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Lithospheric structure in Central Eurasia derived from 3
elevation, geoid anomaly and thermal analysis 4
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Alexandra M. M. Robert1,2*, Manel Fernàndez1, Ivone Jiménez-Munt1 and Jaume 7
Vergés1 8
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* Corresponding author: A. M. M. Robert, Géosciences Environnement Toulouse, 12
UMR CNRS-IRD-Université de Toulouse, 14 avenue Edouard Belin, 31400 Toulouse, 13
France (alexandra.robert@get.obs-mip.fr) 14
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1 Group of Dynamics of the Lithosphere (GDL), Institute of Earth Sciences Jaume 17
Almera, ICTJA-CSIC, Lluís Solé i Sabarís s/n, 08028 Barcelona, Spain.
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2 Géosciences Environnement Toulouse, UMR CNRS-IRD-Université de Toulouse, 19
14 avenue Edouard Belin, 31400 Toulouse, France 20
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Abstract 24
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We present new crustal and lithospheric thickness maps for Central Eurasia from the 26
combination of elevation and geoid anomaly data and thermal analysis. The results are 27
strongly constrained by numerous previous data based on seismological and seismic 28
experiments, tomographic imaging and integrated geophysical studies. Our results indicate 29
that high topography regions are associated with crustal thickening being maximum below 30
the Zagros, Himalaya, Tien Shan and the Tibetan Plateau. The stiffer continental blocks 31
remaining undeformed within the continental collision areas are characterized by a slightly 32
thickened crust and flat topography. Lithospheric thickness and crustal thickness show 33
different patterns that highlight an important strain partitioning within the lithosphere. The 34
Arabia/Eurasia collision zone is characterized by a thick lithosphere underneath the Zagros 35
belt whereas a thin to inexistent lithospheric mantle is observed beneath the Iranian and 36
Anatolian plateaus. On the contrary, the India/Eurasia collision zone is characterized by a 37
very thick lithosphere below its southern part as a consequence of the underplating of the 38
cold and stiff Indian lithosphere. Our new model presents great improvements compared to 39
previous global models available for the region and allows us to discuss major aspects 40
related to the lithospheric structure and acting geodynamic processes in Central Eurasia. 41
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The Central Eurasia region contains two of the most prominent deformed regions on Earth: 43
the Arabia/Eurasia and the India/Eurasia collision zones, forming most part of the well-44
studied Alpine-Himalayan belt. Defining the lithospheric structure of such large region 45
including areas with very scarce information will provide important contributions to 46
understand the relevant geodynamic processes occurring in collisional contexts. 47
During the last decades, the Himalaya-Tibet and Zagros-Iran regions have been the target of 48
numerous geophysical surveys (mostly receiver functions, deep seismic and tomographic 49
studies) to unravel their lithospheric structures. Crustal thickness data include 1D spotted 50
estimations (e.g. Nasrabadi et al., 2008; Chen et al., 2010), 2D transects across the Himalaya-51
Tibet region (e.g. Kind et al., 2002; Tian et al., 2005; Nabelek et al., 2009) or the Zagros-Iran 52
region (e.g. Paul et al., 2006, 2010) and 3D regional studies (e.g. Zör et al., 2003; Pan & Niu, 53
2011). However, some other areas of Central Eurasia are less well known due to the lack of 54
seismological and seismic studies. In fact, less than 19% of the region has crustal thickness 55
data coverage better than 1 measure per 50,000 km2 (i.e. one point every 2x2 arc-degree). 56
Furthermore, results from these geophysical studies show a high scatter related to different 57
surveys, instrumentation, recording periods and methodology. Some global crustal thickness 58
models are proposed and they have provided general features of the crustal structures with a 59
spatial resolution up to 1x1 arc degree (e.g. Nataf & Ricard, 1996; Mooney et al., 1998; 60
Bassin et al., 2000; Laske et al., 2013). Lithospheric thickness estimations are even scarcer 61
and they are highly variable as a function of the used methodology. Regional tomography 62
models have been developed to image the upper mantle structure (e.g. Ritzwoller and 63
Levshin, 1998; Villaseñor et al., 2001) and only few receiver functions studies are able to 64
image the base of the lithosphere (e.g. Ramesh et al., 2010; Zhao et al., 2010). 65
Jiménez-Munt et al. (2012) presented a detailed image of the crust and lithospheric mantle 66
variations in the Arabia-Eurasia collision zone based on the combination of geoid and 67
elevation data and thermal analysis. In the present study we use a similar approach with the 68
aim to map the Crust-Mantle Boundary (CMB) and the Lithosphere-Asthenosphere Boundary 69
(LAB) over the entire Central Eurasia region that we compare with available data. This study 70
therefore incorporates an up-to-date compilation of crustal thickness data from seismological 71
and seismic studies including 1212 estimations coming from more than 100 references used as 72
main constraints. We discuss our resulting crustal and lithospheric thickness maps with 73
previous global crustal and lithosphere thermal models (e.g. Laske et al., 2013, Reguzzoni et 74
al., 2013; Artemieva, 2006; Goutorbe et al., 2011) and we shortly discuss relevant 75
geodynamic processes occurring in the Central Eurasia region. 76
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Geodynamic context 78
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During Cenozoic times, the Central Eurasia region has been affected by the India/Eurasia 80
and Arabia/Eurasia continental collisions. In both cases, collision occurred between the 81
strong and resistant Archean-to-Proterozoic shields of Indian and Arabian plates, and the 82
weaker southern margin of Eurasia. The weakness of the Eurasian margin is interpreted as a 83
result of the presence of major pre-existing structures as suture zones and/or large-scale fault 84
zones between the different accreted Gondwana-derived continental blocks (e.g. Audet & 85
Bürgmann, 2011). As a consequence of these collisions, two major topographic features 86
arised: the Zagros and the Himalaya orogenic belts, which show different stages of 87
development due to differences in amounts of convergence and ages of initiation of the 88
collision (Hatzfeld & Molnar, 2010). 89
For the sake of simplicity, we differentiate the region into 5 major tectonic subdivisions: (1) 90
the Arabian Plate, (2) the Indian Plate, (3) the Arabia/Eurasia collision zone, (4) the 91
India/Eurasia collision zone and (5) the region to the North of the Tethysides (Figure 1). 92
The Arabian Plate 93
The Arabian Plate is composed of the Arabian shield (Late Proterozoic basement) 94
unconformably overlapped to the East by the Phanerozoic Arabian platform and separated 95
from Africa by young spreading centers located within the Red Sea and the Gulf of Aden 96
(Al-Damegh et al., 2005). The Arabian platform consists of thick Paleozoic and Mesozoic 97
sediments dipping NE-ward and reaching more than 10 km thickness (Mokthar et al., 2001). 98
The Red Sea marks the boundary between Africa and Arabia and consists of both oceanic 99
and thinned continental crust (Al-Damegh et al., 2004) covered by up to 5 km thick 100
Cenozoic sediments according to the Exxon Tectonic Map of the World (1994). 101
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The Indian Plate 103
The Indian Plate is composed of several Early to Late Archean cratons delimited by rift 104
zones containing Proterozoic and/or Phanerozoic sediments (Naqvi & Rogers, 1987). The 105
massive basalts of the Deccan volcanic province erupted during Late Cretaceous to Early 106
Tertiary and are considered as a consequence of the separation of India from the Seychelles 107
microcontinent (a Gondwana-derived continental block) under the influence of the Réunion 108
mantle plume (Mahoney, 1988). The break-up of Gondwanaland occurred around 140 Ma 109
(earliest Cretaceous) and corresponds to the beginning of the northern drift of the Indian 110
Plate towards Eurasia (Kumar et al., 2007). 111
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The Arabia-Eurasia collision zone 113
The Zagros orogen is the consequence of the closure of the Neotethys Ocean that was 114
located between Arabia and Eurasia (e.g. Sengör et al., 1988; Agard et al., 2011). Though the 115
timing of the continental collision is still controversial, the final closure of the Neotethys 116
Ocean seems to occur during the Late Oligocene-earliest Miocene (e.g. Agard et al., 2011, 117
McQuarrie & van Hinsbergen, 2013; Mouthereau et al., 2012; Vergés et al., 2011). The 118
Zagros orogenic system continues to the South-East to Makran and its North-West 119
continuation corresponds to South-East Anatolia where the collision started earlier than in 120
the Zagros (Agard et al., 2011). Subduction-related volcanism continues to occur throughout 121
both the Iranian and the Turkish plateaus (Kazmin et al., 1986) suggesting that, despite the 122
change from subduction to continental collision, the tectonics of these plateaus is still driven 123
by subduction processes within the upper mantle (Hearn & Ni, 1994; Agard et al., 2011). 124
The Alborz and Kopet Dagh ranges developed along the Paleotethys suture zone (Sengör et 125
al., 1988; Robert et al., 2014) and they correspond to the northern boundary of the Arabia-126
Eurasia collision zone. The Alborz is bounded to the North by the South Caspian Sea Basin, 127
which is characterized by a sedimentary basin exceeding 20 km in sediment thickness 128
(Brunet et al., 2003) (Figure 2). 129
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The India-Eurasia collision zone 131
The India-Eurasia collision zone hosts the largest and highest topographic feature on Earth: 132
the Tibetan Plateau, which was built by sequential accretion of lithospheric terranes upon the 133
southern margin of Eurasia since the end of the Paleozoic (Yin & Harrison, 2000; Pubellier 134
et al., 2008). From North to South, the Tibetan Plateau is successively formed by (1) the 135
Kunlun terrane, (2) the Songpan Garze flysch complex related to the closure of the 136
Paleotethys, (3) the continental Qiangtang terrane accreted at the end of the closure of the 137
Paleotethys and (4) the Lhasa terrane accreted during Cretaceous times (Figure 1). The Indus 138
Tsangpo suture zone, related to the closure of the Neotethys Ocean, is located to the South of 139
the Lhasa Block and to the North of the Himalaya. The timing of the India/Eurasia collision 140
is still subject to controversies but most authors agree that it occurred between 55 and 45 Ma 141
during the Eocene (e.g. Rowley, 1996; Zhu et al., 2005). 142
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North of the Tethysides 144
The North of the Tethysides is a region composed from East to West by the Ordos block, the 145
Altaid range (Figure 1), the Tien Shan range, the Kazakh terranes, the Turan platform and 146
the North Caspian Basin, the Black Sea and the East European Craton (Figure 1). The Altaid 147
range has a complex geodynamic evolution made of multiple accretions of terranes of 148
different origin, chiefly microcontinents and island arcs occurring from 600 Ma to 250 Ma 149
(latest Precambrian to earliest Triassic times) (Wilhem et al., 2012). Numerous sedimentary 150
basins were developed in this area as the Junggar Basin or the North Caspian Basin, which 151
presents an up to 20 km thick sedimentary sequence. 152
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Methodology and input data 154
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Method 156
We calculate the crustal and the lithospheric mantle thickness by combining elevation and 157
geoid data together with thermal analysis in a 1D approach (Fullea et al., 2007). This 158
method assumes local isostasy with a depth of compensation below the lithosphere-159
asthenosphere boundary (LAB) and considers four layers: sea water, crust, lithospheric 160
mantle and asthenosphere. The crustal density increases linearly with depth between 161
predefined values at surface and at the base of the crust. The lithospheric mantle density 162
(ρm) is considered to be temperature dependent: 163
ρm(z)= ρa(1+α(Ta -T(z))) 164
where ρa is the density of the asthenosphere considered constant everywhere, α is the 165
thermal expansion coefficient, Ta is the temperature at the LAB and T(z) is the temperature 166
of the lithospheric mantle at a given depth z. 167
The geoid anomaly is calculated relative to a reference level (H0), which in turn depends on 168
a selected reference lithospheric column with a known crustal thickness and elevation. It 169
can be demonstrated that knowing the crustal thickness and considering an average crustal 170
density and heat production, we can calculate the lithospheric thickness that fits the 171
elevation corresponding to the selected reference column (see Fullea et al., 2007, for 172
details). The lithospheric reference column for geoid anomaly has been chosen within the 173
Indian shield where consistent crustal and lithospheric thicknesses are obtained from 174
seismic and seismological studies (e.g. Bhattacharya, 2009; Kumar et al., 2007; Ramesh et 175
al., 2010). This method was successfully applied to decipher crustal and lithospheric 176
structures in the Gibraltar arc system (Fullea et al., 2007) and in Arabia/Eurasia collision 177
zone (Jiménez-Munt et al., 2012). 178
Input parameters used in our modeling are summarized in the Tables 1 and 2. According to 179
the large extent of the region considered, lateral variations of the surface density were used 180
in our modeling (Figure 2). This 2D density distribution takes into account the occurrence of 181
large sedimentary basins, as well as small differences between the considered tectonic 182
subdivisions (Table 2). Because the crust is defined as a simple layer with a linear depth-183
density increase, the occurrence of thick sedimentary basins is associated with a decrease of 184
the surface crustal density. First, we obtained the sediment thickness distribution by 185
digitizing the Exxon Tectonic Map of the World (1994), complemented with some local 186
studies (Singh, 1996; Sastri et al., 1971; Rao, 1973; Karunakaran & Rao, 1979) (Figure 2). 187
And then, we calculated the surface density according to: 188
ρsurf = 2700 - 100 × (1-exp (-H/5)) 189
with ρsurf being the crustal density at surface and H the sediment thickness expressed in km. 190
This relation has been chosen in order to simulate the lower average crustal density at the 191
surface with the increase of the sediment thickness. According to this equation, the surface 192
density is equal to 2700 kg.m-3 when there is no sediments and equal to 2602 kg.m-3 when 193
sediments are 20 km thick. 194
Certainly, the absolute density values used for the lithospheric mantle and the asthenosphere 195
are noticeably lower than the actual ones. This is because we are using a ‘pure’ thermal 196
approach in which the lithosphere mantle density is only temperature dependent and the 197
density of the asthenosphere is constant everywhere. In other words, temperature is the 198
dominant effect on density and no pressure effects are considered. Note that this approach is 199
widely used in describing variations of topography and potential fields since they mainly 200
depend on lateral density variations rather than on absolute density values. 201
Because of the occurrence of large magmatic provinces and numerous ophiolitic belts, the 202
crustal density at Moho depth in the Arabia-Eurasia collision zone has been increased to 203
2970 kg.m-3. We use classical thermal conductivities of 2.7 W.m-1.K-1 for the crust and 3.2 204
W.m-1.K-1 for the lithospheric mantle. The thermal conductivities are considered as constant 205
in our modeling. In fact, the larger variability of the crustal thermal conductivity is restricted 206
to the uppermost part of the crust (up to 20 MPa) (Clauser & Huenges, 1995) and then, does 207
not affect much the average crustal thermal conductivity. The average crustal radiogenic heat 208
production is 0.5 μW.m-3 for continental crust and 0.3 μW.m-3 for oceanic crust (Vilà et al., 209
2010). The heat production in the lithospheric mantle is considered as null because of its 210
very low contents in radiogenic elements. 211
Fullea et al. (2007) performed a sensitivity analysis for the calculation of CMB and LAB 212
depths induced by the typical RMS error of elevation and geoid anomaly databases. They 213
found that the inaccuracy for the crustal thickness in emerged regions is less than 2 km 214
whereas maximum error in lithospheric thickness is less than 10 km. The effect of varying 215
crustal and lithospheric mantle conductivity has been tested and the resulting lithospheric 216
thickness is affected up to 6-8 km whereas varying the coefficient of thermal expansion (α) 217
affects the lithospheric thickness up to ~12 km. In addition, the crustal thickness is not much 218
affected (~1 km) by varying thermal parameters. 219
Considering this sensitivity analysis, we estimated the maximum misfits obtained in our 220
modeling result to be up to 5 km for the crustal thickness and up to 15 km for the 221
lithospheric thickness. 222
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Input data 224
Topography and Bathymetry 225
Topographic and bathymetric data are extracted from the ETOPO1 database (Amante & 226
Eakins, 2009) (Figure 3a). The most remarkable geomorphological feature of Central 227
Eurasia is the Tibetan Plateau which covers an area of more than 2.5x106 km2 with a mean 228
elevation of around 4500 m. High elevations are also found to the North of the Tibetan 229
Plateau, along the Tien Shan or the Qilian Shan ranges. The Tarim basin presents a low relief 230
with a mean elevation of around 1000 m. To the West, the highest elevations are located in 231
the Zagros Mountains and they extend towards the Anatolian Plateau. The Alborz, Kopet 232
Dagh, Caucasus and Sistan ranges show also remarkable topographies with elevations up to 233
5000 m. To avoid unrealistic high frequency variations of the Moho and LAB depths we 234
have filtered the elevation using a Gaussian filter with a wavelength of 100 km. 235
Geoid anomaly 236
The geoid anomaly was extracted from the EGM2008 global model (Pavlis et al., 2008). The 237
EGM2008 gravitational model is complete to spherical harmonic degree and order 2159, 238
which determines the resolution of our modeling at 10 arc minute (~18.5 km). In order to 239
avoid the effects of sublithospheric density variations, the geoid was filtered to remove the 240
signature corresponding to the lower spherical harmonics until degree and order 11 (Figure 241
3b). The degree and order of spherical harmonics filtering have been chosen in order to 242
remove very large wavelength anomalies that probably result in sub-lithospheric density 243
variations (see figure S2, supplementary materials). 244
The obtained geoid anomaly shows maximum values of up to 21 m located in two main 245
regions: (1) along the Himalaya and the Tibetan Plateau with a northward decrease of the 246
anomaly, and (2) in the Anatolian Plateau, in the Alborz and in the Sanandaj-Sirjan zones in 247
Iran. Minimum values are down to -21.5 m coinciding with large sedimentary accumulations 248
as occurs in the Tarim, Junggar and Ganga basins, the Caspian Sea, the Nile delta, the 249
Persian Gulf and eastern part of Arabia. 250
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Comparison with seismological and seismic studies 252
We carried out a complete compilation of previously published crustal thickness data (Figure 253
4, Table 3 and supplementary material S1) mainly from receiver function and deep seismic 254
studies. To test the reliability of our modeling approach, we compared this compilation to 255
crustal thicknesses obtained from our modeling (Figures 5, 6 and 7). The resulting CMB and 256
LAB from our modeling approach are downloadable in the supplementary material and they 257
are furnished in .xyz format. In contrast to the CMB, the LAB corresponds to a rheological 258
rather than to a compositional contrast, which explains that its location and properties are 259
more elusive (Eaton et al., 2009). Shear-wave velocity anomaly within the upper mantle 260
computed from global or regional tomography is often considered as correlated with the 261
lithospheric thickness. We compare our results with the S40RTS mantle shear-wave velocity 262
model (Ritsema et al., 2011). The LAB is rarely imaged using receiver function methods but 263
we compared our modeling results with estimations from these studies where they were 264
available (e.g. Kumar et al., 2007; Zhao et al., 2011). 265
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The Arabian Plate 267
The crustal thickness of the Arabian Plate estimated from receiver functions, seismic 268
refraction profiles and surface wave studies (see references in Table 3) ranges between 32 269
and 50 km with an average value of ~40 km, which is consistent with our modeling results 270
(Figure 5a). Our map shows a smooth crustal thinning towards the North-East of the Arabian 271
Plate, which could be interpreted as inherited from the passive margin structure though it is 272
not well imaged by the scarce available seismic data. Both previous seismic studies and our 273
modeling results image an abrupt crustal thinning along the boundaries of the Arabian Plate 274
(the Gulf of Aden to the South-East, the Gulf of Aqaba to the North-West and the Red Sea to 275
the West). Finally, beneath Oman and along the southern part of Red Sea coast, our model 276
indicates that the crustal thickness may reach up to 43 km, whereas Al-Damegh et al. (2005) 277
propose a crustal thickness up to 50 km. 278
The shear-wave velocity anomaly map at 100 km depth shows a large low velocity anomaly 279
in western Arabia and along the Red Sea, whereas a positive anomaly is imaged beneath the 280
Mesopotamian Basin (Villaseñor et al., 2001; Ritsema et al., 2011). Our model exhibits a 281
thicker lithosphere in the eastern part of the Arabian Plate than to the western part, with a 282
maximum thickness exceeding 200-220 km beneath the Persian Gulf and South-East Arabia 283
(Figure 5b), which is in agreement with estimations from tomographic results (Villaseñor et 284
al., 2001; Shomali et al., 2011). According to our results, this thickened lithosphere extends 285
beneath the Zagros Belt, which is also confirmed by shear-wave modeling (Priestley et al., 286
2012). Finally, our model indicates that towards the North-West of the Mesopotamian Basin, 287
the lithosphere thickness decreases reaching minimum values of about ~140 km south of the 288
Anatolian Plateau. 289
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The Indian Plate 291
Previous crustal thickness estimations of the Indian Plate vary from 30 to 44 km but most of 292
these estimations are comprised between 36 and 39 km (see references in Table 3). Our 293
modeling results fit very well with previously obtained crustal thicknesses for the whole 294
Indian Plate except for the Shillong Plateau where our model predicts a thicker crust due to 295
the fact that our approach is not strictly valid to image an uncompensated crustal pop-up 296
structure as proposed by Mitra et al. (2005). 297
Previous studies using surface-wave dispersion data observed a shield-like lithospheric 298
structure in most part of the plate (Singh, 1999; Mitra et al., 2006; Suresh et al., 2008; 299
Bhattacharya, 2009; Prajapati et al., 2011) and there are no major lateral variations in the 300
shear-wave velocity anomalies across the Indian Plate (Villaseñor et al., 2001; Ritsema et 301
al., 2011). According to previous geophysical studies, the lithospheric thickness beneath 302
western India is ~155 km (Mitra et al., 2006) and ~140 km beneath eastern India 303
(Bhattacharya, 2009). Our model images a ~160-190 km thick lithosphere; though a 304
lithosphere thickness reaching up to 230 km is imaged below the central part of the Ganga 305
Basin (Figure 5b). In this region, the possible component of flexural support to the 306
lithospheric bending, evidenced by large sedimentary thickness accumulation, could also 307
affect the validity of our approach resulting in an overestimation of our modeled lithospheric 308
thickness. However, the lack of other geophysical data in this area does not allow for firm 309
conclusions. 310
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The Arabia-Eurasia collision zone 312
In agreement to previous geophysical studies our modeling results show crustal thickness 313
values ranging between 35 and 45 km beneath the Mesopotamian Basin (see references in 314
Table 3). Along the Zagros orogeny, seismic data indicate that the maximum crustal 315
thickness is located below the Sanandaj-Sirjan Zone, whereas our model predicts the 316
maximum crustal thickness values slightly displaced to the South-West coinciding with the 317
highest elevations and directly below the Imbricated Zone and the Zagros Simply Folded 318
Belt (Figure 5a). This shift of the maximum crustal root relative to maximum topography is 319
common in collisional settings where crustal thickening is associated with underthrusting or 320
down-flexuring at crustal scale as it might be the case of the Zagros orogeny (Motavalli-321
Anbaran et al., 2011; Vergés et al., 2011). Our 1D approach, based on the combination of 322
elevation and geoid anomaly, is not able to reproduce these underthrusting structures and 323
therefore to precisely locate the region with maximum crustal thickness. However, our 324
modeling results indicate crustal thickness values between 35 and 54 km beneath the Iranian 325
and Afghan blocks in good agreement with available seismic data in Central Iran (Paul et al., 326
2006; Nasrabadi et al., 2008; Paul et al., 2010; Asfari et al., 2011; Motaghi et al., 2012) 327
suggesting that the stiff lithospheric domains are related to relatively low crustal thickness 328
values. The Alborz, Kopet Dagh and Caucasus mountain ranges are characterized by 329
thickened crust according to both, seismic data and modeling results (Figure 7). In the 330
Anatolian Plateau our model shows crustal thickness values between 36 and 50 km in good 331
agreement with most of the estimations from seismological studies, which in turn show a 332
large scattering (up to ~20 km of difference). 333
Shear-wave velocity studies in the Arabia-Eurasia collision zone show a thick high shear-334
wave velocity anomaly within the upper mantle beneath eastern Arabia and Zagros whereas 335
western Arabia, Central Iran and the Anatolian Plateau domains are characterized by low 336
shear-wave velocity anomalies within the upper mantle (Villaseñor et al., 2001; Priestley et 337
al., 2012). Low shear-wave velocities underneath the eastern part of the Anatolian Plateau 338
are interpreted as a consequence of a very thin or possibly removed lithospheric mantle (Gök 339
et al., 2008; Bakirci et al., 2012), which is confirmed by S-receiver functions studies (Angus 340
et al., 2006) and low Pn wave speeds (Al-Lazki et al., 2004). Mohsen et al. (2006) propose a 341
lithospheric thickness value of 90 km beneath eastern Anatolia, which is in agreement with 342
our model, highlighting a thin lithosphere beneath the whole Anatolian Plateau (from 100 to 343
130 km thick). 344
345
The India-Eurasia collision zone 346
Many geophysical studies have been carried out across the Himalayan range and the Tibetan 347
Plateau to image its complex crustal and lithospheric structure. Most of these studies are 348
based on deep seismic profiles (Zhao et al., 2001; Galvé et al., 2002; Ross et al., 2004; Zhao 349
et al., 2006; Mechie et al., 2012) and on passive seismological data (see the whole set of 350
references in the Table 3). A general result is that, despite its low relief and high elevation, 351
the Tibetan Plateau shows a highly variable crustal thickness ranging from 56 km up to 95 352
km and a heterogeneous lithospheric structure (e.g. Yin & Harrison, 2000; Zhao et al., 2010). 353
Collectively, the results suggest that the crust is thicker in the southern part (between 70 to 354
80 km thick) of the plateau than to the North (between 65 to 70 km thick) except for the 355
easternmost part of the plateau (longitude > 95°E). In contrast, our results indicate that the 356
crust in the southern part of the Tibetan Plateau ranges between 68 and 76 km thick whereas 357
it predicts a thicker crust (more than 75 km thick) in the northern part of the plateau (Figure 358
5a). To the West of the Tibetan Plateau, seismological studies indicate a noticeable crustal 359
thickening from ~50 km below the front of the Himalayan range in Pakistan (Soomro, 2009) 360
to ~78 km below the Pamir (Li & Mashele, 2009; Mechie et al., 2012). Our model is broadly 361
in agreement with these estimations although it slightly underestimates the crustal thickness 362
below southern Pamir. 363
Our model is able to reproduce the lithospheric thickness in the southern and western part of 364
the plateau, in agreement with receiver function data (Owens & Zandt, 1997; Nabelek et al., 365
2009; Zhao et al., 2010, 2011), tomographic data (Villaseñor et al., 2001; Priestley et al., 366
2006; Replumaz et al., 2013) and previous lithospheric modeling (Jiménez-Munt et al., 367
2008) that suggest a ~200 to 240 km thick lithosphere. Most authors agree that this thick 368
Tibetan lithosphere corresponds to the northern boundary of the underthrusted Indian 369
lithosphere below Tibet (e.g., Owens and Zandt, 1997; Nabelek et al., 2009; Zhao et al., 370
2010, 2011; Liang et al., 2012). The main discrepancy concerns the northern part of the 371
Tibetan Plateau where some studies suggest a very thin to inexistent lithospheric mantle (e.g. 372
Meissner et al., 2004; Jiménez-Munt et al., 2008; Liang et al., 2012) whereas other studies 373
suggest a very thick lithosphere (e.g. Priestley et al., 2006; Barron & Priestley, 2009). Our 374
model results indicate a very thick lithosphere, up to 340 km thickness coinciding with the 375
region where our model overestimates the crustal thickness. However, using the same geoid 376
and elevation data plus gravity and thermal data, Jiménez-Munt et al. (2008) proposed a very 377
thin to removed lithosphere. Unfortunately, our results do not allow discerning among the 378
proposed geodynamic processes acting in this peculiar part of the Tibetan Plateau. 379
380
North of the Tethysides 381
According to seismological and seismic studies, the Tarim basin is characterized by a crustal 382
thickness of ~40 km in its central part increasing up to ~60 km in its flanks (Kao et al., 2001; 383
Qiusheng et al., 2002; Wittlinger et al., 2004; Mi et al., 2005; Zhao et al., 2003, 2006; Chen 384
et al., 2010). Our model overestimates by ~ 5 km the crustal thickness in the Central Tarim 385
basin and does not reflect noticeable crustal thickness variations beneath the basin (Figure 386
5a). According to seismological estimations the crustal thickness below the Tien Shan is 55-387
60 km (Bump & Sheehan, 1998; Zhao et al., 2003; Vinnik et al., 2004), whereas our model 388
predicts crustal thickness values between 60 and 68 km. In the Kazakh terranes, North-West 389
from Tien Shan, the small number of available seismological estimations indicate a crustal 390
thickness of about 45 km (Vinnik et al., 2004; Bump & Sheehan, 1998), which is very close 391
to the 40-44 km obtained from our model for the whole region. In the North-East part of the 392
study area, North of the Altaids Ranges, the crustal thickness inferred from receiver 393
functions (Zorin et al., 2002) varies from 44 to 50 km in full agreement with our results. 394
The lithospheric thickness have been estimated from S-wave receiver functions (Kumar et 395
al., 2005) indicating values of 90-120 km beneath the Tien Shan, ~180 km in the Tarim 396
Basin, and ~130 km in the Kazakh Platform. Our results show a thicker lithosphere in all 397
these regions varying from 240 km in the Tien Shan, to 220 km in the Tarim Basin and 180 398
km in the Kazakh Platform, being in agreement with S-wave tomography models, which 399
indicate high velocities below the Tien Shan and the Tarim basin (Ritsema et al., 2011). 400
401
Discussion 402
One of the goals of the present study is to provide new maps of both the Crust-Mantle-403
Boundary and the Lithosphere-Asthenosphere-Boundary in the Central Eurasia region and to 404
compare these results with previous regional/global models. A second related goal is to 405
discuss the implications on the geodynamic evolution of this large continental collision 406
region based on our crustal and lithospheric thickness maps. 407
Our modeling approach is based on the assumptions of local lithospheric isostasy and 408
thermal equilibrium. Whereas local isostasy is an acceptable approximation for wavelengths 409
of hundred kilometers (e.g. McKenzie & Bowin, 1976; England & Molnar, 1997), the 410
assumption of thermal equilibrium is not valid in regions of recent tectonic activity as the 411
Alpine-Himalayan system. Therefore, the results of our model must be interpreted as 412
average physical conditions necessary to produce the required density distribution rather 413
than the actual thermal boundaries. The calculation in steady state minimizes the variations 414
in lithospheric thickness, since the modeling tends to underestimate the lithosphere thickness 415
when thickening processes occur under transient conditions, and overestimate it for thinning 416
processes (see Fullea et al., 2007 and Jiménez-Munt et al., 2011 for more details). 417
418
Comparison with global crustal models 419
The most widely used global datasets related to crustal thickness is the CRUST2.0 (Bassin et 420
al., 2000), which has recently been replaced by the CRUST1.0 that incorporates an updated 421
version of global sediment thickness and presents a better accuracy (Laske et al., 2013). 422
Reguzzoni et al., (2013) combined CRUST2.0 model with gravity observations from the 423
GOCE satellite to propose the new global crustal GEMMA model that we compare with our 424
results (Figure 7). 425
426
The 1x1° grid global crustal model CRUST1.0 is obtained from the analysis of seismic wave 427
travel-times and seismic refraction data (Laske et al., 2013). The resulting crustal thickness 428
from the CRUST1.0 model (Figure 8a) in the study region is relatively smooth with 429
maximum values exceeding 70 km in the Tibetan Plateau and minimum values of <15 km in 430
the oceanic domains of the Red Sea and the Indian Ocean. Most of the continental domains 431
show crustal thickness values of 40±5 km. The excessive smoothness of the resulting map 432
makes difficult to relate crustal thickness variations with surface geological structures. 433
Therefore, the crustal thickness derived from our model exceeds that from CRUST1.0 in 434
more than 10-15 km along the main mountain ranges (i.e., Red Sea rift shoulders, Anatolia, 435
Caucasus, Zagros, Alborz, Makran, Pamir, Qiangtang and Tien Shan). 436
The resulting crustal thickness from GEMMA Moho model (Reguzzoni et al., 2013) ranges 437
from 15 to more than 100 km and presents undulations of ~40 km large over the whole 438
region (Figure 8b). In contrast, our method allows for a better interpolation between regions 439
with reliable measurements and then show a much better correlation with the major 440
geological structures (Figure 8c). 441
442
Comparison with global lithospheric models 443
Concerning the lithospheric thickness, we compare our modeling results with 2 lithospheric 444
models: the TC1 model (Artemieva, 2006) and the thermal lithospheric model presented in 445
Goutorbe et al. (2011). In addition to these two models, there are a number of 446
regional/global seismic tomography models and several regional studies based on the 447
combination of different geological and geophysical data that can also constrain the crust 448
and upper mantle structures as compared below. 449
The resulting lithospheric thickness from our model can be qualitatively compared to 450
regional/global models based on seismic surface waves and thermal approaches. Regional 451
and global S-wave velocity perturbations in the Central Eurasia region at 100 km depth (e.g. 452
Villaseñor et al., 2001; Priestley et al., 2006; Ritsema et al., 2011; Hatzfeld & Molnar, 2010 453
and references therein) show low velocities beneath the Red Sea, Eastern Anatolia, Central 454
Iran and Afghan blocks, Eastern Tibet, and North to the Altaid ranges, with some differences 455
among these models. High velocity perturbations are located beneath North India and West 456
Tibet, and the Northern regions. Major differences between our model and S-wave 457
tomography models are related to the lateral variations in lithospheric thickness from the 458
Zagros to Central Iran and the Mesopotamian Basin. Some tomography models (e.g. 459
Priestley et al., 2006; Priestley et al., 2012) propose a lithosphere thickness exceeding 230 460
km along the so-called Zagros core thinning rapidly towards Central Iran and the 461
Mesopotamian Basin. In contrast, our model shows that the thick lithosphere affects the 462
Mesopotamian Basin, the South-East Arabian Plate and the Zagros until the Arabia-Eurasia 463
continental suture with an abrupt thinning beneath the Sanandaj-Sirjan Zone and the 464
Urumieh-Doktar Magmatic Arc. This lithosphere thinning extends from Anatolia, to Central 465
Iran and to the Afghan Block, which is more in agreement with the tomography models by 466
Villaseñor et al. (2001) and Ritsema et al. (2011) that also show a lithospheric thickening 467
beneath the Mesopotamian Basin as in our model. Another conflictive region is the Eastern 468
Tibet, where all tomography models show a relatively low velocity region at 100-125 km 469
depth vanishing with depth and disappearing at 200 km depth. Finally, the high velocity 470
anomalies occupying most of the northern regions of the study area are in agreement to the 471
lithospheric thickness of 170-200 km proposed by our model. This fitting is not 472
accomplished north to the Altaid ranges where tomography models consistently image a low 473
velocity region in contrast to a rather homogeneous lithospheric thickness of 180-200 km 474
inferred from our calculations. In any case, the interpretation of S-wave tomography in terms 475
of thermal models must take into account that velocity anomalies of non-thermal origin may 476
amount up to ±3% of Vs amplitude and attributed to chemical composition variations and/or 477
to the presence of melts/fluids (Artemieva, 2009). 478
Global thermal models can be directly compared to our results since the approach used in the 479
present work include thermal analysis. TC1 is a global scale continental lithospheric model 480
that has been computed from available heat flow measurements supplemented with 481
electromagnetic and xenolith data in cratonic domains (Artemieva, 2006) (Figure 8a). In 482
Central Eurasia, TC1 model proposes lithospheric thickness ranging from less than 100 km 483
below Anatolia and Iran, up to more than 200 km in the North-East Tibetan Plateau and W-484
Central part of the Indian Plate. Our model presents a roughly similar trend of the 485
distribution of the lithospheric thickness but with a higher resolution and more precise 486
lithospheric thickness estimations in the study region (Figure 8b). The lithospheric 487
thickening affecting the Eastern Arabian Plate and the Turan Platform are not evident in TC1 488
model and, conversely, our model does not image the apparent lithospheric thickening in W-489
Central India. Goutorbe et al. (2011) published a thermal model obtained from heat flow 490
measurements combined with multiple geological and geophysical data and proxies. This 491
model reproduces the lithospheric thickening affecting the eastern Arabian Plate and the 492
Mesopotamian Basin, as well as the Tibetan Plateau and North India Plate. Lithospheric 493
thinning is concentrated along the Red Sea and its shoulders, Anatolia and Central Iran and 494
Afghan blocks. The proposed lithospheric thickness values are however unrealistically low 495
especially in the thinned regions, where the LAB is less than 40 km thick. 496
497
Geodynamic implications 498
The new results indicate that crustal thickening in Central Eurasia is clearly associated with 499
major frontal ranges as the Himalaya and Zagros, but also to distal ranges as the Alborz, 500
Kopet Dagh, and Tien Shan. The more than 1200 km in width of the deformed area resulting 501
from the Arabia/Eurasia and India/Eurasia collision zones is also highlighted by the extent of 502
the area affected by crustal thickening (Figures 1 and 5a). As already suggested by many 503
geophysical studies, the Tibetan Plateau is sustained by a very thick crust with an average 504
thickness of ~75 km according to our model. Furthermore, the tectonic blocks remaining in 505
the deformation zone, as the Central Iran or Tarim blocks, show a slightly thickened crust 506
and uniform topography pointing out their rheological resistance and moderate deformation. 507
A consequence of these features is that tectonic stresses are able to be efficiently transmitted 508
hundreds to thousand kilometers far from the suture zones. 509
Contrarily to the crustal structures that are comparable between the two collision zones, their 510
lithospheric structures differ substantially (Figures 5 and 6). The Arabia/Eurasia collision 511
zone is characterized by a thick lithosphere underneath the Zagros belt and the Arabian 512
platform (between 180 and 220 km thick) whereas the Central Iran, the Anatolian Plateau 513
and the northern part of the Arabian Plate are characterized by thin to very thin lithosphere 514
(90-150 km thickness). In contrast, the India/Eurasia collision zone shows a thick to very 515
thick lithosphere especially in its southern part thus evidencing the plate underthrusting of 516
most part of the Indian lithosphere beneath the Eurasian lithosphere. The large differences in 517
lateral thickness variations between the crustal and the lithospheric mantle indicate a strong 518
strain partitioning with the CMB acting as the preferential detachment level as proved in Iran 519
(Jiménez-Munt et al., 2012). 520
521
Conclusions 522
523
We present a new crustal and lithospheric thickness model assuming thermal and isostatic 524
equilibrium that fits well with existing seismic and seismological data. The method used 525
permits to interpolate crustal and lithospheric thickness values between regions with reliable 526
measurements and provides a full coverage for both the crust and the lithosphere of Central 527
Eurasia. The presented work allows for drawing the following concluding remarks: 528
Crustal thickening is correlated with major mountains ranges (Himalayas, Zagros, 529
Pamir, Caucasus, and Tien Shan) whereas crustal thinning is restricted to the Arabian 530
and Indian oceanic domains and to the South Caspian Sea, the Red Sea and the Black 531
Sea. The Iranian and Anatolian plateaus do not show a significant thick crust whereas 532
the Tibetan Plateau shows an extremely thickened crust up to 78 km thick. At regional 533
scale, our resulting map matches well with geological features. 534
Contrasting lithospheric structures are modeled between the India/Eurasia and the 535
Arabia/Eurasia collision zones. Thickened lithosphere is obtained below the Zagros 536
orogenic belt whereas the Anatolian Plateau and Central Iran are characterized by a 537
thin to very thin lithosphere (~100-130 km thick). In contrast, maximum lithospheric 538
thickness, reaching up to 300 km thick, corresponds to the southern and western 539
Tibetan Plateau and below the Pamir. Our model coincides well with S-wave 540
tomography models but not in Eastern Tibet where our model cannot resolve the 541
existence of a Low Velocity Zone down to 200 km depth, suggesting a lithospheric 542
thinning related to deep geodynamic processes. 543
Our resulting crustal map is roughly in good agreement with the global CRUST1.0 and 544
GEMMA Moho models. However, the crustal thickness derived from our model 545
exceeds that from CRUST1.0 in more than 10-15 km along the main mountain 546
ranges, which better fits the crustal thickness in these tectonic domains based on 547
geophysical studies. Our resulting model presents a similar trend of the distribution of 548
the lithospheric thickness compared to regional and global models but it also presents 549
better resolution and thickness estimations. 550
India-Eurasia and Arabia-Eurasia collisional systems present comparable crustal 551
structure but contrasted lithospheric structure. The Arabia-Eurasia collision zone is 552
characterized by a very thin lithosphere that can be interpreted as the signature of the 553
dominance of subduction processes. Oppositely, the more mature India-Eurasia 554
collision zone shows a thick lithosphere in its western and central parts that highlights 555
the importance of crustal underthrusting processes. 556
Our new results indicate that crustal thickening is not restricted but extends hundreds 557
to thousand kilometers away from the collision front, indicating an effective 558
transmission of tectonic stresses, which is partly related to the presence of stiff 559
lithospheric blocks that remain almost undeformed within the collisional systems. 560
561
562
This research has been funded by DARIUS Consortium and projects ATIZA (CGL2009-09662-BTE), TECLA 563
(CGL2009-09662-BTE) and Consolider-Ingenio 2010 Topo-Iberia (CSD2006-00041). Some figures were 564
generated using GMT (Wessel and Smith, 1991). The authors thank Alexander Koptev, Joerg Ebbing and 565
Marie-Françoise Brunet for their constructive comments on the manuscript that helped to improve this article. 566
567
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Tables captions: 1102
1103
Table 1: Constant parameters used over the whole Central Eurasia region in our 1104
modeling. 1105
1106
Table 2: Parameters used in our modeling, as a function of the zones defined on the 1107
figure 1. 1108
1109
Table 3: List of publications used for the compilation of crustal thickness data that are 1110
represented on the figure 4. Complete dataset of the crustal thicknesses is presented in 1111
the supplementary material (Table S1). 1112
1113
Table S1 (Supplementary Material): Compilation of crustal thickness data from 1114
publications that were used to be compared with our modeling results. These data are 1115
represented on the figure 4 and the Table S3 only indicates references used for this 1116
compilation. 1117
1118
1119
Figure captions: 1120
Figure 1. Simplified structural map of the Central Eurasia region modified from the 1121
Structural map of Eastern Eurasia (Pubellier et al., 2008, modified) and publications 1122
(Sengör and Natal’in, 1996; Frizon de Lamotte et al., 2013). a) this map highlights the 1123
accreted terranes in the region resulting from the long-lived convergence between 1124
Gondwana-derived blocks and the Southern margin of Eurasia since the end of the 1125
Paleozoic. b) Simplified map showing the different zones delimited in this study. The 1126
shaded relief was obtained from the ETOPO1 database (Amante & Eakins, 2009). The 1127
scale bars corresponds to 15° latitude. 1128
1129
Figure 2. Mean crustal density used as an input in our model. Contours denote 1130
sediment thickness in meters taken from EXXON Structural Map of the World (1994) 1131
and publications related in the text. Used parameters for each region are summarized 1132
in the Table 2. 1133
1134
Figure 3. a) Elevation map from ETOPO1 database (Amante & Eakins, 2009). b) 1135
Residual geoid height map resulting from the EGM2008 global model (Pavlis et al., 1136
2008) after removing the lower spherical harmonics until degree and order 11. Main 1137
sutures zones have been represented. The scale bar correponds to 15° latitude. 1138
1139
Figure 4. Compilation of previous crustal thickness estimations from seismological 1140
and seismic experiments (complete dataset in Table S1 and references in Table 3). The 1141
scale bar corresponds to 15° of latitude. 1142
1143
Figure 5: a) Crustal thickness map and b) lithospheric thickness map of the entire 1144
Central Eurasia region calculated from elevation and geoid anomaly inversion. The 1145
scale bar corresponds to 15° of latitude. 1146
Figure 6: a) Topographic map of the study and location of two lithospheric cross-1147
sections, b) Lithospheric cross-section across the India-Eurasia collision zone 1148
representing the topography and the depths of the crust-mantle boundary (CMB) and 1149
the lithosphere-asthenosphere boundary (LAB). c) Same as b) but across the Arabia-1150
Eurasia collision zone. 1151
1152
Figure 7: Comparison between crustal thickness estimations from seismological and 1153
seismic experiments (circles and triangles corresponding to Figure 4) and the 1154
calculated values from inversion of geoid and elevation (background map 1155
corresponding to Figure 5a). Similar colors between filled symbols and background 1156
map indicate coincidence between seismic data and model results. 1157
1158
Figure 8. Crustal thickness maps from a) CRUST 1.0 model (1 degree grid 1159
resolution) (Bassin et al., 2000), b) GEMMA Moho model (Reguzzoni et al., 2013) 1160
and c) our results (0.16 degree grid resolution). Major sutures zones are plotted. 1161
1162
Figure 9: Lithospheric thickness maps from a) TC1 thermal model (Artemieva, 2006) 1163
and from b) our results. Major sutures zones are plotted. 1164
1165
Figure S2 (Supplementary Material): Comparison of the residual geoid anomaly 1166
resulting from filtering at different orders and degrees of spherical harmonics. These 1167
residual geoid anomaly indicates that large wavelength anomalies that probably result 1168
from deep sub-lithospheric density variations are filtered by removing the signal 1169
corresponding to degree and order above 10 of spherical harmonics. 1170
Parameter Va l u e
Asthenosphere density (ρa) 3200 kg.m-3
Sea Water density (ρw) 1031 kg.m-3
Temperature at the LAB (Ta) 1300 °C
Surface temperature (Ts) 15 °C
Thermal expansion coefficient (α) 3.5 10-5K-1
Crustal thermal conductivity (kc) 2.7 W.K-1.m-1
Mantle thermal conductivity (km ) 3.2 W.K-1 .m-1
Parameter Zone 1 Zone 2 Zone 3 Zone 4 Zone 5 Oceanic crust
Crustal density at surface
in kg.m-3 [2608-2700] [2605-2700] [2601-2700] [2605-2700] [2601-2700] [2602-2700]
Crustal density at CMB
in kg.m-3 2910 2910 2970 2910 2910 2980
Crustal radiogenic heat
production
in mW.m-3
0.5 0.5 0.5 0.5 0.5 0.3
Zone References
Zone 1 – Arabian Plate
Al-Damegh et al. [2005]; Al-Lazki [2003]; Badri [1991]; Barton et al. [1990]; Brew [2001]; Gök et al. [2008]; Kumar et al. [2001];
Mohsen et al. [2006]; Sandvol et al. [1998]; Tiberi et al. [2007]
Zone 2 – Indian Plate Gupta et al. [2003]; Kumar et al. [2001, 2004, 2007]; Mitra et al. [2005, 2008]; Radhakrishna et al. [2012]; Soomoro [2009]
Zone 3 – Arabia-Eurasia
collision zone
Al-Damegh et al. [2005]; Angus et al. [2006]; Asfari et al. [2011]; Brew [2001]; Çakir et al. [2000]; Doloei & Roberts [2003]; Gök et al.
[2008, 2011]; Gritto et al. [2008]; Magino & Priestley [1998]; Mellors et al. [2008]; Motaghi et al., [2012]; Nasrabadi et al. [2008];
Nowrouzi et al. [2007]; Paul et al. [2006, 2010]; Radjaee et al. [2010]; Sandvol et al. [1998]; Sodoudi et al. [2009]; Tezel et al.
[2013]; Yaminifard et al. [2012]; Zor et al. [2003]
Zone 4 – India-Eurasia
collision zone
Bump & Sheehan [1998]; Chen et al. [2010]; Galvé́ et al. [2002]; Gao et al. [2013]; He et al. [2014]; Hetenyi [2007]; Jiang et al.
[2006]; Kao et al. [2001]; Kind et al. [2002]; Kumar et al. [2005]; Li et al. [2001,2008]; Li & Mashele [2009] ; Lou et al. [2009];
Makarov et al. [2010]; Mechie et al. [2012]; Mi et al. [2005]; Mitra et al. [2005, 2008]; Nabelek et al. [2009]; Oreshin et al. [2008]; Pan
& Niu [2011]; Qiusheng et al. [2002]; Robert et al. [2010a, b]; Ross et al. [2004]; Schulte-Peklum et al. [2005]; Shi et al. [2004,
2009]; Soomoro [2009]; Tian et al. [2005]; Tilmann [2011]; Tong et al. [2007]; Vergne et al. [2002]; Vinnik et al. [2004]; Wang et al.
[2003,2010]; Wittlinger et al. [2004]; Xu et al., [2013]; Yue et al. [2012]; Zhang & Wang [2009]; Zhao et al. [2001, 2003, 2006]
Zone 5 – North of the
Tethysides
Chen et al. [2010]; Magino & Priestley [1998]; Mellors et al. [2008]; Motaghi et al. [2012]; Pan & Niu [2011]; Radjaee et al. [2010];
Vinnik et al. [2004]; Wang et al. [2004]; Zhao et al. [2003]; Zorin et al. [2002]
