ARTICLE IN PRESS
Geomorphology xx (2004) xxx – xxx
www.elsevier.com/locate/geomorph
Characterization of the aeolian terrain facies in Wadi Araba Desert,
southwestern Jordan
Walid Saqqa *, Mohammad Atallah
Department of Earth and Environmental Sciences, Faculty of Science, Yarmouk University, Irbid, Jordan
Received 6 May 2003; received in revised form 2 February 2004; accepted 17 February 2004
Abstract
The sand dunes in Wadi Araba Desert, southwestern Jordan, conform to the influence of two main wind systems: (1) the
Shamal ‘‘ north wind’’, the main determinant of dune patterns, and (2) the southerly winds caused by the Red Sea Trough and
Khamasin winds. Wadi Araba is a narrow elongated morphotectonic depression bordered by the high eastern and western
mountain ranges that would obstruct most of westerlies in winter and the hot dry easterlies in summer. Wadi Araba Desert is
caused by the rain shadow of the eastern and western topographic highs, with an arid – hyperarid climate and moisture index
between 0.1 and < 0.05. An aeolian terrain occupies about 16% of Wadi Araba Desert divided into four sand dune fields that
contain different dune types, interdunes, and sand sheet facies. The development of the aeolian terrain was more likely in the
interglacial periods of latest Pleistocene – Holocene during which wind deposition and fluvial erosion were prompted. The
variability of wind direction, wind speed and rates of sand supply led to a variety of simple, compound and complex dune
formations. Barchanoid (barchan, barchanoid ridge, transverse dunes) are very common. Linear dunes, nabkhas and climbing
dunes are less common. The dunes in Wadi Araba are either mobile (active) or stabilized. Dune fixation is primarily by desert
shrubs or cementation in case of aeolianites. Interdune troughs vary from dry to damp. Sand sheets are mainly unvegetated or
barren. Dune sands are well sorted to moderately well sorted with more than half the sand falling in fine – medium sand
fraction. Interdune areas and sheet sands are moderately – poorly sorted with grain size between 0.06 – 5 and 0.1 – 0.5 mm,
respectively. Angular-subrounded grain shapes are relatively common on all sides of barchanoid dunes. The sediments of the
aeolian terrain are chiefly composed of quartz, feldspar, mica and kaolinite. Calcite and dolomite are less predominant
minerals.
D 2004 Elsevier B.V. All rights reserved.
Keywords: Wadi Araba Desert; Winds; Sand rose; Drift potential; Barchanoid; Barchan; Transverse; Linear; Nabkha; Climbing dunes; Interdune
trough; Sand sheet
1. Introduction
1.1. Background
* Corresponding author.
E-mail addresses: [email protected] (W. Saqqa),
[email protected] (M. Atallah).
Fascinating geological and geomorphologic features occur within and around Wadi Araba, Jordan.
An aeolian terrain covers about 16% ( c 400 km2)
0169-555X/$ - see front matter D 2004 Elsevier B.V. All rights reserved.
doi:10.1016/j.geomorph.2004.02.002
GEOMOR-01518; No of Pages 25
ARTICLE IN PRESS
2
W. Saqqa, M. Atallah / Geomorphology xx (2004) xxx–xxx
of the Wadi Araba area, extending between a point
c 35 km north of the Gulf of Aqaba to about 22 km
south of southern margin of the Dead Sea basin. The
aeolian terrain contains various types of dunes in
association with deflation zones, sand sheets, aeolianite and interdune troughs. Coalescing alluvial
fans, mud flats (Qa or ‘‘Gaa’’ in Arabic) and inland
sabkhas are common features interrupting the aeolian
terrain.
1.2. Previous studies and aim of study
Previous research investigating the geology, structure, mineral prospecting and the hydrology of the
Wadi Araba area include Quennell (1958, 1983),
Bender (1968, 1974), Farhat (1979), Jarrar (1979,
1984), Atallah (1986, 2002), Abu-Zir (1989), Abed
and Al-Hawari (1991), Abdelhamid and Rabba
(1994), Saqqa (1997, 1998) and Galli (1999). Recently, Abed (2002) wrote a geologic overview
about the Taba inland sabkha in southern Wadi
Araba. A general lack of knowledge pertaining to
the type of movement and mode of accumulation of
the aeolian sands in the study area prompted the
current investigation.
1.3. Study area
The study area is a part of the Dead Sea Transform
Valley (DSTV). It occurs between 29j50VN and 31jN
latitudes (Fig. 1).
1.4. Geomorphology, climate, soil
The DSTV extends for 375 km along the entire
length of Jordan from the Gulf of Aqaba to Lake
Tiberias (Sea of Galilee) (Fig. 1). It consists of
three morphotectonic depressions, the Wadi Araba
in the south, the Dead Sea in the middle and the
Jordan Valley in the north. The DSTV is 9 –25 km.
wide south of the Dead Sea and narrows up to 9
km. near Lake Tiberias in the north. The southern
end of the DSTV lies at sea level at the Gulf of
Aqaba on the shores of the Red Sea; the valley
rises gradually up to 250m asl at Jabal Ar-Risha in
central Wadi Araba. From there, it gradually
decreases up to 412 m bsl to the present Dead
Sea shorelines. The DSTV is bordered by the
Western Mountain Range (in excess of 1000m asl)
and the Eastern Mountain Range (elevation 1200–
1500 m asl) (Bender, 1974). The topographic highs
have a very steep escarpment that is crossed by
deep gorges that drain towards the DSTV from the
east and west. Large alluvial fans built out from
wadi mouths at the base of the escarpment onto
inland sabkhas and/or mud pans in the heart of the
Wadi Araba (Fig. 1). These alluvial fans indicate
periods of heavy rains with resultant debris-flow
and stream-flood deposition.
Desert conditions prevail in Wadi Araba and the
adjacent Negev desert to the west, which are in the
rain shadow of the bordering Eastern and Western
Mountain Ranges. A short dry cool winter and a
very hot dry long summer characterize the modern
climate. The mean air temperature is c 15 jC in
January and c 33 jC in August. The daily maximum temperature reaches 50 jC in summer and
drops to 0 jC in winter. The temperature difference
between daytime and night is large enough (>30 jC)
to restrict the growth of plants. The annual rainfall is
usually < 100 mm. Based on long-term meteorological data (1955 – 2002) obtained from Ghor As-Safi
Meteorological Station (GSMS) and Aqaba Airport
Meteorological Station (AAMS), the annual precipitation decreases from c 80 mm (GSMS) to c 40
mm (AAMS) from north to south of the study area.
The eastern bordering highlands (Jordan Plateau),
however, receive some 200 mm/year of rainfall.
Localized storms occur mainly in conjunction with
the presence of the active Red Sea Trough (RST) or
in conjunction with a low-pressure center over
southern parts of Jordan and the adjacent regions
of Israel, Sinai, and northwestern Saudi Arabia
(Sharon and Kutiel, 1986; Kahana et al., 2002).
The mean annual potential evapotranspiration fluctuates between c 2200 mm (northern Wadi Araba)
and c 2500 mm (southern Wadi Araba). A rapid
loss of soil moisture results from high levels of
evapotranspiration. In winter, the relative humidity
fluctuates between 53% (AAMS) and 62% (GSMS)
but in summer it diminishes to 30% (AAMS) in the
southern parts and 40% (GSMS) in the northern
parts (Fig. 2). Dust storms are common in Wadi
Araba beginning in the spring and continuing
through the early summer (March– June) (Fig. 3).
The DSTV is of a typical Sudanian bioclimate
ARTICLE IN PRESS
W. Saqqa, M. Atallah / Geomorphology xx (2004) xxx–xxx
Fig. 1. (A, B) Distribution of the aeolian terrain facies in Wadi Araba Desert, southwestern Jordan.
3
ARTICLE IN PRESS
4
W. Saqqa, M. Atallah / Geomorphology xx (2004) xxx–xxx
Fig. 1 (continued).
(Long, 1957) with entisol, enceptisol and aridisol
soil types and with tropical trees and shrubs such as
Acacia savanna and Ziziphus spina-christi. (AlQudah, 2000).
1.5. Geology
The DST fault is an active left lateral strike– slip
fault that has undergone tectonic movement since
ARTICLE IN PRESS
W. Saqqa, M. Atallah / Geomorphology xx (2004) xxx–xxx
5
Fig. 2. Plot of the mean temperature (jC), mean rainfall (mm), mean evaporation (mm) and mean relative humidity (%) in Wadi Araba. (Records
for 30 years between 1955 and 1985.)
Miocene– Pliocene time and is still active (Quennell,
1958; Garfunkel et al., 1981; Galli, 1999; Atallah,
2002) The highlands flanking Wadi Araba correspond to the northern end of the Arabo – Nubian
Shield and they resulted from the upheaval of the
eastern Transjordan block during the Tertiary (Bender, 1968). The eastern Transjordan block consists of
a complexity of igneous, volcanic and metasedimentary rocks of Precambrian age (Jarrar, 1984) overlain
by Paleozoic –Tertiary sediments. Three inland sabkhas (Fig. 1) located in southern Wadi Araba are
chiefly composed of detrital mudstone, dolomite and
gypsum (Abed and Al-Hawari, 1991; Abed, 2002).
Lake varved sediments of late Pleistocene age are
exposed in the central parts of the DSTV (Abed,
1982; Abed and Yaghan, 2000). These sediments
have been easily eroded into a badland landscape of
gullies, hills and cliffs near the Dead Sea basin. Late
Fig. 3. Annual recurrence of dust storms in Wadi Araba. (Records
for 30 years between 1955 and 1985.)
Pleistocene– Recent detrital sediments cover the center of Wadi Araba.
2. Materials and methods
The description of the aeolian terrain facies in
Wadi Araba is based on field observations, landsat
imagery (TM, scale 1:150,000) and aerial photographs (1:10,000). The grain size, roundness and
mineralogy of the aeolian sediments were investigated. Grain-size analysis was determined using a
Retsch AS200 Digit Analytical Sieve Shaker and
Retsch-Graintest software. Powers (1953) roundness
chart and binocular microscopy were used to determine the roundness of sand grains. A Phillips PW1729 X-ray Diffractometer was used for the mineralogical analysis. To describe weather conditions in
the study area, mean temperature (jC), total rainfall
(mm), relative humidity (%), mean evaporation
(mm), wind speed (m s 1), wind direction (degrees)
and the time periods of dust storms (days) were
analyzed. Meteorological data (1955 – 2002) were
collected from two stations operated by the Jordan
Meteorological Department (JMD): (1) the Aqaba
Airport Meteorological Station (AAMS) in the south,
and (2) Ghor As-Safi Meteorological Station
(GSMS) in the north. The aridity index is given by
AI = P/T + 10 (De Martonne, 1925; Murai and
Honda, 1991), where P = annual rainfall (mm),
T = sum of monthly mean temperature divided by
12. The moisture index is given by: MI = P/PET
(Pahari and Murai, 1995), where P is the annual
ARTICLE IN PRESS
6
W. Saqqa, M. Atallah / Geomorphology xx (2004) xxx–xxx
rainfall and PET is the potential evapotranspiration.
Sand drift potential DP=(U2(U Ut)/100)t (Fryberger, 1979) was determined, where U is wind
velocity, Ut is threshold velocity, and t is time of
wind blew (%). Sand roses were plotted after simplifying wind speed data and related calculations
using a visual basic application programmed by A.
Saqqa.
3. Results and discussion
3.1. Aridity and moisture indices
In terms of aridity and moisture indices (AI and
MI), the Wadi Araba area is classified as a desert
arid zone or arid – hyperarid dryland. The AI and MI
are respectively between 1.0 – 2.2 and < 0.05 – 0.1
along south –north ascending precipitation gradient.
Usually, a zone of aridity is classified as desert if the
aridity index (AI) is < 5 (Murai and Honda, 1991;
Pahari and Murai, 1995). The terms arid and hyperarid are applied if the moisture index (MI) is
respectively between 0.05 and 2.0 and less than
0.05 (UNEP, 1992; Pahari and Murai, 1995; Middleton and Thomas, 1997).
3.2. Determinant winds and sand drift
Wind speed in the Wadi Araba Desert varies
throughout the year. The variation in wind speed is
more likely to be caused by a large pressure difference
developed between the lowland DSTV and the surrounding eastern and western high elevation mountains, in addition to the influence of desert surface
roughness and local topography. The air above the
DSTV is heated, expands, and rises leaving behind a
low-pressure zone. Cool air flows down from a highpressure zone and is balanced by an outward flow of
the air (Krauskopf and Beiser, 2002). Fig. 4 shows
that wind speed in Wadi Araba Desert reaches up to
36 knots/h (18.5 m s 1) or it may exceed this value
during windstorm events. Monthly wind speed averages 6 – 15 knots/h (3.1 – 7.7 m s 1) but it diminishes
to < 5 knots/h (2.6 m s 1) as a minimum value. The
annual wind speed (Fig. 5) averages 8 – 13 knots/
h (4.1 – 6.7 m s 1). Practically, maximum wind speed
is not restricted in a certain period of year, but it is
rather common at end of spring – early summer
(April– June), end of summer (September), and end
of winter (February –March). These periods are favorable times for the Khamasin sandstorms, monsoons,
and the RST. Undoubtedly, wind speed variation has
its own effect on rates of wind erosion, wind deposition, advance and reshaping of dunes. Moreover, it
may serve as indicator for climatic condition variations in the whole region.
By relating the annual wind speed data to values
z 12 knots/h (6.2 m s 1), where the value 12 knots/
h is taken as the threshold velocity, we found that
the ratios of surface wind energy responsible for
sand drift are between 56% (year 1987) and 25%
(year 2001) (Fig. 6). A remarkable decrease in the
annual wind energy is expected in the ninetieth, as
it drops from 38% (1991) to 25% (2001) (see Fig.
6). It appears from wind speed analysis that only
about half to quarter the potential wind energy is
used as a motive force for wind erosion and
sediment mobilization.
Wind roses (Fig. 7) suggest that four groups of
surface winds blown out over the Wadi Araba Desert
from NW, NE, SW and SE directions. The strongest
and most frequent winds are encompassed within the
first two groups: NW and NE, attributed to the
northern winds, the ‘‘Shamal’’ (Cooke et al., 1993)
drawn in by the Asian monsoon; these northern winds
are the driving force for movement of sand and dune
formation in the Arabian Peninsula (Membery, 1983).
The Shamal wind comprises 70% of the total wind
energy (Abed, 2002). An active Red Sea Trough
(RST) implies southerly winds over southern Israel,
southern Jordan, Sinai, and northern Saudi Arabia and
is capable of transporting moist tropical air in the
middle troposphere causing some rainfall in winter
over such arid regions (Itzigsohn, 1995; Dayan et al.,
2001; Kahana et al., 2002). The southerly winds
which comprise about 8% of the total prevailing
winds (Abed and Al-Hawari, 1991; Abed, 2002),
may also be initiated by the Khamasin winds blowing
in the early and late summer months and lasting for
several days at a time before terminating abruptly as
the wind direction changes and much cooler air
follows. Such winds arise when cold winds from
Europe converge with the hot winds from the Sahara.
The resulting atmospheric depression moves eastward
from Africa off the Arabian Peninsula pushing hot,
ARTICLE IN PRESS
W. Saqqa, M. Atallah / Geomorphology xx (2004) xxx–xxx
Fig. 4. Time series 1975 – 2001, showing monthly minimum, maximum and average wind speed (AAMS).
7
ARTICLE IN PRESS
8
W. Saqqa, M. Atallah / Geomorphology xx (2004) xxx–xxx
Fig. 4 (continued).
ARTICLE IN PRESS
W. Saqqa, M. Atallah / Geomorphology xx (2004) xxx–xxx
9
Fig. 4 (continued).
sand-bearing winds followed by a drop in temperature. Extreme dust storm events in the study area are
simultaneous with the Khamasin (see Fig. 3). A
comparison between the meteorological data from
the two stations AAMS and GSMS showed that
duststorms and wind speeds are more intense in the
southern Wadi Araba Desert. Southwesterly winds
generated by depressions across the Mediterranean
and the Middle East would affect the Sahara and the
Arabia as far as 30jN in winter (Breed et al., 1979; El
Baz, 1986; Jones et al., 1986). Accordingly, the
southwesterly winds do not blow across the whole
region of Wadi Araba, but their zone of influence
extends only up to 55 km north of the Gulf of Aqaba.
The interaction between the westerlies and the hot dry
eastern trade winds (Wilson, 1971; Fryberger, 1980)
or between moist cold air of low pressure and hot dry
air of high pressure (Trenberth, 1992) at the boundary
between winter and summer brings a complex wind
regime to the northern edges of the Sahara and the
Arabian Desert. We believe that this kind of wind
complexity may have little chance of operating in the
Wadi Araba Desert. It is clear that the eastern and
western topographic highs (elevation>1000m asl)
bordering DSTV from east and west serve as natural
obstructions against most easterly and westerly winds.
Fig. 8 shows the annual sand drift potential in Wadi
Araba Desert. The resultant sand drift direction
(RDD) is mainly to the south, southeast or southwest.
Previous studies in many localities of the Arabian
Gulf showed that the season of high drift potential of
dune sands is associated with the Shamal winds at
Fig. 5. Time series 1975 – 2001, showing annual wind speed average (AAMS).
ARTICLE IN PRESS
10
W. Saqqa, M. Atallah / Geomorphology xx (2004) xxx–xxx
Fig. 6. Time series 1975 – 2001, showing the frequency % for wind speed above the threshold velocity.
speeds between 12 and 26 knots/h (6.1 – 13.3 m s 1)
(ARAMCO Survey, 1974; McKee, 1979) similar to
results of the present study. For example, sand drift
shown in Fig. 8 is compatible with the Dhahran sand
rose of the Jafurah sand sea, Saudi Arabia where the
RDD implies drift to the SSE (Fryberger et al., 1983).
The annual sand drift potential (DP) measured in
vector units is between 140 (year 2001) and 1070
(1975), which corresponds to values between 10 and
80 m3/meter-width (Fryberger et al., 1983). The index
of directional variability RDP/DP (RDP: resultant
drift potential) is between 0.5 and 0.9, which is
arbitrarily classified as intermediate-high ratios
(McKee, 1979). This means a narrow –wide unimodal
flow pattern of less directional variability typical for
barchanoid dunes (McKee, 1979).
3.3. Aeolian terrain facies
3.3.1. Sand dunes
Four sand dune fields were mapped in Wadi Araba.
From north to south these are: (1) Salmani sand dune
field, (2) Fidan sand dune field, (3) Qa As –Suaydiyyin sand dune field, and (4) the Gharandal – Taba
sand dune field (Fig. 1). The Gharandal – Taba is the
largest dune field (200 km2). Dune types were determined according to the descriptive typology of dune
and the mode of complexity of pattern displayed by
the groups of dunes (McKee, 1979). Simple and
compound dune types are common in the study area
(Figs. 9 and 10). Complex dunes are occasional and
only observed by aerial photograph analysis in the
mountainous region in the eastern parts of the Gharandal – Taba- and Qa As – Suaydiyyin sand dune fields
(Fig. 11). Barchanoid dunes and intergrading subtypes: barchans (Figs. 9 and 10), barchanoid ridges
(Fig. 12) and transverse dunes (Figs. 12 and 13) are
very common; linear dunes (Fig. 14), nabkhas and
climbing dunes are less common.
The barchan dunes vary in size. They usually occur
as a solitary or as a compound type mainly anchored
by shrubs. Small ephemeral barchans ( < 3m. high)
lasting a few months and mesobarchans (3 –10 m
high) lasting 1 –30 years (Cooke et al., 1993) are
present. Barchan dunes are mainly confined to the
regions where winds allow for crescent forms to grow
and to persist (Bagnold, 1941; Finkel, 1959; Sharp,
1964). Barchan dunes maintain their size by sands
supplied from the upwind side (Hastenrath, 1987).
Isolated barchans tend to develop where limited
amounts of sand are available and almost unidirectional currents (McKee, 1979; Wasson and Hyde,
1983). We suggest that the compound barchan ridges
occur where the large basal mound has a single
proportional slip face and an upwind slope covered
with smaller barchans or smaller barchanoid ridges.
The small dunes are all oriented in the same direction
as the main dune. An oblique sand supply accounts
for some asymmetric barchans.
Barchanoid ridges and transverse dunes form in the
same wind regime as barchan dunes in regions where
more sands are available and winds are variable
(Wasson and Hyde, 1983). Barchans merge into
wave-like shapes to produce barchanoid ridges (Cooke
et al., 1993) or transform into transverse dunes in
regions where greater sand supply is available
(Araya-Vergara, 1987). In the study area, the barchanoid ridges and the transverse dunes do occur as
ARTICLE IN PRESS
W. Saqqa, M. Atallah / Geomorphology xx (2004) xxx–xxx
11
Fig. 7. Annual wind rose patterns showing the vector means of the prevailing wind directions (*: AAMS, **: GSMS).
groups with interdune corridors of different sizes
(tens – hundreds meters). The crestlines are sinuous
and the dune bodies tend to adopt linguoid – barcha-
noid shape with a gentle northern stoss side and a steep
southern lee side. The average dip angle of the windward side (stoss side) is usually < 20j and for the
ARTICLE IN PRESS
12
W. Saqqa, M. Atallah / Geomorphology xx (2004) xxx–xxx
leeside is >30j. As noted by McKee (1966), such dip
angles are typical for barchan dunes early in their
formation in the presence of unidirectional winds.
Active short-lived wind ripples appear on the stoss
sides and crests of dunes but are absent on the slip
faces. Active short-lived ripples tend to develop on
sandy surfaces that are in a state of relative equilibrium
or slow deposition, while surfaces experiencing
marked erosion or vigorous deposition generally do
not display ripples (Sharp, 1963). Seppälä and Lindé
(1978) reported that ripples in well-sorted sands of a
mean size ca. 0.2 mm take less than 10 min to reach
equilibrium under new wind conditions. Field study
showed that the stoss sides of barchanoid dunes are
usually overloaded with coarse-grained sands and the
lee sides are overloaded with fine-grained sands.
Fig. 8. Annual sand roses based on wind data from AAMS station. Sand drift potential (DP) are in solid lines, resultant sand drift direction
(RDD) are in dash lines. Values of DP and resultant drift potential (RDP) are in vector units. RDP/DP is index of directional variability.
ARTICLE IN PRESS
W. Saqqa, M. Atallah / Geomorphology xx (2004) xxx–xxx
13
Fig. 8 (continued).
Bagnold and Barndorff-Nielsen (1980) showed that
the probability that a grain will be eroded or deposited
is likely to be a function of the logarithm of the grain
size. Small grains tend to move away from the windward side (stoss side) and accumulate in the lee side.
Despite the prevalence of barchanoid dunes created
by northerly winds, with their downwind slip faces
directed south, it seems that some wind reversals, i.e.
southerly winds produced another group of barchan
dunes with downwind slip faces directed north (e.g.
Gharandal –Taba sand dune field). A similar case was
mentioned by Fryberger et al. (1983) in the Jafurah
sand sea, Dhahran area in Saudi Arabia.
Barchanoid dunes have tabular cross bedding preserved on slip (downwind) faces with foresets 0.5 – 1.0
m in length and a slope angle around 30j. Tabular
cross bedding is caused by grainfall produced-strata or
straight-crested dune migration (Fryberger et al.,
1983). The crossbeds are truncated at the top by
subhorizontal to slightly oblique reactivation surfaces
as a result of scouring, possibly simultaneous with
wind direction shifts. Cross bedding in dunes is
occasionally disturbed or contorted by plant roots
bioturbation.
Linear dunes (Gharandal – Taba sand dune field)
are straight-crested to slightly sinuous sand ridges
with two symmetrical slip faces. They are typically
much longer than they are wide. The dunes stand a
few meters above the interdune corridors with their
long axes extending more than 1 km. They commonly
ARTICLE IN PRESS
14
W. Saqqa, M. Atallah / Geomorphology xx (2004) xxx–xxx
Fig. 9. Solitary barchan dune (Qa As – Suaydiyyin sand dune field) with two asymmetric arms directed SE. Western mountain range seen in
background of photo.
Fig. 10. Aerial photograph showing compound dunes of barchanoid type (Fidan sand dune field). Arrow indicates sand drift direction.
ARTICLE IN PRESS
W. Saqqa, M. Atallah / Geomorphology xx (2004) xxx–xxx
15
Fig. 11. Aerial photograph showing complex dunes (Gharandal – Taba sand dune field).
occur as isolated ridges or sets of parallel ridges
separated by regularly spaced corridors of widths
< 100 m. The origin of linear dunes is confusing, with
several models suggested in the literature. The barchan-to-linear dune model (Bagnold, 1941) is more
likely to explain the origin of linear dunes in Wadi
Araba Desert for two reasons: (1) availability of wind
strength variation throughout the year, and (2) the
juxtaposition of barchanoid (dominant type) and linear dunes. An asymmetric barchan is converted to a
linear dune by alternating strong winds that build the
dune and gentle persistent winds which elongate the
dune. However, this model has little support from
field evidence (e.g. Lancaster, 1980; Tsoar, 1984;
Hastenrath, 1987). Fig. 14 shows pairs of linear dunes
that merge at low angle Y-junctions. This means that
parallel linear dunes are no longer stable in their
original places, but they are reworked by wind currents. The Y-junctions are open in an east –southeast
direction. They be may produced by side winds at
oblique angles (Bagnold, 1941; Tsoar, 1982; Thomas,
1986). An alternative explanation is that Y-junctions
have been formed by wind flowing in contra-rotating
helical spirals or roll vortex pairs transporting sands
obliquely (Bagnold, 1953; Folk, 1971; Glennie, 1987;
Cooke et al., 1993). The downwind vortices may be
produced by duststorms frequently occurring in parallel lines (e.g. Hastings, 1971; Swift et al., 1978;
Middleton et al., 1986). The subdued topographic
mass of Jabal Um-Nukhyla, located east of the Gharandal – Taba dune field (Fig. 1), may create natural
wind passages or tunnels that would facilitate the
formation of wind vortex pairs.
Nabkha dunes of various sizes develop when
aeolian sands move between growing bushes and
feed sand mounds between them. The nabkhas in
the study area are randomly distributed because of
the scattering of desert plants. Stems of bushes are
still visible in small nabkhas, but in larger ones they
may entirely covered by accumulated sands. Different sizes of the nabkhas are affected by changes in
wind velocity and wind direction, being large when
ARTICLE IN PRESS
16
W. Saqqa, M. Atallah / Geomorphology xx (2004) xxx–xxx
Fig. 12. Aerial photograph showing barchanoid-transverse dune system (Fidan sand dune field). Arrow indicates sand drift direction.
the winds are light, and small when they are strong
(Worral, 1974; Warren, 1988; Hesp, 1989; Cooke et
al., 1993). However, some of the nabkhas from the
study area are built of cohesive, silty-fine sand
when they form near stream channels and around
mudflats.
Fig. 13. Transverse dune in front of basement igneous rock (upper right) and an interdune trough in foreground.
ARTICLE IN PRESS
W. Saqqa, M. Atallah / Geomorphology xx (2004) xxx–xxx
17
Fig. 14. Linear dunes trending NW – SE with prominent low angle bifurcated Y-junctions (Gharandal – Taba sand dune field). Arrow indicates
wind direction.
The anchored dunes (climbing, echo, and flank
dunes) occur within the mountainous regions where
drifted sands climb above the foothills. They may
reach sometimes the summits of the topographic highs
(e.g. Jabal Um-Nukhyla, south Gharandal). These
dunes develop as a result of upwind vortices in front
of the scarp (slope angle is >50j). The vortex developed by the velocity gradients above the ground
(Bowen and Lindley, 1977; Cooke et al., 1993). Echo
dunes develop when the vortex is capable of sweeping
out a corridor in front of the scarp. When winds
accelerate around the flanks of the topographic highs,
they become more erosive and are capable of clearing
corridors between the topographic highs and the
accumulating sand mounds. Flanking dunes develop
in at these locations (Cooke et al., 1993).
3.3.2. Interdune troughs
Interdune troughs which vary in size and location
are mainly at three types: (1) dry depositional
interdunes between barchanoid dunes, which consist
of laterally continuous and flat-gently dipping thin
layers or laminae showing almost a transition upward into dunes. These laminae resulted from downwind migration and deposition of current ripples.
Surfaces of interdunes show current ripples and have
some plant remains and mottled-roots usually around
moist areas. Gravel deposits of intermittent streams
occur in the dry depositional interdunes, (2) dry
erosional interdunes between linear dunes, with
scour surfaces developed as a result of the storage
of drifting sand by the dunes immediately upwind,
despite lower drift rates in interdune areas (Ahlbrandt and Fryberger, 1981; Fryberger et al., 1983),
and (3) damp depositional interdunes with flourishing desert plants ( z 1 m. height) between barchanoid dunes (e.g. Salmani and Fidan sand dune
fields). This type of interdune has no signs of salt
encrustation. Moistening of interdunes and the
growth of plants above the interdune surface are
ARTICLE IN PRESS
18
W. Saqqa, M. Atallah / Geomorphology xx (2004) xxx–xxx
more likely to be related to a saturated water table
of low salinity. The presence of vegetative cover is
indicative of lower wind velocity and lower rates of
sand drift. It seems that this type of interdune was
built up of laterally continuous thin beds of horizontal lamination ( z 0.3 m. total thickness) considered as a subsoil layer for plant growth. Episodic
stream flooding causes some local erosion in interdune troughs that has led to exposure of the horizontally laminated sediments close to the banks of
dry stream channels. Horizontal lamination in sands
is indicative of slow deposition rates. Gravelly sediments occur on the bottom of stream channels.
3.3.3. Sand sheets
Sand sheets in Wadi Araba are of the unvegetated
or barren sand sheet type. The sand sheet surface is
almost smooth, flat and covered with thin mantle of
coarse sands and widely spaced sinuous-crested megaripples (wavelength >1 m). The smoothness of sand
sheet surface is likely to be created by the abrasion of
transported sand. Lian-You et al. (2003) showed that
the abrasion capacity of transporting sand is increased
logarithmically with wind velocity, and the abrasion
itself is affected by sand transport rates. Megaripples
formed at the sand sheet surface were more likely
developed during windstorms. Sharp (1963) and
Sakamoto-Arnold (1981) suggested that the megaripples in the southern San Joaquin Valley, California
were produced in a short period of time, of not more
than hours, under windstorm conditions. However,
some megaripples may take years to develop and may
last for centuries (Cooke et al., 1993). The idea that
ripples tend to flatten out at higher wind strengths and
greater rates of deposition, and can result in sand sheet
formation (Glennie, 1987) is a possible origin for the
unvegetated sand sheets in the study area. Shallow test
pits ( c 0.5 m depth) made into the sand sheets
showed a horizontal lamination facies truncated at
the top by gently dipping ripple cross-lamination (dip
angle c 15j). Glennie (1987) pointed to the possible
change in lamination patterns when the mode of
transport changes from traction to grainfall. Fryberger
et al. (1983) assumed that sand sheets develop independently of dunes particularly when steeply dipping
crossbeds are absent, and, the sand sheets grow
vertically and laterally from ripples and grainfall
deposition during episodes of sandstorms.
3.4. Textural attributes of the aeolian terrain facies
3.4.1. Grain size
Bagnold (1941) defined dune sands as any particle
with a grain size between 0.02 and 1.0 mm. Ahlbrandt
(1979) used the range between 0.1 and 1.6 mm.
Despite the discrepancies in the quantitative definition
of dune particle sizes, dune sands are qualitatively
defined as any particle that is light enough to be
moved by winds but too heavy to be held in suspension in the air. Results of the grain size analyses (Fig.
15a) showed that dune sands fall between 0.05 and 2
mm, which are very, much close to grain size of most
sand sea deposits as mentioned by Cooke and Warren
(1973). The mean grain size is between 0.2 and 0.35
mm. Dune sands are generally well sorted to moderately well sorted (0.35f –0.7f).
The sediments from the interdunes (Fig. 15a) are
more variable in grain size and less well-sorted than
the dune sands. The grain size is between 0.06 and 5
mm with as up to 40% coarser particles (>1 mm). The
mean grain size varies between 0.25 and 0.6 mm.
Sands of sheet-like bodies have grain sizes between 0.1 and 0.5 mm, i.e. towards an excess of fine
sand in comparison with interdune sands (Fig. 15a).
The mean grain size is between 0.15 and 0.4 mm.
Sheet sands are better sorted than the interdune
sands.
Significant differences of grain size were observed
between the sand particles from the stoss side (upwind
side), crest, and the lee side (slip face) of barchanoid
dunes. The results (Fig. 15b) showed that an excess of
fine tailing is in the order: stoss side ! crest ! lee
side. The mean grain size of sand particles in the stoss
side, crest, and lee side is between 1.0 and 1.6 mm,
0.7 and 1.1 mm, and 0.6 and 0.9 mm, respectively.
3.4.2. Roundness
The roundness of 85– 90% of sand populations
from the stoss side, crest, and lee side of barchanoid
dunes is commonly between angular to subrounded,
with little differences in the abundance ratios of the
classes (Fig. 16). The degree of angularity is more
among sand grains from the stoss side, in contrast to
rounding in the crest and lee side samples which tend
to be more rounded, remembering that the fine tailing
of grain size is toward the lee side (Fig. 15b). At first
glance, these findings are incompatible with the
ARTICLE IN PRESS
W. Saqqa, M. Atallah / Geomorphology xx (2004) xxx–xxx
19
Fig. 15. (a) Grain size analyses of sands from barchanoid dune, interdune trough, and sand sheet. (b) Grain size analyses of sands from the stoss
side, crest, and lee side of the barchan sand dune.
ARTICLE IN PRESS
20
W. Saqqa, M. Atallah / Geomorphology xx (2004) xxx–xxx
Fig. 15 (continued).
ARTICLE IN PRESS
W. Saqqa, M. Atallah / Geomorphology xx (2004) xxx–xxx
21
general idea that rounding in aeolian sands increases
with grain size (e.g. Goudie and Watson, 1981; Khalaf
and Gharib, 1985). Similar trends were achieved from
the analysis of aeolian sands from different regions of
the world, where the degree of highest rounding does
not necessarily occur in coarser grains (e.g. 1 – 2 mm),
but instead mean rounding values are relatively higher
in finer sand grains (Thomas, 1984, 1987a; Khalaf
and Gharib, 1985) as later discussed by Thomas
(1987b). This contradiction may resolved by considering that sand grains on the upwind side (stoss side)
as they moved by creep or sliding; rolling and grain
impacts or collisions may be inhibited. In the crest and
lee side, fine – medium sand particles are more likely
to be moved by saltation, involving grain impacts and
collisions. Moreover, strong winds preferentially sort
out more spherical, rounded sand grains (Mazzullo et
al., 1986) more likely near the crest and brink point. In
terms of sand sources, the aeolian sands in the study
area are derived mainly from similar nearby sources.
So, it is not unlikely that the grain shape is an
inherited property (Tucker, 1991). Quartz sand grains
display slight elongation in the direction of their optic
c-axes, a property that is more obvious in coarser
grains in comparison with small grains, which will
inhibit well rounding, and high sphericity of sand
particles.
3.5. Mineralogy of the aeolian terrain facies
Fig. 16. The roundness of the aeolian sand particles from the stoss
side, crest, and lee side of the barchan sand dune.
The sediments of barchanoid dunes, interdunes,
and aeolianite (cemented sand sheet-like bodies) are
composed mainly of quartz, feldspar, mica, calcite,
and dolomite. The dominant clay mineral is kaolinite
with a minor content of smectite. Sands of interdunes
showed a peak at 2h 10.7 (d = 8.26 Å) characteristic
for the hydrated iron-sulphate mineral coquimbite
{Fe2(SO4)39H2O} which forms instead of gypsum
under conditions of low salinity and more oxidizing
nearsurface groundwater conditions.
Quartz, feldspar and mica of detrital origin are
derived from the nearby basement rocks and the
overlying sediments. The dominance of quartz in dune
sands is primarily due to its durability and resistance
to chemical weathering in addition to the source
availability. It is likely that the Paleozoic sandstones
(mainly arkosic- and quartz arenite), exposed along
the eastern shorelines of the Dead Sea, are other
ARTICLE IN PRESS
22
W. Saqqa, M. Atallah / Geomorphology xx (2004) xxx–xxx
suppliers for dune sands in the Salmani field, the
nearest dune field to the southern tip of the Dead Sea.
The nearby Mesozoic –Tertiary carbonate rocks and
inland sabkha sediments in southern Wadi Araba are
possible sources for detrital carbonate minerals in
dune sands. In the interdunes and in the aeolianite,
calcite and dolomite are presumably authigenic and
act as pore-filling cement. The calcite in the aeolianite
is more abundant in the uppermost horizons (peak
intensity f 1500 counts) and tends to decrease in the
intermediate and lower horizons (peak intensity
f 1250 counts). Calcite possibly concentrates in the
capillary fringe zone with the help of evaporation.
Sporadic rainfall may leach calcite downward into the
lower horizons. Results of XRD give unequivocal
evidence of clay illuviation provided by more kaolinite in the lower horizons of the aeolianite where the
peak intensity changes from 1000 to more than 1500
counts.
3.6. Expected age of the aeolian terrain
Speculations about the age of sand dunes and
other associated facies in the aeolian terrain of Wadi
Araba Desert are based on previous research. The
paleoclimate trend in the late Pleistocene Last Glacial Maximum (LGM) ( c 22– 15 ka) showed evidence that it was cold and dry in the Sahara, Arabia
and SE Asia (China) with less rainfall, and expansion of deserts (Gasse et al., 1987; Petit-Maire and
Guo, 1997; El-Baz, 1998; Zhuo et al., 1998). Abed
and Yaghan (2000) suggested a similar cold, dry
paleoclimate trend in the Jordan Valley during the
LGM ( c 22– 15 ka). Glennie (1987) considered the
glacial periods of the Pleistocene as times for the
major erosional features of globally stronger winds
that blew with strength for a much greater part of
the year in modern deserts. He believed also that
fluvial erosion and wind deposition were much more
effective in the interglacial periods of late Pleistocene – early Holocene ‘‘climatic optimum’’ around
8000 – 5000 years BP, when convection-induced
thunderstorms were more likely to be responsible
for building alluvial fans in today’s deserts (Hötzl
and Zötl, 1978; Jado and Zötl, 1984). Similarly,
Gasse et al. (1987), Petit-Maire and Guo (1997),
El-Baz (1998) and Zhuo et al. (1998) believed that
warmer interglacial periods in the late Pleistocene
and the Holocene optimum were associated with
high rainfall and retreat of deserts. Hence, the
climatic optimum of interglacial periods of latest
Pleistocene – early Holocene are more likely time
periods of fluvial erosion, wind deposition and sand
dune formation in the study area, despite the likelihood that deserts shrunk at these times. The eluviation of kaolinite into lower aeolianite sand-sheetlike horizons is evidence of wet pluvial episodes
between dry, cold climate conditions.
4. Conclusions
The principal approaches to dune study in Jordan
and probably elsewhere are practically based on the
classification of dune forms from field study, landsat
imagery, and aerial photographs. Textural studies for
grain size and grain morphology, and understanding
physical processes of wind transport, sand drift, and
deposition of sands are also important. Sand rose
patterns enabled us to draw conclusions on the probable relation of wind strengths and directions to dune
type in the study area. The present study shows that
the types of aeolian terrain facies mainly depend on
the amounts of sand supply, wind variability and wind
speed, and vegetation cover. Local topography
(mountain scarps, inland- or marginal hills), desert
plants, and desert plain roughness are responsible for
obstructing wind forces, allowing sands to pile up in
drifts and ultimately for different shapes and sizes of
dunes to form.
Acknowledgements
The authors are indebted to the Yarmouk University for logistic support and transport facilities during
fieldwork. Dr. Tina Niemi, University of MissouriKansas City (UMKC), USA is especially thanked for
reading the manuscript carefully and for providing
many relevant corrections and advice to improve it.
Thanks are extended to Dr. A. Biss and Dr. M. Taj for
their assistance in map drawing. Mr. I. Zaid, Yarmouk
University, Department of Biology, helped in developing and printing aerial photographs used in this
paper. Comments of anonymous reviewers are so
much appreciated.
ARTICLE IN PRESS
W. Saqqa, M. Atallah / Geomorphology xx (2004) xxx–xxx
References
Abdelhamid, Gh.A., Rabba, I., 1994. An investigation of mineralized zones revealed during geological mapping, Jabal Hamra
Fiddan-Wadi Araba, using landsat TM data. International Journal of Remote Sensing 15, 1495 – 1506.
Abed, A.M., 1982. Geology of Jordan. Al-Nahda Al-Islamiah library, Amman, Jordan. 232 pp. (in Arabic).
Abed, A.M., 2002. An overview of an inland sabkha in Jordan:
the Taba Sabkha, southern Wadi Araba. In: Barth, H.-J., Böer,
B. (Eds.), Sabkha Ecosystems. Kluwer Academic Publishing,
Netherlands, pp. 83 – 97.
Abed, A.M., Al-Hawari, Z.K., 1991. Geology and Sepiolite Formation in the Taba Continental Sabkha, Southern Dead Sea – Araba
Rift, Jordan, vol. 18B. Dirasat, University of Jordan, Amman,
pp. 41 – 65.
Abed, A.M., Yaghan, R., 2000. On the paleoclimate of Jordan
during the last glacial maximum. Paleogeoraphy, Paleoclimatology, Paleoecology 160, 22 – 33.
Abu-Zir, I.M., 1989. Hydrochemistry and Hydrogeology of the
Central Part of the Wadi Araba. Unpubl. MSc Thesis, Dept.
of Geology and Mineralogy, Univ. of Jordan, Amman.
185 pp.
Ahlbrandt, Th.S., 1979. Textural parameters of eolian deposits. In:
McKee, E. (Ed.), A Study of Global Sand Seas. USGS, Professional Paper, vol. 1052, pp. 21 – 52. Washington.
Ahlbrandt, Th.S., Fryberger, S.G., 1981. Sedimentary features and
significance of interdune deposits. In: Ethridge, F.G., Flores,
R.M. (Eds.), Recent and Ancient Nonmarine Depositional Environments: Models for Exploration. SEPM Special Publication,
vol. 31. 349 pp.
Al-Qudah, B., 2000. Soils of Jordan. Draft Report (Unpubl.). Range
Directorate, Ministry of Agriculture, Amman.
ARAMCO Survey, 1974. Projects Construction Services (Shaik
Mahboob compiler), AER-716 southern area climate, estimated
climatic information for Shaybah EPI 151-192 and EPI 169290: Arabian American Oil, Dhahran, SA. 33 pp. and appendix
34 pp.
Araya-Vergara, J.F., 1987. The evolution of the modern coastal
dune systems in central Chile. In: Gardiner, V. (Ed.), International Geomorphology, 1986. Proceedings of the 1st International Geomorphology Conference II. Wiley, Chichester,
pp. 1231 – 1239.
Atallah, M., 1986. Application of remote sensing and field techniques to tectonic problems of the Dead Sea Rift in Jordan.
Unpubl. PhD thesis, der Technischen Universität München.
154 pp.
Atallah, M., 2002. Morphotectonic indices of the Eastern Wadi
Araba (Dead Sea Rift Jordan). Geografia Fisica e Dinamica
Quaternaria 25, 3 – 10.
Bagnold, R.A., 1941. The Physics of Blown Sand and Desert
Dunes. Methuen, London. 265 pp.
Bagnold, R.A., 1953. The surface movement of blown sand in
relation to meteorology. Research Concentrated Israel, Special
Publication 2, 89 – 93.
Bagnold, R.A., Barndorff-Nielsen, O.E., 1980. The pattern of natural size distributions. Sedimentology 27, 199 – 207.
23
Bender, F., 1968. Geologie von Jordanien. Beitr. Region Geol.
Erde, vol. 7. Bornträger, Stuttgart. 230 pp.
Bender, F., 1974. Explanatory notes on the geological map of Wadi
Araba, Jordan (scale 1: 100,000, 3 sheets). Geol. Jb., Reihe B,
Heft, vol. 10. Hannover, 62 pp.
Bowen, A.S., Lindley, D., 1977. A wind tunnel investigation of the
wind speed and turbulence characteristics close to the ground
over various escarpment shapes. Boundary-Layer Meteorology
12, 259 – 272.
Breed, C.S., Fryberger, S.C., Andrews, S., McCauley, C., Lennartz,
F., Gebel, D., Hortsman, K., 1979. Regional studies of sand seas
using LANDSAT (ERTS) Imagery. In: McKee, E. (Ed.), A
Study of Global Sand Seas. Professional Paper, vol. 1052.
USGS, Washington, pp. 305 – 397.
Cooke, R.U., Warren, A., 1973. Geomorphology in Deserts. University of California Press, Beverly. 374 pp.
Cooke, R.U., Warren, A., Goudie, A.S., 1993. Desert Geomorphology. UCL Press, London. 526 pp.
Dayan, U., Ziv, B., Margalit, A., Morin, E., Sharon, D., 2001. A
severe autumn storm over the Middle East: synoptic and mesoscale convection analysis. Theoretical and Applied Climatology
69, 103 – 122.
De Martonne, E., 1925. Traité de Géographie Physique. 3 Vols.,
Paris Armand Colin, 2 volumes In-4 brochés, 1057 P.
El Baz, F., 1986. Monitoring of the sand-moving desert winds. In:
El-Baz, F., Hassan, M.H.A. (Eds.), Physics of Desertification.
Martinus Nijhoff, Dordrecht, pp. 141 – 158.
El-Baz, F., 1998. Sand accumulation and groundwater in the eastern
Sahara. Episodes 21, 147 – 151.
Farhat, J.S., 1979. High grade metamorphic complex in the
basement at Wadi Araba, Jordan. Evolution and Mineralization of the Arabian – Nubian Shield, vol. 2. Pergamon, Oxford, pp. 171 – 178.
Finkel, H.J., 1959. The barchans of southern Beru. Journal of Geology 67, 614 – 647.
Folk, R.L., 1971. Longitudinal dunes of the northwestern edge of
the Simpson Desert, Northern Territory Australia: 1. Geomorphology and grain size relationships. Sedimentology 16, 5 – 54.
Fryberger, S.G., 1979. Dune forms and wind regime. In: McKee,
E.D. (Ed.), A Study of Global Sand Seas. Professional Paper,
vol. 1052. USGS, Washington, pp. 137 – 169.
Fryberger, S.G., 1980. Dune forms and wind regime, Mauritania
West Africa: implications for past climate. Paleoecology of
Africa 12, 79 – 96.
Fryberger, S.G., Al-Sari, A.M., Clisham, Th.J., 1983. Eolian dune,
interdune, sand sheet, and siliciclastic sabkha sediments of an
offshore prograding sand sea, Dhahran area, Saudi Arabia. The
American Association of Petroleum Geologists 67, 280 – 312.
Galli, P., 1999. Active tectonics along the Wadi Araba – Jordan
Valley transform fault. Journal of Geophysical Research 104,
2777 – 2796.
Garfunkel, Z., Zak, I., Freund, R., 1981. Active faulting in the Dead
Sea Rift. Tectonophysics 80, 1 – 26.
Gasse, F., Fontes, J.C., Plaziat, J.C., Carbonet, P., Kaczmarska, I.,
De Deckker, P., Soulie-Marche, I., Callot, Y., Dupeuple, P.,
1987. Biological remains, geochemistry and stable isotopes for
the reconstruction of environmental and hydrological changes in
ARTICLE IN PRESS
24
W. Saqqa, M. Atallah / Geomorphology xx (2004) xxx–xxx
the Holocene lakes from the Northern Sahara. Palaeogeography,
Paleoclimatology, Paleoecology 60, 1 – 46.
Glennie, K.W., 1987. Desert study environments, present and
past—a summary. Sedimentary Geology 50, 136 – 165.
Goudie, A.S., Watson, A., 1981. The shape of desert sand dune
grains. Journal of Arid Environments 4, 185 – 190.
Hastenrath, S.L., 1987. The barchan dunes of southern Peru revisited. Zeitschrift für Geomorphologie 3, 167 – 178.
Hastings, J.D., 1971. Sand sheets. Meteorological Magazine 100,
155 – 159.
Hesp, P.A., 1989. A review of biological and geomorphological
processes involved in the initiation and development of incipient
fordunes. In: Girningham, C.H., Ritche, W., Willetts, B.B., Willis, A.J. (Eds.), Symposium: Coastal Sand Dunes. Proceedings
of the Royal Soc. of Edinburgh, vol. B96, pp. 181 – 201.
Hötzl, H., Zötl, J.G., 1978. Climate change during the Quaternary
period. In: Al-Sayari, S.S., Zötl, J.G. (Eds.), Quaternary Period
in Saudi Arabia. Springer, Berlin, pp. 301 – 311.
Itzigsohn, D., 1995. A climatological and dynamic analysis of the
interaction between the tropics and midlatitudes over the Red
Sea region. Unpubl. MSc Thesis, Tel Aviv Univ. In Hebrew.
Jado, A.R., Zötl, J.G., 1984. Quaternary period in Saudi Arabia,
vol. 2. Springer, Berlin. 360 pp.
Jarrar, Gh., 1979. Petrology and geochemistry of Wadi Um-Saiyala,
Wadi Araba, Jordan. Unpubl. MSc Thesis, Dept. of Geology and
Mineralogy, Univ. of Jordan, Amman. 86 pp.
Jarrar, Gh., 1984. Late Proterozoic crustal evolution of the Arabian – Nubian Shield in the Wadi Araba Area, SW Jordan. Braunschweiger Geologisch – Palaeontologishce Dissertationen 2,
Braunschweig, Germany. 107 pp.
Jones, D.K.C., Cooke, R.U., Warren, A., 1986. Geomorphological
investigation, for engineering purposes, of blowing sand and
dust hazard. Quarterly Journal of Engineering Geology 19,
251 – 270.
Jordan Meteorological Department (JMD). Meteorological data
(1955 – 1985, 1977 – 2001). Aqaba and Ghor Safi Meteorological Stations, Internal Report, Amman.
Kahana, R., Ziv, B., Enzel, Y., Dayan, U., 2002. Synoptic climatology of major floods in the Negev Desert, Israel. International
Journal of Climatology 22, 867 – 882.
Khalaf, F.I., Gharib, I.M., 1985. Roundness parameters of quartz
sand grains of recent aeolian sand deposit in Kuwait. Sedimentary Geology 45, 137 – 158.
Krauskopf, K.B., Beiser, A., 2002. The Physical Universe, 10th
edition. McGraw Hill, New York.
Lancaster, N., 1980. The formation of seif dunes from barchan—
supporting evidence for Bagnold’s model from the Namib Desert. Zeitschrift für Geomorphologie 24, 160 – 167.
Lian-You, L., Shang-Yu, G., Pei-Jun, Sh., Xiao-Yan, L., Zhi-Bao,
D., 2003. Wind tunnel measurements of adobe abrasion by
blown sand: profile characteristics in relation to wind velocity
and sand flux. Journal of Arid Environments 53, 351 – 363.
Long, G., 1957. Bioclimatology and vegetation of Eastern Jordan.
Working paper for FAQ/57/2/1109.
Mazzullo, J., Sims, D., Cunningham, D., 1986. The effects of eolian
sorting and abrasion upon the Shapes of fine quartz sand grains.
Journal of Sedimentary Petrology 56, 45 – 56.
McKee, E., 1966. Structures of dunes at White Sands National
Monument, New Mexico (and a comparison with structures of
dunes from other selected areas). Sedimentology 7, 3 – 69.
McKee, E., 1979. An introduction to the study of global sand seas.
In: McKee, E. (Ed.), A Study of Global Sand Seas. USGS,
Professional Paper, vol. 1052, pp. 1 – 20. Washington.
Membery, D.A., 1983. Low-level wind profiles during the Gulf
Shamal. Weather 38, 18 – 24.
Middleton, N.J., Thomas, D., 1997. World Atlas of Desertification,
2nd edition. Arnold Publication, London. 192 pp.
Middleton, N.J., Goudie, A.S., Wells, G.L., 1986. The frequency
and source area of dune storms. In: Nickling, W.G. (Ed.), Aeolian Geomorphology. Allen and Unwin, Boston, pp. 237 – 260.
Murai, Sh., Honda, Y., 1991. World Vegetation Map from NOAA
GVI Data. In: Murai, Sh. (Ed.), Applications of Remote Sensing
in Asia and Oceania. AARS, Geocartos International, Hong
Kong. 384 pp.
Pahari, K., Murai, Sh., 1995. Modeling for global land degradation
using remote sensing and GIS. Asian Conference on Remote
Sensing (CCRS). Poster Session, vol. 3. Suranaree University of
Technology, Nakorn Ratchasima, Thailand.
Petit-Maire, N., Guo, Z.T., 1997. Holocene paleoprecipitation over
the present-day Sahara desert. Implications for the future. Episodes 20, 232 – 234.
Powers, M.C., 1953. A new roundness scale for sedimentary particles. Journal of Sedimentary Petrology 23, 117 – 119.
Quennell, A.M., 1958. The structural and geomorphic evolution of
the Dead Sea Rift. Quarterly Journal of Geological Society 114,
1 – 24.
Quennell, A.M., 1983. The evolution of the Dead Sea Rift—a
review. Proceedings of the 1st Jordan Geological Conference.
Jordan University, Amman, pp. 462 – 482.
Sakamoto-Arnold, C.M., 1981. Eolian features produced by the
December, 1977 windstorm, southern San Joaquin Valley, California. Journal of Geology 89, 129 – 137.
Saqqa, W., 1997. Heavy mineral study of the recent eolian sands,
Wadi Araba, SW Jordan. Dirasat, Natural and Engineering Sciences 24, 148 – 156.
Saqqa, W., 1998. Surface textures of quartz grains from arid continental environments (Wadi Araba area SW Jordan). Neues
Jahrbuch für Geologie und Paläontologie. Abhandlungen 209,
135 – 144.
Seppälä, M., Lindé, K., 1978. Wind tunnel studies of ripple formation. Geografisca Annaler 60A, 29 – 40.
Sharon, D., Kutiel, H., 1986. The distribution of the rainfall intensity in Israel, its regional and seasonal variations and its climatological evaluation. Journal of Climatology 6, 277 – 291.
Sharp, R.P., 1963. Wind ripples. Journal of Geology 71, 617 – 636.
Sharp, R.P., 1964. Wind-driven sand in the Coachella Valley, California. Geological Society of America Bulletin 75, 785 – 804.
Swift, D.J.P., Molina, B.F., Jackson III, R.G., 1978. Intermittent
structure of the atmospheric boundary layer made visible by
entrained sediment: example from the Copper River delta,
Alaska. Journal of Sedimentary Petrology 48, 897 – 900.
Thomas, D.S.G., 1984. Ancient ergs of the former arid zones of
Zimbabwe, Zambia and Angola. Transactions of the Institute
Geographers 9, 75 – 88.
ARTICLE IN PRESS
W. Saqqa, M. Atallah / Geomorphology xx (2004) xxx–xxx
Thomas, D.S.G., 1986. The response diagram and ancient
desert sands—a note. Zeitschrift funr Geomorphologie 30,
363 – 369.
Thomas, D.S.G., 1987a. Discrimination of depositional environments using sedimentary characteristics in the Mega Kalahari,
Central Southern Africa. In: Frostick, L.E., Reid, I. (Eds.), Desert Sediments, Ancient and Modern. Geol. Soc., London, Spec.
Publ., vol. 35.
Thomas, D.S.G., 1987b. The roundness of aeolian quartz sand
grains. Sedimentary Geology 52, 149 – 153.
Trenberth, K.E., 1992. Climate System Modelling. Cambridge
Univ. Press, Cambridge. 788 pp.
Tsoar, H., 1982. Internal structure and surface geometry of longitudinal (seif) dunes. Journal of Sedimentary Petrology 52,
823 – 831.
Tsoar, H., 1984. The formation of seif dunes from barchans—a
discussion. Zeitschrift für Geomorphologie 28, 99 – 104.
Tucker, M.E., 1991. Sedimentary Petrology, an Introduction to the
25
Origin of Sedimentary Rocks, 2nd edition. Blackwell, Oxford.
260 pp.
United Nations Environment Program (UNEP), 1992. World Atlas
of Desertification.
Warren, A., 1988. A note on vegetation and sand movement in the
Wahiba Sands, Journal of Oman Studies, Spec. Rep.: 3. The
scientific results of the Royal Geographical Society’s Oman
Wahiba Sands Project 1985 – 1987, 251 – 255.
Wasson, R.J., Hyde, R., 1983. Factors determining desert dune
type. Nature 304, 337 – 339.
Wilson, I.G., 1971. Desert sand flow basins and a model for the
development of ergs. Geographical Journal 137, 180 – 199.
Worral, G.A., 1974. Observations of some wind-formed features
in the southern Sahara. Zeitschrift funr Geomorphologie 18,
291 – 302.
Zhuo, Z., Baoyin, Y., Petit-Maire, N., 1998. Paleoenvironments in
China during the last glacial maximum and the Holocene optimum. Episodes 21, 152 – 158.