Seismological constraints on the crustal structure beneath the Zagros Mountain belt (Iran).

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Seismological constraints on the crustal structure beneath the Zagros Mountain belt (Iran).

Denis Hatzfeld, M. Tatar, K. Priestley, M. Ghafori Ashtiany

To cite this version:

Denis Hatzfeld, M. Tatar, K. Priestley, M. Ghafori Ashtiany. Seismological constraints on the crustal

structure beneath the Zagros Mountain belt (Iran).. Geophysical Journal International, Oxford Uni-

versity Press (OUP), 2003, 155 (2), pp.403-410. �10.1046/j.1365-246X.2003.02045.x�. �hal-00110022�

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Seismological constraints on the crustal structure beneath the Zagros Mountain belt (Iran)

Denis Hatzfeld,

¹

Mohammad Tatar,

¹^,²

Keith Priestley

³

and Mohsen Ghafory-Ashtiany

²

1LGIT-Geosciences, CNRS Universite Joseph Fourier, UJF/BP 53, Grenoble Cedex,F-38041,France

2International Institute of Earthquake Engineering and Seismology, PO Box 19395/3913 Tehran, Iran

3Bullard Laboratories, CambridgeCB3 OEZ,UK

Accepted 2003 May 7. Received 2003 April 4; in original form 2002 October 29

S U M M A R Y

The Zagros Mountain belt of western Iran results from the collision of the Arabian and Central Iran continental blocks. The stage of the collision is unclear and the crustal structure of the Zagros is rather poorly known. In this study we investigate the velocity structure of the crust and upper mantle beneath the Ghir region located in the Central Zagros using data collected by a temporary local seismological network including a broad-band instrument. The structures of the sedimentary cover and the upper crystalline crust are estimated from the inversion ofP andStraveltimes of local earthquakes recorded on a dense seismological network. The upper crust consists of an∼11 km thick sedimentary layer (Vp∼4.70 km s⁻¹) above a∼8 km thick upper crystalline crust (Vp ∼5.85 km s⁻¹). The velocity of the lower crust and the depth of the Moho are found using receiver function analysis of teleseismic earthquakes. The lower crystalline crust is unusually slow (Vp∼6.5 km s⁻¹) and∼27 km thick. The upper bound for the total crustal thickness beneath the Ghir region is 46±2 km. A comparison of the thickness of the crystalline crust of the Zagros with available information for the thickness of the crystalline crust of the Arabian Platform shows that, at present, the Zagros has a thinner crust. The current crustal thickness beneath the Zagros is comparable to the pre-collision crustal thickness of the Arabian Platform, suggesting that the Zagros is now in a very early stage of continental collision.

Key words:crust structure, Iran, Zagros.

I N T R O D U C T I O N

The Zagros Mountains (Fig. 1a) are a seismically active fold-and- thrust belt resulting from the collision of the Arabian Plate with the continental crust of Central Iran that began in the Miocene and has continued to the present. At the surface, the Zagros consist of long, linear, asymmetrical folds that form a 200–300 km wide series of ranges extending about 1200 km from eastern Turkey to the Straits of Hormuz. The Zagros contain an almost continuous sequence of shelf sediments, ranging in age from the Paleozoic to the Late Tertiary, deposited on the 1–2 km thick infra-Cambrian Hormuz Salt formation, which lies on a probable Precambrian basement.

Although this basement is not exposed, its age is inferred from

‘exotic’ metamorphic blocks brought to the surface in salt plugs (Haynes & McQuillan 1974).

Geological evidence indicates that the Zagros experienced various tectonic episodes that affected different parts of the belt (Falcon 1974; Stocklin 1977). The building process started during the Upper Cretaceous and was associated with the emplacement of ophiolitic fragments along the line of the Main Zagros Thrust. Throughout the Mesozoic, a very thick layer of sediments was deposited on

the stretched and thinned, subsiding Arabian continental margin (Berberian & King 1981; Stoneley 1981). A second period of com- pressional tectonics started during the Plio-Quaternary and created the uplift and folding of the sedimentary section now observed in the Zagros (Berberian 1995). The sediments are decoupled from the underlying basement at the level of the Hormuz Salt and at higher, younger evaporite horizons (Falcon 1974; Berberian 1981). Folding is more intense toward the Main Zagros Thrust but also affects, in a less intense way, the Persian Gulf sediments (Falcon 1974; Berberian 1981).

Seismicity in the Zagros belt is restricted to the region between the Main Zagros Thrust and the Persian Gulf. Most of the larger earthquakes occur on high-angle reverse planes striking parallel to the trend of the fold axes (Jackson 1980; Jackson & McKenzie 1984;

Ni & Barazangi 1986). Strong earthquakes are thought to occur on ‘blind’ active thrust faults (Berberian 1995) that do not reach the surface. In the Ghir region, microearthquakes primarily occur beneath the sediments in the shallow part of the basement (Tatar et al.2003) as previously suggested by Jackson & Fitch (1981), amongst others. The centroid depths of moderate-size earthquakes throughout the Zagros determined by body wave modelling (Jackson

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404 D. Hatzfeldet al.

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platform Central Iran Zagros

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Figure 1.(a) Map of Iran and surrounding countries showing the main tectonic features. A–B is the location of the cross-section shown in Fig. 4. R is Ryad, S is Shiraz and T is Tabriz. The barbed line denotes the Main Zagros Thrust and the black region shows the location of the volcanic rocks. The Arabian Shield is reported in black and the Arabian interior platform in grey. (b) Map of the seismologic network installed in Zagros Mountains around Ghir. Triangles denote the seismograph stations, the triangle surrounded by a circle denotes the location of the broad-band station. The selected events used for the 1-D inversion for the shallow velocity structure of the crust are shown as open circles.

& Fitch 1981; Bakeret al.1993; Maggiet al.2000) occur between about 8 and 20 km. There is no evidence for any seismicity in the mantle (Maggiet al.2000; Tataret al.2003) and thus no direct evidence of continued present-day subduction.

In this study we describe the first quantitative seismological estimate of the crustal structure of the Zagros. We will use this result and the known crustal structure of the adjacent Arabian Platform to examine the relation between the Quaternary shortening of about 50 km that has folded the sedimentary cover (Falcon 1974) and the amount of shortening that might have affected the crystalline crust.

D AT A A N D R E S U L T S

For 7 weeks between 1997 November and 1998 January, we op- erated a temporary network of 30 portable seismographs around Ghir, a city destroyed by anMs=6.9 earthquake in 1972 (Dewey

& Grantz 1973; Berberian 1995). The seismological network con- sisted of 25 short-period (0.5 s) vertical seismometers, four short- period (0.5 s) three-component seismometers, and a single, three- component, broad-band (20 s) seismometer.

During this period, the network recorded more than 400 local earthquakes. The data logger sampling frequency was 125 Hz; the time was calibrated by a GPS signal every 3 h, allowing an accuracy of better than 0.01 s. The timing uncertainties are estimated to be

Table 1. Velocity structures for the shallow crust obtained with different methods.

Minimum rms 1-D inversion, six layers 1-D inversion, four layers Receiver function

Depth Vp Depth Vp Depth Vp Depth Vp

(km) (km s⁻¹) (km) (km s⁻¹) (km) (km s⁻¹) (km) (km s⁻¹)

−3 4.78±0.44 −3 5.17±0.35

0 5.0 0 4.81±0.35 0 4.66±0.09

9 4.90±0.72

11 6.0 11 5.71±0.41 11 5.84±0.07

13 5.95±0.02

21 6.4 19 5.98±0.37 19 6.13±0.28 19 6.5±0.3

46±2 8.2

about 0.05 s forPand 0.1 s forSarrivals. We took as an initial model (Table 1) the velocity structure of a 5.5 km s⁻¹, 15 km thick, upper crust, overlying a 6.5 km s⁻¹, 20 km thick lower crust by Moazami- Goudarzi (1974). There were 186 events ranging in magnitude from 0.7 to 4.1 which were located using eight or more stations, with rms residual arrival times smaller than 0.1 s, and an azimuthal gap of less than 180^◦ (Fig. 1b). This guarantees an accuracy in epicentre and depth estimates for these events of better than 2 km.

The shallow crustal structure

Most of the 186 well-located microearthquakes occurred between 10 and 14 km depth; none occurred deeper than 20 km (Tataret al.

2003). We used the∼1500S_varrival times recorded on vertical com- ponents to compute a meanVp/Vsratio of 1.77±0.01. The trade- off between the velocity structure and the location of the events is small when both directPandSwaves are recorded on more than eight seismological stations with an azimuthal gap of less than 180^◦, and the earthquake locations are generally relatively stable.

First, we estimate the shallow velocity structure by exploring a wide range of reasonable velocity models for the crust, and mini- mizing the mean rms residual (of bothP- andS-arrival times) for a constant set of earthquakes and arrival times (hereafter called an rms grid search). The thickness of the first layer is allowed to vary from

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6 to 18 km, and its velocity from 4.8 to 5.7 km s⁻¹. The thickness and velocity of the second layer from 13 to 23 km and from 5.5 to 6.4 km s⁻¹, respectively, and the velocity of the third layer from 6.0 to 6.9 km s⁻¹. We obtain a final model (Table 1), which consists of an 11 km thick sedimentary layer withV_p=5.0 km s⁻¹, overlying a 10 km thick layer withVp=6.0 km s⁻¹, and a lower layer with Vp=6.4 km s⁻¹. The residual arrival time is 0.112 s. However, the velocity of the lower crust is not well determined because only a few events were located within or below this layer and very few propagation paths sampled the lower crustal layer. A change in the thickness for one layer of 2 km, or in the velocity of 0.2 km s⁻¹, increases the rms to be greater than 0.120 s. To refine and substantiate the inferred crustal velocity structure, we also perform, using the program VE- LEST (Kissling 1988), a 1-D inversion of the arrival times for the 212 earthquakes that were recorded by a minimum of eight stations with an rms of less than 0.2 s. This routine relocates the earthquakes and inverts for the velocity structure simultaneously. The data set consists of more than 3800P- andS-waves arrival times. We perform several tests to ensure that the 1-D inversion program converge correctly to a unique velocity structure not depending on the starting velocity model. We include several layers, 2 km thick, within the sediments and at the interface between the sediments and the upper crust (at a depth of 11 km). These 2 km thick layers are left free in order to check that our interface located at the base of the sediments does not depend on the starting velocity structure. Starting

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Final models Vp

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Final models Vp

Figure 2.Velocity structure obtained for the shallow crust by inversion of the traveltimes of selected local earthquakes recorded on the seismological network (Fig. 1b). We explore a six-layer velocity structure with 212 events (top) and a four-layer velocity structure with 65 events (bottom), and compute 50 random starting velocity structures from thea priorimodel deduced from rms grid search (left). The resulting models (right) show a good convergence for the velocities of the sedimentary layer (Vp∼4.9±0.15 km s⁻¹) and the upper crust (Vp∼5.85±0.05 km s⁻¹).

with velocity structure obtained by the rms grid search model, we generate a set of 50 different velocity models by introducing random changes of velocities (as large as 0.5 km s⁻¹forVp) in each of the six layers and use these as starting models in the inversion. Not all 1-D inversions converge properly, but 32 inversions converge to a standard deviation rms smaller than 0.5 s, giving an estimate of the uniqueness of the velocity structure (Fig. 2). The mean velocity model of the 32 convergent solutions (Table 1) clearly shows a systematic increase of velocity around 11 km, which we assume to represent the transition between the sediments and the basement.

To reduce the standard deviations and to improve the convergence, we performed a second test restricting the data set to the 65 earthquakes with an rms of less than 0.1 s and merging some of the adjacent layers, at the base of the sediment and at the top of the basement, which were of similar velocities in the previous model. The total number of layers is then reduced to four, and again, we gener- ated 50 random starting models. The 1-D inversion converged much better and, taking into account only the inversions that converge to a standard deviation smaller than 0.3 s, we obtain 34 models. The mean velocity in each layer is given in Table 1 and is very similar to the rms grid search velocity structure and to the six-layer velocity structure. The averaged velocities in the 11 thick sedimentary and upper crystalline layers are about 4.66±0.09 and 5.84±0.07 km s⁻¹, respectively. Both velocities are rather low compared with that observed in other regions.

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Table 2. Parameters of the earthquakes for the receiver function.

Date Time Lat. Long. Depth Mag.

1997 12 5 11 26 54.6 54.841 162.035 33 6.3

1997 12 5 18 48 22.7 53.752 161.746 33 6.2

1997 12 6 10 59 10.0 53.972 161.909 33 5.7

1997 12 7 17 56 18.7 54.658 162.882 33 5.6

1997 12 8 21 6 13.8 53.853 161.778 33 5.4

1997 12 8 22 19 55.7 53.841 161.771 33 5.5

1997 12 20 13 26 31.5 53.424 152.762 614 5.1

Using the same random approach, we also checked for the exis- tence of low-velocity layers (i.e. related to the salt layer at the base of the sediments) and left it free in the 1-D inversion. However, because of the lack of events within or above this layer, we cannot obtain a reliable result for a possible low-velocity layer located at the base of the sediments.

The deep crustal structure

Because we did not record any earthquakes in the lower crust or upper mantle, and because our epicentral distances are generally too small to record refracted waves within the lower crust, we cannot use the arrival times of local or regional earthquakes to infer the thickness and velocity of the lower crust, and therefore define the Moho depth. We use receiver function analysis of seven teleseismic events (Table 2) recorded on the broad-band seismograph to constrain both theS-wave velocity in the lower crust and to deduce the depth to the Moho (e.g. Langston 1979; Owenset al.1984).

The receiver functions are determined from the broad-band seismo- grams using the method of Liggoria & Ammon (1999). Although there is some difference in the sizes, locations and mechanisms of the events, the individual receiver functions are remarkably similar (the observed variability being due to the noise in the data) and ensure that the signal-to-noise ratio is favourable and the signal significant. All individual receiver functions (Fig. 3) show a clear pulse 6–7 s after the directParrival. This pulse is very likely to be due to aPsconversion at the Moho interface. We stack the seven receiver functions to increase the signal-to-noise ratio. Instead of using an inverse method, which could give a biased estimate due to the non-uniqueness of the solution, we use a ‘grid-search inversion’ to determine the structure of the lower crust. We assume the upper crustal velocity structure from the four-layer 1-D inversion discussed above (Table 1) and constrain the mantle velocity to be 8.2 km s⁻¹beneath the Zagros Mountains (Asudeh 1982b). Then we assume a velocity and a thickness for the single-layer lower crust only and compute the misfit (Wallaceet al.1981) between the synthetic receiver function for this crustal model and the stack of the seven observed receiver functions for 10 s duration (Fig. 3). Using this technique, we explore 128 crustal models with a lower crust for which the thickness varies from 21 to 36 km and for which the velocity varies from 6.3 to 7.0 km s⁻¹, assuming aVp/Vsratio of 1.732. There is a clear minimum in the misfit for a 27±2 km thick lower crust of 6.5±0.3 km s⁻¹as shown in Figs 4 and 5.

Fig. 5 shows the effect of a±3 km variation in the velocity of the crustal thickness (Fig. 5a) and±0.2 km s⁻¹ variation in the velocity of the lower crust (Fig. 5b) on the amplitude and timing of thePsphase. Fig. 5 indicates that the lower crustal velocity and Moho depth are reasonably well constrained by thePsarrival time and amplitude. Receiver function analysis is poor in retrieving layer velocities (Ammonet al.1990). The procedure we use does not

Time (sec) Ps

Figure 3. Receiver functions for the 1997 Kamchatka 7 earthquakes (Ta- ble 2). Receiver functions were determined using the iterative deconvolution procedure of Liggoria & Ammon (1999) using 100 iterations and a Gaus- sian smoothing factor of 2.5. All the seven receiver functions show a clear Psconversion 6–7 s behind theParrival, which is probably related to the Moho. The bottom trace is the mean receiver function and the dashed lines enclosing are the±one standard deviation.

provide unique details of the final-scale structure of the lower crust, but it gives the simplest structure for the lower crust consistent with the observed receiver functions. AVp/Vsratio of 1.77 instead of 1.732 would introduce an additional error of 0.5 km in depth and of 0.1 km s⁻¹in velocity.

In summary, the rms grid search and 1-D inversion of local traveltimes first provide an estimate of the upper crustal velocity structure, whereas the receiver function analysis gives an estimate of the Moho depth.

D I S C U S S I O N

Summary of the crustal structure

Our crustal model is for a region located in the Central Zagros near Ghir and may not represent a cross-section of the entire Zagros if important lateral changes exist. The crust beneath this region is 46±2 km thick and consists of an 11 km thick sedimentary layer overlying a 35±2 km thick crystalline layer. The crystalline crust consists of two layers: an upper layer extending from about 11 to about 19 km depth and a lower layer extending from about 19 km to the Moho at 46 km depth. The velocity estimates are likely to be computed with an accuracy of±0.2 km s⁻¹, which ensures that we properly identify the different interfaces between the sediments, the metamorphic crust and the upper mantle. The velocity of the upper crystalline crust (∼5.85 km s⁻¹) is much lower than that of

‘normal’ continental crust (Christensen & Mooney 1995), perhaps indicating that it has a more silica-rich composition. This thickness, given with an uncertainty of about 2 km, is the first estimate of the

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0.09

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0.09

0.1 0.11 0.12 0.13 0.14

Lower crustal Vp (km s

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)

Lower crustal thickness (km)

28

26

24

22 30 32 34 36

30

22 24 26 28 32 34 36

6.30 6.40 6.50 6.60 6.70 6.80 6.90 7.00

3.64 3.70 3.75 3.81 3.87 3.93 3.98 4.04

Lower crustal Vs (km s

⁻¹

)

Figure 4.Grid search of the minimum misfit receiver function. We explore a lower crust with aP-wave velocity ranging from 6.3 to 7.0 km s⁻¹and aS-wave velocity ranging from 3.64 to 4.04 km s⁻¹(assuming aVp/Vsratio of 1.732) and a thickness ranging from 21 to 36 km. The thickness is determined to be 27±2 km (although there is a second minimum at 22 km) and the crustal velocity to be 6.5±0.3 km s⁻¹. The first 10 s of the receiver function waveform were fit in the grid search. This time window contains the main crustal arrivals but not the low-amplitude crustal multiples.

crustal velocity structure beneath the Central Zagros deduced from seismological observations other than an average value from surface waves analysis. Considering the high topography of the Zagros (from 1500 to 2000 m along the cross-section), and the thickness of the sedimentary cover, this suggests a rather thin metamorphic crust.

Comparison with existing data

Very little is known concerning the crustal velocity structure of the Zagros and especially concerning the Moho depth. Pn velocities have been computed between pairs of stations for regional earthquakes (Chenet al.1980; Kadinsky-Cadeet al.1981; Asudeh 1982b; Hearn & Ni 1994). Most of these studies rely on data reported in the ISC bulletins for the path between the seismographs at Tabriz and Shiraz. This path, about 800 km long, is located along the Main Zagros Thrust, and crosses several strike-slip faults. It may represent an average of the velocity structure of the Iranian Plateau and of the Zagros. The meanPnvelocity is about 8.1 km s⁻¹, but there is no information on the crustal thickness. Surface wave analysis (Asudeh 1982a) along the Tabriz and Shiraz path implies a Moho depth of 46 km and a relatively fast upper-mantle velocity of 8.3 km

s⁻¹is observed, but that may not represent the crustal structure of the Central Zagros.

A refraction profile consisting of sparse recordings along a line from Central Iran to the Strait of Hormuz (Gieseet al.1983) shows questionable quality arrivals indicating, if they are Moho reflections, a crustal thickness of 40 km beneath Central Iran. Two reflectors were found at 20–25 km and 60 km depth beneath the Main Za- gros Thrust. However, these features are very poorly constrained, because of uncertainties in the arrival times, because they are not reverse profiles and because they are oblique to the main geological structures. Moreover, they are located beneath the volcanic belt that may be the site of the major discontinuity between the Arabian Plate and Central Iran.

Gravity surveys (Dehghani & Makris 1984; Snyder & Barazangi 1986) provide additional constraints on the Zagros crustal structure. The crustal model, deduced from Bouguer gravity observations (Snyder & Barazangi 1986), is characterized by a Moho at 40 km depth beneath the Persian Gulf, dipping about 1^◦northward beneath the folded belt and 5^◦ beneath the Main Zagros Thrust, reaching a depth of 65 km. The maximum Moho depth is slightly shifted northeastward relative to the highest topography. This crustal model for the Zagros gives a Moho depth of about 45–47 km beneath

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Time (sec)

Time (sec) (a)

(b)

Figure 5. Bounds of the lower crustal structure. (a) The heavy solid line is the minimum misfit synthetic receiver function from the grid search for a lower crustal thickness of 27 km, the two dotted lines are the±one standard deviation for the receiver function stack, and the dashed lines are synthetic receiver functions for a lower crust with the same velocity but a thickness of 27±3 km. (b) Same format but the dashed lines are for a lower crust 27 km thick and a velocityVpof 6.5±0.2 km s⁻¹. The lower crustal thickness is better constrained than the velocity as it was clear from Fig. 4.

Ghir. To explain the negative isostatic residual anomaly, Snyder &

Barazangi (1986) suggest that elastic flexure, due to both vertical and horizontal forces, takes place. However, they assume the upper crustal density to be 2.61 g cm⁻³, an average between the limestone and the salt densities, and the lower crust density to be 2.95 g cm⁻³. The low velocities we obtain from our seismic analysis suggest that the densities may be overestimated. This could partially explain the negative isostatic residual anomaly that was modelled by elastic flexure.

Comparison with surrounding regions

The Zagros Mountain belt lies on the northwest margin of the Ara- bian Platform, which collided with the continental blocks in Central Iran during the Quaternary. There are two important issues concerning the collision process: (1) whether the∼50 km of shortening since the Pliocene, which folded the sedimentary cover, similarly affected the crystalline crust and (2) whether the reverse faulting mechanisms observed for the larger earthquakes are the basement expression of the shortening. We can estimate the amount of total thickening resulting from the collision by comparing the crustal

structure of the Zagros Mountains with what is known of the crustal structure of the surrounding region.

The crustal structure of the Arabian Shield is known from seismic refraction experiments (e.g. Mooneyet al.1985), gravity and aeromagnetic observations (Gettingset al. 1986), receiver function analysis (Sandvolet al.1998; Juliaet al.2000), and surface wave analysis (Rodgerset al.1999). The crustal thickness varies from about 35 km near the Red Sea to 40 km near the NE edge of the shield and this region has only a thin sedimentary cover. The crust consists of an upper layer 16–20 km thick with aP-wave velocity of 6.2–6.4 km s⁻¹and a lower layer of similar thickness with a velocity of 6.9–7.0 km s⁻¹. The upper mantleP_nvelocity is about 7.9 km s⁻¹.

The velocity structure of the undeformed part of the Arabian Platform is less well known. Spectral analysis of long-periodP-wave amplitude ratios for earthquakes from the Eastern Mediterranean recorded in Ryad (Al-Amri 1999), regional waveform modelling of Zagros events recorded at seismological stations located on the Arabian Shield (Rodgers et al.1999), and receiver functions for events recorded at Ryad (Sandvol et al.1998; Julia et al.2000) show the crust to be thicker (∼45 km) than that of the shield and

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overlain by a sedimentary cover that is about 3 km thick. The mean velocity is probably slightly slower than that of the Arabian Shield (Rodgerset al.1999) and the crust overlies a faster upper mantle.

The shallow velocity structure in the Arabian Platform is consistent with results obtained in the Persian Gulf (Rosset al.1986) except that the sediment beneath the Gulf is about 7 km thick.

Tentative interpretation

Comparing the Zagros crustal structure with that of the surrounding regions shows that the sedimentary layer thickens from 3 km beneath the Arabian Platform, to 7 km beneath the Persian Gulf and to 11 km beneath the Zagros fold belt. The depth of the Moho is 36–40 km beneath the Arabian Shield, 45 km beneath the Arabian Platform, 44–48 km beneath the Zagros Mountains and∼40 km beneath Central Iran. Thus, the crystalline crust has a nearly constant thickness of 40–42 km across the Arabian Platform and 35 km thickness beneath the Zagros fold belt (Fig. 6). This result is quite different from the cross-section proposed by Seberet al.(1997) in the sense that we do not observe a significant crustal thickening beneath the Zagros Folded Belt. Our result is similar to the result of Snyder & Barazangi (1986) at Ghir and calibrates their cross-section deduced from gravity observations.

However, the crustal thickness of the leading edge of the Arabian Platform, which now underlies the Zagros, was subjected to rifting and thinning in the Permo-Triassic and again in the Jurassic when the region formed an Atlantic-like passive continental margin. To understand the effect of the continental collision on the crust, it is the thinned pre-collision crustal thickness we should compare with the present-day crustal thickness beneath the Zagros and not the current crustal structure of the undeformed Arabian Platform to the southwest. Trowell (1995) deduced the tectonic history of the Arabian Platform margin from back-stripped and decompacted

Arabian shield Arabian platform

-1000 -500 0 500 km

Moho

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Volcanic belt

Persian Gulf Zagros folded belt Central Iran

Main Front Fault Main Zagros Thrust

NW SE

50

depth

crust sediments 6.2

6.9

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6.4

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6.5

8.2 8.3

6.0 ?

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Figure 6.Synthetic cross-section (see Fig. 1a for location) from the Arabian Shield to Central Iran. We report the mean values for the thickness (hatched segments) and the velocity (numbers) of the sedimentary and the crustal layers from various sources (see text). We report in bold characters, over a grey column, our results in the Central Zagros. All of these values are of various accuracies but are consistent enough to infer a smooth topography of the Moho. Beside the reflectors at 20–25 and 60 km depth beneath the volcanic belt from Gieseet al.(1983), which are not well constrained, the thickness of the crystalline crust is almost constant across the Arabian Platform and Zagros Mountains. The dashed line is a tentative interpretation suggesting that there is no significant shortening and thickening of the crystalline crust to match the shortening observed for the folded belt at the surface in the Zagros Mountains.

stratigraphic data from 82 wells along the southern Zagros margin and found a stretching factor ofβ∼ 1.2 similar to theβ values found across the Tethyan passive margin further to the west (Wooler et al.1992). Using the thickness of∼42 km of the crystalline crust at Ryad (Sandvolet al.1998) as the unstretched value for the Ara- bian Platform, and applying thisβfactor of 1.2, the thickness of the crystalline crust of the leading edge of the Arabian Platform might be∼35 km following the stretching and prior to the collision. Com- paring this thickness for the crystalline crust prior to the collision and the present-day thickness∼35±2 km of the crystalline crust beneath the Zagros suggests that the crust has undergone very little, if any, thickening due to shortening. In comparison, the folded shallow sediments at the surface suggest about∼15 per cent (or

∼50 km) shortening across the Zagros (Falcon 1974).

This is consistent with the hypothesis that the Zagros Mountains are in the very early stage of continental collision because the shortening has yet to thicken the crystalline basement. Future shortening will thicken the crust further, creating a more substantial crustal root similar to that seen in more developed mountain ranges such as the Himalayan Range (Ni & Barazangi 1986). The difference between the shortening of the sediments, as seen in the folding, and that of the basement shortening (related to the seismicity and crustal thickness) suggest that the folding (associated with the 50 km of sediment shortening) could be due to the earlier scrapping of the sediments decoupled from the top of the crystalline crust by the salt layer (the Hormuz sequence) at the base of the sediments.

C O N C L U S I O N

Using arrival times of local events recorded on a dense seismological network, we infer the upper-crust velocity structure to be composed of an 11 km thick sedimentary layer and a 8 km thick upper crystalline crust. Receiver functions analysis of seven teleseismic earthquakes suggests a 27± 2 km thick lower crust of

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6.5 km s⁻¹velocity. These estimates, although obtained with a small amount of data, are the first seismological quantitative estimates in Central Zagros. The total thickness (∼35 km) of the crystalline crust therefore looks similar to the thickness of the stretched margin of the Arabian Platform and suggests that the shortening of the Zagros basement is small and has only started recently, whereas the shortening demonstrated by the folded sediments is due to the long-term scraping of the sediments above the basement.

A C K N O W L E D G M E N T S

We thank the observers and drivers who helped us in the field. KP acknowledges visiting support from INSU-CNRS. We thank J. Jack- son for useful discussions and A. Paul for helping in computation.

This research is supported by the French Embassy in Teheran. M.

Tatar benefited of a fellowship from the French Ministry of Foreign Affairs. We thank D. Snyder, B. Parsons and an anonymous reviewer for helpful reviews.

R E F E R E N C E S

Al-Amri, A.M.S., 1999. The crustal and upper-mantle structure of the interior Arabian platform,Geophys. J. Int.,136,421–430.

Ammon, C.J., Randall, G.E. & Zandt, G., 1990. On the resolution and non- uniqueness of receiver function inversions,J. geophys. Res.,95,15 303–

15 318.

Asudeh, I., 1982a. Seismic structure of Iran from surface and body wave data,Geophys. J. R. astr. Soc.,71,715–730.

Asudeh, I., 1982b.Pnvelocities beneath Iran,Earth planet. Sci. Lett.,61, 136–142.

Baker, C., Jackson, J. & Priestley, K., 1993. Earthquakes on the Kazerun Line in the Zagros Mountains of Iran: strike-slip faulting within a fold- and-thrust belt,Geophys. J. Int.,115,41–61.

Berberian, M., 1981. Active faulting and tectonics of Iran, in,Zagros–Hindu–

Kush–Himalaya Geodynamic evolution,Vol. 3, pp. 33–69, eds Gupta, H.K. & Delany, F.M., Am. Geophys. Union, Geodyn. Ser.

Berberian, M., 1995. Master blind thrust faults hidden under the Zagros folds:

active basement tectonics and surface morphotectonics,Tectonophysics, 241,193–224.

Berberian, M. & King, G.C.P., 1981. Towards a paleogeography and tectonic evolution of Iran,Can. J. Earth Sci.,18,210–265.

Chen, C.-Y., Chen, W.-P. & Molnar, P., 1980. The uppermost mantlePwave velocities beneath Turkey and Iran,Geophys. Res. Lett.,7,77–80.

Christensen, N.I. & Mooney, W.D., 1995. Seismic velocity structure and composition of continental crust: a global view,J. geophys. Res.,100, 9761–9788.

Dehghani, G.A. & Makris, J., 1984. The gravity field and crustal structure of Iran,N. Jb. Geol. Palaont. Agh.,168,215–229.

Dewey, J.W. & Grantz, A., 1973. The Ghir earthquake of April 10, 1972 in the Zagros mountains of southern Iran: seismotectonic aspects and some results of a field reconnaissance,Bull. seism. Soc. Am.,63,2071–2090.

Falcon, N., 1974. Southern Iran: Zagros Mountains, inMesozoic–Cenozoic Orogenic Belts,ed. Spencer, A.M.,Spc. Publ. Geol. Soc. London,4,199–

211.

Gettings, M.E., Blank, H.R., Mooney, W.D. & Healey, J.H., 1986. Crustal structure of Southwestern Saudi Arabia,J. geophys. Res.,91,6491–6512.

Giese, P., Makris, I., Akashe, B., Rower, P., Letz, H. & Mostaanpour, M., 1983. Seismic crustal studies in southern Iran between Central Iran and the Zagros belt,Geol. Surv. Iran, Rep.,51,71–84.

Haynes, S.J. & McQuillan, H., 1974. Evolution of the Zagros suture zone, southern Iran,Bull. geol. Soc. Am.,85,739–744.

Hearn, T.M. & Ni, J.F., 1994.Pn velocities beneath continental collision zones: the Turkish–Iranian Plateau,Geophys. J. Int.,117,273–283.

Jackson, J.A., 1980. Reactivation of basement faults and crustal shortening in orogenic belts,Nature,283,343–346.

Jackson, J. & Fitch, T., 1981. Basement faulting and the focal depths of the larger earthquakes in the Zagros mountains (Iran),Geophys. J. R. astr.

Soc.,64,561–586.

Jackson, J.A. & McKenzie, D., 1984. Active tectonics of the Alpine–

Himalayan Belt between western Turkey and Pakistan,Geophys. J. R.

astr. Soc.,77,185–264.

Julia, J., Ammon, C.J., Herrmann, R.B. & Correig, A.M., 2000. Joint inversion of receiver function and surface wave dispersion observations, Geophys. J. Int.,143,1–19.

Kadinsky-Cade, K., Barazangi, M., Oliver, J. & Isacks, B., 1981. Lateral variations of high-frequency seismic wave propagation at regional distances across the Turkish and Iranian plateau,J. geophys. Res.,86,9377–

9396.

Kissling, E., 1988. Geotomography with local earthquake data,Rev. Geo- phys.,26,659–698.

Langston, C.A., 1979. Structure under Mount Rainier, Washington, inferred from teleseismic body waves,J. geophys. Res.,84,4749–4762.

Liggoria, J.P. & Ammon, C., 1999. Iterative deconvolution and receiver function estimation,Bull. seism. Soc. Am.,89,1395–1400.

Maggi, A., Jackson, J.A., Priestley, K. & Baker, C., 2000. A re-assessment of focal depth distributions in southern Iran, Tien Shan and northern India:

do earthquakes really occur in the continental mantle?,Geophys. J. Int., 143,629–661.

Moazami-Goudarzi, K., 1974. La vitesse de phase des ondes de Rayleigh et la structure de la croˆute et du manteau sup´erieur entre Machad et Chiraz (Iran),Pageophys.,112,675–681.

Mooney, W.D., Gettings, M.E., Blank, H.R. & Healy, J.H., 1985. Saudi- Arabian seismic-refraction profile: a traveltime interpretation of crustal and upper mantle structure,Tectonophysics,111,173–246.

Morris, P., 1977. Basement structure as suggested by aeromagnetic surveys in S.W. Iran,Internal Report,Oil Serv. Co. of Iran, Tehran, Iran.

Ni, J. & Barazangi, M., 1986. Seismotectonics of the Zagros continental collision zone and a comparison with the Himalayas,J. geophys. Res.,91, 8205–8218.

Owens, T.J., Zandt, G. & Taylor, S.R., 1984. Seismic evidence for an ancient rift beneath the Cumberland Plateau, a detailed analysis of broadband teleseismicPwaveforms,J. geophys. Res.,89,7783–7795.

Rodgers, A.J., Walter, W.R., Mellors, R.J., Al-Amri, A.M.S. & Zhang, Y.- S., 1999. Lithospheric structure of the Arabian Shield and Platform from complete regional waveform modelling and surface wave group velocities, Geophys. J. Int.,138,871–878.

Ross, D.A., Uchupi, E. & White, R.S., 1986. The Geology of the Persian Gulf–Gulf of Oman Region: a synthesis,Rev. Geophys.,24,537–556.

Sandvol, E., Seber, D., Barazangi, M., Vernon, F., Mellors, R. & Al-Amri, A., 1998. Lithospheric seismic velocity discontinuities beneath the Arabian Shield,Geophys. Res. Lett.,25,2873–2876.

Seber, D., Vallv´e, M., Sandvol, E., Steer, D. & Barazangi, M., 1997. Middle East tectonics: applications of geographical information systems (GIS), GSA Today,7,1–6.

Snyder, D.B. & Barazangi, M., 1986. Deep crustal structure and flexure of the Arabian Plate beneath the Zagros collisional mountain belt as inferred from gravity observations,Tectonics,5,361–373.

Stocklin, J., 1977. Structural correlation of the Alpine range between Iran and Central Asia,Mém. H. sér. Soc. Géol., Fr.,8,33–353.

Stoneley, R., 1981. The geology of the Kuh-e Dalneshin area of southern Iran, and its bearing on the evolution of southern Tethys,J. geol. Soc.

Lond.,138,509–526.

Tatar, M., Hatzfeld, D. & Ghafory-Ashtiany, M., 2003. Tectonics of the Cen- tral Zagros (Iran) deduced form microearthquakes seismicity,Geophys.

J. Int.,in press.

Trowell, C.G., 1995. Lithospheric stretching, subsidence and magmatism in Oman and the Middle East,PhD thesis,University of Cambridge.

Wallace, T.C., Helmberger, D.V. & Mellman, G.R., 1981. A technique for the inversion of regional data in source parameters studies,J. geophys.

Res.,86,1679–1685.

Wooler, D.A., Smith, A.G. & White, N., 1992. Measuring lithospheric stretching on Tethyan passive margins,J. geol. Soc. Lond.,149,517–

532.

C2003 RAS,GJI,155,403–410