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Graham Brew Ph.D. Dissertation: Chapter 4

Published in modified form by the Journal of the Geological Society, London, 2001. All rights reserved.

Structure and Tectonic Development of the Dead Sea Fault System and Ghab Basin in Syria


We examine the structure and evolution of the Ghab Basin that formed on the active, yet poorly understood, northern Dead Sea transform fault system. Interpretations of seismic reflection and well data, gravity modeling, and surface geology yield a subsurface image of the basin. The basin formed in Plio-Quaternary time at a complex step-over zone on the fault. Subsidence occurred along cross-basin and transform-parallel faults in two asymmetric depocenters. The larger depocenter in the south end of the basin is asymmetric towards the east, the margin along which most active transform displacement is occurring. Our analysis is enhanced by comparison to deformation and deposition observed in other strike-slip basins and physical models.

The topographically and structurally high Syrian Coastal Ranges, located directly west of the Ghab Basin, are a consequence of Late Cretaceous and Cenozoic regional compression, heavily modified by the Plio-Quaternary Dead Sea Fault System and Ghab Basin formation. They are part of the broader scheme of Syrian Arc deformation. Plio-Quaternary uplift of the Coastal Ranges has been preferentially focused west of the Dead Sea Fault, possibly through reverse movement detached along the fault. A Plio-Quaternary age for the development of the Dead Sea Fault System in northwest Syria is consistent with previously proposed models of two-phase Dead Sea Fault System movement and Red Sea spreading.


Continental transform faults, such as the San Andreas Fault in California, the Alpine Fault in New Zealand, the North Anatolian Fault in Turkey, and the Dead Sea Fault System (DSFS), involve complex structural and sedimentary regimes. This complexity relates to the history of displacement along these fundamental components of the global plate tectonic framework. Our work concerns the development of structures and history of deposition along the northern DSFS that is relatively little studied compared to the southern DSFS (i.e., south of the Lebanese restraining bend, Figure 4.1). The evolution of the DSFS remains one of the most contentious issues of Middle Eastern tectonics.

This work begins with a very brief review of the DSFS. After a description of available data, we present our interpretation of the structure of the Ghab Basin that lies along the northern DSFS (Figures 4.1 and 4.2). This interpretation is largely based on high quality seismic reflection profiles from the basin that are published here for the first time. Our analysis, integrated with interpretations of Bouguer gravity anomalies and surface geology, shows the deep, asymmetric, double-depocenter structure of the Plio-Quaternary Ghab Basin. We compare this to other strike-slip basins and models of strike-slip basin formation to provide further insight into the tectonic controls on basin formation and evolution of the basin through time. The gross scale topographic signature of the adjacent Syrian Coastal Ranges is then considered. This prominent topography (Figure 4.2) is shown to be part of the regional Late Cretaceous and Cenozoic Syrian Arc uplift, albeit strong modified by the Plio-Quaternary propagation of the DSFS and development of the Ghab Basin. Our new regional model, founded on the previous work of Hempton (1987) and Chaimov et al. (1990), illustrates how the evolution of the Ghab Basin integrates with theories of Late Cretaceous Cenozoic plate motions in the eastern Mediterranean indicating that the DSFS only propagated through northwest Syria after the Miocene.


The DSFS is a transform fault linking Red Sea / Gulf of Aqaba seafloor-spreading to NeoTethyan collision in Turkey. Most researchers agree that in total ~107 km of sinistral motion has taken place on the 'southern' portion of the fault, south of the Lebanese restraining bend (Figure 4.1) (e.g. Dubertret, 1932). In concert with the episodic rifting in the Red Sea area (Hempton, 1987), many authors have suggested that the lateral motion on the DSFS occurred during two different episodes (e.g. Quennell, 1959; Freund et al., 1970; Beydoun, 1999). In this scenario there was ~65 km of movement during the Early - Middle Miocene, with the remaining ~42 km of from earliest Pliocene until present.

More controversial is the amount of translation experienced by the 'northern' DSFS in Lebanon and farther north in northwest Syria. The controversy arises from limited mapping of the trace of the DSFS, and a lack of piecing points by pre-Pliocene features, making total offset mapping in Lebanon and Syria extremely difficult (Chaimov et al., 1990). Displacement of ophiolite cut by the DSFS in Turkey is open to very broad interpretations, but was used by (Freund et al., 1970) to suggest ~70 km of total sinistral movement on the northern DSFS.

Chaimov et al. (1990), expanding on the ideas of Quennell (1959), suggested that only the second episode of motion on the DSFS (~40 - 45 km since the start of the Pliocene) has affected the northern DSFS. In this scenario shortening in the southwest Palmyride fold and thrust belt (Figure 4.1) accommodated ~20 km of sinistral movement, leaving ~20 - 25 km of movement to be transferred to the DSFS north of the Palmyrides. Supporting evidence for the post-Miocene development of the northern DSFS includes offsets of Pliocene basalt (Quennell, 1984), Quaternary fans, and Mesozoic ophiolite (Freund et al., 1970), although this last interpretation is discounted by many (e.g. Quennell, 1984). The Roum Fault in Lebanon (Figure 4.1) or similar structures, may have translated the ~65 km of pre-Pliocene displacement offshore, hence explaining the absence of a northern DSFS in Miocene time.

Another scenario suggests that the northern DSFS has been inactive since the Miocene (Butler et al., 1997). Given the geomorphic evidence for Pliocene Recent tectonic activity on the fault, however, together with seismicity (Ambraseys and Jackson, 1998) and GPS measurements (McClusky et al., 2000), this inactive northern fault hypothesis seems improbable.

The recent tectonics of the Ghab Basin (Figure 4.2) further attest to the current activity along the northern DSFS. Ponikarov (1966) considered the Ghab Basin to be a Pliocene - Recent feature, and recognized that the basin developed on a left-step in the DSFS (Figure 4.3). These findings were echoed in geomorphic studies by Hricko (1988), Domas (1994), and Devyatkin et al. (1997). Paleostress analysis on faults around the Ghab Basin by Mater and Mascle (1993) further suggest an active step-over geometry. Herein we do not present direct evidence regarding the history of movement on the northern DSFS, however, we suggest that the Ghab Basin formed through left-lateral strike-slip since earliest Pliocene. This is supports the scenario of northern DSFS development in which ~20 - 25 km of sinistral displacement has occurred on the northern DSFS since the earliest Pliocene (Chaimov et al., 1990).


Among the data used for our subsurface analysis of the Ghab Basin are ~260 km of 2-D migrated seismic reflection profiles (Figure 4.2), acquired during 1994 using a Vibroseis source to six seconds two-way time. These data were processed and migrated using standard seismic processing flows. Interpretation utilized the Landmark SeisWorks software package. The sections shown here are in time, rather than depth. Within the basin seismic p-wave velocities are 2.0 0.2 km/s as derived from sonic logs and seismic stacking velocities (Dzhabur, 1985). Hence the two-way time scales in Figures 4.4, 4.5, and 4.6 are a close approximation for depth in kilometers for the basin fill. The data are largely not interpretable past four seconds two-way time.

The one deep well within the basin (Ghab, Figure 4.2) was used, together with seismic signatures, to provide stratigraphic control on the seismic interpretations. There is very limited penetration of the basin fill by drilling (Figure 4.4), so while there is good stratigraphic control of older horizons within and around the basin, age control for the basin fill remains speculative. Regardless, the main objectives of this interpretation the mapping of the structure of the basin are met using the seismic data (Figure 4.7).

Geologic maps (Figure 4.3) and gravity interpretations provide additional information especially where seismic data are lacking. We modeled the Bouguer gravity data from a grid of eight profiles, two of which are presented here (Figures 4.4 and 4.5). The gravity modeling software permitted changing densities and body lengths in the strike direction; hence, the models are sensitive to lateral variations beyond what is usually considered two-dimensional modeling. Along-strike variations, at distances farther than ~5 km from the profile, caused no appreciable impact on the modeled anomaly. Consequently, with the exception of the profile along the axis of the basin (Figure 4.4), the modeling presented here is sufficiently accurate with no along-strike variations. Likewise, faults were not directly incorporated into the density models because of its insignificant effect relative to continuous surfaces. The final models give a reasonable fit (< 3 mGal difference) between calculated and observed anomalies.

During the gravity modeling, density information came from field samples (Hricko, 1988), borehole density logs (this study and Lupa, 1999), and seismic refraction data (Seber et al., 1993). Depth limits came from seismic refraction and reflection data, and well data, as presented in this study (for locations see Figure 4.2). Where density logs from wells were not directly available, sonic log velocities borehole sonic logs (this study and Dzhabur, 1985) were converted to densities using well-established velocity-density curves. Rigorous comparisons between densities obtained directly from density logs, those estimated from sonic logs, and those from direct field samples show differences of less than 0.1 g/cm3 (for further details, see Lupa, 1999). Furthermore, the resulting densities were found to be reasonable according to our knowledge of lithologies derived from drilling information. Given these external controls on densities and depths, this gravity modeling is better-determined and less ambiguous than typical gravity studies.



The surface expression of the Ghab Basin is an extensive, flat plain with almost no topographic relief thus betraying its recent lacustrine history. The plain is ~60 km long and ~15 km wide (Figure 4.2). In the south, between two strands of the DSFS, is the Missyf Graben (Figure 4.3). The eastern fault strand can be traced northward at the surface along the eastern margin of the Ghab Basin before bifurcating to the north-northeast (Figure 4.3). No definitive termination of this eastern fault is observed along the basin margin.

The Syrian Coastal Ranges - that Ponikarov (1966) referred to as Jebel An-Nusseriyeh - rise dramatically by ~1300 m in just four kilometers of distance (Figure 4.2), exposing Jurassic, and even uppermost Triassic, strata (Figure 4.3) directly west of the basin (Mouty, 1997). This steep flank suggests geologically recent uplift along the western margin of the Ghab Basin (the origin of the Coastal Ranges is discussed below). In contrast to the eastern margin, this edge of the basin is poorly defined, obscured by significant mass-wasting and large blocks detached from the Coastal Ranges (Domas, 1994). Faults would be expected along the western margin given a typical fault step-over arrangement for the Ghab Basin. However, no surface expression has been detected along these margins, except in the far north (Figure 4.3), probably owing to burial by mass-wasting.

At the northern end of the Ghab Basin, the surface plain bifurcates and the Balou Trough extends to the north-northeast; Jebel El-Wastani - up to 800 m high - divides this from the northern Ghab (Figure 4.2). Surface observations indicate the Ghab Basin fill is Neogene - Quaternary lacustrine and alluvial deposits, finer grained in the basin center (Domas, 1994; Devyatkin et al., 1997).

Subsurface Analysis


In our seismic interpretations (Figures 4.4, 4.5, and 4.6), tied to the Ghab well, the deepest mapped reflector is a relatively thin bed of Mid-Cambrian age limestone (Best et al., 1993). The unconformity at the top of Paleozoic (generally Upper Ordovician strata) presents a clear reflector where the mainly carbonate Mesozoic section overlies a largely clastic Paleozoic section. Middle Triassic age anhydrite and dolomite form a sequence of strong reflectors, as does Early Cretaceous sandstone. The uppermost mapped reflector, (other than arbitrarily traced horizons within the basin fill shown by dashed lines in Figures 4.4, 4.5, and 4.6), is at the base of basin fill that is Middle Eocene age, or in the south of the basin, Upper Cretaceous (Devyatkin et al., 1997).

The Ghab well penetrates Middle Eocene limestone immediately beneath Pliocene strata. A clear unconformity at this point is expressed by abrupt facies change (clay to limestone), paleontologic evidence, and an absence of volcanic detritus that is found throughout the younger strata. We interpret this unconformity (at a depth of 350 m in the Ghab well, but dropping sharply to the south and north, Figure 4.4) as the base of basin fill. This puts initial Ghab Basin formation, at least at the latitude of the Ghab well, in Pliocene time. Furthermore, we interpret a very thin layer of volcanic rocks encountered at a depth of 200 m within the basin fill in the Ghab well is part of a nearby 1-2 Ma sequence.

Since most of the Ghab Basin fill has not been drilled, directly dating the overall onset of extension and basin formation is not possible. However, when the seismic data are tied to the well data, there is no evidence of basin strata older than earliest Pliocene. Shallow borings (< 500 m) in the main depocenter have also failed to penetrate rocks older than Pliocene, and find Mesozoic strata immediately

Pliocene (Devyatkin et al., 1997). Furthermore, outcrop studies have shown marine Pliocene strata at the northern end of the current Ghab Basin, but continental strata of the same age near the south end of the basin (Ponikarov, 1966). Thus the full extent of the Ghab Basin topographic depression was not fully established until at least after the earliest Pliocene. In summary, the balance of evidence suggests Ghab Basin formation commenced around earliest Pliocene.


The basic structure of the Ghab Basin is a fault-controlled double depocenter. The main depocenter is positioned beneath the southern portion of the surface plain, and slight northward migration of that depocenter with time is clear from the seismic data (Figure 4.4). Also apparent are a mid-basin ridge (on which the Ghab well is drilled) and a second smaller depocenter to the north.

The relatively undeformed nature of much of the basin fill suggests that most subsidence has been accommodated along the major basin-bounding faults. An apparent western bounding fault (marked A on Figure 4.5) when projected to the surface would be close to the foot of the Coastal Ranges. Abrupt sediment thickness changes are very apparent across the eastern bounding fault (marked A in Figure 4.6). Most of the subsidence is clearly asymmetric in the southern depocenter controlled by the more prominent eastern basin-bounding fault (marked F in Figure 4.5). Mesozoic strata encountered by shallow drilling on the western flank (Devyatkin et al., 1997) further support this interpretation. Gravity and seismic interpretations reveal this southern depocenter to be up to ~3400 m deep (Figure 4.7). Assuming basin formation occurred in the last 4.5 Ma, the approximate subsidence rate in the deepest part of the Ghab Basin is ~0.8 m / 1,000 years. This is comparable with similar strike-slip basins elsewhere (Nilsen and Sylvester, 1995).

In the earlier stages of basin formation, accommodation space was created by movement on cross-basin faults that are now internal to the basin, rather than on the flanking faults. This displacement shifted between the faults, with older displacement on the more interior (western) faults (marked B-E in Figure 4.5), and most recent motion accommodated on the eastern basin-bounding fault (marked F in Figure 4.5).

The geometry of the faults can be appreciated from Figures 4.7 and 4.8a. Two depocenters are illustrated - the larger in the south, and the smaller in the north - both formed against the basin-bounding faults. Cross-basin faults are found particularly in the south of the basin, predominately steeply dipping to the northeast (faults B-E on Figure 4.5). These transverse features are generally northwest southeast striking, suggest some extension across the basin. The central region is dominated by acutely striking cross-basin faults that bound a horst extending across the basin (faults A-C on Figure 4.6, feature marked R in Figure 4.8a). Confidence in the interpretation of a second depocenter in the northwest of the basin is improved by analysis of gravity data (Figures 4.4 and 4.7). This second basin is somewhat asymmetric toward the western bounding fault and is up to ~1700 m deep.

North-northeast of the Ghab Basin, faults splay out significantly and several depocenters are present. A gravity low east of the Jebel El-Wastani (Figure 4.7) shows another step-over basin, beneath the Balou Trough (Figure 4.2). Whilst no seismic data have imaged this area, our gravity interpretations and previous work (Hricko, 1988) reports 500 - 1000 m of basin fill in what is apparently another strike-slip basin bounded by left-lateral faults (Ponikarov, 1966). Historical seismicity shows recent activity on some of these fault splays (Ambraseys and Melville, 1995).

East of the south part of the Ghab Basin, (Asharneh Plain, Figure 4.2), there is no significant Bouguer gravity low (Figure 4.6). Seismic interpretations (Figure 4.6) also indicate that there is no significant basin in that area, and this area is not an extension of the Ghab Basin as might be expected from the topographic expression (Figure 4.2). Faults in the Aleppo Plateau area are minor (Figure 4.6), and generally no older than movement on the northern DSFS. Seismic reflection data image deformation associated with the DSFS in this area (labeled A on Figure 4.6). This location corresponds directly with surface faults inferred from topography imagery (Figure 4.7) and Quaternary faults observed in the field. The displacement is distributed among several fault strands that are seen to coalesce at depth. This image is comparable with other examples of continental transform faults (e.g. Ben-Avraham, 1992), and is a typical 'flower structure' such has often been shown to be associated with strike-slip faulting (Harding, 1985).

Comparison with other basins and basin models

Despite out interpretations that have used all available data, several issues regarding the evolution of the Ghab Basin remain unresolved. Furthermore, while the fault geometry controlling the Ghab Basin roughly fits the pattern of a 'step-over' basin (Nilsen and Sylvester, 1995), the Ghab Basin shows several departures from this simple transform-parallel extension case. Below we compare the Ghab Basin with other basin studies and basin models (Figure 4.8), thus shedding some light on many of the second-order complexities we have observed.

Asymmetric basins have been documented along the DSFS (especially in the Gulf of Aqaba), the San Andreas Fault, the North Anatolian Fault, and many other major strike-slip faults (Ben-Avraham, 1992; Ben-Avraham and Zoback, 1992). These asymmetric basins are bound on only one side by a major linear strike-slip fault, against which most deposition always occurs. The opposite side of the basin is bound by predominantly normal faults; thus, the overall fault geometry is distinctly different from the classic step-over. The sense of basin asymmetry commonly changes along strike in these fault systems as strike-slip displacement transfers from one en-echelon strike-slip fault to the next. This geometry could be caused by a reorientation of stresses near a weak fault in an otherwise strong crust, so as to minimize shear stress on the fault, resulting in transform-normal extension (Ben-Avraham and Zoback, 1992).

The asymmetry within the Ghab Basin closely follows this pattern of deformation, with the southern depocenter asymmetric to the east. This suggests that, at the latitude of the Ghab Basin most of the lateral movement on the DSFS is accommodated on the eastern bounding fault of the basin. Some displacement steps over to the western bounding fault farther to the north, and the smaller northern depocenter is slightly asymmetric to the west. This geometry agrees with surface observations and indicates a component of extension across the Ghab Basin.

Transverse structures, such as those found in the Ghab Basin, are also commonly observed in other strike-slip basins. The Dead Sea Basin (Figure 4.8e) is bounded by strike-slip faults on which most of the deformation occurs and transverse structures separate smaller sub-basins there (Garfunkel and Ben-Avraham, 1996). Another analog for the Ghab Basin is the Cariaco Basin, Venezuela (Figure 4.8d), where twin depocenters, asymmetric toward the more active strike-slip and separated by a central sill have developed at a dextral fault step-over (Schubert, 1982).

Physical (e.g., sandbox / clay) models of pull-apart basins can provide insight into strike-slip basin evolution by considering simplistic end-member cases that are rare in nature. For Dooley and McClay (1997), their model with resulting deformation most closely resembling the Ghab Basin was a case of an initial 90 releasing sidestep between the two segments of the strike-slip fault (Figure 4.8b). Strong similarities with the Ghab Basin include: Cross-basin faults (C in Figure 4.8), mid-basin ridge (R in Figure 4.8), strongly terraced sidewalls of basin (T in Figure 4.8), and graben along the principal displacement zone at the basin ends (G in Figure 4.8).

Rahe et al. (1998) used unequal motion on the 'crustal' blocks on opposite sides of the strike-slip fault in their physical models. The results show asymmetric basins, with increased subsidence toward the moving boundary. Commonly observed in these models are intrabasin highs, early opening accommodated on oblique-slip transverse faults, and switching basin asymmetry along strike (associated with 'master fault' step-over). Again, all these features are observed in the Ghab Basin.

Mathematical (finite difference) models for the deformation of a basin under strike-slip conditions were made by Rodgers (1980) and Golke et al. (1994), among others. Rodgers (1980) showed that once the total offset across the bounding strike-slip faults is about equal to the separation between the faults, two distinct depocenters begin to form through normal faulting (Figure 4.8c). If considered analogous to the Ghab Basin (Figure 4.8a), this shows that the northern depocenter in the Ghab developed sometime after the initiation of the southern depocenter owing to increasing displacement on the DSFS. This explains the smaller size of the northern depocenter. Golke et al. (1994) found that two depocenters developed when initial master fault overlap is close to zero - the 90 case of Dooley and McClay (1997). They also saw the formation of asymmetric basins when there is some uneven movement on the master faults. Golke et al. (1994) also observed basin migration, in the same sense as that in the Ghab, because of increasing master fault overlap with time.


Seismic reflection interpretations reveal that the Ghab Basin is not a textbook example of a step-over basin. However, through comparison with other basin studies and models we find that many of the second-order structures within the Ghab Basin are common to other strike-slip basins. The basin asymmetry seen in the Ghab is probably related to the amount, and sense, of relative movement across the bounding lateral faults. The results are consistent with the observed surface faults that show a greater amount of relative motion on the eastern basin-bounding strike-slip fault. Observations from the Ghab are echoed in theoretical models that show cross-basin oblique-slip faults accommodating initial basin opening, but most subsidence on the basin bounding faults. A northward shifting depocenter, and the subsequent development of a second depocenter in the Ghab Basin, are due to increasing fault overlap with time and step-over of the lateral motion from the eastern to the western faults.


To properly discuss the evolution of the Ghab Basin and DSFS in Syria we must also mention the adjacent, topographically prominent Syrian Coastal Ranges (Figure 4.2). Although not well studied, this deformation can illuminate the regional tectonic regime under which the basin and DSFS formed.

Based on stratigraphic evidence, uplift in northwest Syria has been episodic since at least the latest Cretaceous. In the Coastal Ranges Ruske (1981) found tilted and eroded Maastrichtian strata unconformably overlain by Paleogene transgressive deposits that reached a high-stand in Middle Eocene time. The geometry of this latest Cretaceous and Paleogene uplift appears to have been similar to the current Coastal Range topography, albeit without the imposition of the Neogene Ghab Basin.

Middle Eocene limestone was deposited in much of the study area, including some of the Coastal Ranges, indicating that the latest Cretaceous and Paleogene uplift had largely subsided by that time. It is unclear whether absence of the Middle Eocene strata in the southern Coastal Ranges was due to continued emergence and non-deposition, or post-Middle Eocene erosion. In any event, uplift of the Middle Eocene strata on parts of the current Coastal Ranges indicates that most of the uplift has occurred since the Middle Eocene.

Structural relationships and outcrop geology indicate that an anticlinorium, sub-parallel to the present Coastal Ranges, formed at some time since the Middle Eocene. It is this anticlinorium that dominates the current topography. The crest of the anticlinorium forms the current ridge of the Coastal Ranges. The doming clearly narrows towards the north, and all evidence of the upwarping is lost near the present Turkish / Syrian border. The absence of any Late Eocene Miocene strata on or around the Coastal Ranges - or beneath the Ghab Basin - strongly suggests that this second stage of uplift started around Late Eocene time, as suggested by Ruske (1981). Quaternary coastal terraces attest to continued tectonic uplift in this area (Dalongeville et al., 1993).

Clearly the Coastal Range uplift has been very strongly modified by the propagation of the DSFS through northwest Syria, and the related formation of the Ghab Basin. The Plio-Quaternary Ghab Basin formed near what was presumably the crest of the pre-Pliocene Coastal Range uplift. This created the extremely steep scarp on the eastern face of the Coastal Ranges alongside the Ghab Basin. Furthermore, the presence of the DSFS has caused asymmetry in the uplift (Figures 4.2 and 4.9). The Coastal Ranges are topographically and structurally significantly higher directly to the west of the present DSFS. This indicates that some of the post-Middle Eocene uplift has occurred since the DSFS propagated through northwest Syria.

In the remainder of this section we examine two related attributes of the current Coastal Ranges that are presently unexplained. The first is the strong asymmetry of the current uplift. The second issue is the support of the topography. The Bouguer gravity anomalies (Figure 4.6) indicate that the current topography is not locally isostatically compensated, thus an explanation of a regional support mechanism is required.

Superficially, the asymmetry of uplift along the southern DSFS is similar to that near the Ghab Basin (Figure 4.2). On closer inspection, however, the half-width of the uplift is much greater in the southern DSFS (~100 - 125 km) than in the Syrian Coastal Ranges in the north (~15 - 25 km) (Figure 4.9). Even so, we may consider the explanations given for the uplift and asymmetry on the southern DSFS when trying to explain that in the north.

Wdowinski and Zilberman (1996; 1997) concluded that the uplift along the southern DSFS is caused by the isostatic lithospheric response to basin formation along the fault. They suggested the asymmetry along the southern DSFS is caused by deeply detached listric normal faults. ten Brink et al. (1990) also invoked flexure with asymmetric loading, elastic parameters, or thermal effects, to explain the asymmetry.

We have shown that a significant proportion of Coastal Range uplift occurred before the propagation of the DSFS through northwest Syria, hence the fault (and related basin formation) cannot be used to explain all the uplift. However, we can consider the additional Pliocene-Quaternary uplift that may have been caused by the faulting. To test this idea we have examined isostatic uplift of the Coastal Ranges due to Ghab Basin formation by assuming an elastic approximation following Turcotte and Schubert (1982). Using their method the uplift of an assumed elastic lithosphere can be modeled as being due to an upward force on a beam. In our case the upward force is the negative loaded created by basin formation. The uplift is then:

ziU(x) = w0 exp {-x + xl / a} . [sin (x + xl / a) + cos (x + xl / a)].........(1)


w0 = [Lt a^3 / 8 D]..............................................................................(2)

a = [4 D / Dr g]^0.25..........................................................................(3)

D = [E Te / 12 (1-n^2)].......................................................................(4)

If we consider the case of a broken lithosphere (beam), as could be the case along the DSFS, (1) becomes:

ziB(x) = w0 exp {-x + xl / a} . cos (x + xl / a).....................................(5)


w0B = [Lt a^3/ 4 D]...........................................................................(6)

In the above,

ziU,B(x) = flexure of lithosphere as a function of distance for unbroken and broken lithosphere, respectively
x = distance along profile
w0 = maximum amplitude of flexure, unbroken lithosphere
w0B = maximum amplitude of flexure, broken lithosphere
xl = offset distance of point load from center of profile
a = flexure parameter
Lt = 'Negative' Load: force that causes upward flexure The Ghab Basin is approximated with a 30 km^2 cross-sectional area (from seismic data), filled with sediments of density 2200 kg/m^3. The surrounding rock density is assumed to be 2600 kg/m^3, hence the negative load is 1.2x10^11 N/m.
D = flexural rigidity [1.8 x 10^10 N m]
Dr = density change between air and compensating 'fluid' layer [3300 kg/m^3]
g = acceleration due to gravity [9.81 m/s^2]
E = Young's Modulus [6 x 10^10 Pa]
Te = elastic thickness of the lithosphere [15,000 m]
n = Poisson's ratio [0.25]

The numbers in square brackets given next to the terms above are those used by Wdowinski and Zilberman (1996). We initially use these parameters to model the isostatic response due to the formation of the Ghab Basin. The resulting flexures for the case of the unbroken lithosphere, ziU(x), and the broken lithosphere ziB(x), are shown in Figure 4.9. Clearly, these flexures are of too small amplitude, and of too long a wavelength, to explain more than a small fraction of the present topography of the Syrian Coastal Ranges. The result is little changed if regional isostatic compensation occurs in the lower crust, rather than in the upper mantle (e.g. ten Brink et al., 1993). We conclude that the Pliocene-Quaternary Syrian Coastal Range uplift is not simply a consequence of Ghab Basin formation, and a regional isostatic response to Ghab Basin formation is not supporting the topography.

Thus we consider other mechanism for support of the Coastal Range topography. Recent seismological observations (Sandvol et al., 2000) indicate a zone of strong shear wave attenuation in the uppermost mantle beneath western Arabia, especially along the DSFS. This may indicate elevated mantle temperatures that could be supporting the uplift dynamically. However, a mantle driving force seems unlikely given the small wavelength of the uplift (Figure 4.9), and it also fails to explain the asymmetry of the uplift.

A more likely support mechanism for the Coastal Range uplift could be regional compression. We will see in the following section that regional compression caused the initial Coastal Range uplift. Regional plate kinematics from preliminary GPS data permit small convergence across the DSFS plate boundary (McClusky et al., 2000). The DSFS could be acting to decouple this compression by accommodating strike-slip and reverse slip of the crust west of the DSFS. In this scenario the Coastal Ranges west of the DSFS are uplifting though reverse faulting along the predominantly strike-slip DSFS, thus providing a support mechanism and explaining the asymmetry.

In summary, the true cause of the Syrian Coastal Range topographic support and asymmetry remains equivocal given the relatively limited data available. However, we favor a scenario in which the Syrian Coastal Ranges uplift began in the latest Cretaceous with regional compression causing folding and uplift. The area experienced similar compression in Late Eocene time onwards. After propagation of the DSFS through northwest Syria in Pliocene time, the Ghab Basin formed thus causing collapse the eastern flank of the Coastal Ranges. Regional compression continued to drive the uplift through reverse movement along the DSFS until present. This compression is largely detached along the DSFS hence explaining the current asymmetric uplift (Figure 4.10).


The previous sections have discussed our interpreted evolution of the Ghab Basin and Syrian Coastal Ranges. Now we consider these results in the context of the regional tectonic evolution of northwest Syria. As discussed, the timing of DSFS development in Syria is still controversial. Also, previous tectonic models have largely failed to incorporate findings from northwest Syria. Our results, although somewhat speculative, provide insight of this development for Late Cretaceous to Recent.

Late Cretaceous

Deep well data from northwest Syria illustrate some of the Paleozoic and Early Mesozoic history of the area. In general these observations fit previously proposed tectono-stratigraphic models for the region (e.g. Best et al., 1993; Brew et al., 1999). However, the latest Cretaceous period is of most relevance to the current work. The Maastrichtian age initial uplift of the Syrian Coastal Ranges (Figure 4.11a) is coincident with contemporaneous events documented throughout northwestern Arabia (Figure 4.12a). Most notably this time was the first episode in the formation of the 'Syrian Arc'. The Syrian Arc is the swath of folds and structurally inverted faults observed along the Sinai and Levant coasts, sub-parallel to the present shoreline (Figure 4.11a). In the original definition (Krenkel, 1924) the Arc extended northward towards Turkey, although more recent authors have also included Palmyride folds in the definition. The formation of the Syrian Arc is dated as a Maastrichtian phenomenon (Guiraud and Bosworth, 1997), although some subtle precursory compression began earlier in the Late Cretaceous (Bartov et al., 1980; Walley, 1998). Chaimov et al. (1992) considered the initial folding, uplift, and structural inversion in the southwest Palmyride fold and thrust belt to be part of the Syrian Arc and documented this compression as latest Cretaceous. On a more regional scale the cessation of extensional tectonics in eastern Syria is well established as a Maastrichtian phenomenon (Brew et al., 1999).

The Maastrichtian was the time of ophiolite emplacement along the northern Arabian margin, particularly in the Baer-Bassit and Kurd Dagh areas proximal to the present Ghab Basin (Al-Maleh, 1976; Robertson et al., 1991). This emplacement occurred because the north Arabian margin collided with an intra-ocean subduction zone. These collisions can explain the observations of Maastrichtian age compression throughout the northern Arabian platform. Thus, the initial Maastrichtian uplift of the Coastal Ranges fits completely with the previously documented regional plate tectonics. In this scenario, the Syrian Coastal Ranges are considered part of the Syrian Arc folding, as suggested by Walley (1998), in accordance with the original definition of the Arc (Krenkel, 1924).


The uplift that affected the Coastal Ranges in the latest Cretaceous continued into the Paleogene but was subdued during the Eocene. Middle Eocene marine deposits were deposited throughout the studied area, with the possible exception of the crest of the Coastal Range uplift that may have remained emerged. As discussed above, the second episode of Coastal Range uplift was post-Middle Eocene. This corresponds with the second episode of Syrian Arc development (Guiraud and Bosworth, 1997). Middle Eocene was also a time of uplift of the Palmyrides (Chaimov et al., 1992) (Figure 4.12b), and some minor structural inversion in northeast Syria (Kent and Hickman, 1997). Furthermore, through stratigraphic relationships Dubertret (1975) documented how most of the structuration of the Lebanese mountains was emplaced during the Late Eocene and Oligocene, a view supported by the more recent work of Walley (1998).

These periods of renewed compression within the northern Arabian platform are clearly related to the Mid-Late Eocene final collision of Eurasia and Arabia along the northern Arabian margin (Hempton, 1985). This final obliteration of NeoTethys oceanic crust led to the Bitlis suture that still marks the boundary of these plates (Figure 4.12b). From Mid-Late Eocene time until the Middle Miocene, convergence between the Eurasian and Arabian plates was accommodated by continental margin shortening the thickening along this northern margin (Hempton, 1987). Hence the Syrian Coastal Ranges are shown again to be part of the larger Syrian Arc folding coincident with more regional tectonic development.


The first phase of rifting in the Red Sea area saw continental stretching there that probably started in the Oligocene (Hempton, 1987). From Early Miocene time onwards the differential motion between rifting in the Red Sea and the Gulf of Suez began to be accommodated along the newly formed DSFS (Figure 4.12b,c). Thus ~64 km of sinistral motion occurred on the southern DSFS during this first phase of DSFS movement in Early and Middle Miocene time (Figure 4.12c). As discussed above, the balance of evidence suggests that the northern DSFS had not formed at this time, and the motion was perhaps transmitted offshore along a fault or faults in northern Israel / Lebanon (Figure 4.11c).

In his model, Hempton (1987) argues that by Middle Miocene time the northern margin of Arabia had reached full crustal thickness after shortening and thickening in the Eurasia / Arabia collision. Hempton (1987) suggests that this was therefore the terminal suturing of Eurasia/ Arabia, after which Arabia was unable to converge any further on Eurasia, and so spreading in the Red Sea halted. In turn this led to a cessation of movement along the DSFS (Figure 4.12d). Thus, in the model of Hempton (1987) that we support herein, the first phase of motion on the DSFS came to a close during the Middle Miocene and the DSFS was inactive from around 14.5 Ma until about 4.5 Ma. Interestingly, this time also approximately corresponds to a period no volcanic activity in Syria (Mouty et al., 1992).

Pliocene - Recent

Hempton (1987) goes on to argue that activity on the DSFS commenced again in the Early Pliocene (~4.5 Ma). This was due to commencement of Red Sea seafloor spreading as the northward motion of Arabia was accommodated along the newly formed North and East Anatolian Faults. In accordance with this model, we suggest that with the renewed activity and reoriented stress regime, the DSFS formed its current path though Syria beginning in Early Pliocene time (Figure 4.11e). The balance of our evidence indicates that the Ghab Basin only formed during Pliocene time. This strongly suggests that the northern DSFS only formed since the Miocene, as forwarded by the model of Chaimov et al. (1990) that we support here. Further evidence comes from offsets in Pliocene basalt and Quaternary fans (Trifonov et al., 1991; Fleury et al., 1999), and offsets of ophiolites together with GPS current motion vectors. Preliminary GPS measurements suggest roughly 6 mm/year of relative Africa / Arabia motion in the northern Arabian platform (McClusky et al., 2000), in agreement with field studies (Trifonov et al., 1991). If overall this motion has been constant it indicates ~27 km of movement in the last 4.5 Ma, roughly equivalent to previously suggested totals (Quennell, 1984; Trifonov et al., 1991).

We suggest that after the northern propagation of the DSFS, the Ghab Basin formed owing to the complex splaying left step-over in the sinistral fault system. Cross-basin oblique-slip faults appear to have accommodated the initial extension (Figure 4.11d) that was later transferred onto the basin bounding faults that are still prominent today (Figure 4.10 and 4.11e-f). Despite significant topography to the west, surface and subsurface data show the eastern basin-bounding fault to be the more active, and this fault is continuous north of the basin as readily seen in topography and seismicity data. This suggests an incomplete transfer of lateral motion from the eastern to the western strands of the DSFS across the Ghab Basin. Thus north of the Ghab Basin the DSFS splays out into a broad zone of deformation with lateral motion distributed amongst several faults.

We suggest that during the Pliocene - Recent the Ghab Basin and northern DSFS were superposed on the pre-Pliocene Syrian Coastal Range topography. This faulting along the crest of the Coastal Ranges has created the very steep western flank of the uplift that we observe today. Continued compression of northwest Arabian since the propagation of the DSFS through the Coastal Ranges has caused further uplift to the west of the DSFS.


Geomorphology, stratigraphic relationships, and seismicity clearly demonstrate the active deformation of the northern, Syrian segment of the DSFS. Sinistral movement at a left-step and splaying of the fault has resulted in the Ghab Basin that, absence evidence to the contrary, we interpret to have formed since earliest Pliocene time. Cross-basin oblique-slip faults accommodated some initial basin opening, but most subsidence has occurred along the more active eastern basin-bounding fault. The basin exhibits two asymmetric depocenters with geometry suggestive of some transform-normal extension. The timing of Ghab Basin formation strongly supports a model in which the current northern strand of the DSFS (in Lebanon and Syria) has only been active since the latest Miocene / earliest Pliocene to Recent.

Uplift of the Syrian Coastal Ranges has been episodic since latest Cretaceous time. The first episode of uplift, in the Maastrichtian, was clearly related to plate-wide compression and folding caused by collision along the northern Arabian margin. Mid-Late Eocene uplift was again contemporaneous with regionally observed folding due to final continent-continent collision along the northern margin. This uplift is ongoing, and has been strongly influenced by the formation of the DSFS that has delimited the uplift to the east.


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Figure 4.1: Regional shaded relief image of the Eastern Mediterranean illuminated from the northwest. The Dead Sea Fault System (DSFS) extends from the Gulf of Aqaba to Turkey, as highlighted between the two large arrows on this figure. Numerous flat-bottomed step-over basins along the fault, and significant asymmetric topography on either side of the fault, are clear in this image. The proposed location of the Roum fault (e.g. Butler et al., 1997) is also shown. Locations of traverses shown in Figure 4.9 are indicated. The dashed box marks the extents of Figures 4.2, 4.3 and 4.11. Inset illustrates the regional plate tectonic setting (EAF = East Anatolian Fault).

Figure 4.2: Shaded relief image of the area immediately surrounding the Ghab Basin illuminated from the northwest, location shown in Figure 4.1. The very low relief Ghab Basin is at an elevation of ~170 m, and the marked peak of the Coastal Ranges is 1562 m. Locations of other figures are shown. Seismic reflection profile locations are shown as thin black lines and well locations are illustrated with solid circles. (AP = Asharneh Plain, JW = Jebel Wastani, JZ = Jebel Az-Zawieh, BT = Balou Trough.)

Figure 4.3: Generalized geologic and fault map of the study area; location shown in Figure 4.1. Sense of fault movement was not reported by the original author (Ponikarov, 1966).

Figure 4.4: Transect along the Ghab Basin (see Figure 4.2 for location) showing a seismic reflection profile, density model and associated gravity anomalies. Dashed box on density model illustrates extent of seismic reflection data coverage. Intersections with other seismic reflection profiles are shown as small arrows; different line patterns are used to distinguish different reflectors. Densities shown on model are g/cm3, see text for discussion. Quaternary age deposits form the surface layers along the entire length of the transect. Most of the faults shown have components of both normal and strike-slip fault movement. Deeper structure cannot be constrained with current data (see text).

Figure 4.5: Interpreted migrated seismic profile across the Ghab Basin. See Figure 4.2 for location. Intersections with other seismic reflection profiles are shown as small arrows; different line patterns are used to distinguish different reflectors. Quaternary age deposits are the surfical strata along the whole line. Most of the faults shown have components of both normal and strike-slip fault movement. The fault marked F is the main strand of the DSFS and the major eastern bounding fault of the Ghab Basin, see text for discussion.

Figure 4.6: Transect across the Ghab Basin (see Figure 4.2 for location) showing a seismic reflection profile, density model and corresponding gravity anomalies along transect. Dashed box on density model illustrates extent of seismic reflection data coverage. Intersections with other seismic reflection profiles are shown as small arrows; different line patterns are used to distinguish different reflectors. Densities shown on model are g/cm3, see text for discussion. The faults marked A are the along strike continuation of the main strand of the DSFS (as shown in Figure 4.3). The faults marked B and C have surface expression, as documented by Ponikarov (1966). Most of the faults shown have components of both normal and strike-slip fault movement. Note the required thinning of the crust toward the Mediterranean Basin, as commonly observed long the Levantine margin (e.g. ten Brink et al., 1990), although exact lower crustal structure is indeterminate. Step shown in the Moho is not resolvable in the gravity data.

Figure 4.7: Perspective view from the northwest of the Ghab Basin, looking to the southeast. See Figure 4.2 for location. The top layer represents the topography surrounding the Ghab Basin. For this, and the middle layer, darker shades representing higher levels and illumination is from the northeast. Middle layer is a representation of the base of Ghab Basin sedimentary fill; the slightly angular appearance is a consequence of the gridding process. Lowermost layer shows Bouguer gravity contours (BEICIP, 1975). Contour interval is 2 mGal, bolder lines every 10 mGal. Note the large depocenter in the south of the Basin.

Figure 4.8: Comparison of the Ghab Basin structure with physical and mathematical models, and real examples of strike-slip basins. See text for full discussion. Throughout the figure, crosshatched areas indicate major depocenters and bolder lines indicate faults that are more significant. (a) Fault map in the Ghab Basin and immediate surroundings. See Figure 4.2 for location. These faults have been mapped either from surface observations and geomorphology (gray lines) (Ponikarov, 1966), or from seismic reflection and other interpretations (black lines, this study). Letters G, C, T and R correspond to features also observed in (b), see text for discussion. (b) Fault map from sandbox model of step-over basin, after Dooley and McClay (1997). (c) Numerical model of a step-over basin from Rodgers (1999). (d) Simplified structure map of the Cariaco basin, Venezuela, from Schubert (1982). (e) Fault map from Dead Sea Basin, summarized from several sources (Garfunkel and Ben-Avraham, 1996).

Figure 4.9: (a) Graph showing comparison of topographic profiles across the DSFS. These profiles are the average topography across a 20 km wide swath, locations shown on Figure 4.1. The thin black lines are the modeled regional isostatic response of the lithosphere owing to the formation of the Ghab Basin. Each profile has been approximately aligned relative to the fault.

Figure 4.10: Highly schematic, vertically exaggerated, three-dimensional representation of Ghab Basin. Large arrows show approximate relative movements; the Coastal Ranges block is depicted uplifting while translating southwards.

Figure 4.11: Map showing schematic structural evolution of the Ghab Basin and immediately surrounding regions, location shown in Figure 4.2. Extents of zones illustrated outside the Ghab are somewhat speculative. Legend shown in (a) applies to all maps.

Figure 4.12: Map showing schematic structural evolution of the DSFS in a regional setting. Area illustrated is same as Figure 4.1, modern-day geography shown for reference. Bold lines indicate approximate paleo-plate boundaries, and large arrows indicate approximate motion of Arabia relative to Africa. No attempt is made to illustrate all tectonic events on this map; see Figure 4.11 for more detail for the Ghab region in NW Syria. (EAF = East Anatolian Fault).

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Last updated: January 2001