creep movement may become apparent im- 18.6.1 Suspension mediately prior to the onset of saltation (Nickling, 1983; Willetts and Rice, 1985a). The difficult in isolating creep Small particles of less than 62.5 µm (Goudie and Mid- in experimental observations has made the task of definin dleton, 2006) whose settling velocity may be very small the relative importance of creep in terms of other trans- in comparison to the combined effects of wind lift and port mechanisms problematic. Willetts and Rice (1985b) drag might be transported in suspension, with the vertical estimate that creep accounts for approximately one quar- profil of flu decreasing with height and described by ter of the bedload transport rate while Dong et al. (2002) a power function (Nickling, McTainsh and Leys, 1999; found in their wind tunnel tests that the creep transport Wang et al., 2008). The turbulent motion of airfl w can rate increased with wind speed but decreased with grain keep very fin sediment suspended for many days, high size, while the creep fraction varied widely, ranging from in the atmosphere, which may ultimately be deposited as about 4 to 29 %, averaging at about 9 %. P1: TIX/XYZ P2: ABC c18 JWST032-Thomas December 15, 2010 18:2 Printer Name: Yet to Come

472 CH18 SEDIMENT MOBILISATION BY THE WIND

Saltation proper Wind Sand Sheets Bed load (rolling and saltation) t Sand * Outgoing grain U Dune Sands Reptation Incoming grain = 0.1 / U * Approach 0.1 U f = 0.7 angle / U * U f 3m

Figure 18.18 The process of reptation where the impact of a = 0.1 Loess high-velocity saltating grain ejects other grains into the airﬂow / U * Modified saltationShort term U f (after Anderson, 1987). suspension

Grain diameter (mm) diameter Grain tion (Haff and Anderson, 1993). Much more research is

Silt required to understand this process fully, but Anderson, 0.01 Long term suspension Sørensen and Willetts (1991) suggest that it could be extremely important in near-surface aeolian transport and more recent modelling work confirm that transport flu models produce more realistic simulations with the inclusion of a reptation component (Andreotti, 2004; Namikas, 2003). Clay

Dust storms 18.6.4 Saltation 0.001 10 30 50 70 100 200 The most intensively researched mode of aeolian transport -1 U* (cm s ) is that of saltation. This is the characteristic ballistic tra- ≈ Figure 18.17 The relationships between grain diameter, jectory of grains ( 0.06–0.5 mm in diameter) as they are shear velocity and mode of sediment transport showing the ejected from the grain bed, are given horizontal momen- distinction between the suspension of dust-sized sediment and tum by the airfl w, descend to impact the grain bed and the bedload transport of sand-sized sediment; Uf is the particle then continue ‘leaping’ downwind (see Figure 18.19). The fall velocity (after Tsoar and Pye, 1987). impact of saltating grains on the surface often leads to the splash of other grains into the airfl w, which may then un- dergo reptation or, if sufficien momentum is transferred, 18.6.3 Reptation saltation. Hence, saltation is a very efficien transporting mechanism whereby a few saltating grains can rapidly The low hopping of several grains consequent upon the high-velocity impact of a single saltating grain has been Wind termed reptation (Anderson and Haff, 1988) and is an important transitional state between the modes of creep and w1 saltation. On impact and subsequent rebound, a saltat- Initial u u ing grain may lose 40 % of its velocity (Willetts and 1 Final 2 Rice, 1989; Anderson and Haff, 1991; Haff and Ander- w2 son, 1993). This energy imparted to the grain bed re- h α sults in the ejection (or ‘splashing’) of perhaps 10 other grains (Werner and Haff, 1988) with velocities at ap- l proximately 10 % of the impact velocity, often too low to enter into saltation (Willetts and Rice, 1989; Ander- Figure 18.19 The ballistic trajectory of a saltating sand son and Haff, 1991). Hence, each grain takes a single grain; w and u represent vertical and horizontal velocities re- hop (Figure 18.18), the majority in a downwind direc- spectively (after Bagnold, 1941). P1: TIX/XYZ P2: ABC c18 JWST032-Thomas December 15, 2010 18:2 Printer Name: Yet to Come

RIPPLES 473

and longer is its saltating jump. The manner of momentum extraction from the airfl w is demonstrated by the wind tunnel experiments of Rasmussen and Sørensen (2008), who recorded 1:1 ratios between air and grain speed within 80 mm of the surface, increasing to a ratio of 2.0 at 5 mm height. Grain size also has an influenc on trajectory paths, with larger grains rebounding from the surface at lower speeds and following a shorter and lower trajectory (Namikas, 2003). Similarly, Willetts (1983) noted that grain shape influence the trajectory with platy grains tending to have lower and longer paths than spherical grains. The length of jump is thought to be of the order of Figure 18.20 Saltating sand being swept from the crest of 12–15 times the height of bounce (Livingstone and War- an 80 m high linear dune in the Namib Desert. Increasing ren, 1996) and the trajectory is also influence if the grain wind speeds and saltation impacts result in progressively more starts to spin after a glancing impact. Reports of spin intense saltation from dune foot to dune crest (photo: author). rates reaching over 400 r.p.s (White and Schultz, 1977; White, 1982) can induce a lift force (termed the Magnus effect) that may extend saltation trajectories. Particle col- induce mass transport of sediment in a cascading system lisions with the surface effectively convert near-horizontal ◦ (Figure 18.20), often termed the ‘fetch effect’ (Gillette momentum (with angles of descent up to 10 )toaverti- et al., 1996). A limit to the amount of sediment in salta- cal momentum via ejection or rebound. Ungar and Haff tion is set when an equilibrium is established between the (1987) noted that the number of grains splashed up into flu es of sediment and air as a result of the momentum the airstream by an impacting grain was proportional to extracted from the airfl w by the saltating grains. This the square of the impacting grain velocity, which may be momentum extraction causes a considerable reduction in up to 5 times its initial velocity (Anderson and Hallett, near-surface wind speeds during saltation (Owen, 1964) 1986). and results in the majority of grain ejections being intiated The height of the saltation layer is dependent on the by impact from other saltating grains rather than by direct wind speed (Dong et al., 2002), the grain size in transport flui entrainment. and surface characteristics. Bagnold (1941) recognised Bagnold (1936, 1941) was the firs to appreciate that saltation leaps were higher on a pebbly or hard rock the ballistic trajectory of saltating grains and there surface than on a loose sand surface because hard sur- are numerous studies that have investigated the sub- faces are less absorbent of the grain momentum on each processes of grain entrainment, trajectory, bed colli- bounce. Pye and Tsoar (1990) quote a maximum salta- sion and velocity profil modification principally rely- tion height on such surfaces of 3 m, although heights of ing on wind tunnel and numerical modelling approaches ∼0.2 m are more commonly found. However, the mass of (e.g. Willetts and Rice, 1986a, Anderson and Hallett, saltating particles decreases very rapidly with height and 1986; Werner and Haff, 1988; Anderson and Haff, has been described by a declining exponential function 1988; Rasmussen and Mikkelsen, 1991; Werner, 1990; (Rasmussen, Sørensen and Willetts, 1985; Dong et al., McEwan, Willetts and Rice, 1992; McEwan, 1993; Haff 2002; Rasmussen and Sørensen, 2008), with up to 80 % and Anderson, 1993; Dong et al., 2002; Kok and Renno, of all transport taking place within 2 cm of the surface 2009b; Namikas, 2003; Rasmussen and Sørensen, 2008; (Butterfield 1991). Bauer, Houser and Nickling, 2004). Bagnold (1941) and Chepil (1945) both identifie steep ◦ initial take-off angles of grains at angles approaching 90 . 18.7 Ripples However, other studies have recognised a range of lift- ◦ off angles of between 15 and 70 (White and Schultz, The sediment transport mechanisms of creep, reptation 1977; Nalpanis, 1985; Willetts and Rice, 1985a; Ander- and saltation combine to create ripples, a mobile bed- son, 1989; Namikas, 2003; Kok and Renno, 2009b). The form that contributes to bulk sediment transport. Despite actual grain trajectory depends very much on the height of decades of research these features still remain enigmatic bounce of the grain into the boundary layer. Wind speeds and our understanding of them is incomplete. Wind tunnel increase at a logarithmic rate away from the surface (see and numerical modelling investigations on the controls on Figure 18.2); hence, the higher a grain leaps, the more aeolian ripple dynamics and wavelengths was common up momentum it can extract from the airfl w and the faster to the 1990s. More recent modelling work has treated the P1: TIX/XYZ P2: ABC c18 JWST032-Thomas December 15, 2010 18:2 Printer Name: Yet to Come

474 CH18 SEDIMENT MOBILISATION BY THE WIND

problem from a nonlinear perspective (Anderson, 1990; that grain size has a fundamental control on ripple geom- Yizhaq, Balmorth and Provenzale, 2004) or by using a etry (Anderson and Bunas, 1993; Baas, 2007). Theoreti- complex systems approach employing aspects of emer- cal and empirical studies by Ellwood, Evans and Wilson gent behaviour and self-organisation in cellular automa- (1975) led to the conclusion that mega- and normal rip- ton models (Werner and Gillespie, 1993; Anderson and ples do not form distinct populations, but are the upper and Bunas, 1993; Baas, 2002, 2007; Pelletier, 2009). lower bounds of a continuum in wavelengths as a result Bagnold (1941) recognised three categories of ripples: of differences in grain size (Ellwood, Evans and Wilson, normal or ballistic; granule, sand ridge or mega; and 1975) and wind speed, with wavelength responding posi- ﬂuid drag or aerodynamic. They generally form transverse tively to increases in both controls (Sharp, 1963; Claudin to the wind direction in repeated patterns that continu- and Andreotti, 2006; Andreotti, Claudin and Pouliquen, ally adjust in response to variability in windfl w. Ripples 2006). commonly have wavelengths of 1–25 cm and heights of 0.5–1.0 cm (Sharp, 1963) with a asymmetrical profil con- ◦ sisting of windward slopes of about 10 (Mabbutt, 1977). 18.7.1 Ballistic ripples However, megaripples can achieve crest-to-crest wavelengths exceeding 20 m, are characterised by a bimodal Bagnold’s (1941) impact mechanism theory has been distribution of coarse and fin particle sizes (Yizhaq et al., widely used to explain ripple formation. In this theory rip- 2008) and tend to be more symmetrical in profil (Gree- ple wavelength is related to a characteristic or mean salta- ley and Iversen, 1985). This symmetry may be related tion path length. Surface irregularities act as erosional and to shifts in wind direction, suggesting that the form and depositional nuclei for moving particles (Figure 18.21). spacing of larger ripples are less dynamic. However, cel- The windward side of an irregularity is bombarded by lular automaton models of ripple formation also suggest more saltating grains per unit area than the sheltered lee

(a) Wind direction

A C

Descending limb of Rising limb of saltation paths saltation paths

(b) Wind direction

Crestal accumulation of coarser grains

° Veneer of 10 30° finer grains

Core of Foreset bed finer grains

Figure 18.21 (a) Ripple development (after Bagnold, 1941). AB is preferentially bombarded by descending grains compared with BC (see text for details). (b) Laminae and grain-size distribution within a ripple (from diagrams in Sharp 1963). P1: TIX/XYZ P2: ABC c18 JWST032-Thomas December 15, 2010 18:2 Printer Name: Yet to Come

PREDICTION AND MEASUREMENT OF SEDIMENT FLUX 475

side and so is an area of net erosion. The preferential 18.8 Prediction and measurement of loss of particles on the windward side of the emergent sediment ﬂux ripple creates a new zone of bombardment downwind, at a distance equivalent to the mean saltation path length. In order to quantify surface erosion or deposition, geo- Hence, the ripple pattern migrates downwind as a series morphologists are often interested in the changing rate of alternating zones of erosion and protection. Coarser of sediment flu in space and/or time. There are many grains, which are not set into saltation by the impact of equations to calculate mass sand flu from wind velocity bombarding grains, creep forward to accumulate in the data, nearly all derived from theoretical or wind tunnel less bombarded crestal zone. experimental work (see Greeley and Iversen, 1985, for a However, Sharp (1963) notes that ripple spacing in- review). All the expressions tend to the form of creases with time and so it is unlikely to be related to the mean saltation path length under steady wind condi- 3 q = Au∗ (18.9) tions. In contrast, he suggested that the angle of incidence of descending particles and ripple height determined rip- The two types of relationship frequently used to calcu- ple spacing. Height can perhaps be seen as crucial; given late sand flu are typifie by the expressions of Bagnold the narrow range of ripple slope angles that have been (1941) and Kawamura (1951), the latter incorporating a recorded, it must geometrically affect the length of the specifi term for the threshold shear velocity for entrain- windward and lee slopes. Thus, the minimum wavelength ment: must increase with height (Brugmans, 1983) and ripple wavelength must increase correspondingly as ripple 0.5 3 q = C(d/D) u∗ρ/g (Bagnold, 1941) growth proceeds. As the ripple protrudes further into the boundary layer and coarser grains on the crest are more (18.10) readily moved forward, the height of the ripple becomes limited (Greeley and Iversen, 1985). = − − 2ρ/ , Wind tunnel and theoretical studies have indicated q Kk (u∗ u∗ct)(u∗ u∗ct) g (Kawamura 1951) how saltation dynamics might influenc ripple wave- (18.11) lengths. Saltating particles require vertical lift-off momentum while creeping grains require forward momentum. Willetts and Rice (1986a, 1986b) have calculated that where bombarding grains transfer relatively more vertical energy − − on steeper slopes and more horizontal energy on shallower q = sand transport rate (gm 1 s 1) slopes. Thus, saltation is favoured from the middle of the windward side of ripples and creep activity increases C = constant (1.8 for naturally graded dune sand) towards the crest, favouring the accumulation of coarse grains (Figure 18.21). d = grain diameter Anderson (1987) has suggested that while ripple wavelength is affected by grain trajectory lengths, it is not the D = standard grain diameter (0.25 mm) same distance as the mean saltation path length. A crit- icism of Bagnold’s (1941) theory was that it relied on a ρ = air density very narrow range in saltation path lengths, although experimental studies showed that saltation trajectories were Kk = constant (2.78) often widely distributed and variable (Mitha et al., 1986). Anderson’s (1987) model therefore includes a wide range u∗ct = threshold of grain entrainment of saltation trajectory lengths and speeds that drive a reptating population. In this model ripples are formed through Well-known problems with the Bagnold (1941) expres- the accumulation of these reptating grains, with a wave- sion include the fact that it predicts sand movement below length that is a function of the probability distribution of the threshold of entrainment and it commonly predicts the total trajectory population and the ejection rate of the rates that are considered too low at high values of shear reptating grains. This reptation control of ripple develop- velocity (Sarre, 1987). Owing to the inclusion of a thresh- ment has support from the wind tunnel experiments of old term in the Kawamura (1951) calculation, it is more Willets and Rice (1986b) and simulations of Werner and accurate at lower levels of shear velocity. However, this Haff (1988). expression only incorporates the effect of grain size in the P1: TIX/XYZ P2: ABC c18 JWST032-Thomas December 15, 2010 18:2 Printer Name: Yet to Come

476 CH18 SEDIMENT MOBILISATION BY THE WIND

threshold term (u∗ct), despite the fact that it is also likely 160 to have an important effect on the transfer of momentum Measured rate on grain impact. 140 Bagnold (1941) Lettau & Lettau (1978)

Zingg (1953) followed a similar argument to Bagnold )

3/ -1 120 Zingg (1953)

(1941) but used a 4 power function: s -1

3/ 100 = / 4 3ρ/ q C2(d D) u∗ g (18.12) 80 where 60 = C2 constant (0.83) 40 Sand transport rate (g m Owen (1964) found that this expression was more accurate over a wider range of particle sizes than the Bagnold 20 (1941) formula, but it still omits a term for the threshold 0 of entrainment. Such a term was incorporated into a later 0.2 0.3 0.4 0.5 0.6 0.7 0.8 analysis of Bagnold (1956), which was subsequently re- Shear velocity (u*) fine by Lettau and Lettau (1978) into a very commonly used expression: Figure 18.22 Relationship between shear velocity, measured sand transport rate and that calculated from three sand trans- = / 0.5 − 2ρ/ port rate formulae (after Wiggs, 1992). q C3(d D) (u∗ u∗ct)u∗ g (18.13)

where thin vertical design (hence presenting a minimum disturbance to the wind f ow) and incorporate the bleeding of air = C3 constant (4.2) from within the trap to reduce backpressure. The sampling efficiencie of the various types of trap vary considerably Despite this array of transport equations, no specifi model and have been reported to be as low as 20 % (Knott and has proven to be broadly applicable under a variety of nat- Warren, 1981) and as high as 70 % (Gillette and Goodwin, ural environmental conditions. This is partly because the 1974; Marston, 1986). However, the determination of ef- majority of published models have been derived from the- ficiencie is fraught with problems and Jones and Willetts ory or experimentally from wind tunnel observations with (1979) have shown that small differences in the operation homogeneous sediment beds (Spies, McEwan and Butter- of traps in the fiel may give rise to large contrasts in field 1995; Butterfield 1999; Bauer, Houser and Nick- efficien y. ling, 2004), without sufficien testing or development in Substantial improvements in sampling efficien y have the field Field testing that has been accomplished (Sarre, been provided by the incorporation of wings extending 1987; Sherman, 1990; Wiggs, 1992) has shown a great downwind from trap openings, giving the trap a wedge- deal of variation between the observed rates and those shaped appearance (Figure 18.23). This further decreases predicted as a function of u∗ (Figure 18.22). Given the backpressure and allows sand flu collectors to become natural variability and complexity over short spatial and almost isokinetic in terms of their f ow characteristics temporal scales in grain size, moisture status, bedslope (Nickling and McKenna-Neuman, 1997). While these de- and wind conditions, as previously described, such weak signs have achieved near 90 % sampling efficiencie in relationships are not surprising. wind tunnel tests their performance in the fiel is severely Furthermore, fiel calibration of sediment flu equa- restricted by their sensitivity to the incident angle of the tions is severely hampered by the errors inherent in mea- wind. Nickling and McKenna-Neuman (1997) report that ◦ suring sand flu in the field In order to measure the sand incident angles as low as 5 , which may occur frequently transport rate, a collecting device has to be inserted into during periods of measurement, severely reduce the sam- the airfl w for a known period of time. This inevitably pling efficien y of such traps (McKenna-Neuman, Lan- causes airfl w disturbance and hence error in flu mea- caster and Nickling, 2000; Namikas, 2003). surement. Two major problems with sand traps are those More recent attempts to characterise saltation activity of back pressure and scouring of sand around the base of in the fiel have involved the use of saltation impact re- traps. To reduce these errors, the majority of traps are of a sponders. A variety of designs exist but all operate on a P1: TIX/XYZ P2: ABC c18 JWST032-Thomas December 15, 2010 18:2 Printer Name: Yet to Come

PREDICTION AND MEASUREMENT OF SEDIMENT FLUX 477

Figure 18.23 A series of isokinetic wedge-shaped sand traps measuring sand ﬂux in the ﬁeld. For comparison two saltation impact responders can also be seen at the base of the sonic anemometers towards the left of the photograph (photo: author).

similar basis of responding to the individual impacts of ‘snake’ across the surface in an ever-changing weaving saltating grains on a sensitive surface. In the case of the pattern (Figure 18.24). This poses significan measure- Saltiphone, this surface is a highly responsive microphone ment problems, for within a continuously saltating envi- (Arens, 1996; Schonfeldt¨ and von Lowis,¨ 2003; Sterk, Ja- ronment sand traps and saltation impact responders are cobs and van Boxel, 1998; Zobeck et al., 2003), while in sometimes in contact with these streamers but, equally, the case of the Sensit and Safire the surface is a piezo- electric transducer (Gillette and Stockton, 1989; Stout and Contour map Spanwise transect Zobeck, 1997; Baas, 2004). The advantage of such instru-

ments over traditional sand traps is that they are small, streamer family offer little resistance to the f ow and are omnidirectional in response. They also allow high-frequency and instantaneous assessment of saltation activity, something that can only be achieved with a sand trap if a dynamic load cell is incorporated into the design, a technique that is generally restricted to wind tunnel applications (although nested streamers see McKenna-Neuman, Lancaster and Nickling, 2000). However, while such sensors have been shown to be very capable of identifying the onset of saltation and intermittency in saltation dynamics, they are less reliable for quantifying the actual saltation flu (van Pelt, Peters and Visser, 2009). clouds with embedded streamers A considerable difficult in characterising saltation activity using sand flu collectors or impact responders in array (y-axis) the fiel arises from the very nature of saltation. Baas and transport intensity Sherman (2005) and Baas (2008) recognise that saltation time (x-axis) array (y-axis) does not occur as a homogeneous ‘curtain’ across an erod- Figure 18.24 Characteristics of observed sand streamer pat- ing surface; rather it most often occurs in the form of tem- terns in plan view (left) and cross-section (right) (from Baas porally and spatially discontinuous sand streamers that and Sherman, 2005). P1: TIX/XYZ P2: ABC c18 JWST032-Thomas December 15, 2010 18:2 Printer Name: Yet to Come

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often they are not. The spatiotemporal dynamics of these mobile sand dunes (Castro and Wiggs, 1994; Wiggs, Liv- sand streamers in differing wind conditions is not known ingstone and Warren, 1996; Walker and Nickling, 2002, and so the impact they have on measurements of salta- 2003; Weaver and Wiggs, 2010) and sand streamers (Baas, tion activity and flu provided by sand traps and impact 2008). responders is also uncertain. The controls on the forma- The practical difficultie inherent in measuring tion and development of sand streamers is unclear, but it high-frequency turbulence in sand-laden airfl ws has seems likely that they are governed by near-surface tur- resulted in a moderate rate of advance in research. In bulent structures in the wind (Baas, 2008). his wind tunnel experiments, Butterfiel (1991, 1993, 1999) investigated the impact of temporally varying winds (of the order of seconds) on saltation behaviour. He discovered that mass flu and aerodynamic roughness 18.9 The role of turbulence in aeolian responded within about 1 second to an acceleration in sediment transport fl w, but that the response was several seconds longer in a decelerating f ow due to the momentum of the grains. As discussed above, there is an increasing understanding These finding led Butterfiel (1993) to conclude that that sediment entrainment and transport can be highly in- with naturally fluctuatin winds over a sand bed, the termittent under many environmental conditions. While grain-laden boundary layer is in constant adjustment variability in surface conditions can account for some ir- and may rarely achieve an equilibrium state. With the regularity in entrainment, it is now clear that, in a similar increasing application of sonic anemometers in the manner to river f ows (Bennett and Best, 1996; Kostaschuk fiel (Walker, 2005; van Boxel, Sterk and Arens, 2004; and Villard, 1996; Venditti and Bauer, 2005), turbulence Weaver, 2008) and the use of high-frequency grain impact in the boundary layer is also a major driving force behind sensors (Schonfeldt¨ and von Lowis,¨ 2003; Baas, 2004), aeolian sediment entrainment and transport (Livingstone, investigations have now begun to observe the relationship Wiggs and Weaver, 2007; Weaver, 2008; Sterk, Jacobs between saltation dynamics and instantaneous turbulent and van Boxel, 1998; Schonfeldt¨ and von Lowis,¨ 2003). peaks in wind velocities at much higher frequencies (1 to Indeed, it has also been shown that the influenc of peak > 10 Hz). While the relationship between turbulent wind instantaneous turbulent stresses on sand transport dynam- fl w and sand transport is complex at such frequencies ics is required to explain the observed development of (Figure 18.25) the association between the parameters

1600 12

1400 10 1200 8 1000

800 6

600

4 Wind velocity (m/s) 400 Sand transport (pulse count) 2 200

0 0 0 500 1000 1500 2000 2500 3000 3500 4000 4500 5000 5500 6000 6500 Time (10 Hz intervals)

Sand transport Wind velocity

Figure 18.25 Time series of horizontal wind velocity (using a sonic anemometer) and saltation impacts (on a grain impact sensor) measured at 10 Hz. The association is complex but the turbulent frequencies in both series question the application of ‘mean’ measures of ﬂow for the prediction of sediment transport (from Weaver 2008). P1: TIX/XYZ P2: ABC c18 JWST032-Thomas December 15, 2010 18:2 Printer Name: Yet to Come

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has led many to question the effica y of ‘mean’ measures significanc of the vertical fl w component in air (in com- of erosivity, such as mean wind velocity or shear velocity parison to water) is likely to be due to its lower density. (u∗) to predict sediment flu es (Stout, 1998; Baas and Much further research is required into the relation- Sherman, 2005; Weaver, 2008). ships between turbulence structures, sediment transport Research is now focused towards incorporating in- mechanics and evolving bed topographies. In the aeolian stantaneous fl w velocities into assessments of shear case, it seems likely that initial future research directions stress and turbulence structure. By using two- or three- will focus on developing sediment transport equations that dimensional sonic anemometers, the instantaneous veloc- incorporate some aspect of the instantaneous horizontal ity deviations from mean values can be assessed in both wind speed or establishing probabilistic analyses of poten- the horizontal (u component) and vertical (w component). tial sediment flu given wind turbulence characteristics. These data can be used to determine Reynolds shear stress (rather than mean u∗) to assess the magnitude of turbulent shear stresses or in quadrant analysis to assess the exis- 18.10 Conclusions tence of structure in the turbulence signal (Lu and Will- marth, 1973). Quadrant analysis divides the f ow into four Continued investigations over the last decade of aeolian discrete turbulent events depending on the relative signs of sediment mobilisation using field ork, wind tunnels instantaneous velocity deviations away from mean values and mathematical modelling have significantl enhanced in both the horizontal and vertical fl w components (Fig- our understanding. Huge progress has particularly been ure 18.26). Sweeps and ejections contribute to a ‘bursting made regarding the impacts of vegetation on erosivity process’ in the f ow where low-speed ejections of flui and erodibility, intermittency in entrainment thresholds away from the bed are followed by high-speed sweeps of and the role of turbulence in sediment transport rate de- fl w towards the bed, and this process has been found to termination. Much of this work has necessarily involved be important in accounting for sediment transport in rivers a small-scale approach and the continuing challenge is (Best, 1993). However, fiel investigations by Sterk, Ja- to apply our new knowledge at this scale to research cobs and van Boxel (1998), Schonfeldt¨ and von Lowis¨ questions at landscape and global scales. With growing (2003), van Boxel, Sterk and Arens (2004), Leenders, van concerns about the potential impact of global warming Boxel and Sterk (2005) and Weaver (2008) have found that on dryland landscapes, knowledge of aeolian processes, the majority of instances of high aeolian saltation flu are through which much landscape change will occur, will rather associated with sweeps and outward interactions, be of increasing importance. i.e. f ow events that have instantaneous horizontal velocities greater than the mean (Figure 18.26). The reduced References

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(2009) Responses of and dune mobility in the southwest Kalahari. Land Degra- dune activity and desertificatio in China to global warming dation and Rehabilitation, 5, 13–24. in the twenty-firs century. Global and Planetary Change, 67 Wiggs, G.F.S., Thomas, D.S.G., Bullard, J.E. and Livingstone, I. (3–4), 167–185. (1995) Dune mobility and vegetation cover in the southwest Wasson, R.J. and Nanninga, P.M. (1986) Estimating wind trans- Kalahari Desert. Earth Surface Processes and Landforms, port of sand on vegetated surfaces. Earth Surface Processes 20, (6), 515–529. and Landforms, 11, 505–514. Wiggs, G.F.S., Livingstone, I., Thomas, D.S.G. and Bullard, J. E. Weaver, C.M. (2008) Turbulent f ow and sand dune dynamics: (1996) Airfl w and roughness characteristics over partially- identifying controls on aeolian sediment transport, Unpub- vegetated linear dunes in the southwest Kalahari Desert. lished DPhil Thesis, University of Oxford, 228 pp. Earth Surface Processes and Landforms, 21, 19–34. Weaver, C.M. and Wiggs, G.F.S. (2010) Field measurements Wilkinson, R.H. (1983/1984) A method for evaluating statisti- of mean and turbulent airfl w over a barchan sand dune. cal errors associated with logarithmic velocity profiles Geo- Geomorphology (in press). Marine Letters, 3, 49–52. Webb, N.P., McGowan, H.A., Phinn, S.R. et al. (2009) A model Willetts, B.B. (1983) Transport by wind of granular materials to predict land susceptibility to wind erosion in western of different grain shapes and densities. Sedimentology, 30, Queensland, Australia. Environmental Modelling and Soft- 669–679. ware, 24 (2), 214–227. 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19 Desert dune processes and dynamics

Nick Lancaster

19.1 Introduction range of height, width and spacing. Aerial photographs and satellite images show the general regularity of dune Desert sand dunes form part of a self-organised hierar- patterns (manifested by close correlations between dune chical system of aeolian bedforms, which comprises: (a) height, width and spacing) and also indicate that essen- wind ripples, (b) individual simple dunes or superimposed tially similar dune forms occur in widely separated local- dunes on mega dunes and (c) mega dunes (also called ities. Even dunes on other planetry bodies (e.g. Mars, draa or compound and complex dunes) – characterised by Titan) have a simliar morphology to terrestrial dunes. superimposition of simple or elemental dunes on larger These commonalities indicate that there are general phys- forms. The majority of dunes are composed of quartz and ical principles that govern dune form and dynamics. feldspar grains of sand size, although dunes composed of Dunes are created and modifie by interactions be- gypsum, carbonate and volcanic sand also occur. Most tween sand transport rates, dune topography and airfl w sand dunes occur in contiguous areas of dunes known as (Figure 19.2). As the dune projects into the atmospheric sand seas or ergs (with an area of > 100 km2). Smaller boundary layer, the primary air fl w is modifie by in- areas of dunes are called dune field (see Chapter 17). teractions between the dune form and the airfl w, which Dunes occur in self-organised patterns that develop over give rise to modification of the local wind speed, shear time as the response of sand surfaces to the wind regime stress and turbulence intensity, and therefore sand trans- (especially its directional variability) and the supply of port rates. Such interactions also create secondary fl w sand (Werner, 1995). The dune types discussed below circulations, especially in the lee of the dune, as a result of represent the steady state attractors of the aeolian sand fl w separation and fl w diversion. In addition, superim- transport system and can evolve from a wide range of posed dunes on megadunes or draa respond to changes in initial conditions. Key variables that determine dune mor- airfl w and sediment transport on the megadune itself. The phology and dynamics are the spatial and temporal charac- nature of these interactions is discussed in detail below. teristics of: (1) the wind regime (especially its directional variability), (2) sand supply (the amount of sand available for dune building) and (3) vegetation cover. Sand particle 19.3 Dune types and environments size does not appear to be as important as once thought. As dunes and dune patterns evolve over time, the legacy 19.3.1 Crescentic dunes of past climates and wind regimes influence present-day dune morphology and dune patterns in many areas. In wind regimes characterised by a narrow range of wind directions (unidirectional wind regimes), crescentic dunes form with their crests aligned normal to the direction of 19.2 Desert dune morphology sand transport. Where sand supply is low relative to the capacity of Desert dunes occur in a variety of morphologic types (see the wind to transport sand (undersaturated sand transport Box 19.1 and Figure 19.1), each of which displays a conditions), isolated crescentic dunes or barchans form

Arid Zone Geomorphology: Process, Form and Change in Drylands, Third Edition. Edited by David S. G. Thomas © 2011 John Wiley & Sons, Ltd. Published 2011 by John Wiley & Sons, Ltd. ISBN: 978-0-470-51908-0 P1: TIX/XYZ P2: ABC c19 JWST032-Thomas December 15, 2010 18:3 Printer Name: Yet to Come

488 CH19 DESERT DUNE PROCESSES AND DYNAMICS

Box 19.1 Dune classiﬁcations Dunes Mega-dunes

Relative Sediment Thickness Many different classification of types of dunes have been proposed: (1) those based on the external morphol- CRESCENTIC ogy of the dunes (morphological classifications and Barchans Crescentic Compound (2) those that imply some relationship of dune type to ridges crescentic Dome dunes dunes formative winds or sediment supply (dynamic classifi cations). Morphologic classification of dunes facilitate LINEAR mapping of dune types from satellite image data (e.g. Simple Compound Complex - Vegetated Linear dunes Linear dunes Breed et al., 1979) with no assumptions regarding for- - Sinuous mative mechanisms and relations to external variables. Dynamic classification of dunes imply some causal re- REVERSING

lationship between external variables and dune type – so Wind regime complexity STAR that dunes may be identifie as transverse, longitudinal or oblique to the formative winds (Hunter, Richmond and Alpha, 1983). Even though understanding of the coarse SAND conditions in which different types of dunes form has sand SHEETS ZIBARS improved significantl in recent years, this knowledge remains imperfect, so that a largely morphological classificatio scheme is preferred. Figure 19.1 provides a vegetation NEBKHAS BLOW PARABOLIC anchored OUT DUNES framework for classifying dunes based on external mor- DUNES phology, sediment volume and other key parameters. source LUNETTES bordering

topographic ECHO FALLING CLIMBING influence DUNES DUNES DUNES

Figure 19.1 A scheme for the classiﬁcation of desert dunes (modiﬁed from Lancaster, 1995).

(Figure 19.3). Barchans have been described from most seas and dune field as well as in the transport corridors desert regions and have been extensively studied in the extending from sand source areas (Figure 19.3(b)), where field using aerial and satellite images, and most recently a clear size selection process involves merging smaller, via numerical models. Conditions favouring the forma- faster moving dunes with slower moving, larger dunes tion of barchans are encountered on the margins of sand so that the pattern becomes more uniform. In plan view,

Dune morphology

erosion and deposition patterns wind regime form - flow slope effects interactions

Air flow Sand transport rates

momentum exchange

Figure 19.2 A conceptual framework for process–form interactions on aeolian dunes. P1: TIX/XYZ P2: ABC c19 JWST032-Thomas December 15, 2010 18:3 Printer Name: Yet to Come

DUNE TYPES AND ENVIRONMENTS 489

barchans are elliptical in shape with a concave lee face and (Elbelrhiti, Andreotti and Claudin, 2008) and at several horns extending downwind (Figure 19.3). Dune width is localities along the coast of the Namib Desert (Bourke typically about ten times dune height. Good examples of and Goudie, 2009). Most barchans occur at the simple barchan dunes are to be found in the Pampa La Joya, Peru dune scale, but megabarchans are identifie from a few (Finkel, 1959), the Tarfaya–Laayoune area of Morocco areas (e.g. Grolier et al., 1974).

(a) (b)

(e) (f)

Figure 19.3 Crescentic dunes: (a) barchans in transport corridor (Laayoune, Morocco); (b) asymmetric barchan in Skeleton Coast dune ﬁeld, Namibia; (c, d) simple crescentic dunes, White Sands, New Mexico; (e, f) compound crescentic dunes, Algodones Dune Field, California; (g) sand rose for an area of crescentic dunes; (h) anatomy of crescentic dune morphology. P1: TIX/XYZ P2: ABC c19 JWST032-Thomas December 15, 2010 18:3 Printer Name: Yet to Come

490 CH19 DESERT DUNE PROCESSES AND DYNAMICS

Horn (h) Walvis Bay (g)

Avalanche face

Stoss Lee

Brink Crest

Avalanche face RDP/DP = 0.73 To e

Figure 19.3 (Continued)

Where sand supply increases, crescentic ridges or trans- compared to 0.09 m/yr to 0.35–0.40 m/yr for the main verse dunes occur (Figure 19.3). Worldwide, such dunes dunes (Sharp, 1979). occupy about 10 % of the area of sand seas. Good examples of simple crescentic ridges can be found at White Sands, New Mexico (Kocurek et al., 2007; McKee and 19.3.2 Linear dunes Douglass, 1971) and in the Skeleton Coast dune fiel of Namibia (Lancaster, 1982a). At White Sands, the dunes Dunes of linear form (Figure 19.4 and Figure 19.5) are are typically sinuous crescentic ridges and barchanoid widespread in deserts worldwide and dominate in Aus- dunes, 8–12 m high, with a crest length of about 250 m tralia and the Kalahari, as well as in Namibia, Arabia and an average spacing of 136 m. These dunes are migrat- and the southern and western Sahara. Linear dunes oc- ing towards the ENE at 1–7 m/yr. The Skeleton coast dune cur in simple, compound and complex varieties (the latter fiel is characterised by straight to slightly sinuous cres- two at megadune scales). There are two main varieties of centic dunes with a spacing between 150 and 300 m and the simple linear dune: (1) sinuous, relatively short dunes a height ranging between 4 and 30 m, migrating towards (also known as seif dunes) and (2) longer, straight, vege- the north or northeast. Dune size increases northwards tated linear dunes such as those that occur in the Simpson (downwind) and the eastern side of the dune fiel is char- Desert of Australia and the Kalahari. acterised by smaller crescentic dunes that are changed to Sinuous simple linear dunes (seif dunes) have been doc- barchans as the sand supply decreases. umented from many desert regions (Figure 19.4). They Compound or crescentic megadunes are characterised typically consist of a sinuous ridge with alternate peaks by the superimposition of smaller crescentic dunes on and saddles and a more or less triangular cross-section, the their stoss (windward) slopes (Figure 19.4(d) and (e)). sinuosity increasing with dune height. Such dunes are typ- These smaller dunes migrate across the main form at an ically a few hundred metres long. In many areas, it appears order of magnitude greater than that of the main dune. that dunes of this type are very young, comprising dunes Classic examples of crescentic megadunes can be found that have developed in modern wind regimes as a modifi in the Algodones dune fiel of southeast California (De- cation of older patterns, e.g. in the Azefal Sand Sea, Mau- rickson et al., 2008; Havholm and Kocurek, 1988), the ritania (Lancaster et al., 2002), the Wahiba sands (Warren Liwa area of the UAE (El-Sayed, 2000; Stokes and Bray, and Allison, 1998) or representing the extension of larger 2005) and along the coastal areas of the Namib Sand Sea linear dunes (Bristow, Bailey and Lancaster, 2000). (Lancaster, 1989c). In the Algodones, the main dunes are Vegetated linear dunes in the southwest Kalahari and 40–60 m high with a mean spacing of 1025 m, whereas the Simpson–Strzelecki dune field (Figure 19.4(c) and the superimposed coalesced crescentic dunes are 12–20 (d)) are very similar in form and provide good exam- m high with a mean spacing of 93 m (Derickson et al., ples of this dune type. These dunes typically comprise 2008). Superimposed dunes migrate at 0.76–1.73 m/yr, straight subparallel ridges 2–35 m high and 200–450 m P1: TIX/XYZ P2: ABC c19 JWST032-Thomas December 15, 2010 18:3 Printer Name: Yet to Come

DUNE TYPES AND ENVIRONMENTS 491

(a) (b)

(e)

Upington, South Africa

RDP/DP = 0.48

Figure 19.4 Simple linear dunes: (a) Azefal Sand Sea, Mauritania; (b) Northern Namib sand Sea; (c,d) vegetated linear dunes, southwest Kalahari. P1: TIX/XYZ P2: ABC c19 JWST032-Thomas December 15, 2010 18:3 Printer Name: Yet to Come

492 CH19 DESERT DUNE PROCESSES AND DYNAMICS

(a) (b)

(e) (f)

Boutilimit, Mauritania

Figure 19.5 Compound and complex linear dunes: (a) compound linear dunes, Erg Fachi Bilma, Niger; (b) southern Namib Sand Sea; (c, d) complex linear dunes, Namib Sand Sea; (e) wind regimes of compound and complex linear dunes; (f) anatomy of complex linear dunes. P1: TIX/XYZ P2: ABC c19 JWST032-Thomas December 15, 2010 18:3 Printer Name: Yet to Come

DUNE TYPES AND ENVIRONMENTS 493

(a) (b)

Arm

Plinth

Eastern Namib Sand Sea RDP/DP = 0.25

Figure 19.6 Star dunes: (a) clusters of star dunes, eastern Rub al Khali, Saudi Arabia; (b) Silver Peak, Nevada; (c) wind regime of star dunes; (d) anatomy of star dunes.

apart (Bullard et al., 1995; Fitzsimmons, 2007), with a areas (e.g. the eastern Namib Sand Sea and southeastern length of 20–25 km. The ridges may join in Y-junctions. parts of the UAE). Several varieties of this pattern can be identifie (Figure The origins of linear dunes and their relationship to 19.6), related to sand supply, distance from sand sources, formative wind directions have been the subject of con- vegetation cover and persistence of dune-forming winds siderable debate (Lancaster, 1982b; Livingstone, 1988; (Bullard et al., 1995; Wasson et al., 1988). Tsoar, 1989). A substantial body of empirical evidence Compound linear dunes (Figure 19.5(a) and (b)) con- now indicates that linear dunes form in bidirectional ◦ sist of several narrow, closely spaced (100–200 m) seif- wind regimes with the two modes separated by 90 or like ridges on the crest of broad linear plinths, which are more. This includes correlations between the occurrence spaced as much as 2.5 km apart. Good examples occur in of linear dune and information on the directional vari- the southern Namib Sand Sea (Lancaster, 1989c) and the ability of wind regimes in these areas (e.g. Fryberger, Fachi Bilma Erg in Niger (Mainguet and Callot, 1978). 1979): studies of internal sedimentary structures (Bris- Complex linear dunes (Figure 19.5(c) and (d)) are com- tow, Bailey and Lancaster, 2000; McKee, 1982; McKee prised of relatively straight, continuous, linear ridges on and Tibbitts, 1964), experiments in flume (Reffet et al., which other dune types are superimposed. Such dunes 2008; Rubin and Ikeda, 1990) and numerical simulations are typically spaced 1–2.5 km apart and reach heights of (Werner and Kocurek, 1997), as well as detailed process 50–250 m. Complex linear dunes in the Namib Sand Sea studies on linear dunes (Livingstone, 1986, 1988, 1993; and Saudi Arabia consist of a sinuous main crestline with Tsoar, 1983). Models for linear dune formation that invoke superimposed crescentic dunes on the dune flank (Lan- boundary layer roller vortices in which helicoidal fl w caster, 1989c; Livingstone, 1993). In places, the crest line sweeps sand from interdune areas to dunes (Hanna, 1969) may consist of a series of peaks similar to star dunes sep- are not generally supported by empirical data – evidence arated by lower sinuous sharp crest lines. Such complex for the occurrence of such boundary layer structures is linear dunes are transitional to chains of star dunes in some limited. P1: TIX/XYZ P2: ABC c19 JWST032-Thomas December 15, 2010 18:3 Printer Name: Yet to Come

494 CH19 DESERT DUNE PROCESSES AND DYNAMICS

Linear dunes primarily tend to extend downwind, but are characterised by a U or hairpin plan morphology, may also migrate laterally, because sand transport by the with an active ‘nose’ or dune front and trailing partly two modes of a bimodal wind regime are rarely equal to completely vegetated arms that extend upwind. Be- (Rubin, Tsoar and Blumberg, 2008). Evidence for lateral cause a sparse vegetation cover is important to the for- migration has been documented from several areas, in- mation and maintenance of parabolic dunes, they have a cluding northwest China, the Strzelecki Desert, Sinai and restricted distribution in truly arid regions, but are com- the Namib Sand Sea (see Tsoar, Blumberg and Stoler, mon in many semi-arid dune areas, including the Great 2004, and references therein). Rates of lateral variation Plains of the USA (Sun and Muhs, 2007), the Thar Desert vary considerably, ranging from 0.13 m/yr over the past of India (Kar, 1993) and White Sands, New Mexico (Ko- 2400 years in Namibia (Bristow, Duller and Lancaster, curek et al., 2007). 2007) to ∼0.50 m/yr in the Sinai and up to 3 m/year in China (Rubin, Tsoar and Blumberg, 2008). Some linear dunes do apparently form parallel to a 19.3.5 Zibars and sand sheets dominant wind direction. This occurs when the sand is partially stabilised by vegetation or anchored by topo- Sand sheets are common in the upwind areas of many sand graphic obstacles (Hesp, 1981; Tsoar, 1989). Tsoar (1989) seas and dune fields They occur in environments that are suggests that all vegetated linear dunes form parallel unfavourable for the formation of dunes, often in condi- to the dominant wind direction. Recent studies suggest tions in which bypassing or net erosion of sand occurs. that partial induration of sediment may produce simi- Sand sheets may develop in areas of coarse sands, such as lar longitudinally oriented linear dunes (Rubin and Hesp, in the eastern Sahara and southern Namib Sand Sea, the 2009). presence of a high water table and/or periodic flooding as in parts of the White Sands dunefield or a moderate vegetation cover, as in the northern Gran Desierto of Mex- 19.3.3 Star dunes ico (Kocurek and Nielson, 1986; Lancaster, 1992). Many sand sheets as well as interdune areas between linear and Star dunes (Figure 19.6) are the largest and often the most star dunes are organised into low rolling dunes, without complex type of desert sand dunes, with a pyramidal form slipfaces, known as zibars (Holm, 1960; Nielson and Ko- and radiating arms (Lancaster, 1989a). Star dunes rarely curek, 1986; Warren, 1972), with a spacing of 50–400 m occur in isolation and several patterns can be recognised, and a maximum relief of 10 m (see Figure 19.5(a)). including chains and clusters of star dunes. Star dunes occur in the northern Sahara (Mainguet and Chemin, 1984), the northeastern part of the Rub al Khali, the Badain Jaran 19.4 Airﬂow over dunes desert of China (Dong, Wang and Wang, 2004) and in parts of the Namib Sand Sea, and may reach a height of The airfl w around and over dunes is strongly influence 300 m. They occur in areas of complex wind regimes, by the topography of the dune, which projects into the with strong seasonal changes in wind direction, involving atmospheric boundary layer. This results in compression ◦ winds from at least two directions separated by 180 .Al- of streamlines on the upwind side of the dune, leading to though the mode of formation of star dunes is not well acceleration of the wind and an increase in wind shear understood, evidence from several studies indicates that stress towards the dune crest. In the lee of the dune, the modificatio of a pre-existing form (linear or crescen- fl w expands and/or separates from the surface, resulting tic) is important (Nielson and Kocurek, 1987). Thus star in fl w deceleration and secondary f ows; at a point some dunes may form as crescentic dunes migrate into areas of distance downwind of the dune, the f ow reattaches to seasonally reversing wind directions (Lancaster, 1989b), the surface and gradually recovers to its upwind profil or by reworking of large linear or crescentic dunes as a characteristics, assuming that there is no additional dune result of climatic and wind regime changes (Beveridge present. et al., 2004). Boundary layer theory provides a conceptual framework for understanding air f ow over dunes (Nickling and McKenna Neuman, 1999). The fl w can be divided into 19.3.4 Parabolic dunes an outer inviscid region and an inner region, which follows the topography of the dune (Figure 19.8). The inner Parabolic dunes (Figure 19.7) form in areas that are par- layer is further divided into two sublayers: (1) a very thin tially vegetated and experience unidirectional winds. They inner surface layer, in which the shear stress is constant, P1: TIX/XYZ P2: ABC c19 JWST032-Thomas December 15, 2010 18:3 Printer Name: Yet to Come

AIRFLOW OVER DUNES 495

(a) (b)

Figure 19.7 Parabolic dunes: (a, b) White Sands, New Mexico.

with a thickness equivalent to the surface roughness (1/30 19.4.1 The stoss or windward slope of the grain diameter for a sand surface), and (2) a shear stress layer less than 1 m thick in which shear stress effects Winds approaching the upwind toe of a dune stagnate decrease with height above the surface. The shear stress slightly and are reduced in velocity (Livingstone, Wiggs developed in this part of the boundary layer generates sand and Weaver, 2007; Wiggs, 1993). On the stoss, or wind- transport on dunes. The basic features of this conceptual ward slope of the dune, streamlines are compressed and model have been confirme by fiel studies of air f ow winds accelerate up the slope. The magnitude of the ve- over dunes (Livingstone, Wiggs and Weaver, 2007) and locity increase varies with dune height and dune steep- reproduced in numerical models (Parsons, Walker, and ness or aspect ratio (Figure 19.9) and is characterised by Wiggs, 2004; Weng et al., 1991). the speed-up ratio s or amplificatio factor Az (U2/U1,

Outerlayer

Planetaryboundary OuterRegion layer } Middlelayer } } } Innerregion

Innersurfacelayer Shear stress layer

Figure 19.8 Airﬂow over an isolated two-dimensional hill, showing components of the boudary layer (after Nickling and McKenna Neuman, 1999). P1: TIX/XYZ P2: ABC c19 JWST032-Thomas December 15, 2010 18:3 Printer Name: Yet to Come

496 CH19 DESERT DUNE PROCESSES AND DYNAMICS , Flow reattachment Flow and recovery Flow separation and diversion Flow Deposition Lee slope Brink Deposition 0.50 Crest 0.40 Flow expansion and deceleration Flow curvature streamline Convex sand transportDecreased rates 0.30 Aspect ratio 0.20 Wind speed up on stoss slope speed up on stoss Wind 0.10 Linear Dunes Crescentic Dunes Star Dunes Stoss or windward slope Net erosion 0.00

2.5 2.0 1.5 1.0 Amplification Factor (Az) Factor Amplification Flow acceleration Flow curvature streamline Convex Increasing wind shear velocity in sand transportExponential increase rate Toe Net erosion Airﬂow over dunes showing streamline convergence and divergence (after Livingstone, Wiggs and Weaver, 2007) and velocity speedup (after Lancaster Flow stagnation stagnation Flow and deceleration streamline Concave curvature 1995). Figure 19.9 P1: TIX/XYZ P2: ABC c19 JWST032-Thomas December 15, 2010 18:3 Printer Name: Yet to Come

DUNE DYNAMICS 497

where U2 is the velocity at the dune crest and U1 is the ve- When fl w is oblique to the dune crest a helical vor- locity at the same height above the surface at the upwind tex develops in the lee (Walker and Nickling, 2002). The base of the dune). Wind tunnel simulations and numerical oblique fl w is deflecte along the lee slope parallel to the models show that shear stress increases on dune slopes in dune crest, with the degree of deflectio being inversely a manner similar to wind speed, but fiel measurements proportional to the incidence angle between the crestline of wind shear velocity and shear stress on dunes are dif- and the primary wind. When the angle between the dune ◦ ficul to achieve where wind profile are not log-linear as and the wind is less than 40 the velocity of the deflecte a result of f ow acceleration and development of internal wind is greater than that at the crest and sand is transported boundary layers (Frank and Kocurek, 1996a). In addition along the lee side of the dune (Tsoar, 1983). This process to the effects of velocity amplification curvature of wind is especially important on linear dunes, where it leads to ◦ streamlines may play an important role in the pattern of dune extension. When winds are at more than 40 to the wind shear stress on dunes (Figure 19.9). Concave upward crestline the velocity of the deflecte wind is reduced, curvature of streamlines at the toe of the dune enhances giving rise to lee-side deposition. Changes in the local shear stress, whereas convex streamline curvature on the incidence angle between primary winds and a sinuous mid stoss slope and between the crest and the brink acts dune crest result in a spatially varying pattern of deposi- to decrease shear stress (Wiggs, Livingstone and Warren, tion and along-dune transport on the lee face. Deposition 1996). dominates where winds cross the crest line at angles ap- ◦ proaching 90 and erosion or along-dune transport occurs ◦ where incidence angles are <40 (Figure 19.10). 19.4.2 Lee-side ﬂow 19.5 Dune dynamics In the lee of the crest of dunes, wind velocities and transport rates decrease rapidly as a result of fl w expansion between the crest and brink of the lee or avalanche face Dunes may (1) migrate downwind or laterally as a result and fl w separation on the avalanche face itself. There is of erosion of windward slopes and deposition in the lee, a complex pattern of fl w separation, diversion and reat- (2) extend downwind or (3) accrete vertically (Thomas, tachment on the lee slopes of dunes, which is determined 1992). Different dune morphologic types exhibit differ- by the angle between the wind and the dune crest (angle ent degrees of these three processes. Thus, the arms of of attack), atmospheric stability and the dune aspect ratio barchans may extend as small linear features (Tsoar, (Walker and Nickling, 2002; Dong et al., 2009; Sweet and 1984), linear dunes may migrate laterally as well as ex- Kocurek, 1990). Secondary f ows, including lee-side fl w tending down wind, and also accrete vertically (Bristow, diversion, are especially important where winds approach Duller and Lancaster, 2007); star dunes are dominated the dune obliquely, and are an important component of air by vertical acretion, but their arms may migrate laterally fl w on linear and many star dunes (Figure 19.10). and/or extend (Lancaster, 1989b). Conceptual models for airfl w in the lee of fl w- transverse dunes have been developed based on fiel and 19.5.1 Erosion and deposition patterns on dunes wind tunnel studies (Baddock, Livingstone and Wiggs, 2007; Frank and Kocurek, 1996b; Walker, 1999; Walker All dunes exhibit a high degree of interaction between the and Nickling, 2002). These models identify a series of shape of the dune, the amount of change in wind velocities distinct regions of fl w that vary in wind speed, shear and and sand transport rates, and the rate and pattern of erosion turbulence intensity (Figure 19.11(a)), including a sepa- or deposition. These interactions are manifested in several ration cell extending for 4–10 dune heights downwind, ways. two wake regions that merge at 8–10 h and an internal The windward slopes of dunes are characterised by boundary layer that grows downwind of the point at which an exponential increase in sediment transport rates to- the separated fl w reattaches to the surface. According to wards the dune crest (Figure 19.12), as a result of fl w Walker and Nickling (2002), the internal boundary layer acceleration, coupled with effects of stream line curva- is identifiabl at 8–10 h and comes into equilibrium at ture (Lancaster et al., 1996; McKenna Neuman, Lancaster around 25–30 h, downwind of isolated dunes. For mul- and Nickling, 1997, 2000). Although net erosion tends tiple transverse dunes, Baddock, Livingstone and Wiggs to characterise the windward slopes of dunes, the pat- (2007) identify four main zones (Figure 19.11(b)): sep- tern of erosion and deposition varies with overall wind aration, reattachment, recovery and interaction with the speed, so that areas near dune crests may experience ero- next dune downwind. sion at lower wind speeds and deposition at wind speeds P1: TIX/XYZ P2: ABC c19 JWST032-Thomas December 15, 2010 18:3 Printer Name: Yet to Come

498 CH19 DESERT DUNE PROCESSES AND DYNAMICS

Primary flow (a)

Grainflow Flow separation zone

Secondary flow

(b)

SSW - SW winds Easterly winds Crest line Local flow direction

Flow separation Superimposed dune

Figure 19.10 Flow separation and diversion in the lee of a linear dune: (a) field view of flow in the lee of a simple linear dune; (b) schematic of flow over the same dune in different wind regimes (after Bristow, Bailey and Lancaster, 2000). P1: TIX/XYZ P2: ABC c19 JWST032-Thomas December 15, 2010 18:3 Printer Name: Yet to Come

DUNE DYNAMICS 499

significantl above threshold (McKenna Neuman, Lan- a grain fl w or avalanche. The frequency of grain fl ws caster and Nickling, 2000). is dependent on the wind speed and sand transport rate Separation of airfl w at the brink of the lee face leads for the period preceding their initiation. The interval be- to grain fall from the overshoot of saltating grains trans- tween grain f ows decreases exponentially with incident ported up the windward slope. The rate of deposition by wind speed from about 60 minutes at 5 m/s to 8 minutes grainfall decreases exponentially with distance from the at 8 m/s; the magnitude of the f ows is inversely propor- brink of the lee face (Figure 19.13), so that most of the tional to the interval between fl ws (Breton, Lancaster sediment is deposited within 1 m of the brink (Nickling, and Nickling, 2008). McKenna Neuman and Lancaster, 2002). Grain fl ws oc- Seasonal patterns of erosion and deposition measured cur when grainfall deposits build up so that the lee slope on linear and star dunes show that the greatest amount of is steepened above the angle of repose and fails, initiating change occurs when winds change direction seasonally

air flow

Wake region

Turbulent shear zone

I Internal boundary layer

0 5 10 x/h

Separation cell Turbulent shear Turbulent stress (a) layer maximum

Intermittent (b) Erosivity erosivity (a) (b)

Separation

Reattachment Interaction zone (a) no stagnation at dune toe; (b) stagnation at dune toe

Figure 19.11 Conceptual model for ﬂow in the lee of a transverse dune: (a) isolated dune (after Walker and Nickling, 2002); (b) (after Baddock, Livingstone and Wiggs, 2007). P1: TIX/XYZ P2: ABC c19 JWST032-Thomas December 15, 2010 18:3 Printer Name: Yet to Come

500 CH19 DESERT DUNE PROCESSES AND DYNAMICS

wards a convex form similar to that of transverse dunes (Tsoar et al., 1985), and erosion and deposition rates near the crest decline in time as the dune comes into equilibrium with a new wind direction. Erosion and deposition patterns on linear, reversing and star dunes show that they consist of a crestal area where erosion and deposition rates are high and a plinth zone in which there is little surface change but considerable throughput of sand (Figure 19.14). The crest lines of Namibian complex linear dunes migrate over a lateral distance of as much as 14 m over a 12-month period but show little net change over decadal timescales (Living- stone, 1989, 1993, 2003; Sharp, 1966; Wiggs et al., 1995).

19.5.2 Long-term dune dynamics

Insights into dune dynamics over annual to decadal Figure 19.12 Windward slope erosion and deposition pat- timescales can be gained from a variety of sources. A terns (data from Lancaster et al., 1996, and McKenna Neuman, large data set on migration rates of barchans and crescen- Lancaster and Nickling, 1997). tic dunes has been developed by comparing the position of dunes on time series of aerial photographs or satellite images, demonstrating that there is an inverse relationship (Lancaster, 1989b, 1989c). This is the result of winds en- between barchan dune height and migration rate (Figure countering a dune form that is out of equilibrium with a 19.15). This relationship is important to the development new wind direction. Field observations and models show of self-organised dune patterns in that more rapidly mov- that the crestal profile of linear and star dunes tend to- ing smaller dunes tend to catch up and merge with larger

1.00

6 m/sec (lab) 5.8 m/sec (field) 0.80 8 m/sec (lab) 8.1 m/sec (field)

0.60

q/q max 0.40

0.20

0.00 0.00 0.20 0.40 0.60 0.80 1.00 x/L

Figure 19.13 Patterns of deposition in the lee of a transverse dune (ﬁeld data after Nickling, McKenna Neuman and Lancaster, 2002; laboratory wind tunnel data after Cupp et al., 2005). P1: TIX/XYZ P2: ABC c19 JWST032-Thomas December 15, 2010 18:3 Printer Name: Yet to Come

DUNE DYNAMICS 501

W E

20 Meters 20

Winter (E - NE wind) 3000 Deposition

2000

1000

0 (mm)

-1000

-2000 Summer Winter -3000

Erosion -4000

Summer SSW - SW wind

Figure 19.14 Patterns of erosion and deposition cross a Namib linear dune (after Livingstone, 1989).

60 Pampa La Joya Mauritania Kharga 50 Walvis Bay (1961-1975) Walvis Bay (1976-1988) Salton Sea (1941-1956) Salton Sea (1956-193) Salton Sea (1963-1981) 40

Dune celerity (m/yr) 20

0 0 5 10 15 20 Dune Height (m)

Figure 19.15 Relations between barchan migration rates and dune height (data from Finkel, 1959, Slattery, 1990, Ould Ahmedou et al., 2007, Embabi, 1982, and Haff, 1984). P1: TIX/XYZ P2: ABC c19 JWST032-Thomas December 15, 2010 18:3 Printer Name: Yet to Come

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dunes, so promoting an overall increase in dune size and Following Tsoar (1989), vegetated linear dunes are spacing downwind (pattern coarsening). nucleated by vegetation, and extend downwind from At best, these studies span the last few decades. A nebkhas. Vegetated linear dunes can also develop from longer-term (centuries to millennia) perspective on dune source bordering dunes; closely spaced dunes link by dynamics can now be gained through the application of Y-junctions to form fewer, larger dunes downwind high-resolution optically stimulated luminescence (OSL) (Wasson, 1983). dating of dunes, especially when used in combination The formation of sinuous linear dunes (seif dunes) by with ground penetrating radar (GPR) studies of dune sed- elongation of one arm of a barchan as it migrates into a bi- imentary structures. Crescentic dunes superimposed on modal wind regime was f rst proposed by Bagnold (1941) the southern edge of a north–south-oriented linear dune and supported by fiel observations in the Namib (Lan- in the northern part of the Namib Sand Sea have migrated caster, 1980). Strong winds from an oblique direction add to the west at an average rate of 0.12 m/yr over the past sand to one horn and gentler winds from the original direc- 1570 years (Bristow, Duller and Lancaster, 2005), while a tion extend the horn (Figure 19.16(a)). A different model nearby 60 m-high linear dune shows evidence of episodic is suggested by observations in the Sinai (Tsoar, 1984), development and lateral migration at rates of up to 0.13 where winds from both the primary and secondary direc- m/yr to the east over the past 6000 years (Bristow, Duller tions extend the horn on the side of the dune opposite to and Lancaster, 2007) and 20-m high crescentic dunes in the secondary wind direction (Figure 19.15(b)). The linear theLiwaareaoftheUAEhavemigratedtothesouthover element of the dune extends more rapidly than the original the past 320 years at an average rate of 0.78 m/yr, with barchan can migrate, and the dune evolves to a linear form. a more rapid migration of 0.91 m/year between 220 and Linear dunes develop by extension downwind and by 110 years ago (Stokes and Bray, 2005). slow accretion of sand. Bristow, Bailey and Lancaster (2000) present a new model for development of sinuous linear dunes based on GPR visualisation of sedimentary structures. In this model (Figure 19.17), secondary f ows 19.6 Dune development and form–fl w interactions become more important as the dune increases in size, leading ultimately to the establish- The origins and growth of individual dunes and the de- ment of superimposed bedforms on the dune flanks velopment of dune patterns are intimately linked, because Star dunes appear to develop mainly by modificatio of dunes rarely occur in isolation. existing dunes (Lancaster, 1989b; Nielson and Kocurek, Dune initiation is a poorly understood process, but is 1987), either as they migrate or extend into areas ex- likely to involve localised reductions in sand transport periencing multidirectional winds (Figure 19.18) or as a rates by convergence of streamlines, changes in surface result of changes in wind regime as a result of climate roughness (e.g. vegetation cover), surface particle size or change, as in star dunes developed on linear dunes in the by variations in microtopography (slope changes, relict Gran Desierto (Beveridge et al., 2006) and the southeast bedforms), leading to deposition of sand, which forms a UAE. nucleation point for a protodune (Kocurek et al., 1992; Many megadunes possess superimposed dunes, which Lancaster, 1996). There is a minimum size for a proto- are in equilibrium with the current wind regime (e.g. on dune, determined by the saturation length or the distance Algodones and Namib compound crescentic dunes). The required for actual sand mass flu to reach the capacity slopes of these large dunes present an effectively planar of the wind to transport sand. The elementary dune wave- surface on which sand transport takes place. Therefore, length is ∼20 m on Earth and ∼600 m on Mars (Claudin variations in sand transport rates on megadunes in time or and Andreotti, 2006). Once initiated, growth of crescentic space will lead to the formation of superimposed dunes if dunes occurs by merging of smaller, faster-moving dunes the major dune is sufficientl large (Andreotti et al., 2009; with larger, slower-moving dunes, in a series of construc- Elbelrhiti, Claudin and Andreotti, 2005). tive interactions. In some areas, however, the larger primary (mostly The primary response of sand surfaces to bedform de- linear) dunes are clearly a product of past wind regime velopment is to form a asymmetric fl w-transverse dune and sediment supply conditions, as demonstrated by OSL with a convex stoss slope. Crescentic dunes dominate in dating of the major form (Lancaster, 2007). Such dunes unidirectional wind regimes; small dunes in multidirec- have considerable inertia (represented by the reconstitu- tional wind regimes are also of this form, because they tion time, or the time required for the dune to migrate its are small enough to be completely reworked in each wind length in the direction of net transport) and require sig- season. nifican periods of time to adjust to changed conditions of P1: TIX/XYZ P2: ABC c19 JWST032-Thomas December 15, 2010 18:3 Printer Name: Yet to Come

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(a)

B B B g A g A A (a) s (b) (c)

B A g g A s s g (d) g s s s (e)

(b) g

B B s s A A (a) (b)

g g B B g s s

Figure 19.16 Development of linear dunes from barchans: (a) after Bagnold (1941); (b) after Tsoar (1984).

wind regime and sediment supply (Warren and Allison, 19.7.1 Sediment characteristics 1998). Reconstitution time increases by several orders of magnitude from simple to complex dunes. The lifespan of Although sand sheets and zibar are often comprised of crescentic dunes superimposed on Namib linear dunes is coarse, poorly sorted sands, there is no evidence for gen- 30–50 years, compared to a reconstitution time of about eral relationships between sand particle size and sorting 6000 years for the main dune, based upon estimates of and dune morphological types. Although Wilson (1972) migration rates derived from combined GPR and OSL hypothesised that the spacing of aeolian bedforms at all studies (Bristow, Duller and Lancaster, 2007). scales is related to sand particle size, these relationships are not generally supported by empirical data (Wasson and Hyde, 1983a). 19.7 Controls of dune morphology

Empirical studies as well as numerical modelling have es- 19.7.2 Wind regimes tablished that the directional variability of the wind regime is a major determinant of dune type. Wind speed, grain Global (Fryberger, 1979; Wasson and Hyde, 1983b) and size and vegetation play subordinate roles. The effects of regional comparisons (e.g. Lancaster, 1983) between the sand supply are uncertain. The controls of dune size and occurrence of major dune types and the characteristics spacing are, however, less well understood. of local wind regimes support the hypothesis that the P1: TIX/XYZ P2: ABC c19 JWST032-Thomas December 15, 2010 18:3 Printer Name: Yet to Come

504 CH19 DESERT DUNE PROCESSES AND DYNAMICS

Transverse Dune

Seasonal reversal of crest line

Reversing dune with incipient star dune arm

Figure 19.17 Model for linear dune development (after Bris- tow, Bailey and Lancaster, 2000).

Development of arm major control on dune type is the directional variability of by secondary flow the sand-transporting winds, as characterised by the ratio between total potential sand transport (drift potential, or DP) and resultant (vector sum) potential sand transport (resultant drift potential, or RDP) (Figure 19.19). The directional variability or complexity of the wind regime increases from environments in which crescentic dunes are Accentuation of found to those where star dunes occur (Fryberger, 1979). arms by third wind Crescentic dunes occur in areas where RDP/DP ratios direction and secondary exceed 0.50 (mean RDP/DP ratio 0.68) and frequently flow occur in unimodal wind regimes, often of high or moderate energy. Linear dunes develop in wind regimes with a much greater degree of directional variability and com- Figure 19.18 Formation of star dunes by modiﬁcation of cres- monly form in wide unimodal or bimodal wind regimes centic dunes (after Lancaster, 1989b). with mean RDP/DP ratios of 0.45. Star dunes occur in areas of complex wind regimes with RDP/DP ratios less Wasson and Hyde (1983a) confirme that it was possi- than 0.35. Such a model is further supported by labora- ble to discriminate between dune types on the basis of tory experiments and numerical simulations of bedforms the complexity of the wind regime (the RDP/DP ratio) in directionally varying f ows (Reffet et al., 2008; Rubin and the equivalent (or spread-out) sand. Barchans oc- and Ikeda, 1990; Werner, 1995), which clearly show that cur where sand supply is low and winds are unidirec- crescentic dunes are not stable in directionally variable tional, crescentic dunes are located where sand is more fl ws and that linear dunes are the product of bimodal abundant. linear dunes develop where sand supply is re- wind regimes in which the two modes are separated by at ◦ lateively low but winds are more variable and star dunes least 90 . form in complex wind regimes with abundant sand supply. Similar relationships are evident in the Namib Sand 19.7.3 Sand supply Sea, although the range of directional variability is lower (Figure 19.19). The supply of sand has long been considered a factor in- EST is, however, a measure of the volume of sand fluencin dune morphology (Hack, 1941; Wilson, 1972). contained in the dunes and may be a product of dune type, P1: TIX/XYZ P2: ABC c19 JWST032-Thomas December 15, 2010 18:3 Printer Name: Yet to Come

CONTROLS OF DUNE MORPHOLOGY 505

Namib crests and braided patterns from vegetated linear dunes. Star Crescentic Sand Sea Linear Recent work (Tsoar, 1989) has suggested that there is a Fryberger distinct class of linear dunes that are partially vegetated. (1979) Their staight crests and orientation parallel to the dominant wind result from the anchoring of dunes by vegetation, as well as the deposition of sand downwind of 50 Star vegetation. Crescentic 40 19.7.5 Controls of dune size and spacing 30

20 Dunes exhibit a wide range of sizes, with close relationships between dune height and crest-to-crest spacing, indi- 10 cating that the patterns are self-organised (Figure 19.20). Linear Barchan In many sand seas, there is also a spatial pattern of dune Equivalent sand thickness (m) 0 size and spacing, such that small, closely spaced dunes 000.1 0.30.2 0.80.70.60.50.4 .9occur on the margins, with large dunes in areas of net Directional variability of wind regime (RDP / DP) sand accumulation. The controls on dune size and spacing have been de- Figure 19.19 Relations between dune type, wind regime bated for many years, but recent work has clarifie some variability (RDP/DP ratio) and sand thickness (RDP/DP data key issues. For example, the hypothesis of Wilson (1972) from Fryberger, 1979, Wasson and Hyde, 1983a, and Lancaster, 1989). that sand particle size controls dune spacing is not supported by empirical data (Wasson and Hyde, 1983a). Likewise, there is no simple relationship between dune with the dune type being influence by the other factors, size and sand transport rates (Elbelrhiti, Andreotti and especially the wind regime (Rubin, 1984). In the Namib Claudin, 2008; Lancaster, 1988). In the Namib Sand Sea, Sand Sea, dune types clearly occupy separate wind regime small dunes occur in areas of high sand transport rates domains and the EST data suggest that there is more sand (as indicated by wind data), whereas large dunes occur in complex linear and star dunes than in compound linear in areas of low net sand transport, as a result of their and all types of crescentic dunes. growth in conditions of abundant sand supply and a wind regime that promotes vertical accretion, rather than migration or extension. The size of superimposed dunes on 19.7.4 Vegetation larger forms, however, scales with the local sand transport rate (Havholm and Kocurek, 1988; Lancaster, 1988). The Vegetation affects the rate of sand transport via direct minimum size of such simple dunes is a function of the protection of the surface, absorption of momentum and sand transport saturation length (Claudin and Andreotti, partitioning of shear stress between the surface and plants 2006), which on Earth is around 20 m. (Wolfe and Nickling, 1993). As a result, sand transport Studies of dune patterns and numerical models indicate rates decrease exponentially with vegetation cover and that the spacing of crescentic dunes is a dynamic property sand surfaces are effectively stabilised when it exceeds of the pattern (Elbelrhiti, Andreotti and Claudin, 2008). about 15 % (Lancaster and Baas, 1998; Wiggs et al., The developed pattern results from merging and linking of 1994). dunes with a highly variable spacing, size and migration The effects of vegetation on dune morphology are less rate into fewer dunes with a much more regular pattern of well known. Hack (1941) suggested that in northeastern size and spacing (Figure 19.21), in a similar way to the Arizona crescentic dunes changed to parabolic dunes with development of patterns in wind ripples and subaqueous increased vegetation cover and that linear dunes occurred environments. Such ‘pattern coarsening’ leads to a more in areas with less sand and lower vegetation cover than ordered and stable pattern with fewer interactions between parabolic dunes. In the Negev Desert, Tsoar and Møller dunes. Similar downwind pattern development has been (1986) documented changes in the morphology of linear documented from crescentic dunes at White Sands (Ew- dunes as vegetation cover decreased as a result of graz- ing and Kocurek, 2010b), as well as linear dunes in the ing, including development of sharp sinuous ‘seif’ dune Simpson Desert (Fitzsimmons, 2007). P1: TIX/XYZ P2: ABC c19 JWST032-Thomas December 15, 2010 18:3 Printer Name: Yet to Come

506 CH19 DESERT DUNE PROCESSES AND DYNAMICS

100 Crescentic Dunes 1000 Linear Dunes

100

10 Dune height (m) Dune Height (m) 10

Namib Sand Sea (complex) Namib Sand Sea Namib Sand Sea (compound) Skeleton Coast South West Kalahari Gran Desierto Simpson Strzelecki Great Sandy Desert

1 1 10 100 1000 10000 100 1000 10000 Dune Spacing (m) Dune spacing (m)

Star dunes 1000

100 Dune height (m) 10

Namib Sand Sea Gran Desierto

1 100 1000 10000 Dune Spacing (m)

Figure 19.20 Dune height and spacing: (a) crescentic dunes, (b) linear dunes, (c) star dunes.

Fundamentally, the limits to dune size are set by the A further limiting factor on dune size may be the depth long-term supply of sand from sources external and/or of the atmospheric boundary layer in a region, which is internal to the sand sea or dune field This is a function a function of the annual variation of surface tempera- of the regional setting of the sand sea and also its size, so ture (Andreotti et al., 2009). The depth of the boundary that large dunes almost always occur in relatively large layer sets a limit to megadune growth by amalgamation sand seas. Ewing and Kocurek (2010a) have documented and nonlinear interactions between simple or elemental an increase in average dune spacing with dune fiel size, dunes. GPR and OSL dating studies show that large dunes which may reflec the degree to which the dune pattern are the product of multiple periods of dune construction, can evolve. in some cases spanning at least the last 25,000 years P1: TIX/XYZ P2: ABC c19 JWST032-Thomas December 15, 2010 18:3 Printer Name: Yet to Come

CONTROLS OF DUNE MORPHOLOGY 507

(a)

(b)

Figure 19.21 Dune pattern development in a downwind direction, White Sands, New Mexico (from Ewing and Kocurek, 2010b). P1: TIX/XYZ P2: ABC c19 JWST032-Thomas December 15, 2010 18:3 Printer Name: Yet to Come

508 CH19 DESERT DUNE PROCESSES AND DYNAMICS

(a) Dominant transport

Gross bedform normal transport Resultant transport

Dune trend

Subordinate transport

(b)

180

160

140

120

100

60 Gross bedform normal trend Crescentic 40 Linear Star 20

0 0 20 40 60 80 100 120 140 160 180 Dune trend

Figure 19.22 Gross bedform normal (GBN) transport direction and dune trends: (a) explanation of the gross bedform normal model (after Rubin and Ikeda, 1990); (b) relations between dune trend and GBN direction (after Lancaster, 1991, with additional data).

(Bristow, Duller and Lancaster, 2007; Kocurek et al., relative to the resultant or vector sum of sand transport 1991; Lancaster et al., 2002), so age of the megadunes (Fryberger, 1979). Thus dunes can be classifie as trans- is also an important factor. verse (strike of crestline approximately normal to resultant), longitudinal (crest parallel to resultant) or oblique ◦ (15–75 to resultant direction) (Bagnold, 1953; Glennie, 19.7.6 Dune trends 1970; Hunter, Richmond and Alpha, 1983; Mainguet and Callot, 1978). The factors that determine the alignment of dunes with re- Field and laboratory experiments with wind ripples and spect to the wind have long been a subject for speculation. subaqueous dunes (Rubin and Hunter, 1987; Rubin and Many workers have concluded that dunes are oriented Ikeda, 1990) suggest that all bedforms are oriented in the P1: TIX/XYZ P2: ABC c19 JWST032-Thomas December 15, 2010 18:3 Printer Name: Yet to Come

DUNE PATTERNS 509

direction subject to the maximum gross bedform normal (a) transport across the crest (Figure 19.22(a)), so that sand 30° Atlas Mountains transport from all directions contributes to bedform development. These experiments and subsequent numerical simulations show that the type of dune that forms is determined by the divergence angle between the dominant 20° and subordinate transport vectors and the ratio between S en eg al Lake the two primary transport directions (transport ratio). A R Nile R. . Chad

trend parallel to or normal to the resultant direction of Niger R. 10° sand transport is purely coincidental. This model is supported by empirical data on dune trends and winds. There is a close agreement between ob- 0 1000 km Sand sea Sand transport served and predicted gross bedform normal orientations in 0° vector the case of many barchans, crescentic dunes, simple linear 10° 0° 10 ° 20° 30° Ε dunes and star dunes (Figure 19.22(b)), suggesting that all 0 600 km 48¡ E (b) major dune types are oriented to maximise gross bedform normal sediment transport and therefore are dynamically Nafud

Ar similar (Lancaster, 1991). This approach can also be used abia n to identify dune trends that are out of equilibrium with the Dahna Gulf

Jafu

modern sand transporting wind regime, as demonstrated rah 24¡N for Mauritania, and to suggest wind regimes that could have produced the trends of dunes constructed in the past. Liwa

(Lancaster et al., 2002). Wahiba

i hal l K b'a Ru

19.8 Dune patterns Arabian Sea Sand Seas Dune systems exhibit clear patterns of dune morphology, size, spacing and crestline alignment on several spa- (c) tial scales: (1) regional, (2) within a sand sea and (3) within an area of dunes of similar morphology. The pattern of dunes is a reflectio of the external and internal factors that determine the present and past dynamics of the 20° Great Sandy Tanami Desert system. Desert On a regional scale, mapping of dune trends shows a pattern that is determined by regional atmospheric circulations (Figure 19.23), especially the pattern of winds Simpson Desert

outblowing around anticyclonic cells, as in the Kalahari Gibson Strzelecki Desert (Lancaster, 1981) and Australia (Wasson et al., 1988). Desert In the Sahara, the pattern of dune trends corresponds 30° Great Victoria Desert broadly to the of the Trade Wind circulation (Mainguet, Mallee 1984), modulated in some areas by topography and the Dunefield

development of low-level jets (Washington et al., 2006), 0 500 km while in Arabia, dune trends follow the Shamal circulation 118 ° 130° 142° (Glennie, 1998). Within individual sand seas and dune fields mapping of Figure 19.23 Dune morphology patterns – regional: (a) sand dune patterns from aerial photographs and satellite images transport patterns in the Sahara; (b) dune trends in the Arabian shows distinct spatial patterns of dune type, size and spac- Peninsula; (c) linear dune trends in Australia. ing (Breed and Grow, 1979; Ewing, Kocurek and Lake, 2006). In part, these patterns are the product of regional changes in wind regimes – with star dunes in areas of P1: TIX/XYZ P2: ABC c19 JWST032-Thomas December 15, 2010 18:3 Printer Name: Yet to Come

510 CH19 DESERT DUNE PROCESSES AND DYNAMICS

Box 19.2 Dune generations in Mauritania

Linear dunes dominate sand seas in the western Sahara. In the Azefal, Agneitir and Akchar sand seas of western Mauritania, three distinct crossing trends of linear dunes can be identifie on satellite images (Figure 19.24). From oldest to youngest they are: (1) a northeast–southwest trending class of large, degraded red brown linear ridges, (2) a north–northeast trending class of moderate-sized linear dunes, commonly with active crests, that appear and (3) a north trending class of small simple linear dunes. Field studies of dune sediments and stratigraphy show that the surface of the two oldest dune generations is littered with Neolithic stone artefacts and pottery and is underlain by a bioturbated, slightly indurated red-brown sand enriched in silt and clay, which is interpreted as a paleosol. OSL ages obtained from dunes of the different generations cluster between 25–15 ka (spanning the Last Glacial Maximum), 10–13 ka (during the Younger Dryas event) and after 5 ka. Modelling of the wind regimes that produced these dunes using the gross bedform normal (GBN) approach shows that the wind regimes that occurred during each of these periods were significantl different, leading to the formation of dunes on three superimposed trends – northeast–southwest, north–northeast to south–southwest and north to south. The trend of the N–S dunes corresponds closely to the GBN trend for modern winds near the coast, whereas the NE–SW and NNE–SSW dune trends imply increased sand transport from the east and north to northeast directions, corresponding to enhanced trade wind circulations (Lancaster et al., 2002).

020 km

19 30'N 19 30'N

Dune trends

< 5 ka

10 - 12 ka

15 - 25 ka 16W

Figure 19.24 Generations of linear dunes: western Mauritania (after Lancaster et al., 2002). P1: TIX/XYZ P2: ABC c19 JWST032-Thomas December 15, 2010 18:3 Printer Name: Yet to Come

REFERENCES 511

multidirectional winds, linear dunes in areas of bimodal creases as dunes become more widely spaced and the pat- winds and crescentic dunes where winds are unimodal tern becomes more organised away from sediment source (Figure 19.19). Such paradigms were used to explain the areas (Fitzsimmons, 2007; Mabbutt and Wooding, 1983). distribution of dune types in the Namib Sand Sea (Lan- By contrast, in the southwestern Kalahari dune field the caster, 1983) (Figure 19.21) and elsewhere (Breed and regional scale pattern of linear dunes becomes more com- Grow, 1979). Closer examination of dune patterns re- plex downwind and also from southwest to northeast veals sharp transitions in dune type, sediment compo- (Bullard et al., 1995). These patterns probably reflec de- sition, dune spacing and dune trend, which cannot be creased aridity to the south and east of the dune field attributed to changes in wind regime (Lancaster, 1999). as well as local sources of sediment in river valleys and These patterns are better explained by the concept of dune pans. At a local scale, however, self-organisation of dunes generations, in which dune field and sand seas are com- downwind of river valleys and pans is expressed by joining prised of multiple genetically distinct groups of dunes, of smaller dunes in Y-junctions to create a more organised each formed in a different set of initial conditions of wind pattern (Bullard and Nash, 1998). regime and sediment configuratio (Kocurek and Ewing, 2005; Lancaster, 1999). Dune generations can be recognised by differences in 19.9 Conclusions dune morphology, which are manifested by statistically distinct populations of dunes with a different crest orien- Significan advances in understanding of the dynamics tation, spacing, length and defect density (Beveridge et al., of desert dunes and the factors that determine their mor- 2006; Derickson et al., 2008), by differences in sediment phology have occurred in the past three decades. This has composition and colour, by chronometric and relative age come about as a result of several intersecting method- of different elements of a dune pattern and by geomorphic ological and conceptual advances, which include, but are relations between dunes of different characteristics, e.g. not limited to, the application of remote sensing im- crossing patterns (Lancaster et al., 2002) (see Box 19.2). agery to understand dune patterns, fiel studies of key Because only one set of dunes can form at a time in a dune-forming processes, OSL dating of periods of dune given set of boundary conditions, it follows that dune formation and GPR investigations of dune sedimentary field with simple patterns of dune morphology most structures. A rapidly evolving new development is the likely represent relatively young systems, whereas com- numerical modelling of dunes and dune systems, facil- plex patterns are the product of multiple periods of dune itated by vastly increased computing power, as well as construction. Such information can be used to infer the better understanding of fundamental dune processes. age of a dune fiel as well as the degree to which it has As a result, there is now a general understanding of the been subject to changes in the factors that control its de- processes that form or at least maintain most major dune velopment and current dynamics. types. The importance of scale effects is recognised, so Within a group of dunes, patterns of dune size and that the processes that form simple dunes need to be con- spacing and the parameters of dunefiel organisation vary sidered at different temporal and spatial scales from those systematically. In areas of crescentic dunes, such as at that form compound and complex dunes and sand seas. White Sands dune field dunes become larger and more The importance of past conditions of sediment supply, widely spaced in the direction of migration away from availability and mobility, determined by climatic and sea the linear source area (Figure 19.21). Dunes close to the level changes, in the formation of mega dunes and sand source of sand are smaller, migrate more rapidly and have seas and dune field is now well recognised. a high defect density. Downwind, these dunes interact with each other to produce a self-organisation of the pattern in which the dunes are larger, with longer crestlines and References fewer defects (Ewing and Kocurek, 2010b). Similar patterns of dune pattern evolution have been noted in barchan et al. dune corridors (Elbelrhiti, Andreotti and Claudin, 2008). Andreotti, B., Fourriere, A., Ould-Kaddour, F. (2009) Giant aeolian dune size determined by the average depth of the In areas of linear dunes, such as the Simpson–Strzelecki atmospheric boundary layer. Nature, 457 (7233), 1120–1123. Deserts of Australia, dune spacing is apparently related Baddock, M.C., Livingstone, I. and Wiggs, G.F.S. (2007) The to substrate type, reflectin sediment supply from lo- geomorphological significanc of airfl w patterns in trans- cal sources (Fitzsimmons, 2007; Mabbutt and Wooding, verse dune interdunes. Geomorphology, 87, 322–336. 1983; Wasson et al., 1988). Within this pattern, dune spac- Bagnold, R.A. (1941) The Physics of Blown Sand and Desert ing increases downwind and the number of Y-junctions de- Dunes, Chapman & Hall, London, 265 pp. P1: TIX/XYZ P2: ABC c19 JWST032-Thomas December 15, 2010 18:3 Printer Name: Yet to Come

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