Arab J Geosci (2013) 6:2343–2350 DOI 10.1007/s12517-011-0503-4
ORIGINAL PAPER
The lithosphere–asthenosphere boundary in the eastern part of the Dead Sea Basin (DSB) from S-to-P receiver functions Ayman Mohsen & Guenter Asch & Rainer Kind & James Mechie & Michael Weber
Received: 28 June 2011 / Accepted: 4 December 2011 / Published online: 3 January 2012 # Saudi Society for Geosciences 2011
Abstract Clear S-to-P converted waves from the crust–mantle boundary (Moho) and lithosphere–asthenosphere boundary (LAB) have been observed on the eastern part of the Dead Sea Basin (DSB), and are used for the determination of the depth of the Moho and the LAB. A temporary network consisting of 18 seismic broad-band stations was operated in the DSB region as part of the DEad Sea Integrated REsearch project for 1.5 years beginning in September 2006. The obtained Moho depth (∼35 km) from S-to-P receiver functions agrees well with the results from P-to-S receiver functions and other geophysical data. The thickness of the lithosphere on the eastern part of the DSB is about 75 km. The results obtained here support and confirm previous studies, based on xenolith data, geodynamic modeling, heat flow observations, and S-to-P receiver functions. Therefore, the lithosphere on the eastern part of the DSB and along Wadi Araba has been thinned in the Late Cenozoic, following rifting and spreading of the Red Sea. The thinning of the lithosphere occurred without a concomitant change in the crustal thickness and thus an upwelling of the asthenosphere in the study area is invoked as the cause of the lithosphere thinning. A. Mohsen (*) : G. Asch : R. Kind : J. Mechie : M. Weber Deutches GeoForschungsZentrum—GFZ, 14473 Potsdam, Germany e-mail:
[email protected] A. Mohsen An-Najah National University, Nablus, Palestine G. Asch : R. Kind Freie Universität, Berlin, Germany M. Weber Institut für Geowissenschaften, Universität Potsdam, Potsdam, Germany
Keywords Dead Sea basin . S receiver functions . Lithosphere
Introduction The Dead Sea transform (DST; Fig. 1) is a structure of vital importance and has been studied by several geological and multi-disciplinary geophysical projects. It plays a significant role in the Middle East region in the sense that it demarcates the boundary between the Arabian plate and the Sinai subplate (Quennell 1958; McKenzie et al. 1970; McKenzie 1972; Freund et al. 1970; Garfunkel 1981). The Dead Sea region has remained a stable platform almost since its formation in the late Proterozoic. This tectonic stability was only recently (∼20 My ago) interrupted by the formation of a transform with a left lateral motion of about 107 km as of today (Quennell 1958; Garfunkel 1981). The DST strikes in a north-northeast direction and extends over some 1,000 km from the active spreading center of the Red Sea to the continental collision zone in the Taurus–Zagros mountain belt (McKenzie et al. 1970; Garfunkel 1981; Girdler 1990; El-Isa 1990). The estimated slip rate varies between 1 and 10 mm/year (Garfunkel et al. 1981; Klinger et al. 2000; Meghraoui et al. 2003; Le Beon et al. 2008). The DST is the main source of earthquakes in the region, where large earthquakes are known to have occurred along this structure during the historical period (Abou Karaki 1987; Ambraseys et al. 1994). The internal structure of the transform is dominated by left-stepping enechelon strike-slip faults (Garfunkel et al. 1981; Garfunkel 1981). This fault arrangement produced several pull-apart structures (rhomb-shaped grabens) which form deep basins. The largest of these structures is the DSB with a total length of more than 150 km (e.g. Garfunkel and Ben-Avraham 1996; Al-Zoubi et al. 2002; ten Brink et al. 2006; Smit et al. 2008;
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Fig. 1 Tectonic map, showing the area studied represented by the green box. The blue box shows the area studied during the DESERT project (DESERT Group 2004; Mohsen et al. 2005). The black box shows the area studied during the DESIRE project (Mohsen et al. 2011). The black line shows the wide-angle reflection/refraction
(WRR) profile across the Dead Sea Basin (Mechie et al. 2009). The inset shows the regional tectonic setting of the Dead Sea Transform. Closed squares represent the broad-band seismic stations used in this study together with station names
Mechie et al. 2009; Mohsen et al. 2011). The DSB is located between two principal longitudinal faults which are the continuation of the major strike-slip faults north and south of the basin (Garfunkel and Ben-Avraham 1996). The eastern part of these faults, the Araba fault, runs from the Gulf of Aqaba to the north over some 190 km, whereas the western fault, the Jericho fault, seems to start some 20 km west of the northern end of the Araba fault in the western part of the Dead Sea and runs in a NNE direction to the Tiberias lake over a length of ca 150 km. The DSB is composed of two sub-basins, the southern basin and the northern basin, which are separated by the Lisan salt diapir (Quennell 1958; Neev and Hall 1979). The northern sub-basin is occupied by the deep lake, the Dead Sea, whereas the southern sub-basin extends southwards to about the middle of the Araba valley. The basin reaches a depth of about 9 km in its central part under the Lisan diapir, as indicated by depth conversion of the seismic lines and gravity data (Al-Zoubi and ten Brink 2001; ten Brink et al. 1999). Al-Zoubi et al. (2002) have suggested a 10-km thick basin under the Lisan peninsula. The most recent seismic reflection/refraction profile, the DEad Sea Integrated REsearch (DESIRE) profile, estimates the depth to the Precambrian crystalline basement to be 11 km below sea level where it crosses the southern DSB (Mechie et al. 2009). From a receiver function study across the DSB, Mohsen et al. (2011) give a depth of 8 to 10 km for the top of the crystalline basement beneath the DSB.
Until recently, the coverage of seismic stations in this area was rather sparse. However, the recently installed DESERT temporary network (1 year deployment), and the DESIRE temporary network (1.5 years deployment), provide a wealth of data. These data improved the knowledge of the crustal structure in some segments of the DST, such as Wadi Araba (DESERT Group 2004; Mohsen et al. 2005) and the Dead Sea (Mechie et al. 2009; Mohsen et al. 2011). The DESIRE project is a multi-national, inter-disciplinary geophysical initiative aimed at determining the crustal and upper mantle structure beneath the DSB. It consists of five sub-projects, namely plate movement, crustal structure that includes two controlled source seismic experiments (Mechie et al. 2009) and a natural source magnetotelluric experiment, aero-gravity, earthquakes and radon, and geodynamic modeling. In a previous study by Mohsen et al. (2011), a P-to-S receiver function analysis has been used to study the crustal structure of the DSB. Here, we describe an S-to-P receiver function analysis of teleseismic records in terms of lithospheric structure on the eastern part of the DSB.
Data and methods The data used in this study consist of teleseismic recordings from 18 broad-band seismic stations in the Dead Sea Basin (Fig. 1). The seismic stations were part of the DESIRE
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project that was launched in March 2006. Table 1 lists the locations of stations used in this study. The data were continuously recorded with a sampling rate of 100 samples/s. The data which have been used for this study were selected according to the following criteria: 1 Only teleseismic earthquakes with epicentral distances ranging from 60° to 85° have been used. Good S-to-P conversion from the S phase can be observed in this distance range (Yuan et al. 2006). 2 Earthquakes with magnitudes greater than 5.6 and clear records of S onset with high signal-to-noise ratio. 3 Earthquakes within all depth ranges were selected. The epicenters of the selected earthquakes are shown in Fig. 2. Most of the events are in the northeast quadrant. The S-to-P receiver function technique has been used to identify the LAB (e.g., Farra and Vinnik 2000; Li et al. 2004; Kumar et al. 2005; Yuan et al. 2006; Mohsen et al. 2006; Sodoudi et al. 2006; Hansen et al. 2007). The S receiver function method is an extension of the conventional P receiver function method. In the S receiver function method, the S-to-P converted waves at seismic discontinuities in the crust and mantle beneath a seismic station are analyzed. The low velocity zone in the upper mantle usually derived with surface waves is frequently interpreted as the asthenosphere (e.g., Romanowicz 2009). The directly converted S-to-P waves arrive at the station earlier than the incident S wave, and thus are naturally separated from the multiple reverberations that appear later than the S arrival, thus making this method suitable for investigating the mantle structure of the lithosphere and asthenosphere. Similar to P receiver Table 1 Stations codes (ID) and coordinates of stations used ID
Sensor
LAT.
LONG.
JB01 JB04 JB07 JB10 JB12 JB15 JB18 JB19
STS2 STS2 STS2 40-T STS2 40-T 40-T 40-T
31.23710 31.24997 31.29146 31.33855 31.37247 31.24745 31.25148 31.23319
35.47915 35.42810 35.47960 35.49455 35.50018 35.54355 35.62162 35.64560
JB37 JB41 IB09 IB11 IB13 IB15 IB17 IB21 IB31 IB38
STS2 STS2 STS2 40-T 40-T STS2 STS2 STS2 3-T 3-T
31.38895 31.27262 31.23122 31.20105 31.17122 31.13636 31.10100 31.02075 31.31658 31.44987
35.67463 35.54250 35.39659 35.42019 35.44404 35.45527 35.45228 35.42053 35.34742 35.38094
Fig. 2 Distribution of teleseismic events with magnitude greater than 5.6 used in this study for S receiver function analysis. The equidistant circles show distances in degrees to the DESIRE network in the Dead Sea Basin (DSB)
functions, the signal of the S receiver function is related to the velocity contrast across a discontinuity. Positive converted phases from the Moho indicate velocity increase with depth, whereas negative converted phases indicate velocity decrease with depth. S receiver function observations of a negative discontinuity in the upper mantle are frequently explained as observations of the LAB. We follow this interpretation. In this study, we apply the S receiver function technique to look at the base of the lithosphere in the Dead Sea Basin area. The usual processing starts with rotation of the three components Z–N–E into the P–SV–SH (LQT) coordinate system (Li et al. 2004). After rotation, the converted S-to-P phases are primarily on the L component. Deconvolution of the P component with the S signals on the SV component is then carried out to equalize effects of different S signal forms. Therefore, the resulting P components, which contain only the converted phases are called the S receiver functions. The incidence angle of an S-to-P converted phase is usually larger than that of the incident S wave. Here, the incidence angle was determined by minimizing the S wave energy on the P component. A bandpass filter of 5–20 s has been applied. To make the S-to-P receiver functions directly comparable with the P receiver functions, the polarity of the S receiver functions and also the time axis have been reversed. Since conversions are weak phases, we summed S receiver functions of many earthquakes. Prior to summation, as in the P receiver function method, the moveout correction for a reference slowness of 6.4 s/degree was applied.
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Crust–mantle boundary Although the crust–mantle boundary (Moho) is not the main target in this study, it is noteworthy that conversions from it are observed with an average delay time of about 4.5 s using the P and S receiver function method (Fig. 3a and b). To transform the time domain into the depth domain for each seismic station, the P and S wave velocity models for the DESIRE wide-angle reflection/refraction profile (Mechie et al. 2009; Mohsen et al. 2011) have been used. The corresponding Moho depths are between 27 and 37 km (Fig. 4) which is in agreement with other geophysical studies carried out in this part of the DST region, such as seismic reflection/refraction profiles (Makris et al. 1983; El-Isa et al. 1987a, 1987b; DESERT Group 2004; Mechie et al. 2005; 2009), gravity studies (Al-Zoubi and Ben-Avraham 2002; Götze et al. 2006), local source and teleseismic tomography (Koulakov and Sobolev 2006; Koulakov et al. 2006), and receiver functions (Hofstetter and Bock 2004; Mohsen et al. 2005; 2011). Most of the mentioned studies including this study suggest that the crustal thickness varies from about 30 km in the west to about 38 km in the east. This study and the complementary P receiver function study in the same area show a crustal thickness which is shallower by about 2–4 km beneath the DSB compared to the crustal thickness under
Fig. 3 Plot of all P-to-S and S-to-P receiver functions of the broadband stations from the DESIRE temporary network. In a (left), the traces are plotted according to the slowness. The traces are moveoutcorrected for PS conversions underneath the stations. The sum trace is displayed at the top. The y-axis represents the number of traces. The
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Jordan due to 2-D/3-D effects, especially on and in the immediate vicinity of the Lisan Peninsula (Mohsen et al. 2011; Weber et al. 2011). A lower crustal discontinuity (LCD) has been found on the eastern part of the DST, around a depth of 30 km (DESERT Group 2004; Mohsen et al. 2005; Götze et al. 2006). A similar structure (LCD) has also been found beneath the DSB and to the west of it (Mohsen et al. 2011).
Lithosphere–asthenosphere boundary The thickness of the lithosphere is one of the most crucial parameters to address the problems of plate tectonic processes. The transition from the lithosphere to the asthenosphere in the DST region is still poorly sampled by most seismic data. Beneath the DESERT temporary broad-band seismic network a first order seismic discontinuity or a mantle lid appears at approximately 6 to 8 s, that is, at depths of approximately 60 to 80 km (see Fig. 5, left). Ginzburg et al. (1979b), 1981 have reported a reflected P phase from the top of an upper mantle layer with a velocity of 8.6 km/s at a depth of 55 km along a profile from the southernmost part of the Gulf of Aqaba to the DSB. Using data of the station UNJ (located at about 32°N and 36°E), El-Isa (1990) has noted upper mantle phases with apparent velocities of 8.4 and 8.69 km/s. These phases have
crustal phases that can be easily seen especially in the summation trace are: S sediments, M Moho, LCM lower crustal multiple, MM Moho multiples. In b (right), the traces are ordered according to the latitude of their piercing points (from south to north). The Moho and the LAB are marked
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Fig. 4 Map of Moho depths derived with the receiver function method (modified from Weber et al. 2011 and Mohsen et al. 2011). The circles indicate the conversion points at the Moho depths; see color code for Moho depths. White circles are locations where no Moho depth could be determined
Fig. 5 (Left) Distribution of S-to-P piercing points at 80 km depth from the DESERT temporary network. The area has been divided into seven non-overlapping boxes denoted by numbers. The average thickness of the lithosphere is given in kilometers beside each box (Mohsen
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been interpreted to arise from discontinuities at 55 km depth with a thickness of approximately 2 to 3 km and at 105 km depth. In between these refractors, El-Isa (1990) inferred the presence of a low velocity zone reaching its minimum velocity at a depth of 80 km. From a receiver function analysis, Hofstetter and Bock (2004) studied the shear wave velocity structure of the Sinai subplate and reported a velocity decrease beginning at about 60–70 km depth beneath the EIL and JER seismic stations. Park et al. (2008) has reported a broad low velocity region to depths of about 100 km in the mantle across the Arabian shield and a narrower low velocity region at depths of more than 150 km localized along the Red Sea coast. From a joint inversion of receiver functions and surface wave group velocities, Julia et al. (2003) reported upper mantle shear velocities ranging from 4.3 to 4.6 km/s in the Arabian shield. They suggest that the lithosphere beneath one station in Saudi Arabia (TAIF) that is located at 21.4°N and 40.30°E could be as thin as 50 to 60 km. McGuire and Bohannon (1989) and Stein et al. (1993), from a study of mantle xenoliths in western Saudi Arabia (220 km east of the Red Sea), suggest that the lithosphere has been thinned by several tens of kilometers to its present calculated thickness of about 80 km or less. Hansen et al. (2007) estimated lithospheric thickness beneath the Arabian shield using S wave receiver functions. They reported a thin (∼50 km) lithosphere under the Red Sea thickening toward the Arabian Platform. Geothermometric data from xenoliths also suggests that the lithosphere is less than 80 km thick under the Arabian rift shoulder. On the
et al. 2006). (Right) Distribution of S-to-P piercing points from the DESIRE temporary network together with the average lithospheric thickness
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other hand, the crust in this area has a thickness of about 40 km (McGuire and Bohannon 1989). This would give an indication that the lithosphere was thinned (compared to the Arabian platform) without concomitant crustal thinning in the area beneath the Arabian rift shoulder, where the volcanic eruptions directly relate to the formation of the Red Sea commenced in Oligocene times (Camp and Roobol 1989). In Fig. 5, we calculate the distribution of the piercing points of S-to-P conversions at 80 km depth. Due to the location of usable earthquakes (see Fig. 2) according to the criteria mentioned earlier, all piercing points are located east of the DSB. This distribution prevents getting a complete picture of the lithospheric thickness across the DSB. Figure 3b shows the S-to-P receiver functions for all the DESIRE temporary network in a time window between 0 and 30 s. The traces are sorted according to the latitude. The x-axis represents the conversion time of S-to-P phases with respect to S arrival time, termed as “delay time” and the y-axis represents the number of traces. The sum trace is displayed on the top of the panel. Two phases are easily identified in this figure. The first phase is a conversion from the Moho boundary labeled 'M' at 4.5 s delay time with respect to the S arrival time. The second stable and coherent phase is a negative phase that gives a clear image of the LAB. The arrival time of this phase varies between 6.5 and about 8 s for the individual traces. From the sum trace at the top, the arrival time is about 7.2 s. To convert the arrival times of the LAB to depth, the IASP91 reference model has been used (Kennett and Engdahl 1991). Each second in time corresponds to about 10 km in depth. This conversion method may result in about 3 km maximum error in the LAB depth estimation, due to uncertainty of up to 5% in the velocity structure in the crust and mantle lithosphere. The second possible error occurs in the selection of the arrival times of the converted phases due to noise, which is about 0.2 s estimated from boot strapping of the data. The estimated maximum error due to the abovementioned uncertainties is about 5 km. The average thickness of the lithosphere on the eastern side of the DSB is about 75 km. This would indicate that the lithosphere is thin on the eastern side of the DSB in comparison to the Arabian interior as reported by Hansen et al. (2007) using S wave receiver functions and suggests thinner lithosphere under the western side of the shield and thicker (∼100–150 km thick) under the rest of the shield and platform. Our determinations correlate well with previous estimates from P and S receiver functions (Hofstetter and Bock 2004; Mohsen et al. 2006; Hansen et al. 2007).
Discussion and conclusion We have derived the thickness of the lithosphere east of the DSB using S-to-P receiver functions. We analyzed data from
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18 broad-band seismic stations that have been set up in the area for one and a half years beginning in September 2006. The results presented in this study and those of earlier studies suggest that the crustal thickness is about 35 km in the eastern part of the DSB region (Mechie et al. 2009; Mohsen et al. 2011). Refraction, reflection, receiver functions, and gravimetric studies show the Moho is shallower near the Mediterranean Sea and increases in depth towards the highlands in Jordan (e.g., DESERT Group 2004; Mohsen et al. 2005; Al-Zoubi and Ben-Avraham 2002; Götze et al. 2006; Mechie et al. 2009). The thickness of the lithosphere is about 75 km on the eastern part of the DSB region. Mohsen et al. (2006) found lithospheric thicknesses of 67 km near the Gulf of Aqaba, 80 km south of the Dead Sea and 73 km east of the Dead Sea (see Fig. 5, left). A recent analysis of S wave receiver functions reported thinner (∼70 km thick) lithosphere under the western side of the Arabian shield and about 100 to 150 km thick lithosphere under the rest of the platform. The data presented here indicate a 75-km thickness of the lithosphere east of the Dead Sea. The piercing point map presented in Mohsen et al. (2006) had less than 50 traces east of the Dead Sea shown in box 5 on the left side of Fig. 5, compared to more than 350 traces obtained by the dense DESIRE temporary network. In this study, we confirm the results obtained by Mohsen et al. (2006). The LAB is located at a shallow depth of about 75 km in the eastern part of the Dead Sea region. There is an apparent inconsistency between the xenolith data in this area, suggesting a hot mantle and a thin (∼75 km) lithosphere, and the low heat flow (40–50 mWm−2 (Eckstein and Simmons 1979; Ben-Avraham et al. 1978). However, revision of some of the heat flow data in Jordan (Förster et al. 2004, 2007) and modeling results (Petrunin and Sobolev 2006) suggest that the surface heat flow in the region may have been previously underestimated. Nevertheless, the proposed elevated values of 50–60 mWm−2 are still inconsistent with a lithospheric thickness of 75 km east of the Dead Sea and 70–80 km between the Dead Sea and Red Sea. The explanation of this inconsistency according to Stein et al. (1993), and Sobolev et al. (2005), is that the lithosphere of the Arabian shield has been significantly thinned recently enough (10–30 Ma ago) not to allow elevated heat flow to reach the surface. The thin lithosphere between the Dead Sea and the Red Sea obtained by Mohsen et al. (2006), and in this study east of the Dead Sea supports the explanation given by Stein et al. (1993), and Sobolev et al. (2005), and is also in line with the history of the regional late Cenozoic uplift (Steinitz and Bartov 1991). Based on the tectonic history of the region and modeling results, Sobolev et al. (2005) suggest that a big portion of the Arabian shield and the adjacent Mediterranean was tectonically stagnant following the Mesozoic, probably due to its location far from active mantle convection flows. This might have resulted in cooling and over thickening of the lithosphere as well as in reduced temperature of the sub-
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lithospheric mantle in the entire region. Due to the initiation of the Afar plume and the rifting and spreading of the Red Sea at about 20–30 Ma, the rejuvenation of the asthenosphere, the destabilization and thermal erosion of the over thickened mantle lithosphere began, resulting in the thinned lithosphere and a peak of surface uplift 5–10 Ma ago. Therefore, the lithosphere on the eastern part of the Dead Sea as well as along Wadi Araba is thinned without a concomitant thinning of the crust which is taken as an indication of an upwelling of the asthenosphere in the area covered in this study.
Acknowledgments The DESIRE project was funded by the Deutsche Forschungsgemeinschaft. The instruments were provided by the Geophysical Instrument Pool of the Deutsches GeoForschungsZentrum (GFZ) Potsdam, Germany. The first author A. Mohsen was funded by the DFG and the GFZ.
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