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 Originaally published as: Förster, H.‐J., Förster, A., Oberhä
änsli, R., Stroomeyer, D. (2
2010): Lithosspheric compposition and thermal structuree of the Arab
bian Shield in
n Jordan. ‐ Teectonophysics, 481, 1‐4, 29‐37 DOI: 10.1016/j.tecto
o.2008.11.014 Tectonophysics 481 (2010) 39‐37 Lithospheric composition and thermal structure of the Arabian Shield in Jordan H.‐J. Förster a,*, A. Försterb, R. Oberhänsli a, D. Stromeyer b Institute of Earth Sciences, University of Potsdam, P.O. Box 601553, D­14415 Potsdam, Germany Deutsches GeoForschungsZentrum, Telegrafenberg, D­14473 Potsdam, Germany a
b
A R T I C L E I N F O Article history: Received 27 May 2008 Received in revised form 7 November 2008 Accepted 18 November 2008 Available online 27 November 2008 Keywords: Geotherms Surface heat flow Xenolith Peridotite Late Proterozoic terrane Arabian Shield A B S T R A C T In this paper, a unique set of samples from the uppermost crust down to the lithospheric mantle of Jordan is ana‐
lyzed for composition and petrophysical properties (density, thermal conductivity, radiogenic heat production).
These data, covering a vertical section of almost 65 km, are used in conjunction with surface heat flow to generate a detailed and comprehensive lithospheric thermal model that reflects the conditions of the Arabian Shield (AS) prior
to the post‐Oligocene onset of lithosphere thinning and voluminous basaltic volcanism. The pre‐Miocene model geotherms, based on conductive surface heat flows of 55 and 60 , (a) meet the range of lithosphere‐
asthenosphere boundary depths of 110‐160 km known from seismology, (b) conform to results of thermomechani‐
cal models on the origin of the Dead Sea basin that started in Miocene time, and (c) are consistent with typical xenolith‐derived geotherms for terranes of similar age and lithospheric thickness. Moho temperatures (at depths
between 35 and 40 km) of the AS in pre‐Miocene times were most likely in the order of 530‐650 °C, with mantle heat flows averaging between 24 and 29 . Results contradict former views of the late Proterozoic/early Cambrian‐stabilized AS being an anomalously cold terrane. A “cold” thermal structure inferred from previously measured low surface heat flows (generally ≤45 ) is inconsistent with the thickness, composition, and petrophysical properties of the stable lithosphere of the shield. © 2008 Published by Elsevier B.V.
1. Introduction Earlier studies have characterized the late Proterozoic Ara‐
bian Shield (AS) as a terrane of generally low conductive surface heat flow of 36‐45 (Gettings, 1982; Gettings and Sho­
wail, 1982; Galanis et al., 1986). The implication was that the AS constitutes a cold continental lithosphere, thermally similar to many terranes consolidated in the Archean, averaging to 41±11 (e.g., Nyblade and Pollack, 1993). This widely accepted view of a cold lithosphere was recently questioned by a heat‐
flow study in southeastern Jordan, which yielded a substantially higher surface heat flow of 60±3 (Förster et al., 2007). This higher heat flow is in satisfactorily agreement with the global heat‐flow average of stable late Proterozoic terranes far removed from Archean cratons (55±17 ; Rudnick et al., 1998). Starting from the new heat‐flow determination in southeas‐
tern Jordan, this paper is aimed to elaborate a consistent model of the thermal structure of the Arabian Shield in Jordan. The approach is multidisciplinary by comprehensively crosscheck‐
ing results from the fields of seismics, seismology, gravity, pe‐
trology, geochemistry, petrophysics, and geothermics. It ex‐
amines a unique set of samples from the uppermost crust down to the lithospheric mantle, providing a full coverage of composi‐
tional, density, thermal conductivity, and radiogenic heat pro‐
duction data over a vertical section of almost 65 km. Previous inferences on the thermal state of the AS (e.g., Mc­
Guire, 1988; McGuire and Bohannon, 1989; Stein et al., 1993; Medaris and Syada, 1998) are based on the low surface heat‐
flow concept and considered extrapolated or supposed rather than real values for petrophysical properties and radiogenic heat production, disregarding the temperature and pressure dependence of rock thermal conductivity. * Corresponding author. Tel.: +49 331 288 1242. E­mail address: forhj@gfz‐potsdam.de (H.‐J. Förster) doi: 10.1016/j.tecto.2008.11.014 2. Evolution and regional geology of the Arabian Shield The Arabian Peninsula (Plate) is composed of the (western) Arabian Shield (AS) and the (eastern) Arabian Platform (Fig. 1). The AS constitutes a region of Precambrian/early Cambrian basement that became exposed as result of uplift associated with rifting along the Red Sea. The shield contains the best‐
preserved and most widely exposed juvenile continental crust of Neoproterozoic age on Earth (e.g., Stern and Kröner, 1993; Stern and Abdelsalam, 1998), although some juvenile magmas contain older continental material of early Neoproterozoic to Archean age (Hargrove et al., 2006). Most of the AS formed during the Pan‐African orogenic cycle (~1000‐530 Ma) by suturing of juve‐
nile intra‐oceanic arc terrains followed by magmatic thickening, giving rise to regional metamorphism and the generation of batholiths of predominantly granitic to granodioritic composi‐
tion between 630 and 610 Ma (Stoeser and Camp, 1985; Cosca et al., 1999, Katz et al., 2003; Jarrar et al., 2003). Formation of the AS was likely associated with the rise of a plume head to the upper mantle, leading to the formation of an enriched “plume mantle” and oceanic plateaus (Stein, 2003). During the last stag‐
es of its early evolution (~610‐530 Ma), the shield was sub‐
jected to strong extension. This extension was accompanied by intrusion of abundant within‐plate A‐type granites, alkali gra‐
nites, and syenites, sometime associated with gabbros, by the emplacement of composite dykes, and by the eruption of bimod‐
al volcanic rocks (Beyth et al., 1994; Mushkin et al., 2003; Katzir et al., 2007; Jarrar et al., 2008). Owing to its position as a passive continental margin since early Paleozoic times, the region was the place of sediment deposition repeatedly interrupted by times of uplift and denu‐
dation. After a long period of magmatic quiescence, several minor episodes of Mesozoic intra‐plate magmatic activity have been recognized in the area around the DST (Dead Sea Trans‐
form), which are thought to be associated with the evolution of the eastern Mediterranean Neo‐Tethyan passive continental margin along the northern edge of Gondwana (e.g., Laws and Wilson, 1997; Wilson et al., 2000). The main events occurred 29 H.­J. Förster et al. / Tectonophysics 481 (2010) 29­37 try. Thorium and U were analyzed by inductively coupled plas‐
ma‐mass spectrometry (ICP‐MS). Averages and standard devia‐
tions for the three parameters given in Table 1 involve between 3 and 20 single measurements. For lower crustal rocks, density, compressional wave velocity ( ), and shear wave velocity ( ) were calculated with the algorithm of Sobolev and Babeyko (1994), which translate bulk‐rock compositions (averages of Al­
Mishwat and Nasir, 2004) into modal mineralogy as function of temperature and pressure. 4. Architecture of the lithosphere Refraction/reflection seismic soundings (El­Isa et al., 1987; Batayneh and Al Zoubi, 2001; DESERT Group et al., 2004; Mechie et al., 2005), receiver‐function analyses of teleseismic records (Hofstetter and Bock, 2004; Al­Damegh et al., 2005; Mohsen et al., 2005, 2006; Hansen et al., 2007), teleseismic tomographic inver‐
sion (Koulakov and Sobolev, 2006; Laske et al., 2008), and gravi‐
ty studies (Götze et al., 2007) demonstrate that four major dis‐
continuities determine the structure of the lithosphere beneath Jordan. These features are associated with changes in lithology and correspond to discontinuities recognized elsewhere in the AS (i.e., Nasir and Safarjalani, 2000; AlMishwat and Nasir, 2004). The first discontinuity at 19‐20 km, which is associated with a change in from 6.1‐6.3 to 6.7 and in from 3.553.65 to 3.75 (Mechie et al., 2005), corres‐
ponds to the boundary between the upper and lower crust. The second reflector, marked by a jump in from 6.7 to 7.2 , is located within the lower crust, at about 29 km. The third discontinuity, the Moho, is located at depths between 32 (im‐
mediately east of the DST) and 39 km. The crustal thickness in southern Jordan is about 37 km and, thus, closely approximates the average thickness of the AS crust, calculated between 36 km (Rodgers et al., 1999) and 39 km (Al­Damegh et al., 2005). The fourth reflector, a low‐velocity zone underlying a high‐
velocity mantle lid, marks the lithosphere‐asthenosphere boun‐
dary (LAB). There is ample evidence that the lithosphere below Jordan is tectonically active and thinned, with LAB depths vary‐
ing between about 55 and 80 km. Lithosphere thinning is ob‐
served throughout the AS along the Red Sea rift and the DST and beneath areas of Cenozoic basaltic volcanism (e.g., Julia et al., 2003; Hansen et al., 2007). Seismic stations define the LAB depth beneath the heat‐flow area to ~65 km (Mohsen et al., 2006; Hansen et al., 2007). 5. Composition of the lithosphere Surface geology and extensive geophysical data indicate that the upper crust is felsic to intermediate in composition. The bulk of the upper crust is composed of Pan‐African metamorphic and igneous rocks, which are locally overlain by a thin, up to several kilometers thick sedimentary cover. The rock types considered for fine‐tuning the model are listed in Table 1. The proportions of rock types are constrained by regional geology (relative vo‐
lume at surface exposures), gravity modeling (bulk density), and seismic soundings (low abundance of foliated rocks). Our pres‐
sure‐uncorrected mean density of the upper crust of 2.67×103 corresponds with the regional gravity model of the study area (e.g., Götze et al., 2007). Geophysical information (e.g., Rodgers et al., 1999) and xeno‐
lith data emphasize that the lower crust is two‐layered and predominantly mafic. Accordingly, it is mainly composed of two types of metamorphosed rocks of broadly noritic composition (Al­Mishwat and Nasir, 2004), i.e., plagioclase‐rich two‐pyroxene granulites (PLMG) and two‐pyroxene granulites (PYMG). Rocks of this chemical affiliation are compatible with the island‐arc Fig. 1. Schematic geological map of the western Arabian Plate, showing
the location of the heat‐flow site and the distribution of Cenozoic basal‐
tic volcanic fields. during the late Triassic to early Jurassic (‐200 Ma), the late Ju‐
rassic to early Cretaceous (146‐115 Ma), and in the late Creta‐
ceous (‐80 Ma). During Miocene to Pliocene times, mafic volcan‐
ic activity resumed throughout the region related to the complex tectonics associated with the opening of the Red Sea at around 30 Ma and the activity of the Afar mantle plume (Wilson et al., 2000). The younger event of this two‐phase continental mag‐
matism (<14 Ma) is generally attributed to lithosphere thinning owing to asthenosphere upwelling, possibly involving a mantle‐
plume component (Camp and Roobol, 1992; Shaw et al., 2003; Krienitz et al., 2007; and references therein). The regional‐scale heating of the upper mantle is well imaged by seismological data (Phillips et al., 2007), but is not yet recorded in the surface heat flow. 3. Sampling sites and analytical techniques A total of 155 samples were collected from various regions of Jordan, encompassing Paleozoic sediments, Pan‐African magmatic and metamorphic rocks, lower crustal granulites, and upper mantle spinel lherzolites (Fig. 2). Rock density, thermal conductivity, and radiogenic heat production were determined on unaltered specimens devoid of fractures. Density was meas‐
ured at standard conditions on water‐saturated samples by weighting. Thermal conductivity was measured by optical scan‐
ning (Popov et al., 1999) with the apparatus used by Norden and Förster (2006). The optical‐scanning method has a high preci‐
sion (1.5%) and accuracy (1.5%). Radiogenic heat production was calculated using the equation of Rybach (1976). The concen‐
tration of K was determined by X‐ray fluorescence spectrome‐
30 H.­J. Förster et al. / Tectonophysics 481 (2010) 29­37 Fig. 2. Simplified geological map of Jordan. The rectangles schematically enclose the areas from which the samples analyzed in this study were col‐
lected: a = upper crustal Paleozoic sedimentary rocks, b = upper crustal late Proterozoic to early Cambrian magmatic and metamorphic rocks, c =
lower crustal granulites and upper mantle peridotites. Also shown are the locations of the boreholes used for heat‐flow calculation. The broken line represents the approximate axis of the Dead Sea Transform Fault. 6. Thermal parameters of the lithosphere 6.1. Radiogenic heat production Heat production ( ) of the sedimentary veneer was deter‐
mined to 0.8 on average. The average of the upper crustal metamorphic and igneous basement is variable (Förster et al., 2007). The highest mean values of refer to alkali‐
feldspar granites (3.55 ) and rhyolites/rhyodacites (2.52 ). Quartz monzodiorites (0.92 ) and mon‐
zogabbros (0.73 ) display the lowest values of . The bulk heat production of our model upper crust of Jordan amounts to 1.48 , which is slightly less than of the average global upper crust (1.8 ) according to the model of Taylor and McLennan (1985). Granulite xenoliths from northeastern Jordan permitted cal‐
culation of first data for the Arabian lower crust east of the DST (Table 2). Accordingly, the two types of mafic granulites have virtually the same low mean value of 0.06 (cf. Table 1), corresponding to the global median of mafic granulite (Rudnick and Fountain, 1995). Heat production of the lower crust of the AS is lower than is implied from the estimated present‐day average of global model lower crust (0.18 ; Rudnick and Fountain, 1995). This discrepancy is caused by the tectonic setting inferred for the AS (e.g., McGuire and Stern, 1993). Densities, ‐ and ‐wave velocities calculated separately for PLMG and PYMG (Table 1) are in excellent agreement with geophysical constraints provided by the seismic (Mechie et al., 2005) and gravity models (Götze et al., 2007). These correspon‐
dences and the calculated temperatures and pressures of last re‐
equilibration of the xenoliths (McGuire,1988; Nasir, 1995; Al­
Mishwhat and Nasir, 2004) imply that the upper part of the lower crust is dominantly composed of PLMG, whereas PYMG±garnet granulites (GMG) constitute the lower part down to the Moho (cf. McGuire and Stern, 1993; Al­Mishwat and Nasir, 2004). These constraints result for the lower crust in a density of 3.04×103 , and for the bulk crust in a density of (uncorrected for ). This value is consistent 2.83×103 with the most recent gravity model of the Arabian Plate, imply‐
ing a mean density of the AS crust of 2.85×10‐3 (Hansen et al., 2007). Mantle xenoliths entrained in late Cenozoic basalts provide direct probes of the composition of the lithospheric mantle below Jordan and elsewhere in the AS. Accordingly, the shallow upper mantle is composed of peridotites (spinel lherzolites and harzburgites) and minor pyroxenites, encompassing garnet, spinel, and plagioclase websterites (e.g., Nasir, 1992; Medaris and Syada, 1998; Nasir and Safarjalani, 2000; Shaw et al., 2007). 31 H.­J. Förster et al. / Tectonophysics 481 (2010) 29­37 Table 1 Structure, composition, and petrophysical properties of the lithosphere in southeastern Jordan Age Depth Rock type (km) Upper crust Lower crust 20 29 Paleozoic Cambrium‐late Proterozoic Late Proterozoic Lith. mantle 37 65 Sandstone/siltstone Metamorphic rocks Rhyolite/rhyodacite Alkali‐feldspar granite Monzogranite Granodiorite/qtz‐monzonite Qtz‐monzodiorite Monzogabbro Plag‐rich two‐pyroxene mafic granulite Two‐pyroxene mafic granulite Spinel lherzolite % 10 15 5 5 15 30 10 10 50 50 Density ( 10
2.42±0.11 2.71±0.06 2.63±0.02 2.58±0.01 2.61±0.02 2.69±0.02 2.77±0.04 2.83±0.01 2.95a (6.91, 3.84)a 3.15a (7.28, 4.09)a 3.36±0.02 ) Proport. Thermal Conductivity conductivity (
) Proport. Heat production ( ) Proport. 0.24 0.41 0.13 0.13 0.39 0.81 0.28 0.28 ∑ 2.67 1.47 3.8±0.50 2.79±0.13 2.63±0.21 2.93±0.07 2.60±0.14 2.30±0.09 2.17±0.02 2.04±0.02 2.83±0.12 0.38 0.42 0.13 0.15 0.39 0.69 0.22 0.20 ∑ 2.58 1.41 0.80±0.15 1.19±0.17 2.52±0.93 3.55±1.55 1.47±0.68 1.73±0.40 0.95±0.17 0.73±0.06 0.06±0.03 0.08 0.18 0.13 0.18 0.22 0.52 0.10 0.07 ∑ 1.48 0.03 1.57 3.01±0.15 1.50 0.06±0.02 0.03 ∑ 3.04 ∑ 2.91 4.5b 0.011±0.002c 0.025±0.009d 0.036±0.005e ∑ 0.06 0.02f a and ( , ) calculated for 600 °C and 1 GPa according to the formalism of Sobolev and Babeyko (1994), using the bulk‐rock averages of Al­Mishwat and Nasir
(2004). b After Kukkonen & Peltonen (1999) and Tommasi et al. (2001). c NE Jordan, this study. d NE Jordan (including pyroxenites), Shaw (2003). e Dead Sea area, this study. f Mean of the lithospheric mantle (see text). bulk chemical composition of the lower crust beneath the AS, which is more mafic than average lower crust (Al­Mishwat and Nasir, 2004). On the other hand, our data are fully compatible with recent xenolith‐derived values for chemically similar lower crustal rocks worldwide, in particular those from the well‐investigated Sino‐Korean tectonic blocks (0.03‐0.12 (0.20) ; Liu et al., 2001; Yu et al., 2003; Huang et al., 2004; Dai et al., 2008). U‐Th data for garnet‐bearing mafic granulites from Arabia are available from N Israel (Gazit, 2005), indicating heat‐
production values maximizing to 0.02 . Values of obtained from xenoliths of comparable composition from else‐
where tend to be similar or slightly lower (0.002‐0.10 ; e.g., Yu et al., 2003; Garrido et al., 2006) compared to garnet‐free varieties. However, ambiguities in this parameter have no im‐
pact on the results of the thermal modeling. Heat production of our Jordan model bulk crust totals to 0.83 , i.e., is in harmony with reliable global bulk‐crust estimates, ranging (Rudnick et al., 1998). between 0.70 and 0.93 For the lithospheric mantle beneath the AS, xenoliths from Jordan are yet the sole source of information on . Spinel‐
lherzolite xenoliths examined in this study display values between 0.011 and 0.036 (Table 2). This range in A is in excellent agreement with heat‐production data calcu‐
lated from bulk‐rock chemical analyses (Shaw, 2003) of both peridotites and pyroxenites from northeastern Jordan (0.010‐
0.038 ). Heat production values typical for lithospheric mantle rocks are in the range 0.001 to 0.05 (e.g., Lenoir et al., 2000; Beccaluva et al., 2001; Ionov et al., 2005; Xu et al., 2008; Bodinier et al., 2008). For the mantle, the mean of up‐
permantle xenoliths (0.02 ) from Jordan is implemented in our model, which corresponds to the pre‐
ferred estimate of in the cratonic mantle (0.019 ; Rudnick et al., 1998). No indication exists for the presence of a metasomatic, high‐heat‐production layer in the lithospheric mantle of the AS, which may display of up to 0.4 (Neves et al., 2008). 6.2. Thermal conductivity Thermal conductivity ( ) of the quartz‐dominated, clastic Paleo‐
zoic sedimentary rocks in southeastern Jordan averages to 3.8 (Förster et al., 2007). The upper crustal igneous and metamorphic basement displays a range in between ~2.9 in the most quartz‐rich alkali‐feldspar granites and in the most quartz‐poor monzogabbros. The ~2.0 mean of the upper crust amounts to 2.58 . Thermal conductivity for the lower crustal PLMG approximates to 2.8 ; that for the PYMG is only insignificantly higher (3.0 ). The bulk of the 37‐km‐thick crustal segment yields 2.73 . A thermal‐conductivity value of 4.5 is assigned to the lithospheric mantle, in correspon‐
dence to the well‐constrained global average of spinel lherzolite (e.g., Kukkonen and Peltonen, 1999; Tommasi et al., 2001). 6.3. Surface heat flow and lithospheric temperatures Our study considers the average surface heat flow ( ) of recently obtained in southeastern Jordan 60±3 (Förster et al., 2007). This average refers to five boreholes pene‐
trating consolidated Paleozoic sediments (cf. Fig. 2). The bore‐
holes are 550‐950 m deep, free of artesian flow, and in thermal equilibrium. Heat flow was determined for discrete depth inter‐
vals from temperature gradient and thermal conductivity meas‐
ured on water‐saturated samples. For each borehole, the inter‐
val heat‐flow values were finally converted into a heat‐flow average. The average of the five boreholes ranged from 56 to 66 . The determined by Förster et al. (2007) is at the upper end of range previously measured in Jordan (42‐65 ; Galanis et al., 1986). However, it deviates significantly from values reported for the Saudi Arabian part of the shield, outside of the Red Sea rift (36‐45 ; Gettings, 1982; Gettings and Showail, 1982). This discrepancy is probably related to the poor data quality that characterizes the heatflow estimates from Saudi Arabia. Problems encountered comprise (a) the difficulty of relating the thermal conductivity from drill cuttings to depth and to the sedimentary formations in the boreholes, (b) the measurement of thermal conductivity on dry samples, (c) the use of small depth intervals for the determination of interval temperature gradients, and (d) the use of shallow boreholes (usually ≤70 m deep), in which erroneously low temperature gradients were recorded, reflecting perturbed shallow thermal conditions (e.g., Förster et al., 2007). The higher used in this study is unlikely being related to a local positive anomaly, caused by heat refraction associated with high‐heat‐production (HHP) rocks underlying the sedi‐
ments. HHP rates of 4–6.8 are confined to rhyolites and alkali‐feldspar granites that form aplitic dikes or minor intru‐
32 H.­J. Förster et al. / Tectonophysics 481 (2010) 29­37 Table 2 Chemical composition of mafic granulites (lower crust) and spinel lherzolites (lithospheric mantle) from Jordan Origin Region Sample number SiO2 (wt.%) TiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K2O P2O5 H2O+ CO2 Total Li (ppm) Sc V Cr Co Ni Cu Ge Zn Ga Rb Sr Y Zr Nb Mo Cd Sn Sb Cs Ba La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Hf Ta Tl Pb Bi Th U A Lower Crust NE Jordan DP‐G13 50.42 1.11 21.73 5.98 0.07 3.71 9.88 4.73 0.43 0.27 0.96 0.20 99.48 5.5 11 85 21 23 28 19 1.1 37 13 1.6 2210 7.0 20.1 2.0 0.01 0.05 0.26 0.07 0.01 382 8.51 17.0 2.24 9.98 2.10 1.52 2.08 0.27 1.51 0.27 0.69 0.09 0.48 0.07 0.58 0.15 <0.01 0.79 0.01 0.06 0.06 0.06 DP‐G2 48.03 2.54 17.61 10.37 0.16 5.79 9.34 3.56 0.50 0.08 1.00 0.99 99.98 6.6 27 258 63 37 37 39 1.6 59 19 3.0 449 4.1 14.8 2.9 0.15 0.21 0.51 0.06 0.06 170 2.68 5.53 0.64 2.97 0.72 0.97 0.83 0.12 0.77 0.15 0.44 0.05 0.40 0.06 0.43 0.17 <0.01 1.6 0.01 0.19 0.09 0.09 DP‐G3 53.89 0.56 18.31 7.11 0.12 5.38 8.56 4.26 0.32 0.19 0.79 0.25 99.73 4.4 18 140 189 24 41 12 1.3 71 20 0.90 526 9.0 10.9 0.98 0.02 0.12 0.33 0.04 0.02 127 4.22 8.40 1.32 6.66 1.78 1.03 1.86 0.28 1.78 0.35 1.00 0.13 0.91 0.13 0.44 0.08 <0.01 2.3 0.03 0.06 0.07 0.05 DP‐G21 54.36 0.26 18.50 7.69 0.13 5.74 7.90 4.26 0.27 0.03 0.79 0.18 100.10 6.5 18 111 167 29 49 18 1.4 74 21 1.5 458 4.1 8.7 0.62 0.05 0.07 0.49 0.06 0.02 83 1.93 4.22 0.51 2.35 0.68 0.53 0.80 0.12 0.81 0.17 0.49 0.07 0.49 0.07 0.37 0.05 0.01 1.1 <0.01 0.05 0.05 0.04 DP‐G8 45.67 1.97 17.35 13.34 0.19 8.08 9.19 3.18 0.43 0.12 0.46 0.08 100.06 5.1 20 238 <10 44 51 18 1.4 105 17 2.4 865 15.3 34.5 1.7 0.01 0.09 0.47 0.04 0.04 420 6.75 16.1 2.47 12.8 3.35 1.76 3.69 0.54 3.24 0.62 1.62 0.21 1.32 0.19 1.21 0.17 <0.01 0.48 0.02 0.02 0.04 0.03 DP‐G5 50.28 0.25 21.06 5.34 0.09 10.24 9.10 2.18 0.09 0.04 0.97 0.19 99.82 4.4 12 76 221 42 86 36 1.3 29 14 0.77 585 1.4 10.6 1.6 0.53 0.04 0.19 0.04 0.02 53 2.07 4.99 0.43 1.66 0.34 0.31 0.33 0.05 0.30 0.06 0.17 0.02 0.16 0.02 0.28 0.10 <0.01 0.51 <0.01 0.10 0.04 0.03 DP‐G4 47.88 0.47 14.52 6.92 0.14 11.40 15.05 1.72 0.06 0.07 0.85 0.90 99.99 4.7 41 164 965 51 180 27 1.6 26 13 0.66 332 7.8 16.4 0.46 <0.01 0.12 0.35 0.05 0.02 50 1.68 4.26 0.70 3.79 1.26 0.67 1.61 0.25 1.58 0.31 0.84 0.11 0.70 0.10 0.65 0.06 <0.01 0.55 <0.01 0.09 0.05 0.05 Lithospheric mantle NE Jordan DP‐G 15 DP‐G 17 43.22 42.76 0.02 0.02 1.33 1.41 8.93 9.11 0.13 0.13 43.76 43.99 0.91 0.83 0.10 0.08 0.02 0.02 0.04 0.04 0.49 0.37 0.17 0.11 99.10 98.87 0.47 0.29 5.3 5.1 37 40 4120 5190 112 116 3900 4110 5.5 6.3 0.93 0.74 51 58 1.5 1.5 0.52 0.74 21 20 0.59 0.42 6.9 5.1 0.93 0.74 1.7 2.5 0.02 0.03 0.09 0.07 0.13 0.04 0.01 0.01 6.2 5.4 1.08 0.91 2.16 1.80 0.23 0.19 0.79 0.62 0.11 0.09 0.04 0.03 0.10 0.07 0.02 0.01 0.08 0.07 0.02 0.02 0.05 0.04 0.01 0.01 0.06 0.05 0.01 0.01 0.10 0.05 0.08 0.04 0.01 0.01 0.16 0.29 0.01 0.01 0.05 0.04 0.02 0.01 0.01 0.09 Dead Sea area DP‐38aI(II) 40.33 0.08 1.38 9.24 0.13 40.75 4.10 0.93 0.14 0.04 0.81 2.16 99.34 0.54 6.6 42 1470 132 3380 8.4 0.79 39 1.4 0.85 35 1.0 4.9 1.4 0.08 0.03 0.04 0.06 0.01 14 1.00 1.83 0.21 0.79 0.14 0.05 0.16 0.02 0.17 0.04 0.11 0.01 0.11 0.02 0.10 0.45 0.01 0.28 <0.01 0.05 0.07 0.04 DP‐38aI(III) 41.89 0.08 1.26 8.79 0.13 42.33 3.24 0.22 0.15 0.04 0.39 1.70 99.46 0.74 6.8 43 2760 135 3320 7.7 0.97 62 1.2 1.1 27 0.69 7.4 1.2 0.04 0.02 0.07 0.06 0.01 8.3 0.41 1.08 0.17 0.85 0.23 0.08 0.23 0.03 0.17 0.02 0.07 0.01 0.06 0.01 0.13 0.53 0.01 0.23 <0.01 0.03 0.04 0.03 A – radiogenic heat production (µW m‐3). surface. Thus, it is reasonable to assume that the surface heat flow determined for southeastern Jordan represents the stable thermal state of the shield before lithosphere thinning and sub‐
sequent thermal perturbations. 7. Thermal model Calculation of steady‐state geotherms is based on a three‐
layered lithosphere and the parameters given in bold in Table 1. The heat production distribution is applied as a step model. Thermal conductivity is corrected for temperature and pres‐
sure. Temperature dependency of considers the experimental results of Seipold (1998), using the following relations fitted for single rock types (Seipold, 2001): Upper crust: 1/ 0.156 5.45 10
0.763 10
(granites) 1/ 0.191 5.25 10
0.670 10
(gneisses) 1/ 0.315 3.04 10
0.326 10
(amphibolites) Lower crust: 1/ 0.344 3.27 10
0.445 10
(mafic granulites) Mantle: / 42.9 0.389
7.20 10
(peridotites) sions of only several square kilometers of surface exposure. The heat‐flow site of Förster et al. (2007) encloses an area of about 150 km2, thus significantly exceeding the dimension of HHP rocks. Ample evidence exists for presently elevated temperatures in the deep lithosphere in Jordan and elsewhere in the AS, where the LAB is uprised. Mantle xenoliths entrained in post‐Miocene basalts record lithospheric mantle temperatures clustering between 950 and 1150 °C (Fig. 3; Al­Mishwat and Nasir, 2004; Kaliwoda et al., 2007, and references therein). Moreover, xeno‐
lith‐derived probably still significantly underestimate the mantle temperatures, considering that temperatures at 90–
100 km (in the asthenosphere), inferred from the composition of young basalts from northeastern Jordan/southern Syria (Har‐
rat Ash Shamah, cf. Fig. 1) may already approach 1500 °C (Krie­
nitz et al., 2007). Basalt melting in this area proposed to have occurred shallower, near the LAB (Shaw et al., 2003), require a minimum of ~1300 °C. Active mantle upwelling, contemporaneous with crustal up‐
lift, is constrained to have started in the middle Miocene (15‐12 Ma, Camp and Roobol, 1992; Ilani et al., 2001). Taking into ac‐
count the time/length scale of thermal diffusion through the lithosphere (e.g., Turcotte and Schubert, 2002), heat flow asso‐
ciated with this young event should not have yet reached the 33 H.­J. Förster et al. / Tectonophysics 481 (2010) 29­37 not necessarily true for the xenoliths from the lower crust. The apparent inconsistency of granulite ‐ data with the „present‐
day“ pressures and temperatures in the lower crust has been demonstrated elsewhere for xenoliths in Cenozoic basalts (e.g., Embey­Isztin et al., 2003; Yu et al., 2003). Instead, we suggest that the ‐ data recorded in the lower‐crustal xenoliths reflect frozen‐in mineral equilibria, portraying the conditions of granu‐
lite‐facies metamorphism in Pan‐African time. In this instance, this geotherm would be geologically meaningless. 8. Discussion Several constraints exist permitting to evaluate the validity of the thermal models, and specifically, to infer those geotherms most accurately reflecting the past thermal structure (>15 Ma) in Jordan and elsewhere in the AS. 8.1. Depth of the LAB before lithosphere thinning Stable regions of the shield not affected by upper‐mantle upwelling display LAB depths between about 110 km (Angus et al., 2006) and 160 km (Mohsen et al., 2006; Hansen et al., 2007). This depth range corresponds to the mean thickness of conti‐
nental lithosphere in middle and late Proterozoic cratons (140±50 km; Artemieva and Mooney, 2001). Both geotherms constructed from the 55 and 60 surface heat‐flow scenarios meet these estimates, whereas a of 65 appears to slightly underestimate the LAB depth (e.g., Table 3). The 45 geotherm is incompatible with either the past and present lithospheric thickness of the AS. 8.2. Focal depth of earthquakes Temperature and lithology are among the major factors con‐
trolling the thickness of the seismogenic crust. Unambiguous evidence exists that the present‐day LAB of the AS in the region around the Dead Sea is shallow (65‐75 km, Mohsen et al., 2006). Consequently, lithospheric temperatures are elevated, and exceeding 900 °C may exist beneath the Moho located at ~33 km depth. Heating associated with post‐Oligocene upwelling of the mantle has certainly affected the lower crust, perhaps moving the brittle/ductile transition zone only a few kilometers upward compared to its position prior to the onset of lithosphere thin‐
ning. If so, the focal depths of earthquakes can still be consi‐
dered for validation of the crustal section of our pre‐Miocene geotherms. Feldspar‐dominated rocks, typical for the lower crust, de‐
form at temperatures between 380 °C and 650 °C (Tullis and Yund, 1985). Our 45 heat‐flow scenario predicts the minimum temperature for the brittle‐ductile transition at a depth of 44 km, i.e., in the upper mantle. All three scenarios of higher (55, 60, and 65 ) predict the generation of micro‐seismicity in the transition zone between the upper and lower crust, at minimum depths between 22 and 16 km, respec‐
tively. These depths are slightly lower than the focal depths for the Dead Sea basin calculated by Ten Brink et al. (2006), imply‐
ing earthquake clustering in the upper crust (between 8‐10 and 12‐16 km). The shallower observed focal depths relative to what our geotherms imply may be related to the previously discussed slight uprise of the brittle/ductile transition zone, triggered by Fig. 3. Model geotherms calculated for surface heat flows of 45, 55, 60,
and 65 , respectively, using the model parameters listed in
Table 1 (see text). Shown are the depths of major discontinuities rele‐
vant to the heat‐flow site in southeastern Jordan. The hatched area
represents the approximate ‐ range inferred from upper mantle
xenoliths entrained in post‐Miocene basaltic rocks from Jordan and
elsewhere in the AS, indicating thermal perturbations associated with
asthenosphere upwelling and lithosphere thinning. The field shown in
light gray encompasses the depth range known for the LAB in stable
portions of the shield. Abbreviations: UC ‐ upper crust, LC ‐ lower crust, LM ‐ lithospheric mantle, AM ‐ asthenospheric mantle, PLMG ‐ plagioc‐
lase‐rich mafic granulites, PYMG ‐ two‐pyroxene mafic granulites, GMG ‐
garnet‐bearing mafic granulites (cf. Table 1). The mean xenolith‐derived geotherm constructed for terranes of 100‐140 km lithospheric thickness
and tectonothermal ages <1 GPa (O'Reilly and Griffin, 2006) is shown as
dotted line. Also displayed is the approximate post‐Miocene xenolith's
geotherm of the AS (dashed line) reported by Al­Mishwat and Nasir
(2004). For the upper crust, an average fitting relation is applied, de‐
scribing this layer as a mixture of 60% granite, 20% gneiss, and 20% amphibolite (e.g., Table 1). Pressure dependence of is included in the model using the correction factor (1
). The values are 0.11 GPa‐1 for upper crust (derived from single‐
rock‐type values of Seipold, 2001), 0.03 GPa‐1 for lower crust (average of mafic granulite; Kukkonen et al., 1999), and 0.04 GPa‐1 for mantle (peridotite values of Katsura, 1995). The cor‐
rection was made up to of 1273 K and of 1.2 GPa, corres‐
ponding to the approximate upper limits of laboratory control. Four different geotherm scenarios are investigated reflecting the thermal conditions at pre‐Miocene time (Fig. 3). The scena‐
rios differ in : 45 (reflecting the former view of a cold Pan‐African lithosphere) and 55, 60, and 65 (the approximate range obtained for southeastern Jordan; Förster et al., 2007). Predicted depths of the brittle/ductile transition in the crust, Moho temperatures, LAB depths (corresponding to the depth, where the ‐ array intersects the mantle adiabat at ~1320 °C), and mantle heat flows ( ) are listed in Table 3. Geotherms not accounting for the and dependency of (not shown) predict lower Moho temperatures and greater depths for the brittle‐ductile transition and the LAB. Fig. 3 shows for comparison the approximated post‐Miocene xenolith's geotherm of the AS constructed by Al­Mishwat and Nasir (2004). The justification to elaborate this type of geo‐
therm, connecting ‐ data obtained from lower‐crustal and mantle rocks, is questioned. Whereas the ‐ information rec‐
orded in the mantle xenoliths is generally thought to reflect the temperatures and pressures at the time of entrapment, this is Table 3 Results of thermal modeling Surface heat flow (
) Brittle‐ductile transition (at ~400 °C) (km) Moho temperature (°C)
LAB depth (km)
Mantle heat flow (
) 34 45 44 375 ~300 14 55
22
550
150
24
60
18
630
120
29
65
16
720
95
34
H.­J. Förster et al. / Tectonophysics 481 (2010) 29­37 lith‐based geotherms based on model‐independent temperature markers constructed for terranes that have lithospheric thick‐
nesses between 100 and 140 km and experienced their last major tectonothermal event less than 1 Ga ago (O'Reilly and Griffin, 2006), is quasi‐identical with our 60 geotherm in the high‐ range (cf. Fig. 3). Moreover, this geotherm also accommodates the lithospheric temperatures of 1100‐1200 °C at 100 km depth predicted by Artemieva and Mooney (2001) in their global study of the thermal thickness of Precambrian cra‐
tons. The lithospheric model established in this paper produces a crustal heat flow of 31 . Removing the heat flow gener‐
ated in the crust from the three values, the pre‐Miocene man‐
tle (Moho) heat flow ( ) amounts to 24, 29, and 34 , respectively (e.g., Table 3). The lowest value, representing the 55 geotherm, is closest to the average range ob‐
tained for late Proterozoic orogens (20‐25 ; Rudnick et al.,1998). A of 65 however, implies a mantle heat flow that would be atypically high for this type of stabilized lithosphere. Altogether, we conclude that the thermal structure of the Pan‐African consolidated Arabian Shield in pre‐Miocene time is imaged most reliably by the 55 and 60 geotherms. Accordingly, we expect Moho temperatures as high as 530‐
650 °C (for Moho depths ranging from 35 to 40 km) in those regions of the shield that were not affected by lithosphere thin‐
ning. Acknowledgements This work was performed within the multinational and in‐
terdisciplinary projects DESERT and DESIRE, in part funded by the German Science Foundation, which unified scientists from Germany, Jordan, Israel, and Palestine. We want to thank espe‐
cially M. Weber, S. V. Sobolev, A. Petrunin, and Z. Garfunkel for their continuous interest in this study and valuable discussions. S. Nasir kindly provided lower‐crustal and mantle xenoliths from the Harrat Ash Shamah volcanic field in northeastern Jor‐
dan, and J.E. Shaw provided a copy of her Ph.D thesis. R. Nau‐
mann, P. Dulski, and M. Zimmer conducted the whole‐rock ana‐
lyses. Constructive reviews of A.J. Rodgers and an anonymous referee helped to strengthen the message of this paper. References recent geodynamic processes. Our geotherms based on 55 contradict earlier estimates of focal depths made by Aldersons et al. (2003), ac‐
cording to which the bulk of micro‐earthquakes in the area of the Dead Sea basin was generated at depths between 20 and 32 km, i.e., in the lower crust. These authors determined the brit‐
tle/ductile transition (at 380 °C) at a depth of 31 km, assuming a of 40 and a quartz‐depleted lower crust. Incidental‐
ly, the elevated temperatures from lithosphere thinning make significant present‐day seismic activities in the (deeper) lower crust and the lithospheric mantle unlikely, if no abnormally high rigidity is assumed, for which there is no evidence. 8.3. Formation of the DST Three‐dimensional thermo‐mechanical modeling of pull‐
apart basin formation along a transform fault (Petrunin and Sobolev, 2006) has shown that for a given strike‐slip displace‐
ment and friction of faults, the major parameter that controls basin length, thickness of sediments, and deformation pattern beneath a basin is the thickness of the brittle layer. For the Dead Sea basin case, a thick (20‐22 km, up to 27 km locally) brittle part of the cold lithosphere is required. It is also inferred that a pull‐apart basin may only form if a several‐kilometer‐thick ductile detachment zone exists between the brittle crust and the upper mantle. This all would not be the case for the Dead Sea basin if were as low as 40 as previously assumed and used in thermal modeling (e.g., Aldersons et al., 2003; and references therein). A second strong argument for a higher pre‐Miocene steady‐
state value beneath the AS was supplied by thermo‐
mechanical modeling of the origin of the Arava segment of the DST (Sobolev et al., 2005). Here, the initial conditions for repro‐
ducing the present‐day conditions at this plate boundary re‐
quired a from the Mediterranean margin eastward thickening lithosphere, whose temperatures reflect a steady‐state conduc‐
tive geotherm of at least 50‐60 . 9. Conclusions and implications Although the depths of the Moho and the LAB are variable and the composition of the upper crust may vary from place to place, the AS constitutes a plate, which is relatively uniform in structure and composition of the lithosphere (cf. Al­Mishwat and Nasir, 2004; and references therein). Taking this into considera‐
tion, we suggest that the general implications derived from our thermal model for southeastern Jordan generally hold for the entire AS. Certainly, the of the shield is variable to some ex‐
tent, mainly reflecting the differences in composition of the upper crust. For instance, it will be higher in areas of more vo‐
luminous and radiogenic monzogranites and alkali‐feldspar granites in the uppermost crust, and comparatively lower where more mafic igneous or metamorphic rocks predominate or the sedimentary cover is thicker. Likewise, the Moho temperatures vary slightly, primarily depending on the thickness of the crust. Most importantly, our thermal models contradict the pre‐
vious view of the Arabian Shield as an anomalously cold terrain, characterized by a low of generally ≤45 . A surface heat flow on this order implies large lithospheric thicknesses (>280 km) and low temperature conditions, which typify Arc‐
hean to early Proterozoic cratons mostly from the Northern Hemisphere (e.g., Jaupart and Mareschal, 1999; Artemieva and Mooney, 2001), but are opposed to what has been established for the stable AS. The 55 and 60 model geotherms appear the most reliable. Both correctly reflect the range of lithospheric thick‐
ness in stable parts of the AS of 110 to 160 km. The typical xeno‐
Al‐Damegh, K., Sandvol, E., Barazangi, M., 2005. Crustal structure of the Arabian plate: new constraints from the analysis of teleseismic receiver functions. Earth Planet. Sci. Lett. 231, 177‐196. Aldersons, F., Ben‐Avraham, Z., Hofstetter, A., Kissling, E., Al‐Yazjeen, T., 2003. Lowercrustal strength under the Dead Sea basin from local earthquake data and rheological modeling. Earth Planet. Sci. Lett. 214, 129‐142. Al‐Mishwat, A.T., Nasir, S.J., 2004. Composition of the lower crust of the Arabian Plate: a xenolith perspective. Lithos 72, 45‐72. Angus, D.A., Wilson, D.C., Sandvol, E., Ni, J.F., 2006. Lithospheric structure of the Arabian and Eurasian collision zone in eastern Turkey from S‐wave receiver functions. Geophys. J. Int. 166, 1335‐1346. Artemieva, I.M., Mooney, W.D., 2001. Thermal thickness and evolution of Precam‐
brian lithosphere: a global study. J. Geophys. Res. B106, 16,387‐16,414. Batayneh, A.T., Al Zoubi, A.S., 2001. The gravity field and crustal structure of the northwestern Arabian Platform in Jordan. J. Afr. Earth Sci. 32, 141‐148. Beccaluva, L., Bianchi, G., Coltorti, M., Perkins, W.T., Siena, F., Vaccaio, C., Wilson, M., 2001. Multistage evolution of the European lithospheric mantle: new evidence from Sardinian peridotite xenoliths. Contrib. Mineral. Petrol. 142, 284‐297. Beyth, M., Stern, R.J., Altherr, R., Kröner, A., 1994. The Late Precambrian Timna igneous complex, Southern Israel: Evidence for comagmatic‐type sanukitoid monzodiorite and alkali granite magma. Lithos 31, 103‐124. Bodinier, J.‐L., Garrido, C.J., Chanefo, I., Bruguier, O., Gervilla, F., 2008. Origin of pyroxenite‐peridotite veined mantle by refertilization reactions: evidence from the Ronda peridotite (southern Spain). J. Petrol. 49, 999‐1025. Cosca, M.A., Shimron, A., Caby, R., 1999. Late Precambrian metamorphism and cooling of the Arabian‐Nubian Shield: petrology and 40Ar/39Ar geochronolo‐
gy of metamorphic rocks of the Elat area (southern Israel). Precambr. Res. 98, 107‐127. 35 H.­J. Förster et al. / Tectonophysics 481 (2010) 29­37 and A‐type silicic magmas from composite dykes in a bimodal suite of south‐
ern Israel, northernmost Arabian‐Nubian Shield: geochemical and isotope constraints. Lithos 97, 336‐364. Koulakov, I., Sobolev, S.V., 2006. Moho depth and three‐dimensional P and S struc‐
ture of the crust and the uppermost mantle in the Eastern Mediterranean and Middle East derived from tomographic inversion of local ISC data. Geophys. J. Int. 164, 218‐235. Krienitz, M.‐S., Haase, K.M., Mezger, K., Shaikh‐Mashail, M.A., 2007. Magma genesis and mantle dynamics at the Harrat Ash Shamah volcanic field (southern Sy‐
ria). J. Petrol. 48, 1513‐1542. Kukkonen, I.T., Peltonen, P., 1999. Xenolith‐controlled geotherm for the central Fennoscandian Shield: implications for lithosphere‐asthenosphere relations. Tectonophysics 304, 301‐315. Kukkonen, I.T., Jokinen, J., Seipold, U., 1999. Temperature and pressure depen‐
dence of thermal transport properties of rocks: Implications for uncertainties in thermal lithosphere models and new laboratory measurements of high‐
grade rocks in the central Fennoscandian Shield. Surveys in Geophysics 20, 33‐59. Laske, G., Weber, M., DESERT Working Group, 2008. Lithosphere structure across the Dead Sea Transform as constrained by Rayleigh waves observed during the DESERT experiment. Geophys. J. Int. 173, 593‐610. Laws, E.D., Wilson, M., 1997. Tectonics and magmatism associated with Mesozoic passive continental margin development in the Middle East. J. Geol. Soc. Lon‐
don 154, 459‐464. Lenoir, X., Garrido, C.J., Bodinier, J.‐L., Dautria, J.‐M., 2000. Contrasting lithospheric mantle domains beneath the Massif Central (France) revealed by geochemi‐
stry of peridotite xenoliths. Earth Planet. Sci. Lett. 181, 359‐375. Liu, Y.‐S., Gao, S., Jin, S.‐Y., Hu, S.‐H., Sun, M., Zhao, Z.‐B., Feng, J.‐L., 2001. Geochemi‐
stry of lower crustal xenoliths from Neogene Hannuoba basalt, North China Craton: implications for petrogenesis and lower crustal composition. Geochim. Cosmochim. Acta 65, 2589‐2604. McGuire, A.V., 1988. The mantle beneath the Red Sea margin: xenoliths from western Saudi Arabia. Tectonophysics 150, 101‐115. McGuire, A.V., Bohannon, R.G., 1989. Timing of mantle upwelling: evidence for a passive origin for the Red Sea Rift. J. Geophys. Res. B94, 1677‐1682. McGuire, A.V., Stern, R.J.,1993. Granulite xenoliths from western Saudi Arabia: the lower crust of the late Precambrian Arabian‐Nubian Shield. Contrib. Mineral. Petrol. 114, 395‐408. Mechie, J., Abu‐Ayyash, K., Ben‐Avraham, Z., El‐Kelani, R., Mohsen, A., Rümpker, G., Saul, J., Weber, M., 2005. Crustal shear velocity structure across the Dead Sea Transform from 2‐D modelling of project DESERT explosion seismic data, Geophys. J. Int.160, 910‐924. Medaris, L.G., Syada, G.,1998. Spinel peridotite xenoliths from the Al Ashaer volca‐
no, Syria: a contribution to the elemental composition and thermal state of subcontinental Arabian lithosphere. Inter. Geol. Rev. 40, 305‐324. Mohsen, A., Hofstetter, R., Bock, G., Kind, R., Weber, M., Wylegalla, K., DESERT Group, 2005. A receiver function study across the Dead Sea Transform. Geo‐
phys. J. Int. 160, 948‐960. Mohsen, A., Kind, R., Sobolev, S.V., Weber, M., DESERT Group, 2006. Thickness of the lithosphere east of the Dead Sea Transform. Geophys. J. Int. 167, 845‐852. Mushkin, A., Navon, O., Halicz, L., Hartmann, G., Stein, M., 2003. The petrogenesis of A type magmas from the Amram Massif, southern Israel. J. Petrol. 44, 815‐832. Nasir, S., 1992. The lithosphere beneath the north‐western part of the Arabian plate (Jordan): evidence from xenoliths and geophysics. Tectonophysics 200, 357‐370. Nasir, S., 1995. Mafic lower crustal xenoliths from the northwestern part of the Arabian Plate. Eur. J. Mineral. 7, 217‐230. Nasir, S., Safarjalani, A., 2000. Lithospheric petrology beneath the northern part of the Arabian Plate in Syria: evidence from xenoliths in alkali basalts. J. Afr. Earth Sci. 30, 149‐168. Neves, S.P., Tommasi, A., Vauchez, A., Hassani, R., 2008. Intraplate continental deformation: influence of a heat‐producing layer in the lithospheric mantle. Tectonophysics 274, 392‐400. Norden, B., Förster, A., 2006. Thermal conductivity and radiogenic heat production of sedimentary and magmatic rocks in the Northeast German Basin. AAPG Bull. 90, 6, 939‐962. Nyblade, A.N., Pollack, H.N., 1993. A global analysis of heat flow from Precambrian terranes: implications for the thermal structure of Archean and Proterozoic li‐
thosphere. J. Geophys. Res. B98, 12,207‐12,218. O'Reilly, S.Y., Griffin, W.L., 2006. Imaging global chemical and thermal heterogenei‐
ty in the subcontinental lithospheric mantle with garnets and xenoliths: geo‐
physical implications. Tectonophysics 416, 289‐309. Petrunin, A., Sobolev, S.V., 2006. What controls thickness of sediments and lithos‐
pheric deformation at a pull‐apart basin? Geology 34, 389‐392. Phillips, W.S., Begnaud, M.L., Rowe, C.A., Steck, L.K., Myers, S.C., Pasyanos, M.E., Ballard, S., 2007. Accounting for lateral variations of the upper mantle gra‐
dient in Pn tomography studies. Geophys. Res. Lett. 34, L14312. doi:10.1029/2007GL029338. Popov, Y.A., Pribnow, D.F.C., Sass, J.H., Williams, C.F., Burkhardt, H., 1999. Characte‐
rization of rock thermal conductivity by high‐resolution optical scanning. Geo‐
thermics 28, 253‐276. Rodgers, A.J., Walter, W.R., Mellors, R.J., Al‐Amri, A.M.S., Zhang, Y.‐S., 1999. Lithos‐
pheric structure of the Arabian Shield and Platform from complete regional waveform modelling and surface wave group velocities. Geophys. J. Int. 138, 871‐878. Rudnick, R.L., Fountain, D.M., 1995. Nature and composition of the continental crust: a lower crustal perspective. Rev. Geophys. 33, 267‐309. Rudnick, R.L., McDonough, W.F., O'Connell, R.J., 1998. Thermal structure, thickness Camp, V.E., Roobol, M.J., 1992. Upwelling asthenosphere beneath western Arabia and its regional implications. J. Geophys. Res. B97, 15,255‐15,271. Dai, B.‐Z., Jiang, S.‐Y., Jiang, Y.‐H., Zhao, K.‐D., Liu, D.‐Y., 2008. Geochronology, geochemistry and Hf‐Sr‐Nd isotopic compositions of Huziyan mafic xenoliths, southern Hunan Province, South China: petrogenesis and implications for lower crust evolution. Lithos 102, 65‐87. DESERT Group, Weber, M., Abu‐Ayyash, K., Abueladas, A., Agnon, A., Al‐Amoush, H., Babeyko, A., Bartov, Y., Baumann, M., Ben‐Avraham, Z., Bock, G., Bribach, J., El‐
Kelani, R., Förster, A., Förster, H.‐J., Frieslander, U., Garfunkel, Z., Grunewald, S., Götze, H.J., Haak, V., Haberland, Ch., Hassouneh, M., Helwig, S., Hofstetter, A., Jäckel, K.‐H., Kesten, D., Kind, R., Maercklin, N., Mechie, J., Mohsen, A., Neubau‐
er, F.M., Oberhänsli, R., Qabbani, I., Ritter, O., Rümpker, G., Rybakov, M., Ry‐
berg, T., Scherbaum, F., Schmidt, J., Schulze, A., Sobolev, S., Stiller, M., Thoss, H., Weckmann, U., Wylegalla, K., 2004. The crustal structure of the Dead Sea Transform. Geophys. J. Intern. 156, 655‐681. El‐Isa, Z.H., Mechie, J., Prodehl, C., Makris, J., Khim, R., 1987. A crustal structure study of Jordan derived from seismic refraction data. Tectonophysics 138, 235‐253. Embey‐Isztin, A., Downes, H., Kempton, P.D., Dobosi, G., Thirlwall, M., 2003. Lower crustal granulite xenoliths from the Pannonian Basin, Hungary. Part 1: Miner‐
al chemistry, thermobarometry and petrology. Contrib. Mineral. Petrol. 144, 652‐670. Förster, A., Förster, H.‐J., Masarweh, R., Masri, A., Tarawneh, K., DESERT Group, 2007. The surface heat flow of the Arabian Shield in Jordan. J. Asian Earth Sci. 30, 271‐284. Galanis, S.P., Sass, J.H., Munroe, R.J., Abu‐Ajamieh, M.,1986. Heat flow at Zerqa Ma'in and Zara and a geothermal reconnaissance of Jordan. USGS Open File Report 86‐631, 2‐110. Garrido, C.J., Bodinier, J.‐L., Burg, J.‐P., Zeilinger, G., Hussain, S.S., Dawood, H., Chaudhry, M.N., Gervilla, F., 2006. Petrogenesis of mafic garnet granulite in the lower crust of the Kohistan Paleo‐arc Complex (Northern Pakistan): implica‐
tions for intracrustal differentiation of island arcs and generation of continen‐
tal crust. J. Petrol. 47, 1873‐1914. Gazit, O., 2005. Mafic granulite xenoliths from Qarnei Hittin and the evolution of the lower crust beneath northern Israel. Geol. Surv. of Israel, Report GSI/11/05 (in Hebrew). Gettings, M.E., 1982. Heat‐flow measurements at shot points along the 1978 Saudi Arabian seismic deep‐refraction line, Part II. Discussion and interpretation. Saudi Arabia Deputy Min. for Mineral Res., Open File Report USGS‐OF‐02‐40. 40 pp. Gettings, M.E., Showail, A., 1982. Heat‐flow measurements at shot points along the 1978 Saudi Arabian deep‐refraction line, Part I: Results of measurements. Saudi Arabia Deputy Min. for Mineral Res., Open File Report USGS‐OF‐02‐39. 98 pp. Götze, H.‐J., El‐Kelani, R., Schmidt, S., Rybakov, M., Hassouneh, M., Förster, H.‐J., Ebbing, J., DESERT Group, 2007. Integrated 3‐D density modeling and segmen‐
tation of the Dead Sea Transform. Int. J. Earth Sci. 96, 289‐302. Hansen, S., Rodgers, A.J., Schwartz, S.Y., Al‐Amri, A.M.S., 2007. Imaging ruptured lithosphere beneath the Red Sea and Arabian Peninsula. Earth Planet. Sci. Lett. 259, 256‐265. Hargrove, U.S., Stern, R.J., Kimura, J.‐L., Manton, W.I., Johnson, P.R., 2006. How juvenile is the Arabian‐Nubian Shield? Evidence from Nd isotopes and pre‐
Neoproterozoic inherited zircon in the Bi'r Umq suture zone, Saudi Arabia. Earth Planet. Sci. Lett. 252, 308‐326. Hofstetter, A., Bock, G., 2004. Shear‐wave velocity structure of the Sinai subplate from receiver function analysis. Geophys. J. Int. 158, 67‐84. Huang, X.‐L., Xu, Y.‐G., Liu, D.‐Y., 2004. Geochronology, petrology and geochemistry of the granulite xenoliths from Nushan, East China: implication for a hetero‐
geneous lower crust beneath the Sino‐Korean Craton. Geochim. Cosmochim. Acta 68, 127‐149. Ilani, S., Harlavan, Y., Tarawneh, K., Rabba, I., Weinberger, R., Ibrahim, K., Peltz, S., Steinitz, G., 2001. New K‐Ar ages of basalts from the Harrat Ash Shaam volcan‐
ic field in Jordan: implications for the span and duration of the upper‐mantle upwelling beneath the western Arabian plate. Geology 29, 171‐174. Ionov, D.A., Ashchepkov, I., Jagoutz, E., 2005. The provenance of fertile off‐craton lithospheric mantle: Sr‐Nd isotope and chemical composition of garnet and spinel peridotite xenoliths from Vitim, Siberia. Chem. Geol. 217, 41‐75. Jarrar, G., Stern, R.J., Saffarini, G., Al‐Zubi, H., 2003. Late‐ and post‐orogenic Neopro‐
terozoic intrusions of Jordan: implications for crustal growth in the nor‐
thernmost segment of the East African Orogen. Precambr. Res. 123, 295‐319. Jarrar, G.H., Manton, W.I., Stern, R.J., Zachmann, D., 2008. Late Neoproterozoic A‐
type granites in the northernmost Arabian‐Nubian Shield formed by fractiona‐
tion of basaltic melts. Chem. Erde 68, 295‐312. Jaupart, C., Mareschal, J.C., 1999. The thermal structure and thickness of continen‐
tal roots. Lithos 48, 94‐114. Julia, J., Ammon, C.J., Herrmann, R.B., 2003. Lithospheric structure of the Arabian Shield from the joint inversion of receiver functions and surface‐wave group velocities. Tectonophysics 371, 1‐21. Kaliwoda, M., Altherr, R., Meyer, H.‐P., 2007. Composition and thermal evolution of the lithospheric mantle beneath the Harrat Uwayrid, eastern flank of the Red Sea rift (Saudi Arabia). Lithos 99, 105‐120. Katsura, T.,1995. Thermal diffusivity of olivine under upper mantle conditions. Geophys. J. Intern. 122, 1, 63‐69. Katz, O., Beyth, M., Miller, N., Stern, R., Avigad, D., Basu, A., Anbar, A., 2003. A Late Neoproterozoic (~630 Ma) high‐magnesium andesite suite from southern Israel: implications for the consolidation of Gondwanaland. Earth Planet. Sci. Lett. 218, 475‐490. Katzir, Y., Litvinovsky, B.A., Jahn, B.M., Eyal, M., Zanvilevich, A.N., Valley, J.W., Vapnick, Ye, Beeri, Y., Spicuzza, M.J., 2007. Interrelations between coeval mafic 36 H.­J. Förster et al. / Tectonophysics 481 (2010) 29­37 Arabian‐Nubian shield: evidence from granitic rocks of the Nakasib suture, NE Sudan. Geol. Rundsch. 76, 150‐160. Stern, R.J., Kröner, A.,1993. Late Precambrian crustal evolution in NE Sudan: isotopic and geochronological constraints. J. Geol. 101, 555‐574. Stoeser, D.B., Camp, V.E., 1985. Pan‐African microplate accretion of the Arabian shield. Geol. Soc. Amer. Bull. 96, 817‐826. Taylor, S.R., McLennan, S.M., 1985. The Continental Crust: Its Composition and Evolution. Blackwell, Cambridge, Mass. 312 pp. Ten Brink, U.S., Al‐Zoubi, A.S., Flores, C.H., Rotstein, Y., Qabbani, I., Harder, S.H., Keller, G.R., 2006. Seismic imaging of deep low‐velocity zone beneath the Dead Sea basin and transform fault: implications for strain localization and crustal rigidity. Geophys. Res. Lett. 33, L24314. doi:10.1029/2006GL027890. Tommasi, A., Gibert, B., Seipold, U., Mainprice, D., 2001. Anisotropy of thermal diffusivity in the upper mantle. Nature 411, 783‐786. Tullis, J., Yund, R.A.,1985. Dynamic recrystallization of feldspar: a mechanism for ductile shear zone formation. Geology 13, 238‐241. Turcotte, D.L., Schubert, G., 2002. Geodynamics. Cambridge Univ. Press, Cambridge, UK. 456 pp. Wilson, M., Shimron, A.E., Rosenbaum, J.M., Preston, J., 2000. Early Cretaceous magmatism of Mount Hermon, Northern Israel. Contrib. Mineral. Petrol.139, 54‐67. Xu, Y.‐G., Blusztajn, J., Ma, J.‐L., Suzuki, K., Liu, J.‐F., Hart, S.R., 2008. Late Archean and Early Proterozoic lithospheric mantle beneath the western North China craton: SrNd‐Os isotopes of peridotite xenoliths from Yangyuan and Fansi. Li‐
thos 102, 25‐42. Yu, J.‐H., Xu, X., O'Reilly, S.Y., Griffin, W.L., Zhang, M., 2003. Granulite xenoliths from Cenozoic basalts in SE China provide geochemical fingerprints to distinguish lower crust terranes from the North and South China tectonic blocks. Lithos 67, 77‐102. and composition of continental lithosphere. Chem. Geol. 145, 395‐411. Rybach, L., 1976. Radioactive heat production; a physical property determined by the chemistry of rocks. In: Strens, R.G.J. (Ed.), The Physics and Chemistry of Minerals and Rocks. J. Wiley & Sons, London, pp. 309‐318. Seipold, U., 1998. Temperature dependence of thermal transport properties of crystalline rocks ‐ a general law. Tectonophysics 291, 161‐171. Seipold, U., 2001. Der Wärmetransport in kristallinen Gesteinen unter den Bedin‐
gungen der kontinentalen Erdkruste. Scientific Technical Report STR01/13, GeoForschungsZentrum Potsdam, Potsdam. 142 pp. Shaw, J., 2003, Geochemistry of Cenozoic volcanism and Arabian lithospheric mantle in Jordan. PhD thesis, Royal Holloway University of London, London, 267 pp. Shaw, J.E., Baker, J.A., Menzies, M.A., Thirlwall, M.F., Ibrahim, K.M., 2003. Petroge‐
nesis of the largest intraplate volcanic field on the Arabian plate (Jordan): a mixed lithosphere‐asthenosphere source activated by lithosphere extension. J. Petrol. 44, 1657‐1679. Shaw, J.E., Baker, J.A., Kent, A.J.R., Ibrahim, K.M., Menzies, M.A., 2007. The geoche‐
mistry of the Arabian lithospheric mantle ‐ a source for intraplate volcanism? J. Petrol. 48, 1495‐1512. Sobolev, S.V., Babeyko, A.Yu., 1994. Modeling of mineralogical composition, density and elastic wave velocities in anhydrous magmatic rocks. Surv. in Geophys. 15, 515‐544. Sobolev, S.V., Petrunin, A., Garfunkel, Z., Babeyko, A.Y., DESERT Group, 2005. Thermo‐mechanical model of the Dead Sea Transform. Earth Planet. Sci. Lett. 238, 78‐95. Stein, M., 2003. Tracing the plume material in the Arabian‐Nubian Shield. Precambr. Res. 123, 223‐234. Stein, M., Garfunkel, Z., Jagoutz, E.,1993. Chronothermometryof peridotitic and pyroxenitic xenoliths; implications for the thermal evolution of the Arabian li‐
thosphere. Geochim. Cosmochim. Acta 57, 1325‐1337. Stern, R.J., Abdelsalam, M.G., 1998. Formation of juvenile continental crust in the 37