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1 Mantle convection in the Middle East:
2 Reconciling Afar plume, Arabia collision and Aegean trench rollback
3
4 Claudio Faccenna1, Thorsten W. Becker2, Laurent Jolivet3 and Mehmet Keskin4
5
6 1 LET, Laboratory of Experimental Tectonics, Università Roma Tre, Rome, Italy
7 2 Department of Earth Sciences, University of Southern California, Los Angeles CA 90089,
8 USA
9 3Institut des Sciences de la Terre d'Orléans - Université Orleans
10 4 Dept. of Geological Engineering, Istanbul University, 34320 Avcılar, Istanbul, Turkey
11
12 Abstract
13 The Middle East region represents a key site within the Tethyan domain where continental
14 break-up, collision, backarc extension and escape tectonics are linked together. We perform global
15 mantle circulation computations to test the role of slab pull and mantle upwellings as driving forces
16 for the kinematics of the Arabia-Anatolia-Aegean (AAA) system, evaluating different boundary
17 conditions and mantle density distributions as inferred from seismic tomography or slab models.
18 Model results are compared with geodesy, residual topography, and shear wave splitting
19 observations. The AAA velocity field with respect to Eurasia shows an anti-clockwise toroidal
20 pattern, with increasing velocities toward the Aegean trench. The best match to these crustal
21 motions can be obtained by combining the effect of slab pull exerted in the Aegean with a mantle
22 upwelling underneath Afar and, more generally, with the large-scale flow associated with a whole
23 mantle, Tethyan convection cell. Neogene volcanism for AAA is widespread, not only in the
24 extensional or subduction settings, but also within plates, such as in Syria-Jordan-Israel and in
25 Turkey, with geochemical fingerprints similar of those of the Afar lava. In addition, morphological
26 features show large uplifting domains far from plate boundaries. We speculate that the tectonic
1 27 evolution of AAA is related to the progressive northward entrainment of upwelling mantle material
28 which is itself associated with the establishment of the downwelling branch of a convection cell
29 through the segmented Tethyan slab below the northern Zagros and Bitlis collision zone. The newly
30 established westward flow dragged Anatolia and pushed the Aegean slab south-westward, thus
31 accelerating backarc extension. Our model reconciles Afar plume volcanism, the collision in the
32 Bitlis mountains and northern Zagros, and the rapid increase of Aegean trench rollback in a single
33 coherent frame of large scale mantle convection, initiated during the last ~40 Ma.
34
35
36
37
38 1. Introduction
39 The Arabia-Anatolia-Aegean (AAA) system represents a key regional setting within the Tethyan
40 domain where continental break-up, collision, backarc extension, and escape tectonics are
41 intricately linked. This interacting plate system with a wealth of available geological and
42 geophysical constraints offers a chance to evaluate the dynamics behind the motions and
43 deformation of the plates.
44 The geodetically imaged crustal motions are spectacular; they show marked, counter-clockwise
45 rotation of the velocity field with respect to Eurasia (Fig. 1a; Kreemer & Chamot-Rooke 2004,
46 Reilinger et al. 2006; Le Pichon and Kreemer, 2010). Arabia behaves as a rigid plate moving NE to
47 NNE along a rotation pole positioned in Libya, producing spreading in the Red Sea and Gulf of
48 Aden and collision against Eurasia along the Bitlis-Zagros. While convergence in Iran is
49 accommodated by distributed northward decrease in the relative velocity, toward the west, the
50 Anatolian plate moves coherently at an average rate of ~2 cm/yr, separated from Eurasia by the
51 North Anatolian Fault (Reilinger et al., 2006). In the west, the Aegean plate is moving SW at
52 ~3.3 cm/yr following the retreat of the Aegean trench (Kreemer & Chamot-Rooke 2004, Reilinger
2 53 et al. 2006). Overall, the collisional AAA zone shows a circular velocity pattern, accelerating
54 toward the Aegean trench, with a rotation pole for Anatolia-Eurasia positioned north of the Nile
55 delta (Fig. 1a, see also Le Pichon and Kreemer, 2010).
56 The dynamics of the AAA system are not only reflected in horizontal motions, but also seen in
57 the remarkable geomorphic features indicating rapid and recent vertical motions. High plateaux are
58 diffuse within the system from the Ethiopia-Arabian dome to the Turkish-Iranian plateau, both
59 standing at ~2 km elevation (Sengör et al., 2003). In contrast, subsiding regions are found behind
60 the Aegean trench and in the Makran-Persian Gulf. The vertical motion of these areas is Neogene to
61 Recent and its origins are still poorly understood. Residual topography, computed by removing the
62 crustal isostatic topography component from the actual present-day topography, shows positive
63 features over the Ethiopia - Kenya swells extending northward along the Red Sea and Anatolia-Iran,
64 while a negative signal is seen along the Hellenic trench-Persian Gulf (Figs. 1b and c). This
65 indicates that at least part of the topography could be related to mantle dynamics (Daradich et al.,
66 2003; Moucha and Forte, 2011; Becker and Faccenna, 2011).
67 Volcanism is diffuse over the region and it is mainly focused on the large Ethiopian-Yemen
68 igneous province (Fig. 1b). Intraplate volcanism is also present in Syria-Jordan and Israel far from
69 the plate boundary with large basaltic effusions and within the collisional zone, i.e. Eastern
70 Anatolia, where the volcanic signature gets complicated by the mixed contributions of different
71 geochemical sources. Figure 1b shows a general agreement between volcanism and positive
72 residual topography, suggesting a link between mantle dynamics, volcanism, and dynamic
73 topography.
74 Figures 1c and d show shear-wave (mainly SKS) splitting based estimates of azimuthal seismic
75 anisotropy of the upper mantle (from the compilation of Wuestefeld et al., 2009, 05/2012 update),
76 on top of dynamic topography, now as estimated directly from free air gravity data using the same
77 conversion and filtering as in Craig et al. (2011). The fast polarization axes of SKS splitting
78 correlate poorly with the mantle shear that may be inferred from plate motions both in Anatolia and
3 79 in Arabia (e.g. Hansen et al., 2006, 2012), illustrating the complexity of the system. The large-scale
80 mantle structure as imaged by seismic tomography shows low velocity anomalies rooted beneath
81 the Afar system as well as in northern Arabia-Syria-Jordan, and high velocity anomalies beneath the
82 Aegean and the Zagros-Makran trenches (Fig. 1d, also see Hansen et al., 2012). This pattern
83 suggests that the deeper mantle component of the AAA system is dominated by strong seismic
84 velocity contrasts and velocity anomalies that are indicative of large temperature differences, and
85 hence vigorous convection.
86 Several prior studies analyzed the kinematics and the dynamics of the region, addressing such
87 problems as to why Anatolia is moving westward, what is driving Arabia toward the north, or why
88 the Hellenic trench retreats so fast, and different models have been proposed. The presence of an
89 active subduction system and the westward acceleration of the geodetic field argue in favour of the
90 action of the pull of the Aegean slab system as the main engine, sucking Anatolia - Aegea toward
91 the trench (Le Pichon, 1992; Faccenna et al., 2006; Le Pichon and Kreemer, 2010). The motion of
92 Arabia is more problematic. It has been related to the slab-pull along the Makran trench (Bellahsen
93 et al., 2003; Reilinger and McKlusky, 2011), even if there is clear tomographic evidence that the
94 slab is broken beneath the Bitlis suture zone (Keskin, 2003, 2007; Şengör et al., 2003, 2008;
95 Faccenna et al., 2006). Another possible explanation is the Afar plume pushing Arabia northward,
96 even if the constant velocity attained by Arabia during the last ~30 Ma (Reilinger and McKlusky,
97 2011) contrasts with the expected progressive decrease in the velocity from the main upwelling
98 plume phase, ~32 Ma (van Hinsbergen et al., 2011b). More recently, it has been proposed that
99 mantle drag related to a large-scale convection cell could efficiently pull continental plates toward
100 collisional zones, thus favouring indentation (Alvarez, 2010; Becker and Faccenna, 2011). This
101 model also agrees with previous studies which suggested that Arabia is dynamically tilted by mantle
102 convection toward the Persian Gulf (Daradich et al., 2003; Moucha and Forte, 2011).
103 However, basic questions still remain unsolved. For example, it is unclear how the large-scale
104 convection cell associated with the African deep mantle plume relates to the smaller-scale one
4 105 associated with the pull of the Aegean slab, and what causes intraplate volcanism in the Middle
106 East. Moreover, it is also unclear if the high plateaux in the collisional belt are caused by active
107 mantle upwellings, lithospheric delamination, or by crustal thickening during collision.
108 The aim of this paper is to test different hypotheses about the driving forces of the Arabia-
109 Anatolia-Aegean system. Our results show the importance of mantle flow for driving the system,
110 reconciling Afar plume volcanism, the collision within the Bitlis, and the rapid backward
111 acceleration of the Aegean trench.
112
113 2. Mantle dynamics inferred from tomography and anisotropy
114 Several structural models have been proposed for the mantle underneath the Africa-Middle East-
115 Aegean region. Most seismological models consistently show a marked low velocity region over the
116 Ethiopian-Yemen dome and expanding north toward the Red Sea up to Arabia and Middle East
117 (Figs. 2 and S1). Whether the Afar and Arabian low velocity zone is rooted or not in the lower
118 mantle, and, more specifically, whether it is connected or not to the African large shear velocity
119 zone are debated questions. 3D shear wave tomography first showed the linkage between the deep
120 mantle structure and the Ethiopian province (Ritsema et al., 1999; Mégnin and Romanowicz, 2000;
121 Grand, 2002; Simmons et al., 2009; 2010; Ritsema et al. 2011). Similarly, global P-wave
122 tomography models explored the connection between the Afar hot region and deep structures
123 (Montelli et al., 2006; Li et al., 2008; Hansen et al., 2012). Regional models explored the presence
124 and the position of the low velocity anomalies in some detail (Benoit et al., 2006; Bastow et al.,
125 2005; Bastow, 2008) also showing a deep deflection of the low velocity anomaly towards the lower
126 mantle one (Park et al., 2007, 2008; Hansen et al., 2012).
127 In northern Arabia, upper mantle, low velocity anomalies have been imaged by both P-wave and
128 S-wave models (Benoit et al., 2006; Park et al., 2007, 2008). Recently, regional S-velocity structure
129 for Arabia and East Africa has been inferred by jointly inverting regional waveforms, group
130 velocities, and travel times (Chang et al., 2010; Chang and van der Lee, 2011), indicating that the
5 131 low velocity zone beneath Afar is perhaps distinct from the low velocity zone beneath northern
132 Arabia, implying a two plume origin. Conversely, other recent models favour a shallow upper
133 mantle connection of the Arabia low velocity anomaly to the Afar plume, i.e. a single plume,
134 perhaps dragged northward (Figure 2; Hansen et al., 2012).
135 Regional and global studies also show low velocity anomalies in the uppermost parts of the
136 upper mantle below Anatolia, above what can be inferred to be a deformed slab that is stagnating
137 over the transition zone, partially detached from the lithosphere (Faccenna et al., 2006; Lei and
138 Zhao, 2007; van Hinsbergen et al., 2010; Biryol et al., 2011). The Hellenic subduction zone shows a
139 40°-50° north-dipping, high velocity anomaly marked by seismicity down to intermediate depth (~
140 200 km). Lower mantle, high velocity anomalies have also been detected beneath the Hellenic
141 trench down to ~1400 km (Piromallo and Morelli, 2003; Faccenna et al., 2003; Bijward et al., 1998;
142 van Hinsbergen et al., 2005). High velocity anomalies are present also beneath the Zagros-Makran
143 collision zone (Agard et al., 2011) but the resolution over this area is relatively low.
144 Figure 2 shows a comparison between structure seen in two global S-wave tomography
145 approaches, the composite model SMEAN (Becker and Boschi, 2002), and the more recent model
146 TX2008 (Simmons et al., 2009), as well as a P-wave model, HNB12 (Hansen et al., 2012), which
147 includes more regional data than the global study by Li et al. (2008), on which it is based. It can be
148 seen that: i) Low-velocity anomalies are well defined beneath Afar, connect to the lower mantle,
149 and extend north toward the Red Sea; ii) low-velocity anomalies are also found beneath the
150 northern Arabia-Middle East region but there, they do not seem to connect to the lower mantle. The
151 recent P-wave model (Hansen et al., 2012) shows a more pronounced slow anomaly compared to
152 the other global models, widespread at shallow levels also beneath Anatolia, and Iran; iii) high
153 velocity anomalies are located beneath the Aegean and the Persian Gulf; they get more pronounced
154 and enlarge downward at mid-mantle level, as part of the larger Tethyan anomaly.
155 Mantle anisotropy as reflected in shear wave splitting shows the following features (Fig. 1d):
156 roughly NE-SW trending apparent fast polarization axes north of the East African rift in Ethiopia
6 157 (e.g. Gashawbeza et al., 2004) which rotate to mainly N-S “fast orientations” in Arabia and Israel
158 (e.g. Hansen et al., 2006; Kaviani et al., 2009), a mainly NE-SW trending in Anatolia and in central
159 Iran (e.g. Biryol et al., 2011), turning to a more northerly trend in the Aegean, and a more NW-SE
160 direction below the Zagros. Regional-scale patterns of relatively simple azimuthal anisotropy
161 alternate with more complex signatures, such as a mix of null observations and large spatial
162 gradients in fast orientations within Iran (e.g. Kaviani et al., 2009).
163 The interpretation of anisotropy in terms of mantle flow is not straightforward (see, e.g., Long
164 and Becker, 2010, for a recent review), and both lithospheric and a mantle origins of the observed
165 patterns have been suggested for the region (e.g., Hansen et al., 2006; Biryol et al., 2011). Below,
166 we will explore the predictions for anisotropy from mantle flow modeling; those can provide a
167 baseline to evaluate the feasibility of suggested asthenospheric flow scenarios such as northward
168 transport of a plume.
169
170 3. Quantitative mantle flow and anisotropy modeling
171 To quantify the contribution of mantle heterogeneity such as due to plumes and slabs to mantle
172 circulation and plate motions, we compute instantaneous mantle flow from seismic tomography by
173 converting velocity anomalies to temperature (e.g., Hager and Clayton, 1989; Lithgow-Bertelloni
174 and Silver, 1998; Moucha and Forte, 2011). We solve the Stokes equation for incompressible flow
175 in a global, spherical shell using a finite-element approach (Zhong et al., 2000). The basic approach
176 is discussed in Becker and Faccenna (2011) and Miller and Becker (2012). We test several,
177 geologically-motivated models of lateral viscosity variations as well as several density models to
178 explore the effects of the different components of the AAA convection system. The goal is not so
179 much to find the best model that matches all data at this point, but rather to identify some of key
180 regional convective processes as reflected in geological and geophysical observables.
181 The boundary conditions of the mantle flow computations are shear-stress free at the surface,
182 allowing for dynamically-consistent plate motions guided by weak zones prescribed at plate
7 183 boundaries (Ricard and Vigny, 1989; Zhong et al., 2000). Results are presented for a regional
184 zoom-in of the global models around the Arabia-Middle East, with all velocities rotated into a
185 Eurasia-fixed reference frame whose corresponding best-fit Euler pole is obtained from the surface
186 velocities within that plate.
187 The reference density structure for the mantle is based on SMEAN, with fast anomalies within
188 inferred continental keels removed above 300 km to correct for shallow compositional effects. We
189 use a Rayleigh number of 3.4·108 (defined as in Zhong et al., 2000) with temperature scaled such
190 density anomalies go with shear wave velocity anomalies as dln ρ /dln VS = 0.2. This constant
191 scaling is adopted for simplicity; it does not take depth-dependent mineral physics relationships into
192 account (e.g., Forte, 2007), nor do we allow for compositional anomalies besides keels. However,
193 such complications should mainly affect the amplitude but not the geometry of mantle flow (cf.
194 Gurnis et al., 2000).
195 Density models can also be inferred for slabs alone, based on seismicity in Wadati–Benioff
196 zones and/or inferred past subduction (e.g. Hager, 1984; Ricard et al., 1993; Lithgow-Bertelloni and
197 Richards, 1998), or by mixed models that explore the respective roles of slabs versus tomography
198 and edge forces (e.g. Becker and O'Connell, 2001; Conrad and Lithgow-Bertelloni, 2002; Ghosh et
199 al., 2010; Stadler et al., 2010). Here, we consider upper mantle models that are entirely or partially
200 based on slabs inferred from the Engdahl et al. (1998) seismicity based, RUM slab geometry of
201 Gudmundsson and Sambridge (1998), following Ghosh et al. (2010). No crustal density anomalies
202 or other effects leading to lithospheric driving forces due to gravitational potential energy
203 variations, besides those imaged by tomography, are considered.
204 We use a Newtonian viscous rheology with background viscosities for the lithosphere (depth z ≤
205 100 km), upper mantle (100 < z ≤ 660 km), and lower mantle (z > 660 km) of 5 ⋅ 1022, 1021, and
206 5 ⋅1022 Pas, respectively. Within the lithosphere, relative viscosities are 0.01 for weak zones along
207 plate boundaries, defined as in Becker and Faccenna (2011) and outlined in Fig. 1 for the reference
208 model). This simple model achieves a good general match to global plate velocities and the geoid
8 209 (Becker, 2006; Miller and Becker, 2012), and we explore additional lateral viscosity variations as
210 described below.
211 The mantle flow model results for the region are then evaluated in the light of the present-day
212 geodetic velocity fields by comparing predicted horizontal surface motions. Additionally, we
213 consider geomorphological information concerning the vertical motion and predictions of “dynamic
214 topography”. For the latter, the normal stresses generated by viscous flow at the surface are used to
215 infer the instantaneous surface deflection that would be expected because of convection (Ricard et
216 al., 1984; Richards and Hager, 1984). Assuming that all non-crustal topography is due to active
217 flow, this signal can be compared to observed topography once the effect of crustal isostatic
218 adjustment is removed to arrive at “residual topography” as shown in Fig. 1b (e.g. Daradich et al.,
219 2003; Gurnis et al., 2000).
220 Another constraint for mantle flow is provided by seismic anisotropy which in the upper mantle
221 most likely reflects the alignment of intrinsically-anisotropic, single crystals of olivine into lattice-
222 preferred orientation (LPO) under mantle flow (e.g., Long and Becker, 2010). Our approach
223 consists of assuming that flow is steady-state for the few Myr needed to achieve saturated textures,
224 following tracer particles in the mantle circulation until stable LPOs, as predicted by the mineral
225 physics theory of Kaminski et al. (2004) are reached at each desired location. Then, we assign
226 pressure and temperature-dependent elastic tensors to the resulting 3D model of anisotropy from 50
227 to 410 km depth (Becker et al., 2006a). From this elastic anisotropy, we compute synthetic
228 seismograms assuming lateral homogeneity on Fresnel zone lengths (~100 km), from which the
229 expected splitting can be computed for arbitrary back-azimuths (Becker et al., 2006b). However,
230 here we will only consider station-averaged splits (cf. Miller and Becker, 2012) and mainly discuss
231 our models in terms of the average angular misfit, 〈Δα〉, of apparent fast axes for all available non-
232 null splitting measurements (334 for the study region; Fig. 1d).
233
234 4. Results
9 235 We next illustrate the dynamics of the system using a few selected models to highlight the main
236 driving processes. To isolate the role of suction produced by the subducting slab, we first show
237 results from a model where a negative (slab) density anomaly is localized along the Aegean Wadati-
238 Benioff subduction zone only (Fig. 3a). The effect of the subducting Aegean slab produces mantle
239 suction, driving the Aegean and Anatolia toward SW and WSW, respectively. The Aegean is
240 moving rapidly, similar to what is observed, and Anatolia is moving at ~1 cm/yr, accounting for
241 ~half of its geodetic motion. The suction exerted by the slab produces an unrealistic westward
242 motion of both Arabia and Africa, giving an overall poor regional correlation of plate velocities.
243 The pattern of motion produced by the slab produces a rather good fit with the anisotropy in
244 Anatolia and eastern Aegean, but not in Arabia nor in the central Aegean. This result shows that
245 part of the mantle motion beneath Aegea and Anatolia is likely related to the subduction of the
246 Aegean slab.
247 The influence of the large-scale convection patterns induced by lower mantle density anomalies
248 as inferred from tomography when added to that of the Aegean upper mantle slab is shown on
249 Fig. 3b. Compared to the previous model, the velocity of Arabia is faster, on average ~1 cm/yr,
250 matching about half of its actual speed, and moving in the right direction. Africa's motion is also
251 faster, faster than it should be, but also with a more northerly direction. In addition, the velocity
252 field of Fig. 3b matches the gradual velocity decrease observed across the Zagros and northern
253 Iranian plateau, then fading out north of Kopet Dagh and Alborz. This model also provides a rather
254 good match for the mantle anisotropy (〈Δα〉 ~ 30°) over the entire area, with the exception of the
255 Zagros and south Arabia. The results from the model of Fig. 3b illustrate how superimposing a
256 large-scale convection cell, driven by the superswell beneath Africa and the downwelling over the
257 Tethyan slab (Becker and Faccenna, 2011) to the small scale convection related to the Aegean slab
258 improves the solution both in terms of plate velocity direction and mantle flow, although Arabia and
259 Anatolia velocities are underestimated. Dynamic topography shows a large scale tilting of Arabia
1 0 260 towards the NE, in agreement with previous predictions (Daradich et al., 2003) with maximum
261 uplift centred in southern Ethiopia.
262 To evaluate the influence of the upper mantle related to the Afar-Red Sea-Anatolia system, the
263 slow (“hot”) upper mantle anomalies are added to the previous model in Fig. 3c. This leads to an
264 overall increase of the plate velocities, as expected, produced by the vigorous upwelling centred on
265 the Afar and flowing toward the Aegean subduction zone. In this model, Arabia and Anatolia both
266 move in the right direction, while Arabia moves faster and Anatolia slightly slower than observed
267 and overall correlation for regional velocities is quite good. As expected, the contribution of the
268 upper mantle anomaly within the system increases the dynamic topography showing a very strong
269 signature in the Afar and Carlsberg Ridge (~2 km of topography) and some hundreds of meters
270 positive elevation in eastern Anatolia and the Caucasus, as discussed earlier by Boschi et al. (2010).
271 The faster Arabia motion degrades slightly the correlation of the mantle anisotropy (〈Δα〉 = 35°)
272 with respect to the model of Fig. 3b. This is particularly evident in southern Arabia, where
273 upwelling produces radial mantle flow striking perpendicularly to the observed NS trending
274 anisotropy.
275 The effect of lateral strength variations in modifying seismic anisotropy and flow is explored by
276 adding a low viscosity channel to the model of Fig. 3b. The low viscosity channel runs from the
277 Afar to the Red Sea, where mantle lithosphere is imposed to be thinner than the surrounding. This
278 modification increases the match with anisotropy 〈Δα〉 to 28°, and particularly improves the match
279 in Arabia (Fig. 3d). This exploration of the low viscosity channel provides a quantitative test of
280 previous suggestions that the Afar upwelling is guided within a low viscosity mantle north beneath
281 Red Sea and western Arabia (Hansen et al., 2006) or towards the north beneath Arabia (Chang et
282 al., 2011).
283 The geometry of typical mantle convection patterns for the model of Fig. 3c can be better shown
284 on map view at different depths and cross sections (Fig. 4). At shallow upper mantle depth, the
285 sinking of the Aegean slab and the upwelling of the mantle in Afar generates a large-scale current,
1 1 286 upwelling in the Afar and downwelling in the Aegean, leading to some toroidal motion when
287 interacting with the plate geometry. The convection cell is well represented by the cross sections
288 running from Afar to the Persian Gulf and to the Aegean trench (Figs. 4e, 4f). In addition, a shallow
289 upwelling in eastern Anatolia generates a smaller-scale convection feature from the Aegean trench
290 to the Anatolian plateau. Much of these convection patterns are related to upper mantle motions,
291 which efficiently drag plates and produce pronounced positive dynamic topography at the surface.
292 At the base of the upper mantle, these perturbations flatten out, generating a more regular pattern of
293 mantle flow trending toward WNW.
294 We computed mantle flow and seismic anisotropy for a large number of models, and a
295 representative selection of results for different density models and different choices for lateral
296 viscosity variations are shown in Fig. 5. Such a representation in misfit matrix form allows
297 assessing the combined effects of the different parameters that were illustrated in Fig. 3 (Miller and
298 Becker, 2012). Moreover, the role of density models and the relevance of different model
299 assumptions, such as on the geometry of weak zones, and regional plate coupling, can be assessed.
300 Fig. 5 shows misfits as measured by linear correlation coefficient for global plate velocities (a),
301 regional plate velocities (b), and for the average misfit of the fast axes from seismic anisotropy (c).
302 At a regional scale, the best predictions of plate velocities when measured in terms of correlation
303 are obtained by adopting the SMEAN composite model with upper mantle slabs added. Overall, the
304 presence of a sub-oceanic plate asthenosphere improves the fit for velocity (e.g. Becker 2006). This
305 is also true for the presence of a low viscosity channel beneath Red Sea (cf. Hansen et al., 2006),
306 and Fig. 5 provides an estimate of the sensitivity of details of the viscosity variations underneath
307 Arabia on shear wave splitting estimates.
308 The delay time amplitudes of splitting measurements are a measure of the strength of anisotropy,
309 yet typically less stable than the “fast azimuth”. In our modeling approach, delay times are mainly
310 controlled by the degree of LPO saturation. Guided by the good fit to SKS splitting delay times
311 achieved in other regions (Becker et al., 2006b, Miller and Becker, 2012), we here chose a
1 2 312 logarithmic strain of 0.75. For the AAA system, this choice leads to an overprediction of delay
313 times by ~0.4 s (Fig. 3). We mainly take this an indication for the heterogeneity and complexity of
314 seismic anisotropy in the study region (Figs. 1c and d). This may indicate the role of lithospheric,
315 frozen-in anisotropy, in obscuring the asthenospheric signal in some regions such as in Iran or the
316 Hellenic, but delay times will not be considered further in this study.
317 Overall, we find that the match to the shear wave splitting fast axes is very sensitive to details of
318 the lateral viscosity variations, and in regions such as Iran we are close to the limit of what can
319 usefully be explored with our global mantle flow models with homogeneous lithospheric rheology.
320 Further model refinement may thus allow a better simultaneous fit to crustal velocities, dynamic
321 topography, and seismic anisotropy. However, for our purposes, the results in terms of the major
322 drivers of horizontal and vertical motions of the AAA plate system dynamics are robust.
323 In summary, the global circulation models show the importance of the combined effect of slab
324 pull and plume push to match the velocity field of microplates in the Middle East. Model results
325 indicate that slab pull is essential to drive Anatolia and the Aegean microplate toward the west and
326 southwest, respectively. A large-scale upwelling in the Afar can be channelled into the upper mantle
327 beneath the Red Sea, and such flow is necessary to match the motion of Arabia correctly. The
328 overall convection patterns are dominated on different scales by convection systems that link the
329 upwelling areas in Afar up to northern Arabia, to the shallow Aegean slab and, at larger depth, with
330 the Tethyan slab.
331
332 5. Geological constraints on mantle convection
333 Geological evidence implies that the convection patterns reconstructed from our circulation
334 models for the present-day were activated in the Oligocene-late Eocene. The first trace of volcanism
335 in Ethiopia dates back to ~45 Ma (Ebinger et al., 1993). It is probably at that time that a large
336 volume of mantle emplaced at the base of the lithosphere (Sleep and Ebinger, 1998) while Africa
337 was still a coherent plate, slowly moving NE toward, Eurasia consuming the Tethyan ocean along
1 3 338 the Aegean-Bitlis trench system (Dercourt et al., 1986; Barrier and Vrielynck, 2010). The deep
339 mantle structure associated with the evolution of the Aegean subduction has previously been
340 reconstructed from the seismically imaged high velocity anomalies, restored accounting for time-
341 dependent convergence kinematics (Faccenna et al., 2003; van Hinsbergen et al., 2005). The match
342 between the absolute motion of the trench and the deep position of the high velocity anomalies into
343 the lower mantle gives a maximum age for the entrance of the Aegean slab into the lower mantle of
344 ~50+10 Ma, probably after a stagnation phase of ~10-20 Myr duration on the 660 km discontinuity
345 (Faccenna et al., 2003; 2012). This tectonic episode is extending regionally toward the Himalayan
346 trench (Replumaz et al., 2004). The episode of slab avalanching at ~45-50 Ma in the Hellenic
347 domain has been interpreted as the cause of stretching of the Sirte basin in north Africa (Capitanio
348 et al., 2009). We infer that the large-scale, Tethyan convective cell that is active today (Becker and
349 Faccenna, 2011) between Afar and the Tethyan subduction zone was already operating at ~45 Ma.
350 The most apparent expression of the deep mantle upwelling is the outpouring of the Ethiopian
351 large igneous province at ~33-30 Ma (Hoffman et al. 1997, Fig. 6a). The Red Sea and the Gulf of
352 Suez (Fig. 1) rifted at the beginning of the Oligocene (Ghebread, 1998). In the Red Sea, dykes and
353 basaltic floods were emplaced during the Oligocene (Ghebread, 1998; Bosworth et al., 2005) and
354 syn-rift deposits accumulated from the Oligocene-Miocene (~30-25 Ma) to Miocene (14 Ma)
355 (Bayer et al., 1988; Garfunkel and Beyth 2006; ArRajehi et al. 2010). Thermochronology shows
356 that uplift of the Red Sea rift shoulders commenced at ~27 Ma and peaked at ~20 Ma (Bosworth et
357 al., 2005) and the erosion of the Blue Nile gorges commenced at 25-29 Ma (Pik et al., 2003).
358 Seismic data show a marked change in sedimentation in the Nile delta both offshore and onshore
359 and both in the Levantine and Herodotus Basin at the Eocene-Oligocene boundary. This probably
360 marks the onset of the proto-Nile drainage (Steinberg et al., 2011). Stratigraphic and morphological
361 studies show the presence of a regional truncation surface developed during the Oligocene from the
362 northern Red Sea to southern Levant related to the formation of an ~3000 km × 1500 km crustal
363 dome above the Afar plume, with a widespread fluvial erosion that truncated the northern edge of
1 4 364 this dome (Avni et al., 2012). Evidence of episodic Miocene to Pliocene regional uplift are also
365 found in Isreal (Gvirtzman et al., 2011). First syn-rift sediments in the Red Sea are dated there to
366 the late Eocene-early Oligocene, ~35 Ma (Roger et al, 1989; Watchorn et al. 1998). The eastern
367 arm of the rift system is represented by the Gulf of Aden (Fig. 6b). Extension occurred in two
368 directions, first N20°E and then N160°E (Huchon et al., 1991; Lepvrier er al., 2002; Huchon and
369 Khanbari, 2003). The absolute northward motion of Africa progressively slowed down after 35 Ma
370 (Jolivet and Faccenna 2000; Reilinger and McKlusky, 2011) while Arabia-Eurasia convergence
371 kept a constant velocity of ~2-3 cm/yr with respect to Eurasia from 56 Ma to the present, and
372 Africa-Eurasia convergence slowed down to ~1 cm/yr after ~25 Ma once the opening of the Red
373 Sea had started (McQuarrie et al., 2003).
374 The timing of the onset of collision in the Bitlis-Zagros is debated. Some authors proposed that
375 continental subduction initiated around the Oligocene (Jolivet and Faccenna, 2000; Allen and
376 Armstrong 2008; Agard et al. 2005; 2011) whereas others argue that orogenic structures were built
377 in the Early Miocene (Berberian and Berberian, 1981; Okay et al., 2010). This latter hypothesis is
378 supported by the timing of the closure of the Tethyan gateway ~18 Ma (Reuter et al., 2009),
379 allowing the deposition of detritus issued from the Eurasian margin onto Arabia. In addition, in
380 Eastern Anatolia last marine deposits date back to the Early Burdigalian (~12 Ma, Gelati, 1975)
381 whereas in the central Anatolia plateau they post-date the Messinian (Cosentino et al., 2012). The
382 stratigraphic record thus indicates that thrust systems in the Bitlis emerged from shallow water ~16
383 Ma. Fission track data indicate a period of increased exhumation around the mid-late Miocene
384 (Okay et al., 2010). Most probably, the onset of continental subduction occurred in the Oligocene,
385 while the final closure of the Tethyan gateway and creation of an orogenic relief later, around the
386 Middle Miocene. Volcanism initiated at HAS and in Northwest Syria ~ 25 Ma ago. The occurrence
387 of volcanism at HAS and in NW Syria has been interpreted as due to the arrival of hot material with
388 melting and magma ascent, enhanced by or enhancing crustal extension (Krienitz et al., 2007;
389 2009).
1 5 390 Around 11+2 Ma the velocity of Africa slowed down by ~50% for most velocity models
391 (Reilinger and McKlusky, 2011) while the motion of Arabia was not strongly perturbed keeping a
392 constant velocity at ~2 cm/yr (Burke, 1996; McQuarrie et al., 2003; Muller̈ et al., 2008). McQuarrie
393 et al. (2003) and Reilinger and McKlusky (2011) described a major re-organization of plate
394 boundaries at 9+4 Ma (Fig. 6c). At that time, Arabia was fully separated from Africa with the
395 complete opening of the Gulf of Aden (Reilinger and McKlusky, 2011). During Miocene times, the
396 activity of the Gulf of Suez decreased, transferred to the Gulf of Aqaba with the consequent
397 formation of the Levant fault, while collisional processes in the Bitlis- Pötürge massif are already
398 well developed. Contemporaneously, a surge of volcanism occurred in NW Syria in the Aleppo
399 region (Krienits et al., 2007; 2009) which has been related to asthenosphere upwelling during slab
400 break-off (Keskin, 2003, 2007; Krienitz et al., 2009; Şengör et al., 2003, 2008).
401 In Eastern Anatolia, the geochemical character of volcanism changed at around 10 Ma from an
402 original calc-alkaline, subduction-related type towards a more alkaline one (Keskin et al., 2007). In
403 particular, the the presence of calc-alkaline volcanism with a distinct subduction signature across
404 Eastern Anatolia has been related to a first episode of slab-steepening event between ~15 and 10
405 Ma, before the slab break off at ~10 Ma (Keskin, 2003). This slab steepening episode likely
406 exposed the base of the Eastern Anatolian accretionary complex to asthenospheric temperatures and
407 dragged subduction-related mantle material southward from beneath the Pontides toward the mantle
408 wedge (Keskin, 2003, 2007). Petrological studies also show that the slab steepening episode was
409 likely followed by slab rupture at ~10 Ma, resulting in the formation of a slab window. This is
410 attested to the eruption of the first alkaline basanitic lavas in the Eastern Anatolian volcanic
411 province at ~10 Ma (Lebedev et al., 2010) in the north of Lake Van (Keskin et al., 2012c). This
412 implies that enriched and fertile asthenospheric, hotter material once located beneath the slab (i.e.,
413 underneath the Arabian continent) moved northward to shallow depths through a slab window
414 (Keskin, 2007). I.e., the convective flow possibly coming from the Afar plume may be the source of
415 the volcanism across SE Anatolia (Keskin et al, 2012b) (Fig. 5). The North Anatolian Fault
1 6 416 probably initiated in east Anatolia (~10-12 Ma) (Sengör, 1979), propagating westward at ~7-6 Ma
417 to central Anatolia, and arriving in the Aegean in the Pliocene. Keskin et al. (2012a, b) show that
418 the extensive volcanic field of the Karacadağ volcanic area of SE Anatolia was sourced by a garnet-
419 bearing, deep asthenospheric mantle which is similar to that of Afar in terms of its Sr, Nd, and Pb
420 isotopic ratios (Fig. 7).
421 The evolution of the volcanic system displays a progressive decrease in the degree of
422 involvement of a shallow, subcontinental lithospheric mantle source toward a deeper,
423 asthenospheric one, with isotopic composition similar to Afar lavas. From the Pliocene onward,
424 widespread volcanism occurred in Syria along the Dead Sea Fault but also in Harrat Ash Shamah
425 (HAS), NW and NE Syria, and SE Turkey (e.g. Karacadağ: Lustrino et al., 2010, 2012; Keskin, et
426 al., 2012a, 2012b; Ekici et al., 2012) contemporaneous with the activity of large volcanic fields
427 along the Red Sea coast (e.g., Giannerini et al., 1988) and during the emplacement of ocean floor in
428 the Red Sea (Girdler, 1990). In the HAS, effusion rates are similar to Ethiopia (Hofmann et al.,
429 1997), with ~700 m thick of plateau basalts erupted in less than 1 Ma (Krienitz et al., 2009). As for
430 Karacadağ, the HAS intraplate volcanic is also derived from a mixed lithosphere-asthenosphere
431 source, activated by lithosphere extension (Shaw et al., 2003).
432 The similarities in isotopic composition between the lavas of the Afar region and those of the
433 Arabian province and the expected increase of the mantle temperature during the diffuse phase of
434 volcanism has been related to a sub-lithosphere source related to the Afar plume (Camp and
435 Roobol, 1992; Krienitz et al., 2007, 2009), supporting a model of northward directed mantle flow
436 (Ershov and Nikishin, 2004; Yarmolyuk et al., 2004) that at ~10 Ma reached the Eastern Anatolian
437 Collision Zone after the slab steepening and breakoff (Keskin, 2003, 2007; Keskin et al., 2012c).
438 The temporal increase in the alkalinity coupled with a gradual decrease in the subduction
439 component starting at around 10 Ma in the collision-related volcanic units of Eastern Anatolia
440 supports this interpretation (Keskin et al., 2012c).
441
1 7 442 6. Discussion: Geodynamic evolution of the AAA system
443 In the context of geological and geophysical constraints on the tectonic history of the region, our
444 geodynamic modeling results indicate that the upwelling mantle underneath Afar, the northward
445 motion of Arabia with consequent collision in the Bitlis, the westward escape of Anatolia, and
446 backarc extension in the Aegean may all be the result of the progressive evolution of the Tethyan
447 convection cell (Figs. 6 and 8).
448 Combined effects of an upwelling in Afar and subduction in the Hellenic drag Arabia toward
449 collision. The upwelling in the Afar likely commenced at ~40 Ma but massive volcanic outpouring
450 and the formation of a large igneous province is dated at ~33 - 30 Ma. The volcanic record shows
451 that magmatism propagated northward from Ethiopia-Yemen toward the Arabia-Anatolia
452 collisional zone. The spatio-temporal distribution of volcanism, and the similarities of the isotopic
453 compositions of the late stage lavas, suggest that the origin of these melts is most likely related to a
454 sub-lithosphere source resembling the one erupted in Arabia, and connected with the increased
455 temperature due to the emplacement of the Afar plume (Krienitz et al., 2009). The isotopic
456 signature of the Afar plume can be traced up to the Karacadağ neovolcanic area on the northernmost
457 part of the Arabian continent in Turkey (Fig. 7).
458 Those data then support a model of northward directed mantle flow (Ershov and Nikishin, 2004;
459 Yarmolyuk et al., 2004), channelling the Afar plume head beneath Arabia (Camp and Roobol, 1992;
460 Ebinger and Sleep, 1998; Hansen et al., 2006) and further north beneath Eastern Anatolia after slab
461 breakoff (Keskin, 2007). Ebinger and Sleep (1998) proposed that rifts may represent natural
462 channels for hot material from mantle plumes to flow horizontally beneath thin lithosphere. Ershov
463 and Nikishin (2004) argue that the lithosphere is relatively thin along a south-north line extending
464 from Afar to Anatolia due to a previous orogenic event. Alternatively, Chang and van der Lee
465 (2011) suggested that the low velocity anomaly beneath the Middle East is distinct from the Afar
466 one and is rooted in the lower mantle. However, other tomography models indicate that the slow
467 velocity beneath Arabia is limited to the upper mantle (Figs. 2, S1). For example, Hansen et al.
1 8 468 (2012) find that the Arabia low velocity zone is connected to the Afar plume in the upper mantle,
469 eventually connecting down to the lower mantle African low shear wave velocity anomaly. We
470 consider this single plume scenario the simpler explanation of deep mantle structure, on which local
471 complexities and volcanic trends may well be superimposed.
472 Shear-wave splitting and tomographic results of Hansen et al. (2006) and Park et al. (2007, 2008)
473 and the simplified models of Ershov and Nikishin (2004) suggest channelled, horizontal mantle
474 flow from the Afar hotspot. Our modelling substantiates these suggestions and shows that shear
475 wave splitting and mantle anisotropy below Arabia can be fit best when introducing a low viscosity
476 zone, running along the Red Sea which would channel hot mantle material northward. The spatio-
477 temporal distribution of within plate alkaline volcanism in Arabia also suggests that the deep source
478 migrated northward from Yemen (31 Ma) to northern Syria (19 Ma) at a high, average speed of ~
479 22 cm/yr (Krienitz et al., 2009) even though this estimate could be significantly reduced by
480 considering a larger initial extent of the Afar plume (Ebinger and Sleep, 1998). Our models show
481 that the velocity of sub-lithosphere material can reach velocity around 10 to 20 cm/yr for models
482 accounting for a low viscosity asthenosphere, or temperature dependent viscosity, and this
483 asthenospheric flow could be even faster if channelled along the NS corridor.
484 Our flow computations also show that the Afar - Arabia upwellings drive Arabia northward,
485 while dynamically tilting the entire plate toward the Persian Gulf (Daradich et al., 2003). The
486 asthenospheric mantle is “leading the plate”, and the present collision with Eurasia is inferred to be
487 driven by a mantle “conveyor belt” (Becker and Facenna, 2011). The steady, ~2 cm/yr northward
488 motion of Arabia prior and after its separation with Africa has been taken as an indication of the
489 role of slab-pull from the Zagros-Bitlis to move the plate northward (Reilinger and McKlusky,
490 2011). Conversely, we take the relatively constant plate speed, irrespectively of the collision and
491 break-off episodes (e.g. Keskin, 2003, 2007; Şengör et al., 2003, 2008Faccenna et al., 2006), as a
492 sign that other forces rather than slab-pull drive Arabia northward. We conclude that mantle
1 9 493 upwellings in the Afar-Arabia first separated Arabia from Africa, and then propelled Arabia
494 northward.
495 Following Burke (1996) and Jolivet and Faccenna (2000) we suggest that the reduction of the
496 absolute velocity of Africa after ~30-35 Ma to the present is a consequence of the collision in the
497 northern Zagros and Bitlis suture zone and that Arabia was detached from the slowing main African
498 plate driven by the continuous basal drag imposed by the asthenospheric flow underneath. Arabia
499 then rotated counterclockwise as a consequence of the asymmetric conditions imposed by the new
500 plate boundary geometry, as suggested in Jolivet and Faccenna (2000) and Bellahsen et al. (2003).
501 The massive outpouring of volcanic material in the Afar triple junction area is then allowed by the
502 extension recorded in this area.
503 Downwelling system in the Aegean and the extensional system. Subduction in the Aegean-
504 Tethyan system extended from the Mediterranean to Asia (van der Voo et al., 1999), but the timing
505 of the tectonic evolution of the Tethyan margin appears to also reflect temporal fluctuations in
506 mantle upwellings, encouraging the view of a coupled system. In the Aegean, the penetration of a
507 slab into the lower mantle probably occurred at ~50-40 Ma (Faccenna et al., 2003; Capitanio et al.,
508 2009), contemporaneously with the onset of Ethiopian volcanism. This indicates that those two
509 processes overlap in time, producing the onset of a large-scale, whole mantle convection cell.
510 Subduction in the Aegean exerts a first order control on the evolution of the AAA system. The
511 efficiency of slab pull on the Middle east system is illustrated by the model where only slab pull is
512 at work, matching shearing deformation in the mantle as revealed by mantle anisotropy data
513 (Fig. 3a). Geological data show that fast back-arc extension in the area commenced at ~30-35 Ma
514 (Jolivet and Faccenna, 2000; Jolivet et al., 2012) even though slower extension in the Rhodope
515 region could have started earlier (Brun and Sokoutis, 2007). The total amount of extension in the
516 Aegean has been estimated to be of the order of 300-400 km. But paleomagnetic data show that the
517 present day curvature of the Aegean arc occurred mainly after 15 Ma, indicating that extensional
518 rate accelerates with time.
2 0 519 The interaction between mantle upwelling and collisional processes. Efficient subduction, under
520 the influence of vigorous slab pull, does not favour collision when the trench is moving backward
521 (Molnar and Atwater, 1971; Faccenna et al., 2012). The trench northeast of Arabia is, however,
522 moving northward, indenting the overriding plate. This advancing motion of the trench can be
523 favoured by the entrance of continental lithosphere in the subduction zone (Bellahsen et al., 2003),
524 although numerical models suggested that this motion is limited to less than 100 km (Magni et al.,
525 2012). The Bitlis convergence zone advanced toward the upper plate by 400-500 km during the last
526 25 Ma. This cannot be easily reconciled with the entrance of continental material alone, and an
527 extra push force related to mantle drag is thus required. The onset of continental subduction
528 probably occurred at ~30 Ma. Remarkably, only a few hundreds of kilometers south of the collision
529 zone, effusion of alkali basalts was taking place some 10 Ma later while the collision process was
530 ongoing (Fig. 6). The proximity of these two contrasting processes suggests local interactions
531 between mantle upwellings and collision. The orogenic build up as revealed by rapid exhumation
532 and continental stratigraphy (Okay et al., 2010) occurred in the late early Miocene (~18 to 13 Ma).
533 Tomographic images and volcanological evolution reveals that the slab is broken beneath the Bitlis
534 suture zone (Faccenna et al., 2006). It is then possible that slab first acted as a “sail”, advancing
535 northward under the action of the flowing mantle and was then eroded and perhaps segmented,
536 facilitated by a thermal, weakening effect of the plume, which then allowed for the entrance of
537 mantle material within the Mediterranean region (Fig. 6) (Keskin, 2007).
538 Collision, Anatolia motion and Aegean extension. One of the main geodynamic problems of the
539 Middle East is represented by the escape of Anatolia. This motion probably initiated during the Late
540 Miocene, from ~ 10-12 Ma with a progressive localisation of the North Anatolian Fault at 5-6 Ma
541 and onward (Sengör et al., 2005; Faccenna et al., 2006), when both collision in the Bitlis mountains
542 and backarc extension in the Aegean were well under way. The lateral escape of Anatolia probably
543 occurred contemporaneously with the rise up of the plateau, which in Central Anatolia occurred
544 after ~6 Ma (Cosentino et al., 2012), while in eastern Anatolia it could have started earlier after
2 1 545 ~12 Ma (Gelati, 1972). Our results show that more than ~50% of the motion of Anatolia is
546 explained by the upwelling beneath Arabia-Middle East. As shown in the cross sections of Fig. 6,
547 for this flow to be efficient in Anatolia or Aegean, it has to bypass the subduction system plunging
548 northward. The formation of a slab window, probably at ~8-10 Ma and progressively enlarging
549 westward (Faccenna et a., 1996; Cosentino et al., 2012) could have triggered and favoured by the
550 arrival of hot material from Arabia towards the collision zone (Keskin, 2007). The onset of a surge
551 of volcanism over the central Anatolia plateau attests to the arrival of hot material in the system at
552 ~5 Ma. We may further speculate that the hot, northward flowing material could have even
553 triggered and favoured the formation of this slab window, progressively pushing against the slab
554 itself. If this is true, we are facing the evolution of a convection cell that was invaded by the hot-
555 upwelling material because the downwelling branch of the cell below Bitlis vanished due to the
556 closure of the oceanic basin.
557 None of our models succeeds to fully match the complexity of the region as attested by non-
558 isostatic topography, surface velocities, and the complex patterns of seismic anisotropy (Fig. 5), and
559 further adjustments of models may yield locally refined dynamic scenarios. However, our models
560 give a robust, first order view on the dynamics of the system, where the coupling between
561 upwelling mantle in the Afar-Arabia system and the downwelling of the Aegean - Tethyan slab can
562 generate the general patterns of deformation observed at the surface.
563
564 Conclusion
565 We reconstruct the patterns of mantle convection beneath the Afar-Anatolia-Aegean system and
566 compare our present-day results to geodetically measured crustal motions and constraints from
567 seismic anisotropy. Our mantle flow models highlight the importance of coupling the pull exerted
568 by the Hellenic slab with the upwelling below Afar-Arabia; this explains the northward indenter
569 motion of Arabia, the northward progression of volcanism along Arabia, the dynamic uplift of
2 2 570 Anatolia, its westward motion, and the Aegean extension in a single convection system, established
571 episodically over the last ~40 Ma.
572 Jointly interpreting the present-day geodynamic models with geological constraints, we suggest
573 that it may be possible to backtrack the geometry and interactions of the AAA lithosphere-mantle
574 system throughout the Tertiary. The upwelling material associated with the deep mantle Africa-Afar
575 plume flowed northward in the asthenosphere during the closure of the ocean, assisted in
576 continental breakup, and propelled Arabia to induce collision. Several slab segmentation episodes
577 might have opened mantle flow windows, facilitating plume northward transport. Eventually, this
578 lead to asthenospheric flow inside the Aegean region, and accelerating of the Hellenic backarc
579 extension to the rates we see today. Our findings emphasize the role of the deep mantle in affecting
580 microplate motions and controlling the geological expression of upper mantle convection.
581
582 Acknowledgments
583
584 Discussions with Zohar Gvirtzman and Jean Pierre Brun contributed to clarify some of the
585 concepts expressed here. We thank CitcomS authors including S. Zhong, E. Tan, A. McNamara,
586 and L. Moresi for their contributions, CIG (geodynamics.org), E. Kaminski for sharing DREX, and
587 all seismologists who share their tomographic models and shear wave splitting measurements in
588 electronic form. Computations were performed on USC's High Performance Computing Center, and
589 research was partially supported by NSF grants EAR-0809023 and EAR-0643365. M. Keskin
590 thanks to TUBITAK for their support (project no. 108Y222).
591
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944
945
946
3 7 946 References
947 Figure and Table Captions
948
949 Figure 1: Topography and deformation indicators for the Arabia-Anatolia-Aegean system. a)
950 Geodetic velocity field (ArRajehi et al., 2010) in a Eurasia fixed reference frame (orange vectors)
951 compared to geological plate motions from NUVEL1A (de Mets et al., 1994; dark gray vectors),
952 topography from ETOPO1, and seismicity (Engdahl et al., 1998), color coded by depth, z, with
953 major geographic features labelled. b) Residual topography, zres, estimated by correcting smoothed
954 actual topography for crustal isostatic adjustment using the CRUST2.0 model (Bassin et al., 2000)
955 using the same parameters as in Becker and Faccenna (2011). Open triangles are Holocene
956 volcanoes from Siebert and Simpkin (2002), and colored inverted triangles show all volcanic rocks
957 younger than 50 Ma in the EarthChem database compilation available at www.earthchem.org as of
958 05/2012. Plate boundaries (dark green) are modified from NUVEL1 (de Mets et al., 1994) to
959 include an Anatolian plate, as in Becker and Faccenna (2011). c) Dynamic topography estimated by
960 converting free-air gravity anomalies using a 50 mGal/km conversion, assuming a mantle density of
961 3300 kg/m3 (Craig et al., 2011). d) SMEAN composite tomography model (Becker and Boschi,
962 2002) averaged over uppermost mantle depths from 150 to 400 km (see also Figs. 2 and S1). Both
963 c) and d) are superimposed with SKS splitting observations (from the compilation of Wuestefeld et
964 al., 2009, updated 05/2012) where yellow stick orientation shows the “fast axes” of azimuthal
965 anisotropy, c) for all available splits, and d) for station-averaged estimates.
966
967 Figure 2: Comparison of tomographically imaged structure underneath the study region (also see
968 Fig. S1). Maps a-c show 250 km depth layer of models HNB12 (Hansen et al., 2012), SMEAN
969 (Becker and Boschi, 2002), and TX2008 (Simmons et al., 2007), respectively. Smoothed elevation
970 contours are overlaing in 2000 m contours for positive topography, plate boundaries (dark green)
971 from Bird (2002). Dark cyan profiles are plotted in the underneath (d-f), correponding to the model
3 8 972 shown in map above; dashed line indicates 660 km.
973
974 Figure 3: Mantle flow solutions for four, illustrative example models for the AAA system. Left
975 column: surface velocity field (white arrows) compared with geodetic velocity (orange arrows,
976 averaged from the original vectors shown in Fig. 1a) with respect to Eurasia reference frame, with
977 background indicating dynamic topography estimates. Right column: Shear wave splitting
978 measurements indicated by sticks, colored by angular misfit of “fast axes” with model predictions
979 (gray) whose back-azimuthal variations are denoted by wedges and two different stick lengths for
980 fast axes and delay times, respectively (see Becker et al., 2006b, and Miller and Becker, 2012, for
981 details). Inset shows the regional arithmetic averages for the angular misfit, <Δα>, and the delay
982 time misfit, <Δ(δt)>, for all 334 non-zero measurements of Fig. 1d. a) Flow produced by upper
983 mantle density anomaly assigned to the Wadati-Benioff zones only (slab driven flow); b) Flow
984 produced by density anomaly within Wadati-Benioff zones and lower mantle convection produced
985 by density anomaly scaled from the SMEAN model; c) Flow produced by density anomaly within
986 the Wadati-Benioff zones, lower mantle anomalies, and convection produced by slow velocity
987 anomaly only (from SMEAN) in the upper mantle. d) As in b) but with a low viscosity channel
988 extending northward from Afar to Israel, sensu Hansen et al. (2006).
989
990 Figure 4: Mantle flow predictions at three different depths: a) 210, b) 345, and, c), 570 km for the
991 model of Fig. 3c, with vectors and background shading denoting horizontal and radial velocities,
992 respectively, and velocities within three cross sections (d-f) as marked on a), with background
993 showing non-dimensional temperature.
994
995 Figure 5: Misfit matrices for a subset of all global flow models considered (cf. Miller and Becker,
996 2012), illustrating the effect of different choices of density (columns) and lateral viscosity variation
997 (rows) models for, a), global plate motion correlations, b), regional plate motion correlations, and,
3 9 998 c), average angular misfits of shear wave splitting predictions . White boxes show the models
999 illustrated in Fig. 3. For legend, see Table 1.
1000
1001 Figure 6: Tectonic evolution of the AAA system at four stages at 30-33 (a), 20-23 (b), 10-12 (c),
1002 and 3-5 Ma (d) and relative, idealized cross sections (e-h), cf. Fig. 4 for the present-day. Plate
1003 motions are from GPlates (Muller et al., 2008; Gurnis et al., 2012). Geological features from Barrier
1004 and Vrielynck (2010). Red indicates active volcanic fields (undistinguished); blue indicates deep
1005 basins; red lines indicate active subduction zone; black lines indicate collisional thrust system;
1006 yellowish background area indicate regions under influence of Afar plume; arrows indicate
1007 direction of mantle flow.
1008
1009 Figure 7: Geochemical isotope Diagram for 206Pb/204Pb vs. 207Pb/204Pb for the lavas of the
1010 Karacadağ neovolcanic area (data from Keskin et al., 2012b). Data source for the fields drawn in
1011 this diagram are listed in Ma et al. (2011), except for two basaltic samples from the Erzurum-Kars
1012 Plateau (i.e. EKP in the legend) which are from Keskin et al. (1998).
1013
1014 Figure 8: Conceptual cartoon of mantle flow in the AAA system as inferred from our analysis of
1015 geological and geophysical models.
1016
1017 Table 1: Legend for density and viscosity models considered in the misfit matrix plots of Fig. 5.
1018
4 0 1018 Table 1
1019
1020 Mantle density models.
1021 Columns in Fig. 5 denote density models tested (details of model parameters are described in
1022 Becker and Faccenna, 2011, and Miller and Becker, 2012):
MITP08 Li et al. (2008) (MIT gorup) global P wave model, scaled to temperature in whole mantle HNB12 Hansen et al. (2012) P wave model in whole mantle, building on MITP08, highlighting effect of regional resolution improvement. LH08 Lebedev and van der Hilst (2008) SV wave model in upper mantle, SMEAN (Becker and Boschi, 2002) in lower mantle SMEAN Composite S wave model with anomalies underneath continental keels down to 300 km removed RUM Density anomalies in upper mantle only, as inferred from Wadati-Benioff zone geometry as assigned from the RUM model (Gudmundsson and Sambridge, 1998). RUM+SMEAN_lm RUM plus density anomalies in the lower mantle inferred from SMEAN tomography RUM+SMEAN_neg Like RUM+SMEAN_lm, but additionally slow (“hot”) velocity anomalies in the upper mantle from SMEAN, as in Ghosh et al. (2010) RUM+SMEAN_neg+SMEAN_100 Like RUM+SMEAN_neg, but additionally all velocity anomalies in the uppermost 100 km (“lithospheric anomalies”) from SMEAN TX2008 Global mantle temperature anomalies inferred from the TX2008 S wave tomography model of Simmons et al. (2007) TX2008+RUM RUM upper mantle slabs as in RUM, lower mantle density anomalies from TX2008. 1023 1024 Models of lateral variations in viscosity as shown in different rows of the misfit matrices of Fig. 5 reference Only radial viscosity variations besides weak zone geometry prescribed in the lithosphere with modified NUVEL1 geometry as in Becker and Faccenna (2011), as in Fig. 4. Bird_boundaries Same width and viscosity reduction for weak zones as in reference, but geometry taken from Bird (2003), as in Fig. 2. Bird_wo_Arabia Same as Bird_boundaries, but no weak zone between NW Arabia and Eurasia keels Same as reference, but stiff (factor 500 viscosity increase) continental keels down to 300 km depth as inferred from the
4 1 3SMAC model (Nataf and Ricard, 1996) asthenosphere Sub-oceanic plate reduction of viscosity by a factor of 0.01 in the depths between 100 and 300 km keels+asthenosphere Combined effect of keels and asthenosphere. simple_channel Reference model with factor 0.01 viscosity reduction along a channel from Afar to Isreal between 100 and 300 km depth. channel-LVV Reference model with lateral viscosity variations inferred from scaling the depth-averaged velocity anomalies in LH08 between 100 and 300 km depth to ~four orders of magnitude lateral variations of viscosity (cf. Fig. 1d), yielding a low viscosity channel similar to simple_channel. channel-LVV(z) As channel-LVV but allowing for depth-variations in lateral viscosity variations from 100 to 300 km depth. 1025 1026
4 2 *Highlights (for review)
-The Middle East region represents a key site within the Tethyan domain where continental
break-up, collision, backarc extension and escape tectonics are linked together.
-We perform global mantle circulation computations to test the role of slab pull and mantle
upwellings as driving forces for the kinematics of the Arabia-Anatolia-Aegean (AAA) system and
we teste our result with geodesy, residual topography, and shear wave splitting observations.
- The best match to these crustal motions can be obtained by combining the effect of slab pull
exerted in the Aegean with a mantle upwelling underneath Afar and, more generally, with the large-
scale flow associated with a whole mantle;
-We suggest that the tectonic evolution is related to the progressive northward entrainment of
upwelling mantle material from Afar associated with the downwelling Tethyan slab;
-Our model reconciles Afar plume volcanism, the collision in the Bitlis mountains and northern
Zagros, and the rapid increase of Aegean trench rollback in a single coherent frame of large scale
mantle convection, initiated during the last ~40 Ma.
20˚ 30˚ 40˚ 50˚ 60˚ 20˚ 30˚ 40˚ 50˚ 60˚ Figure1z [km] zres [km] Click here to downloadAnatolia Figure: figure1.pdf 2 2 40˚ 40˚ 1 0 Aegea Bitlis 0 -1 -2 -2
30˚ 30˚ Arabia
d [km] age [Ma] 200 50
40
20˚ 30 20˚ 100 Africa 20 10
0 0 Gulf of Aden Afar 10˚ 10˚ Ethiopian a) 4 cm/yr Rift b) 20˚ 30˚ 40˚ 50˚ 60˚ 20˚ 30˚ 40˚ 50˚ 60˚ zeq [km] δv [%] 2 2
40˚ 40˚ 1 1
0 0
-1 -1
-2 -2
30˚ 30˚
20˚ 20˚
10˚ 10˚ δt = 2 s δ c) SKS d) tSKS = 2 s
Figure 1 a)Figure220˚ 30˚ A 40˚ 50˚ 60˚ b) 20˚ 30˚ A 40˚ 50˚ 60˚ c) 20˚ 30˚ A 40˚ 50˚ 60˚ hnb12Click here to download Figure: figure2.pdf40˚ smean 40˚ tx2008 40˚
δvP [%] δvS [%] δvS [%] 2
2 2 1 30˚ 30˚ 30˚
0 0 0
-1 -2 -2
-2
20˚ 20˚ 20˚
10˚ 10˚ 10˚
B B B d) e) f) x [km] x [km] x [km] 0 0 0 -1000 1000 -1000 1000 -1000 1000 A 0 2000 B A 2000 B A 2000 B -200 -2000 -2000
] 0 6000 600 6000
r [km] r [km r [km]
0 0 550 550 5500
δvS δvS δvS -1 0 1 -2 -1 0 1 2 -2 -1 0 1 2
Figure 2 Figure3 Click here to download Figure: figure3.pdf
20˚ 30˚ 40˚ 50˚ 60˚ δzdyn [km] 2
40˚ 1 2 cm/yr
0
-1
-2
30˚
20˚
a) 10˚ 20˚ 30˚ 40˚ 50˚ 60˚ δzdyn [km] 2
40˚ 1 2 cm/yr
0
-1
-2
30˚
20˚
b) 10˚
20˚ 30˚ 40˚ 50˚ 60˚ δzdyn [km] 2
40˚ 1 2 cm/yr
0
-1
-2
30˚
20˚
c) 10˚ 20˚ 30˚ 40˚ 50˚ 60˚ δzdyn [km] 2
40˚ 1 2 cm/yr
0
-1
-2
30˚
20˚
d) 10˚
Figure 3 Figure4 Click here to download Figure: figure4.pdf 20˚ 30˚ 40˚ 50˚ 60˚ 20˚ 30˚ 40˚ 50˚ 60˚ vr [cm/yr] v [cm/yr] B r 4 4 40˚ 40˚ 2 cm/yr 2 cm/yr 0 z = 210 km z = 345 km F E 0
-4 D -4
30˚ 30˚
20˚ 20˚
smean_al+RUM smean_al+RUM ref ref C 10˚ 10˚ a) A b)
20˚ 30˚ 40˚ 50˚ 60˚ vr [cm/yr]
4 40˚ 2 cm/yr distance along track [km] 0 z = 570 km F E -1000 -500 0 500 1000
-4 0
30˚ -200
-400
-600
20˚ -800 depth [km]
smean_al+RUM -1000
ref -1200 10˚ T c) d) 0.3 0.4 0.5 0.6 0.7 B distance along track [km] A -2000 -1500 -1000 -500 0 500 1000 1500 2000 0
-200
-400
-600 depth [km] -800
-1000 e) -1200 T 0.3 0.4 0.5 0.6 0.7
C distance along track [km] D -1500 -1000 -500 0 500 1000 1500 0
-200
-400
-600 depth [km] -800
-1000 f) -1200 T 0.3 0.4 0.5 0.6 0.7 Figure 4 Figure5 Click here to download Figure: figure5.pdf
MITP08HNB12 LH08 SMEANRUM RUM+SMEAN_lmRUM+SMEAN_negRUM+SMEAN_neg+SMEAN_100TX2008TX2008+RUM a) 0.2 0.3 0.7 0.8 0.7 0.7 0.8 0.8 0.7 0.7 reference 0.6
Bird_boundaries 0.6 0.9 0.8 Bird_w/o_Arabia 0.6 0.7 keels 0.6 0.6 vel
0.5 r asthenosphere 0.8 0.4 keels+asthenosphere 0.8 0.3 0.2 simple_channel 0.8
channel-LVV 0.6
channel-LVV(z) 0.6
0.4 0.7 0.7 0.7 0.7 0.6 0.7 0.7 0.6 0.6 b) reference 0.6
Bird_boundaries 0.6 0.7 Bird_w/o_Arabia 0.6
0.6 keels 0.7
0.5 vel reg
asthenosphere 0.6 r
0.4 keels+asthenosphere 0.6
0.3 simple_channel 0.6
channel-LVV 0.6
channel-LVV(z) 0.6
54.8 44.7 47.6 34.7 40.2 34.2 36.4 39.4 44.5 43.5 c) reference 40.1
Bird_boundaries 46.2 55 Bird_w/o_Arabia 45.3 50 ] keels 41.0 o 45 [
asthenosphere 41.5 40 ∆α keels+asthenosphere 41.7 35 30 simple_channel 39.2
channel-LVV 42.0
channel-LVV(z) 41.0
LH08 RUM MITP08 HNB12 SMEAN TX2008
TX2008+RUM RUM+SMEAN_lm RUM+SMEAN_neg
RUM+SMEAN_neg+SMEAN_100
Figure 5 Figure6 Click here to download Figure: figure6.pdf about 30 Ma Cyprus A a) b) subduction Arabia Red Sea Nubia MER B 0 0
-200 -200
-400 -400
-600 -600
-800 -800
depth [km] -1000 -1000
-1200 -1200 -2000 -1500 -1000 -500 0 500 1000 1500 2000 distance along track [km] e) upper about 20 Ma A Cyprus B mantle subduction Arabia Red Sea Nubia MER ow 0 0 -200 -200
-400 -400
-600 -600
-800 -800
depth [km] -1000 -1000 depth [km]
-1200 -1200 -2000 -1500 -1000 -500 0 500 1000 1500 2000 distance along track [km] 30 - 33 Ma 20 - 23 Ma about 10 Ma Collision - slab break o f) Dead Sea Cyprus Transform A NAF B subduction Arabia Red Sea Nubia MER 0 0 c) d) -200 -200 -400 -400
-600 -600
-800 -800
depth [km] -1000 -1000 depth [km]
-1200 -1200 -2000 -1500 -1000 -500 0 500 1000 1500 2000 distance along track [km] g) present day Dead Sea Cyprus Transform A NAF B Arabia subduction Arabia Red Sea Nubia MER Nubia Arabia 0 0 Nubia -200 -200
-400 -400
-600 -600
-800 -800
depth [km] -1000 -1000 depth [km]
-1200 -1200 -2000 -1500 -1000 -500 0 500 1000 1500 2000 T Somalia 0.3 0.4 0.5 0.6 0.7 10 Ma 3-5 Ma distance along track [km] h)
Figure 6 Figure7 Click here to download Figure: figure7.pdf
15,8 Crust or metasomatised ? mantle Saudi Arabian intraplate 15,7 volcanism NHRL 207Pb/204Pb 15,6
S Syria, Jordan & Israel Red Sea MORB 15,5 Afar plume
206Pb/204Pb 15,4 18,0 18,3 18,6 18,9 19,2 19,5
Figure8 Click here to download Figure: figure8.pdf
10 Ma
20 Ma
35 Ma
Figure 8