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Slab pull and indentation tectonics: insights from 3D laboratory experiments
Physics of the Earth and Planetary Interiors 149 (2005) 99–113
Slab pull and indentation tectonics: insights from 3D laboratory experiments Vincent Regardc,∗ , Claudio Faccennab,1 , Joseph Martinodc,2 , Olivier Belliera,3 a
CEREGE, UMR CNRS 6635, Universit´e Aix-Marseille III, Europˆole de l’Arbois, 13545 Aix-en-Provence Cedes 4, France b Dipartimento di Scienze Geologiche, Universit` a Roma tre, Largo S. L. Murialdo 1, 00146 Rome, Italy c LMTG, UMR CNRS/IRD 5563, Observatoire Midi-Pyr´ en´ee´ s, 14 av. E. Berlin, 31400 Toulouse, France Received 28 July 2003; received in revised form 15 December 2003; accepted 26 August 2004
Abstract We investigate, using 3D laboratory experiments, how the dynamics of indentation process are affected by the evolution at depth of the oceanic and continental subductions. Lithospheric plates are modelled by sand–silicone plates floating on glucose syrup, and the density contrast between oceanic and continental lithospheric plates and asthenosphere is reproduced. Analogue experiments model the convergence between two lithospheric plates, a small continent indenting a large continental plate. We show that the surface deformation in front of the indenter and above the oceanic subduction zone depends on the behaviour of the slab below the collision zone. Slab break-off following the subduction of the small continent favours the indentation process, because it results in an increasing compression in front of the indenter, and extension above the neighbouring oceanic subduction, both of them being responsible for the appearance of the indenter-like geometry of the plate boundary. When the slab does not deform significantly at depth, in contrast, the closure of the oceanic domain in front of the indenter is followed by a longer period of continental subduction, during which the tectonic regime within the wide continent remains quite homogeneous. Comparing the presented analogue experiments with the subductions both part of the Arabian indenter within Eurasia, our results suggest that the different tectonic regime on both sides of the Arabia indenter may partly result from the probable occurrence of a detachment at depth under eastern Anatolia. © 2004 Elsevier B.V. All rights reserved. Keywords: Indentation; Collision; Subduction; Analogue modelling; Slab; Break-off
1. Introduction ∗
Corresponding author. Tel.: +33 5 61332656; fax: +33 5 61332560. E-mail addresses: [email protected] (V. Regard), [email protected] (C. Faccenna), [email protected] (J. Martinod), [email protected] (O. Bellier). 1 Tel.: +39 06 54888029; fax: +39 06 54888201. 2 Tel.: +33 5 61332666; fax: +33 5 61558250. 3 Tel.: +33 4 42971660; fax: +33 4 42971559. 0031-9201/$ – see front matter © 2004 Elsevier B.V. All rights reserved. doi:10.1016/j.pepi.2004.08.011
The term “indentation” has been coined to define the collision between India and Asia, and is classically intended to define the ploughing up of a small continental plate inside a larger one. By definition, indentation is characterised by collision occurring in front of the “indenter”, which is generally flanked on the sides
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by subduction zones. The tectonic process driving and accompanying indentation has been investigated especially for those aspects concerning the way the upper plate deforms. In particular, it has been pointed out that a major role is played by the strength of the upper plate, providing the possibilities to squeeze and thicken the crust or producing the lateral escape of crustal block along strike-slip faults (e.g., Molnar and Tapponnier, 1975; Peltzer et al., 1984; Vialon et al., 1984; Shen et al., 2001). The indentation mechanisms are frequently evoked to explain how crustal or lithospheric deformation in collision zones is associated with vertical or horizontal extrusion processes (Argand, 1924; Molnar and Tapponnier, 1975; Johnson, 2002). Presently, many studies argue in favour of combined or successive vertical/horizontal extrusion to explain deformation related to collision (Johnson, 2002). Consequently, to understand collision processes it appears necessary to look at the mechanisms of indentation in the 3D. An interesting example of indenter is Arabia, since the indentation of Arabian within Eurasia is differentially expressed to the East (Iran) than to the West (Aegean–Anatolian domain). We present 3D laboratory experiments performed to define the role of subduction in the tectonics of indentation. These experiments are part of an experimental program devoted to the investigation of the physical parameters that drive collision processes. Previous study from Regard et al. (2003) described 2D illustrating the way continental lithosphere subducts prior to collision, following an oceanic subduction. This set of experiments shows that the possibility for the continent to subduct is related to the pull exerted by the previously subducted oceanic lithosphere, and therefore to the way the slab deforms at depth. We express our results by means of two dimensionless numbers (F) where the slab pull force is normalised to the strength of the slab: favourable conditions for continental subduction correspond to low values of F, or, in other words, to situations in which the oceanic slab is more prone to maintain its integrity and to exert an higher pull on the subducting continent. Here, we extend this analysis to a more complex 3D configuration with a lateral transition from oceanic to continental subduction. In particular, we show two representative experiments out of 8, where an oceanic lithosphere neighbours a continental one. Our results
highlight that the subduction history prior to the collision and indentation exerts a primary role in the style of deformation of the upper plate. Our work has been inspired by the tectonic regime that governs the indentation of Arabia within Eurasia; a comparison with this natural example will be discussed.
2. Laboratory experiments Many 3D laboratory experiments have been constructed to analyse the tectonic processes that accompany indentation. Vertically constrained paraffin wax experiments first showed that lateral extrusion might be a fundamental process accommodating indentation (Peltzer et al., 1984). Sand–silicone–honey experiments also addressed the same issue but allowed crustal thickening, hence obtaining a more realistic partitioning between extrusion and thickening (cf. Davy and Cobbold, 1988; Martinod et al., 2000). Those experiments were all performed considering the lateral boundary of the upper plate above the subduction zone as a free boundary where the continental lithosphere can eventually spread. In this view the role of subduction process is subordinated to indentation or is simulated by the spreading of the continental lithosphere. Faccenna et al. (1996), using a similar technique, shows that the upper plate deformation depends on the way the subduction process operates on the sides of the upper plate. It is possible to obtain the whole range of deformation spanning from rigid extrusion to extension parallel to indentation by increasing the efficiency of the subducted lithosphere, i.e., of the slab pull. The problem of what does happen at depth in the collisional zone and how deep structures transfer laterally on its sides has been totally unexplored. The closure of an oceanic basin and the beginning of the collisional process itself has also been reproduced in wax experiments (Chemenda et al., 1995, 1996, 2000), or using sand–silicone–honey experiments (Regard et al., 2003). Experimental results agree in considering the possibilities that slab deforms at depth after the entrance of buoyant continental material. In particular, Regard et al. (2003) show that the slab can loose its integrity deforming at depth by ductile delamination (Rayleigh–Taylor instability) or by brittle rupture (break-off), depending on the initial imposed boundary condition and on the rheology of the slab. In
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turn, the onset of collision and the amount of continental subduction depend on the level of slab pull exerted by the previously subducted oceanic lithosphere. The current work explores the effects of a lateral transition between subduction and collision, emphasising the role of the subducting plate deformation. We show that the way the subducting plate indents the upper plate depends upon the deformation of the slab at depth. 2.1. Materials and experimental technique Following an experimental design adopted by previous authors (e.g., Davy and Cobbold, 1988, 1991), we simulate the stratified lithospheric rheological profile (e.g., Ranalli and Murphy, 1987) constructing a brittle–ductile layered model (Fig. 1). We use dry sand to model the brittle behaviour of the upper lithosphere and silicone putty to model the ductile behaviour of
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the lower lithosphere. Dry sand is nearly cohesionless, and its internal friction angle is close to 30◦ (Mandl et al., 1977). Silicone putty is a quasi-Newtonian fluid (Weijermars and Schmeling, 1986), and its viscosity is measured in laboratory before each experiment at room temperature (25 ◦ C). The density of silicone can be varied mixing it with iron fillers, which preserves its Newtonian rheology. The density of the brittle layer can be diminished mixing the sand with lighter granular materials. The sand–silicone layering rests on glucose syrup, which simulates the asthenosphere. We use two different kinds of sand–silicone layers, either lighter or denser than the glucose syrup, to represent continental or oceanic lithospheres, respectively. Two plates are constructed. One is constituted by light silicone and sand simulating a large continental plate, whereas the other is constituted by a heavy silicone–sand layer simulating a dense oceanic plate with a small continent on one side (indenter) (Fig. 1).
Fig. 1. Experimental set up: (a) upper view and (b) cross-section of the western part of the model. Sand–silicone plates modelling the lithosphere are lying above honey that represents the asthenosphere. These plates are lighter or denser than honey, to represent continental or oceanic lithospheres, respectively. The piston is pushing northward at a constant velocity of 4.4 mm/h scaled to represent 2.3 cm/year in nature. The oceanic plate will subduct below the northern continent, the subduction place being controlled by the removal of a 1-cm wide band of sand that weakens the northern margin of the oceanic plate.
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Table 1 Physical parameters and some results of experiments Experiment number
1
2
3
4
5
6
7
8
Hsil Hosand Hcsand Hm
9 4 2/4 109
9 4 2/2 195
9 3 3/3 186
9 3 3/3 186
9 3 3/3 173
9 4 2/2 103
9 4 2/4 189
9 4 1.5/4 105
Length (10−3 m) Southern continent Ocean (western part) Northern continent
Lcs Lo Lcn
190 30 240
150 70 280
220 70 210
205 70 225
210 70 220
200 70 230
190 70 240
210 70 220
Viscosity (Pa s) Oceanic silicone Continental silicone Honey
ηo ηc ηm
3.1 × 105 4.6 × 105 4.6 × 102
2.1 × 105 1.3 × 105 4.6 × 102
1.4 × 105 1.0 × 105 4.6 × 102
1.4 × 105 1.0 × 105 4.6 × 102
1.4 × 105 1.6 × 105 4.6 × 102
3.1 × 105 2.0 × 105 4.6 × 102
3.1 × 105 3.0 × 105 4.6 × 102
3.1 × 105 4.0 × 105 4.6 × 102
Density (kg m−3 ) Oceanic silicone Oceanic sand Continental silicone Continental sand Honey
ρosil ρosand ρcsil ρcsand ρm
1484 1500 1350 1300 1422
1470 1500 1310 1250 1425
1465 1500 1325 1300 1425
1465 1500 1325 1300 1425
1465 1500 1380 1300 1425
1484 1500 1355 1300 1422
1484 1500 1355 1300 1422
1484 1500 1390 1300 1422
u
1.2 × 10−6 1.2 × 10−6 1.2 × 10−6 1.8 × 10−6 1.2 × 10−6 1.2 × 10−6 1.2 × 10−6 1.2 × 10−6 63 42 72 49 72 73 70 68
Thickness (10−3 m) Silicon layers Oceanic sand layer Continental sand layers (N/S) Honey
Time Piston velocity (m/s) Total duration (h) Dimensionless numbers F1 F2
0.5 1.5
1.7 3.0
Continental and oceanic plates differ in thickness, density and viscosity of the upper silicone layer, and in thickness and density of the sand layer (Table 1). The oceanic plate strength (mainly controlled by brittle sand layer thickness) is usually larger or similar to that of continent. Sand is sieved over the entire model, with the exception of a 1-cm wide strip located at one boundary between the two plates (no sand above silicone, Fig. 1). The shortening imposed in the experiment will thus preferentially localise on this side of the oceanic plate, and the thrust fault accommodating subduction will localize there. The lower boundary of the box approximates a high gradient viscosity transition within the deep mantle (Funiciello et al., 2003). The box depth varies from 11 to 19 cm. Models are constructed inside a rectangular Plexiglas tank (50 cm long and 30 cm wide; Fig. 1). The lateral walls of the tank are lubricated with wax and the sand layer is removed near the boundaries of the
1.0 3.7
1.0 2.5
1.0 3.5
1.0 1.5
1.1 2.8
1.3 1.6
box to reduce friction at plate–walls interfaces. In some experiments (cf. experiment 2) such a lubrication is stopped 5 cm before the back wall in order to avoid the development of structures against the back wall; this difference does not modify significantly the stress and strain regime in the centre of the model, and it does not influence the conclusions presented in this paper. Horizontal shortening is imposed displacing a rigid piston at constant velocity perpendicular to the plate boundary. The piston is confined in the upper part of the tank and the glucose syrup is free to move underneath. The motion of the piston mimics the effect of an external push related to plate motion. Constant velocity is imposed, and we have no precise idea of the magnitude of the forces acting on the system. For convenience, we shall refer thereafter to regions of the experiment in terms of geographical directions, the southern boundary corresponding to the piston. A squared grid of passive sand markers enables the visu-
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alisation of the surface deformation. Photographs are taken at regular time intervals in three orientations, top, oblique, and lateral view through the transparent tank walls. The materials selected to model the stratified oceanic and continental lithosphere have been chosen to fulfil stress similarity criteria scaled for gravity (see Davy and Cobbold, 1991). Material parameters for each experiment are listed in Table 1. The scaling of our experiments is detailed in Regard et al. (2003) (see Table 2 therein); 1 cm in these experiments represents 70 km in nature and 1 h represents about 1 My in nature. The convergence velocity is scaled to represent the Arabia–Eurasia convergence velocity, the 4.4 mm/h of our model corresponding to about 2.3 cm/year in nature. Previous experiments (Regard et al., 2003) show that the shortening of the continental plate (that can be both the indenter or the indented plate) only begins after a certain amount of continental material has subducted into the mantle. This amount of subducted continental material, in turn, depends on the rheology of the slab, on the depth of the low-viscosity mantle, and on the amount of oceanic lithosphere subducted before oceanic closure. In Regard et al. (2003), these results have been expressed using two dimensionless numbers, F1 and F2 . F1 characterises the possibility for the slab to break under its own weight: F1 = FT /Frupture
(1)
where FT is the maximum tension force within the slab and Frupture , the force necessary to achieve the rupture of the slab: FT = ρo−m g Z H
(2)
Frupture = α ρm g H Hsand + ηo u Hsil /H
(3)
where ρo−m is the density contrast between the oceanic lithosphere and asthenosphere, ρm is the asthenosphere density, g is the gravity, Z is the depth of the high-gradient viscosity increase within the deep mantle (modelled by the bottom of the Plexiglas tank), H is the thickness of the slab, Hsand is the sand thickness, Hsil is the silicone thickness, ηo is the viscosity of the oceanic plate silicone, u is the convergence velocity, and α is a parameter related to the friction coefficient µ of the sand (µ = tan 30◦ ; α = 2µ/(µ + (1 + µ2 )1/2 )).
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F2 characterises the possibility for a viscous instability affecting the oceanic slab to develop before the slab reaches the bottom of the convective layer: F2 =
t1 Hsil = Zρo−m g t2 u ηo
(4)
where ρo−m is the density contrast between the oceanic lithosphere and the asthenosphere. F2 is the ratio between two characteristic times t1 and t2 , t1 = Z/u is the characteristic time for the tip of the slab to reach the bottom of the low-viscosity mantle (the slab vertical velocity is assumed to be equal to the convergence velocity, u), t2 = ηo /ρo−m × g × Hsil is the time necessary to achieve the formation of a viscous instability. In particular, we observe that slabs deform for high F2 (>3), and the possibilities to have break off is limited to the case where also F1 is high (>1.2). In contrast, low F2 values correspond to models in which the slab does not deform significantly. In 2D models the possible deformation of the slab at depth governs the delay between the oceanic closure and the beginning of collision. This delay (and consequently the amount of continental subduction) is long if the slab does not deform and decreases if it can deform. In case of break-off it can be significantly low. We focus here on the influence of F1 and F2 on the behaviour of the 3D experiments. We show that the process of indentation is sharply influenced by variations of these parameters.
3. Experimental results Experiments have been designed to test the influence of the slab pull force resulting from the subduction of the oceanic lithosphere. This parameter has been investigated varying the length of the oceanic plate prior to closure (Table 1). In the following, we describe the evolution of two selected experiments (see Table 1 for a complete description of the physical parameters of the experiments). 3.1. Experiment 1 In experiment 1 (Fig. 2), a narrow 3-cm large ocean (equivalent to roughly 200 km in nature) initially separates the indenter from the main continent. F1 is about
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Fig. 2. Top views (a), line drawings (b), shortening vs. time (c) of experiment 1. We report in (c) the total shortening occurring within the plates, excluding the shortening accommodated at the subduction zone. Two curves are reported for the eastern domain (continuous line) and for the western domain (dashed line), at 8 cm, respectively, of the eastern and western side of the box.
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0.5, and F2 is about 1.5. Under this condition the subducting plate does not deform much at depth. During the first 12 h of experiment, E–W folds, i.e., compressional structures, localise at the weak oceanic plate margin and, subordinately, near the boundaries of the box. The total shortening rate accommodated outside the young oceanic subduction zone is evaluated to around 1 mm/h (Fig. 2c). This compressional deformation characterises the initiation of subduction (e.g. Faccenna et al., 1999; Regard et al., 2003). After 12 h, the oceanic basin closure is completed, causing an increase of the shortening rate to 2 mm/h, demonstrating the low efficiency of the subducting system during the first phase of subduction in both the eastern and western side of the experiment. Compressive deformation concentrates on a linear E-trending fold belt that develops in the middle of the plate in correspondence of the end of the model/box boundary sand-less strip (Fig. 2). Only after 29 h of deformation, the shortening rate in front of the indenter becomes larger than that accommodated in the area where oceanic subduction takes place. In the eastern part of the experiment, in correspondence of the oceanic subduction, the shortening rate diminishes on average to 1.4 mm/h with a tendency to further decrease after 51 h. The western domain is more shortened, and the difference in shortening rate is accommodated by a N-trending strike-slip zone that links the trench with the main intracontinental zone of deformation, located far from the trench at the centre of the overriding plate. This strike-slip zone accommodates about 20 mm between 30 and 50 h, i.e. 1 mm/h on average. At 51 h, the eastern sector undergoes extensional deformation and the former folds are unfolded and inverted into E-trending extensional structures (normal faults and grabens). From this moment, the upper plate rotates clockwise to accommodate the different tectonic regime. 3.2. Experiment 2 Experiment 2 differs from experiment 1 as the oceanic domain, i.e. the distance separating the indenter and the target continents, is wider (7 cm). In this manner we obtain a longer oceanic slab that exerts a larger slab pull force. To obtain a maximum level of slab pull we also increase the depth of the box. Hence, F1 and F2 both increase to 1.7 and 3.0, respectively.
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Under this condition, the experimental slab must be prone to deform at depth. The initial stage of the experiment is similar to what observed in previous experiments, with some E–W compressional folds characterising the subduction initiation occurring in the experiment outside the trench (Faccenna et al., 1999; Regard et al., 2003). This phase lasts for around 10 h producing a shortening of about 10–15 mm (Fig. 3). Afterwards, the subduction zone accommodates efficiently the imposed convergence and the oceanic domain located in front of the indenter is entirely consumed after 18 h of deformation. During the final stage of oceanic closure, extension appears within the upper plate: E-trending normal faults are observed within the overriding plate with an average N–S extensional rate comprised between 1 and 1.6 mm/h (Fig. 4). The beginning of continental subduction, as observed in 2D experiments (Regard et al., 2003), does not change significantly the subduction dynamics. Extension lasts up to 26 h of deformation (Fig. 3). The extensional phase of deformation is related to the slab growth at depth. The slab pull resulting from the previous entrance of the dense oceanic plate counterbalances both the effect of the positive buoyancy of the continental part of the slab, and the effect of the piston-imposed shortening. After about 30 h of model run, under the collisional zone (western part of the experiment), we observe that the slab breaks and that the subducting oceanic lithosphere detaches from the continental part of the slab (Fig. 3). Afterwards, the detachment progresses eastward under the boundary between the two continents. It is difficult to establish the rate of lateral propagation of the detachment. However, it seems to have occurred quite rapidly. As a matter of fact, we note that in surface, on the western side of the experiment, E-striking folds and thrust-faults develop just after the slab detachment reactivating and inverting normal faults related to the previous stage (Figs. 3 and 4). Surface deformation is then strongly influenced by the detachment process and it indirectly evidences the eastward progression of the detachment (Fig. 5): the shortening domain observed during the period 33–36 h is offset of 3.3 cm with respect to that of period 30–33 h, suggesting that the detachment has progressed at a rate of about 1.1 cm/h (equivalent to roughly 8 cm/year in nature).
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Fig. 3. Top views (a) and the corresponding line drawings (c), lateral view taken from the western side of the experiment at stage D, 33 h (b), and Shortening vs. time (d) of experiment 2. The two curves reported in (d) correspond to the shortening occurring within the plates in the eastern domain (continuous line) and in the western domain (dashed line), at 8 cm, respectively, of the eastern and western side of the box.
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Fig. 4. Deformation maps of the upper continental plate in experiment 2. Graduations are in centimetres. We report the velocity fields (arrows) for the periods 20–28 h (a) and 28–32 h (b), and the rate of surface change for the same periods. Blue surfaces correspond to increasing surfaces (dilatation), and red ones to decreasing surfaces (contraction).
After this break-off, the shortening in front of the indenter reaches a rate of 2.8 mm/h, accommodating 65% of the total amount of convergence (Fig. 4). On the eastern side, conversely, the speed of the oceanic subduction increases with a maximum velocity of about 10 mm/h, twice the convergence velocity. It results in a rapid back-arc extension (Fig. 6). We observe that till the end of the experiment, the break-off of the slab remains confined to the continental collision zone, in front of the indenter. This deep behaviour increases dramatically the pull exerted by the rest of the oceanic slab on the active oceanic subduction, explaining the large amount of back-arc extension in the western part of the experiment. The contrast between shortening in front of the indenter, and extension above the oceanic subduction zone, is responsible for the appearance of the indenter-like geometry of the plate boundary.
4. Interpretation of the experimental results and application to natural examples Our results show that the way the upper plate deforms during indentation depends much on the characteristics of the subduction system and on the history of subduction. Comparison between experiments 1 and 2 highlights the dependence of the pattern of collision and deformation in 3D on the way the subduction process develops. We check that the behaviour of the slab in 3D experiments fits predictions based on the F1 − F2 diagram constructed using the 2D experiments (Fig. 6). Viscous blubs only appear when F2 is larger than about 2.5, and both F1 and F2 must be large (larger than 1.2 and 2.5, respectively) to have break-off. Experiment 1 demonstrates that the plate boundary is likely to absorb several hundreds of kilometres of continent–continent convergence before an indenter-
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Fig. 5. (a) Total N–S shortening between the marker line close to the trench and the boundary of the box, for experiment 2. This value is different from the shortening shown in Figs. 2 and 3 in which we reported the sum of all the shortening that occurred once within the plate, including within the already subducted parts of the plate. (b) Shortening rate vs. E–W position in experiment 2. In the western part of the experiment, the shortening occurs as a consequence of the continental collision. In the eastern part of the experiment, extension occurs above the oceanic subduction zone. Before break-off, shortening varies smoothly from east to west. After break-off, all the western part of the subducted slab is pulling the eastern subduction zone, and provokes the sharp variation of the shortening rate. Arrows highlight the 3.3 cm offset of the shortening domain between periods 30–33 h and 33–36 h. This offset results from the lateral migration of the slab break-off.
like geometry appears. In fact, previously published 2D experiments already evidenced that the oceanic subduction is followed by an episode of continental subduction, that may absorb large amounts of continental plate, particularly when the slab does not break or stretch at depth (Chemenda et al., 1996; Regard et al., 2003). In the 3D configuration, the subduction of the indenting continent is favoured both by the pull of the previously subducted oceanic plate, and by the pull resulting from the lateral active oceanic subduction.
Only during the last phases of this experiment, the different efficiency of the subduction system is accommodated in the surface with internal deformation or via the rigid rotation of the upper continental block. If we allow for a larger subduction system (larger initial ocean width, experiment 2), the slab pull level on the subducting slab can increase up to the point that slab breaks (Regard et al., 2003). The lateral propagation of this process (several cm/year) results in a tremendous concentration of the slab pull effect at the tip of the propagating tear (Yoshioka and Wortel, 1995). This, in turn, may induce a large-scale back-arc extension above the oceanic subducting plate (Wortel and Spakman, 1992; Buiter et al., 2002). This process favours the appearance of an indentation-like geometry between the two lithospheric plates. Then, the entrance of the indenter within the large continent results on the one hand from the shortening accommodated in front of the indenter, but also from the southern motion of the large continent on the sides of the indenter. Most of the experimental sets previously built to study the indentation tectonics focus on the rheology of the upper plate, without taking into account the rheology of the colliding/subducting plate. The indenter and the oceanic plate, however, generally correspond to the same advancing plate, and their motion is not independent. To permit a different behaviour of the large continent, in front of the indenter on the one hand, and above the subduction zone on the other hand, it is necessary to deform both lithospheric plates. This is much easier when a slab break-off narrows the contact between the subducting oceanic plate and the collision zone. In this case (experiment 2), we observe that the tectonic regime within the large continent passes from shortening in front of the indenter, to back-arc extension above the oceanic subduction. Moreover, back-arc extension above the oceanic subduction sharply accelerates as soon as the slab below the collision zone breaks, partly decoupling the motion of the indenter from that of the ocean. Before the onset of break-off, in contrast, the tectonic regime in the large continent remains rather homogeneous (e.g. Fig. 4a). In experiment 1, the slab below the collision zone and below the oceanic subduction preserves its integrity, smoothing the tectonic regime within the large continental plate, and only small amounts of back-arc extension are observed above the subduction zone. This extension is smaller and appears later than in experiment 2, because
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Fig. 6. F1 and F2 dimensionless parameters for experiments. Symbols are filled for 2D experiments (Regard et al., 2003) and empty for 3D experiments (this paper); bold symbols correspond to the two experiments described in this paper.
the deformation of the continental plate above the subduction zone is still largely influenced, 1000 km East of the indenter, and 50 My after the oceanic closure, by the shortening that is developing in front of the indenter. Anyway, a slab break-off below the collision zone only partly decouples the behaviour of the main continent in front of the indenter from that prevailing above the subduction zone. In experiment 2, we observe a roughly 800 km transition zone from the region undergoing pure collision to the region where back-arc extension occurs perpendicularly to the subduction zone. In fact, the transition zone located between collision and back-arc extension has already been observed in analogue experiments, like those of Martinod et al. (2000), which simulated the interaction between Anatolia and Eurasia using a piston indenting collapsing sand–silicone plate. These experiments did not take into account the rheology of the indenter-bearing plate, and the width of the transition zone was only controlled by the rheology of the upper continent. Comparison of experiments 1 and 2 clearly shows that the geometry and rheology of the subducting plate also influence the characteristics of the transition between collision and subduction. Comparison between experimental results and natural systems is limited by the inherent (over)-
simplification of our experiments. However, our results first underline that the way continental lithosphere deforms during indentation depends not only on the upper plate strength or the lateral boundary conditions. A strong and rigid cratonic area should be more prone to escape laterally as a rigid block than a softer, postcollisional lithosphere. The lateral escape mechanism can be very pronounced if a pre-existing weakness zone exists within the upper continent. However, the rheology of the indenter-bearing plate must also be considered. In this sense, our experiments can provide insights into natural situations where the indenting plate shows a transition from ocean to continent, as for the case of the Arabia–Eurasia convergence. 4.1. Comparison with the Arabia–Eurasia collision The Arabia–Eurasia collision zone of Bitlis (Fig. 7) is flanked on both sides by active subduction at the Hellenic and Makran trenches. The complete consumption of the (Neo-) Tethyan ocean and related collision of the Arabian plate with Eurasia occurred only in the Cenozoic in correspondence of the Middle East, with a progressive migration of the collision towards the east from 40 Ma onward (Dercourt et al., 1986). The
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Fig. 7. Map of the present-day tectonics of the Middle East. The main tectonic structures are modified after Dercourt et al. (1986). White arrows indicate local deformation and solid arrows indicate velocities with respect to stable Eurasia (DeMets et al., 1990; McClusky et al., 2000; Vernant et al., 2004). KF: Kephalonia fault; NAF: North-Anatolian fault; EAF: East-Anatolian fault (McClusky et al., 2000); NGF: Nayband and Gowk faults (Walker and Jackson, 2002); NZF: Neh and Zahedan faults (Freund, 1970; Tirrul et al., 1983; Walker and Jackson, 2002); CF: Chaman fault (Lawrence et al., 1992).
African–Arabian platform collided with Eurasia, first in Turkey around 40 Ma (Hempton, 1985) and later in Iran from 20 Ma (Berberian and Berberian, 1981). Beneath Makran and beneath the Hellenic trench, deep earthquakes distribution and arc volcanism show that oceanic subduction has been, and is still active today (e.g. Fytikas et al., 1984; Byrne et al., 1992; Carbon, 1996). Such a similar setting, however, corresponds to quite different tectonic regime on the surface (Fig. 7). The Hellenic arc formed as a result of extension acting in the Aegean Domain from about 25–30 Ma onwards (Jolivet et al., 1994; Gautier et al., 1999; Faccenna et al., 2003). The reason why the Hellenic trench retreats so fast during the Neogene has been the subject of a long-standing debate. Several possible mechanisms have been proposed to explain the origin of extension in the Aegean. Most of the models agree to consider gravitational forces, such as those related to slab pull (Meijer and Wortel, 1997) or related to the collapse of the pre-existing mountain belt (Jolivet et al., 1994; Le Pichon et al., 1995; Gautier et al., 1999), as the dominant factor controlling large-scale extension in the Aegean. Analysing the evolutionary pattern of the eastern Mediterranean region, one can speculate on the way experiment 2 resembles the pattern of defor-
mation observed moving from Anatolia to the Hellenic arc. A simple alternative solution, indeed, would be to consider that the extensional process in the Aegean could be also driven by a rapid increase in the slab pull force due to slab break-off occurring below eastern Anatolia. The traction exerted by the detached portion of the slab may be transferred further to the west, below the Hellenic trench. The absence of actively subducting slab below Bitlis favours this hypothesis. Tomographic images, moreover, show a large-scale high velocity anomaly below Anatolia at transition zone depth (Piromallo and Morelli, 2003). This would confirm that there, the slab detached from the upper plate whereas, below the central portion of the Hellenic trench, the subduction zone is still active. The present-day dynamics of the Aegean is close to what observed in the subduction part of experiment 2, even if length scale is approximate. At the junction between the Hellenic and Cyprus arcs a tip of a break-off may accelerate the southeastward motion of the trench. A similar hypothesis has already been formulated by Wortel and Spakman (1992) and Meijer and Wortel (1996). They propose that the evolution of the Neogene dynamics of Aegea result from slab detachment that would have initiated in the Northwest (under Dinarids) and then
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migrated towards the Aegea. In fact, new published tomographic images do not confirm slab detachment below Yugoslavia, nor below the Hellenic arc (Piromallo and Morelli, 2003). The situation is very different on the eastern side of the Arabia indenter, in Iran. In western Iran, the convergence between the Arabian and Asian continental plates is accommodated by several thrust-belts (mainly the Zagros and Alborz ranges, e.g. Vernant et al., 2004; Alavi, 1994; Berberian, 1995) while to the East, most of the shortening appears to be accommodated in the Makran accretionary prism (Byrne et al., 1992; Kopp et al., 2000, Vernant et al., 2004). Two main N–S strikeslip zones accommodate the transition between continental collision and oceanic subduction (Regard et al., 2004): the Nayband–Gowk domain located North of the Hormuz strait and moving at around 2 mm/year (Walker and Jackson, 2002), and the East Iranian ranges where the Neh and Zahedan faults appear to accommodate, 12 and 2.5 mm/year, respectively (Freund, 1970; Tirrul et al., 1983; Walker and Jackson, 2002). This pattern of deformation resembles more the situation of experiment 1. The strong development and the high elevation of the Makran accretionary prism, although partly explained by the large sediment supply, suggest that the slab-pull force is moderate. Moreover, the flatlying shallow Wadati–Benioff zone corresponding to the Makran oceanic subduction (Byrne et al., 1992) suggests that the old and heavy Neo–Tethys oceanic lithosphere is laterally supported by the continental subductions, and that the behaviour of the collision zone is strongly coupled with that of the subduction zone via the subducting plate. The geology of the collision–subduction transition on both sides of the Arabian plate differs on many points, and probably many parameters explain the different dynamics of the Makran and Aegean subduction zones. For example, the convergence velocity between Africa and Eurasia in Aegea (∼10 mm/year, McClusky et al., 2000) differs from that between Arabia and Eurasia through eastern Iran (∼25 mm/year, Sella et al., 2002; Vernant et al., 2004). Moreover, the India–Asia continental collision, East of the Makran, probably exerts a much stronger influence on the Makran tectonics than the Apulia–Europe collision zone does on the Aegean area. However, the comparison of the dynamics and active tectonics prevailing on both sides of the Arabia–Eurasia collision zone, with the presented ex-
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periments suggests that deep processes affecting the slab may partly explain the different behaviour reported on the two sides of the Arabia indenter. Slab detachment below the western part of the Arabia–Eurasia collision zone could actually increase the slab-pull below the Hellenic arc and favour extension in the Aegean domain.
5. Conclusion We explore, using laboratory experiments, how the deep behaviour of the subducting plates affect the phenomenon of indentation. These simple experiments highlight that the tectonics within the large continent does not only depend on the rheology of this continental plate, as often assumed in analogue or numerical models, but is also strongly dependant on the behaviour of the small indenting continent and of the neighbouring oceanic plate. As a matter of fact, it is both necessary to deform the upper plate, and the subducting/colliding oceanic/continental plate, to obtain the formation of an indenter-like geometry of the plate boundary. In particular, we observe that slab break-off following the continental subduction of the small continent favours the indentation process, because it increases compression in front of the indenter, and extension above the neighbouring oceanic subduction. We check that behaviour of the slab at depth within the mantle is controlled by the F1 and F2 parameters defined in Regard et al. (2003). The larger are F1 and F2 , the easier is the slab deformation, and the sharper is the transition between subduction and collision in the upper plate. Acknowledgments V. Regard’s modelling is part of his Ph.D. thesis collaborative project between the CEREGE (UMRCNRS 6635) and the Laboratory of Experimental Tectonics (University Roma Tre) supported by the French Research Ministry (Action sp´ecifique, Cotutelle de th`ese) and the “Galil´ee” project (Minist`ere des Affaires Etrang`eres). This study was developed in the framework of a wide co-operative French–Iranian project, supervised by D. Hatzfeld and M.G. Ashtiany. A portion of the program was set apart for studying the “modelling of young collision process”, research funding being provided by the program Int´erieur de la
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Terre (INSU-CNRS). Special thanks are due to J.-C. Thomas for helpful discussions and to Chad Hall, Susanne Buiter and Peter Molnar for reviews and criticism which significantly improved this paper.
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laboratory experiments. J. Geophys. Res. 104 (B2), 2749– 2766. Faccenna, C., Jolivet, L., Piromallo, C., Morelli, A., 2003. Subduction and the depth of convection in the Mediterranean mantle. J. Geophys. Res. 108, 2099, doi:10.1029/2001JB001690. Freund, R., 1970. Rotation of strike-slip faults in Sistan, southeast Iran. J. Geol. 78, 188–200. Funiciello, F., Faccenna, C., Giardini, D., Regenauer-Lieb, K., 2003. Dynamics of retreating slabs (part 2): insights from 3-D laboratory experiments. J. Geophys. Res. 108, 2207, doi:10.1029/2001JB000896. Fytikas, M., et al., 1984. Tertiary to Quaternary evolution of volcanism in the Aegean region. Geol. Soc. Spec. Publ. 17, 687–699. Gautier, P., et al., 1999. Timing, kinematics and cause of Aegean extension: A scenario based on a comparison with simple analogue experiments. Tectonophysics 315, 31–72. Hempton, M.R., 1985. Structure and deformation history of the Bitlis suture near Lake Hazar, southeastern Turkey. Geol. Soc. Am. Bull. 96, 233–243. Johnson, M.R.W., 2002. Shortening budgets and the role of continental subduction during the India–Asia collision. Earth Sci. Rev. 59, 101–123. Jolivet, L., Brun, J.-P., Gautier, P., Lallemant, S., Patriat, M., 1994. 3D kinematics of extension in the Aegean from the Early Miocene to the Present, insight from the ductile crust. Bull. Soc. G´eol. France 165, 195–209. Kopp, C., et al., 2000. Structure of the Makran Subduction zone from wide angle and reflection seismic data. Tectonophysics 329, 171–191. Lawrence, R.D., Khan, S.H., Nakata, T., 1992. Chaman Fault, Pakistan-Afghanistan. Ann. Tectonicae Spec. Iss.-Suppl. VI, 196–223. Le Pichon, X., Chamot-Rooke, N., Lallemant, S., Noomen, R., Veis, G., 1995. Geodetic determination of the kinematics of central Greece with respect to Europe: Implications for eastern Mediterranean tectonics. J. Geophys. Res. 100, 12675–12690. Mandl, G., De Jong, L.N.J., Maltha, A., 1977. Shear zones in granular material. Rock Mech. 9, 95–144. Martinod, J., Hatzfeld, D., Brun, J.-P., Davy, P., Gautier, P., 2000. Continental collision, gravity spreading, and kinematics of Aegea and Anatolia. Tectonics 19 (2), 290–299. McClusky, S., et al., 2000. Global positionning system contraints on plate kinematics and dynamics in the eastern Mediterranean and Caucasus. J. Geophys. Res. 105 (B3), 5695–5719. Meijer, P.T., Wortel, M., 1996. Temporal variation in the stress field of the Aegean region. Geophys. Res. Lett. 23, 439–442. Meijer, P., Wortel, M., 1997. Present-day dynamics of the Aegean region: a model analysis of the horizontal pattern of stress and deformation. Tectonics 16, 879–895. Molnar, P., Tapponnier, P., 1975. Cenozoic tectonics of Asia: effects of a Continental Collision. Science 189, 419–426. Peltzer, G., Gillet, P., Tapponnier, P., 1984. Formation des failles dans un mat´eriau mod`ele: la plasticine. Bull. Soc. G´eol. France 26, 161–168. Piromallo, C., Morelli, A., 2003. P wave tomography of the mantle under the Alpine-Mediterranean area. J. Geophys. Res. 108 (B2), doi:10.1029/2002JB001757.
V. Regard et al. / Physics of the Earth and Planetary Interiors 149 (2005) 99–113 Ranalli, G., Murphy, D.C., 1987. Rheological stratification of the lithosphere. Tectonophysics 132, 281–296. Regard, V., Faccenna, C., Martinod, J., Bellier, O., Thomas, J.-C., 2003. From Subduction to Collision: control of deep processes on the evolution of convergent plate boundary. J. Geophys. Res. 108, 2208, doi:10.1029/2002JB001943. Regard, V., et al., 2004. The accommodation of Arabia-Eurasia convergence in the Zagros-Makran transfer zone, SE Iran: a transition between collision and subduction through a young deforming system, Tectonics 157, in press. Sella, G.F., Dixon, T.H., Mao, A., 2002. REVEL: A model for Recent plate velocities from space geodesy. J. Geophys. Res. 107, doi:10.1029/2000JB000033. Shen, F., Royden, L.H., Burchfiel, B.C., 2001. Large-scale crustal deformation of the Tibetan Plateau. J. Geophys. Res. 106 (B4), 6793–6816. Tirrul, R., Bell, I.R., Griffith, R.J., Camp, V.E., 1983. The Sistan suture zone of eastern Iran. Geol. Soc. Am. Bull. 94, 134–150. Vernant, P., et al., 2004. Contemporary Crustal Deformation and Plate Kinematics in Middle East Constrained by GPS
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Measurements in Iran and Northern Oman. Geophys. J. Int. 157, 381– 398. Vialon, P., Bonnet, J.-L., Gamond, J.-F., Mugnier, J.-L., 1984. Mod´elisation des d´eformations d’une s´erie stratifi´ee par le d´eplacement horizontal d’un poinc¸on. App. au Jura, Bull. Soc. G´eol. France 26, 139–150. Walker, R., Jackson, J., 2002. Offset and evolution of the Gowk fault, SE Iran; a major intra-continental strike-slip system. J. Struct. Geol. 24, 1677–1698. Weijermars, R., Schmeling, H., 1986. Scaling of Newtonian and non Newtonian fluid dynamics without inertia for quantitative modelling of rock flow due to gravity (including the concept of rheological similarity). Phys. Earth Planet. Inter. 43, 316– 330. Wortel, M.J.R., Spakman, W., 1992. Structure and dynamics of subducted lithosphere in the Mediterranean region. Proc. Kon. Ned. Akad. v. Wetensch. 95, 325–347. Yoshioka, S., Wortel, M.J.R., 1995. Three-dimensional numerical modelling of detachment of subducted lithosphere. J. Geophys. Res. 100, 20223–20244.
Slab pull and indentation tectonics: insights from 3D laboratory experiments Vincent Regardc,∗ , Claudio Faccennab,1 , Joseph Martinodc,2 , Olivier Belliera,3 a
CEREGE, UMR CNRS 6635, Universit´e Aix-Marseille III, Europˆole de l’Arbois, 13545 Aix-en-Provence Cedes 4, France b Dipartimento di Scienze Geologiche, Universit` a Roma tre, Largo S. L. Murialdo 1, 00146 Rome, Italy c LMTG, UMR CNRS/IRD 5563, Observatoire Midi-Pyr´ en´ee´ s, 14 av. E. Berlin, 31400 Toulouse, France Received 28 July 2003; received in revised form 15 December 2003; accepted 26 August 2004
Abstract We investigate, using 3D laboratory experiments, how the dynamics of indentation process are affected by the evolution at depth of the oceanic and continental subductions. Lithospheric plates are modelled by sand–silicone plates floating on glucose syrup, and the density contrast between oceanic and continental lithospheric plates and asthenosphere is reproduced. Analogue experiments model the convergence between two lithospheric plates, a small continent indenting a large continental plate. We show that the surface deformation in front of the indenter and above the oceanic subduction zone depends on the behaviour of the slab below the collision zone. Slab break-off following the subduction of the small continent favours the indentation process, because it results in an increasing compression in front of the indenter, and extension above the neighbouring oceanic subduction, both of them being responsible for the appearance of the indenter-like geometry of the plate boundary. When the slab does not deform significantly at depth, in contrast, the closure of the oceanic domain in front of the indenter is followed by a longer period of continental subduction, during which the tectonic regime within the wide continent remains quite homogeneous. Comparing the presented analogue experiments with the subductions both part of the Arabian indenter within Eurasia, our results suggest that the different tectonic regime on both sides of the Arabia indenter may partly result from the probable occurrence of a detachment at depth under eastern Anatolia. © 2004 Elsevier B.V. All rights reserved. Keywords: Indentation; Collision; Subduction; Analogue modelling; Slab; Break-off
1. Introduction ∗
Corresponding author. Tel.: +33 5 61332656; fax: +33 5 61332560. E-mail addresses: [email protected] (V. Regard), [email protected] (C. Faccenna), [email protected] (J. Martinod), [email protected] (O. Bellier). 1 Tel.: +39 06 54888029; fax: +39 06 54888201. 2 Tel.: +33 5 61332666; fax: +33 5 61558250. 3 Tel.: +33 4 42971660; fax: +33 4 42971559. 0031-9201/$ – see front matter © 2004 Elsevier B.V. All rights reserved. doi:10.1016/j.pepi.2004.08.011
The term “indentation” has been coined to define the collision between India and Asia, and is classically intended to define the ploughing up of a small continental plate inside a larger one. By definition, indentation is characterised by collision occurring in front of the “indenter”, which is generally flanked on the sides
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by subduction zones. The tectonic process driving and accompanying indentation has been investigated especially for those aspects concerning the way the upper plate deforms. In particular, it has been pointed out that a major role is played by the strength of the upper plate, providing the possibilities to squeeze and thicken the crust or producing the lateral escape of crustal block along strike-slip faults (e.g., Molnar and Tapponnier, 1975; Peltzer et al., 1984; Vialon et al., 1984; Shen et al., 2001). The indentation mechanisms are frequently evoked to explain how crustal or lithospheric deformation in collision zones is associated with vertical or horizontal extrusion processes (Argand, 1924; Molnar and Tapponnier, 1975; Johnson, 2002). Presently, many studies argue in favour of combined or successive vertical/horizontal extrusion to explain deformation related to collision (Johnson, 2002). Consequently, to understand collision processes it appears necessary to look at the mechanisms of indentation in the 3D. An interesting example of indenter is Arabia, since the indentation of Arabian within Eurasia is differentially expressed to the East (Iran) than to the West (Aegean–Anatolian domain). We present 3D laboratory experiments performed to define the role of subduction in the tectonics of indentation. These experiments are part of an experimental program devoted to the investigation of the physical parameters that drive collision processes. Previous study from Regard et al. (2003) described 2D illustrating the way continental lithosphere subducts prior to collision, following an oceanic subduction. This set of experiments shows that the possibility for the continent to subduct is related to the pull exerted by the previously subducted oceanic lithosphere, and therefore to the way the slab deforms at depth. We express our results by means of two dimensionless numbers (F) where the slab pull force is normalised to the strength of the slab: favourable conditions for continental subduction correspond to low values of F, or, in other words, to situations in which the oceanic slab is more prone to maintain its integrity and to exert an higher pull on the subducting continent. Here, we extend this analysis to a more complex 3D configuration with a lateral transition from oceanic to continental subduction. In particular, we show two representative experiments out of 8, where an oceanic lithosphere neighbours a continental one. Our results
highlight that the subduction history prior to the collision and indentation exerts a primary role in the style of deformation of the upper plate. Our work has been inspired by the tectonic regime that governs the indentation of Arabia within Eurasia; a comparison with this natural example will be discussed.
2. Laboratory experiments Many 3D laboratory experiments have been constructed to analyse the tectonic processes that accompany indentation. Vertically constrained paraffin wax experiments first showed that lateral extrusion might be a fundamental process accommodating indentation (Peltzer et al., 1984). Sand–silicone–honey experiments also addressed the same issue but allowed crustal thickening, hence obtaining a more realistic partitioning between extrusion and thickening (cf. Davy and Cobbold, 1988; Martinod et al., 2000). Those experiments were all performed considering the lateral boundary of the upper plate above the subduction zone as a free boundary where the continental lithosphere can eventually spread. In this view the role of subduction process is subordinated to indentation or is simulated by the spreading of the continental lithosphere. Faccenna et al. (1996), using a similar technique, shows that the upper plate deformation depends on the way the subduction process operates on the sides of the upper plate. It is possible to obtain the whole range of deformation spanning from rigid extrusion to extension parallel to indentation by increasing the efficiency of the subducted lithosphere, i.e., of the slab pull. The problem of what does happen at depth in the collisional zone and how deep structures transfer laterally on its sides has been totally unexplored. The closure of an oceanic basin and the beginning of the collisional process itself has also been reproduced in wax experiments (Chemenda et al., 1995, 1996, 2000), or using sand–silicone–honey experiments (Regard et al., 2003). Experimental results agree in considering the possibilities that slab deforms at depth after the entrance of buoyant continental material. In particular, Regard et al. (2003) show that the slab can loose its integrity deforming at depth by ductile delamination (Rayleigh–Taylor instability) or by brittle rupture (break-off), depending on the initial imposed boundary condition and on the rheology of the slab. In
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turn, the onset of collision and the amount of continental subduction depend on the level of slab pull exerted by the previously subducted oceanic lithosphere. The current work explores the effects of a lateral transition between subduction and collision, emphasising the role of the subducting plate deformation. We show that the way the subducting plate indents the upper plate depends upon the deformation of the slab at depth. 2.1. Materials and experimental technique Following an experimental design adopted by previous authors (e.g., Davy and Cobbold, 1988, 1991), we simulate the stratified lithospheric rheological profile (e.g., Ranalli and Murphy, 1987) constructing a brittle–ductile layered model (Fig. 1). We use dry sand to model the brittle behaviour of the upper lithosphere and silicone putty to model the ductile behaviour of
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the lower lithosphere. Dry sand is nearly cohesionless, and its internal friction angle is close to 30◦ (Mandl et al., 1977). Silicone putty is a quasi-Newtonian fluid (Weijermars and Schmeling, 1986), and its viscosity is measured in laboratory before each experiment at room temperature (25 ◦ C). The density of silicone can be varied mixing it with iron fillers, which preserves its Newtonian rheology. The density of the brittle layer can be diminished mixing the sand with lighter granular materials. The sand–silicone layering rests on glucose syrup, which simulates the asthenosphere. We use two different kinds of sand–silicone layers, either lighter or denser than the glucose syrup, to represent continental or oceanic lithospheres, respectively. Two plates are constructed. One is constituted by light silicone and sand simulating a large continental plate, whereas the other is constituted by a heavy silicone–sand layer simulating a dense oceanic plate with a small continent on one side (indenter) (Fig. 1).
Fig. 1. Experimental set up: (a) upper view and (b) cross-section of the western part of the model. Sand–silicone plates modelling the lithosphere are lying above honey that represents the asthenosphere. These plates are lighter or denser than honey, to represent continental or oceanic lithospheres, respectively. The piston is pushing northward at a constant velocity of 4.4 mm/h scaled to represent 2.3 cm/year in nature. The oceanic plate will subduct below the northern continent, the subduction place being controlled by the removal of a 1-cm wide band of sand that weakens the northern margin of the oceanic plate.
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Table 1 Physical parameters and some results of experiments Experiment number
1
2
3
4
5
6
7
8
Hsil Hosand Hcsand Hm
9 4 2/4 109
9 4 2/2 195
9 3 3/3 186
9 3 3/3 186
9 3 3/3 173
9 4 2/2 103
9 4 2/4 189
9 4 1.5/4 105
Length (10−3 m) Southern continent Ocean (western part) Northern continent
Lcs Lo Lcn
190 30 240
150 70 280
220 70 210
205 70 225
210 70 220
200 70 230
190 70 240
210 70 220
Viscosity (Pa s) Oceanic silicone Continental silicone Honey
ηo ηc ηm
3.1 × 105 4.6 × 105 4.6 × 102
2.1 × 105 1.3 × 105 4.6 × 102
1.4 × 105 1.0 × 105 4.6 × 102
1.4 × 105 1.0 × 105 4.6 × 102
1.4 × 105 1.6 × 105 4.6 × 102
3.1 × 105 2.0 × 105 4.6 × 102
3.1 × 105 3.0 × 105 4.6 × 102
3.1 × 105 4.0 × 105 4.6 × 102
Density (kg m−3 ) Oceanic silicone Oceanic sand Continental silicone Continental sand Honey
ρosil ρosand ρcsil ρcsand ρm
1484 1500 1350 1300 1422
1470 1500 1310 1250 1425
1465 1500 1325 1300 1425
1465 1500 1325 1300 1425
1465 1500 1380 1300 1425
1484 1500 1355 1300 1422
1484 1500 1355 1300 1422
1484 1500 1390 1300 1422
u
1.2 × 10−6 1.2 × 10−6 1.2 × 10−6 1.8 × 10−6 1.2 × 10−6 1.2 × 10−6 1.2 × 10−6 1.2 × 10−6 63 42 72 49 72 73 70 68
Thickness (10−3 m) Silicon layers Oceanic sand layer Continental sand layers (N/S) Honey
Time Piston velocity (m/s) Total duration (h) Dimensionless numbers F1 F2
0.5 1.5
1.7 3.0
Continental and oceanic plates differ in thickness, density and viscosity of the upper silicone layer, and in thickness and density of the sand layer (Table 1). The oceanic plate strength (mainly controlled by brittle sand layer thickness) is usually larger or similar to that of continent. Sand is sieved over the entire model, with the exception of a 1-cm wide strip located at one boundary between the two plates (no sand above silicone, Fig. 1). The shortening imposed in the experiment will thus preferentially localise on this side of the oceanic plate, and the thrust fault accommodating subduction will localize there. The lower boundary of the box approximates a high gradient viscosity transition within the deep mantle (Funiciello et al., 2003). The box depth varies from 11 to 19 cm. Models are constructed inside a rectangular Plexiglas tank (50 cm long and 30 cm wide; Fig. 1). The lateral walls of the tank are lubricated with wax and the sand layer is removed near the boundaries of the
1.0 3.7
1.0 2.5
1.0 3.5
1.0 1.5
1.1 2.8
1.3 1.6
box to reduce friction at plate–walls interfaces. In some experiments (cf. experiment 2) such a lubrication is stopped 5 cm before the back wall in order to avoid the development of structures against the back wall; this difference does not modify significantly the stress and strain regime in the centre of the model, and it does not influence the conclusions presented in this paper. Horizontal shortening is imposed displacing a rigid piston at constant velocity perpendicular to the plate boundary. The piston is confined in the upper part of the tank and the glucose syrup is free to move underneath. The motion of the piston mimics the effect of an external push related to plate motion. Constant velocity is imposed, and we have no precise idea of the magnitude of the forces acting on the system. For convenience, we shall refer thereafter to regions of the experiment in terms of geographical directions, the southern boundary corresponding to the piston. A squared grid of passive sand markers enables the visu-
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alisation of the surface deformation. Photographs are taken at regular time intervals in three orientations, top, oblique, and lateral view through the transparent tank walls. The materials selected to model the stratified oceanic and continental lithosphere have been chosen to fulfil stress similarity criteria scaled for gravity (see Davy and Cobbold, 1991). Material parameters for each experiment are listed in Table 1. The scaling of our experiments is detailed in Regard et al. (2003) (see Table 2 therein); 1 cm in these experiments represents 70 km in nature and 1 h represents about 1 My in nature. The convergence velocity is scaled to represent the Arabia–Eurasia convergence velocity, the 4.4 mm/h of our model corresponding to about 2.3 cm/year in nature. Previous experiments (Regard et al., 2003) show that the shortening of the continental plate (that can be both the indenter or the indented plate) only begins after a certain amount of continental material has subducted into the mantle. This amount of subducted continental material, in turn, depends on the rheology of the slab, on the depth of the low-viscosity mantle, and on the amount of oceanic lithosphere subducted before oceanic closure. In Regard et al. (2003), these results have been expressed using two dimensionless numbers, F1 and F2 . F1 characterises the possibility for the slab to break under its own weight: F1 = FT /Frupture
(1)
where FT is the maximum tension force within the slab and Frupture , the force necessary to achieve the rupture of the slab: FT = ρo−m g Z H
(2)
Frupture = α ρm g H Hsand + ηo u Hsil /H
(3)
where ρo−m is the density contrast between the oceanic lithosphere and asthenosphere, ρm is the asthenosphere density, g is the gravity, Z is the depth of the high-gradient viscosity increase within the deep mantle (modelled by the bottom of the Plexiglas tank), H is the thickness of the slab, Hsand is the sand thickness, Hsil is the silicone thickness, ηo is the viscosity of the oceanic plate silicone, u is the convergence velocity, and α is a parameter related to the friction coefficient µ of the sand (µ = tan 30◦ ; α = 2µ/(µ + (1 + µ2 )1/2 )).
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F2 characterises the possibility for a viscous instability affecting the oceanic slab to develop before the slab reaches the bottom of the convective layer: F2 =
t1 Hsil = Zρo−m g t2 u ηo
(4)
where ρo−m is the density contrast between the oceanic lithosphere and the asthenosphere. F2 is the ratio between two characteristic times t1 and t2 , t1 = Z/u is the characteristic time for the tip of the slab to reach the bottom of the low-viscosity mantle (the slab vertical velocity is assumed to be equal to the convergence velocity, u), t2 = ηo /ρo−m × g × Hsil is the time necessary to achieve the formation of a viscous instability. In particular, we observe that slabs deform for high F2 (>3), and the possibilities to have break off is limited to the case where also F1 is high (>1.2). In contrast, low F2 values correspond to models in which the slab does not deform significantly. In 2D models the possible deformation of the slab at depth governs the delay between the oceanic closure and the beginning of collision. This delay (and consequently the amount of continental subduction) is long if the slab does not deform and decreases if it can deform. In case of break-off it can be significantly low. We focus here on the influence of F1 and F2 on the behaviour of the 3D experiments. We show that the process of indentation is sharply influenced by variations of these parameters.
3. Experimental results Experiments have been designed to test the influence of the slab pull force resulting from the subduction of the oceanic lithosphere. This parameter has been investigated varying the length of the oceanic plate prior to closure (Table 1). In the following, we describe the evolution of two selected experiments (see Table 1 for a complete description of the physical parameters of the experiments). 3.1. Experiment 1 In experiment 1 (Fig. 2), a narrow 3-cm large ocean (equivalent to roughly 200 km in nature) initially separates the indenter from the main continent. F1 is about
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Fig. 2. Top views (a), line drawings (b), shortening vs. time (c) of experiment 1. We report in (c) the total shortening occurring within the plates, excluding the shortening accommodated at the subduction zone. Two curves are reported for the eastern domain (continuous line) and for the western domain (dashed line), at 8 cm, respectively, of the eastern and western side of the box.
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0.5, and F2 is about 1.5. Under this condition the subducting plate does not deform much at depth. During the first 12 h of experiment, E–W folds, i.e., compressional structures, localise at the weak oceanic plate margin and, subordinately, near the boundaries of the box. The total shortening rate accommodated outside the young oceanic subduction zone is evaluated to around 1 mm/h (Fig. 2c). This compressional deformation characterises the initiation of subduction (e.g. Faccenna et al., 1999; Regard et al., 2003). After 12 h, the oceanic basin closure is completed, causing an increase of the shortening rate to 2 mm/h, demonstrating the low efficiency of the subducting system during the first phase of subduction in both the eastern and western side of the experiment. Compressive deformation concentrates on a linear E-trending fold belt that develops in the middle of the plate in correspondence of the end of the model/box boundary sand-less strip (Fig. 2). Only after 29 h of deformation, the shortening rate in front of the indenter becomes larger than that accommodated in the area where oceanic subduction takes place. In the eastern part of the experiment, in correspondence of the oceanic subduction, the shortening rate diminishes on average to 1.4 mm/h with a tendency to further decrease after 51 h. The western domain is more shortened, and the difference in shortening rate is accommodated by a N-trending strike-slip zone that links the trench with the main intracontinental zone of deformation, located far from the trench at the centre of the overriding plate. This strike-slip zone accommodates about 20 mm between 30 and 50 h, i.e. 1 mm/h on average. At 51 h, the eastern sector undergoes extensional deformation and the former folds are unfolded and inverted into E-trending extensional structures (normal faults and grabens). From this moment, the upper plate rotates clockwise to accommodate the different tectonic regime. 3.2. Experiment 2 Experiment 2 differs from experiment 1 as the oceanic domain, i.e. the distance separating the indenter and the target continents, is wider (7 cm). In this manner we obtain a longer oceanic slab that exerts a larger slab pull force. To obtain a maximum level of slab pull we also increase the depth of the box. Hence, F1 and F2 both increase to 1.7 and 3.0, respectively.
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Under this condition, the experimental slab must be prone to deform at depth. The initial stage of the experiment is similar to what observed in previous experiments, with some E–W compressional folds characterising the subduction initiation occurring in the experiment outside the trench (Faccenna et al., 1999; Regard et al., 2003). This phase lasts for around 10 h producing a shortening of about 10–15 mm (Fig. 3). Afterwards, the subduction zone accommodates efficiently the imposed convergence and the oceanic domain located in front of the indenter is entirely consumed after 18 h of deformation. During the final stage of oceanic closure, extension appears within the upper plate: E-trending normal faults are observed within the overriding plate with an average N–S extensional rate comprised between 1 and 1.6 mm/h (Fig. 4). The beginning of continental subduction, as observed in 2D experiments (Regard et al., 2003), does not change significantly the subduction dynamics. Extension lasts up to 26 h of deformation (Fig. 3). The extensional phase of deformation is related to the slab growth at depth. The slab pull resulting from the previous entrance of the dense oceanic plate counterbalances both the effect of the positive buoyancy of the continental part of the slab, and the effect of the piston-imposed shortening. After about 30 h of model run, under the collisional zone (western part of the experiment), we observe that the slab breaks and that the subducting oceanic lithosphere detaches from the continental part of the slab (Fig. 3). Afterwards, the detachment progresses eastward under the boundary between the two continents. It is difficult to establish the rate of lateral propagation of the detachment. However, it seems to have occurred quite rapidly. As a matter of fact, we note that in surface, on the western side of the experiment, E-striking folds and thrust-faults develop just after the slab detachment reactivating and inverting normal faults related to the previous stage (Figs. 3 and 4). Surface deformation is then strongly influenced by the detachment process and it indirectly evidences the eastward progression of the detachment (Fig. 5): the shortening domain observed during the period 33–36 h is offset of 3.3 cm with respect to that of period 30–33 h, suggesting that the detachment has progressed at a rate of about 1.1 cm/h (equivalent to roughly 8 cm/year in nature).
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Fig. 3. Top views (a) and the corresponding line drawings (c), lateral view taken from the western side of the experiment at stage D, 33 h (b), and Shortening vs. time (d) of experiment 2. The two curves reported in (d) correspond to the shortening occurring within the plates in the eastern domain (continuous line) and in the western domain (dashed line), at 8 cm, respectively, of the eastern and western side of the box.
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Fig. 4. Deformation maps of the upper continental plate in experiment 2. Graduations are in centimetres. We report the velocity fields (arrows) for the periods 20–28 h (a) and 28–32 h (b), and the rate of surface change for the same periods. Blue surfaces correspond to increasing surfaces (dilatation), and red ones to decreasing surfaces (contraction).
After this break-off, the shortening in front of the indenter reaches a rate of 2.8 mm/h, accommodating 65% of the total amount of convergence (Fig. 4). On the eastern side, conversely, the speed of the oceanic subduction increases with a maximum velocity of about 10 mm/h, twice the convergence velocity. It results in a rapid back-arc extension (Fig. 6). We observe that till the end of the experiment, the break-off of the slab remains confined to the continental collision zone, in front of the indenter. This deep behaviour increases dramatically the pull exerted by the rest of the oceanic slab on the active oceanic subduction, explaining the large amount of back-arc extension in the western part of the experiment. The contrast between shortening in front of the indenter, and extension above the oceanic subduction zone, is responsible for the appearance of the indenter-like geometry of the plate boundary.
4. Interpretation of the experimental results and application to natural examples Our results show that the way the upper plate deforms during indentation depends much on the characteristics of the subduction system and on the history of subduction. Comparison between experiments 1 and 2 highlights the dependence of the pattern of collision and deformation in 3D on the way the subduction process develops. We check that the behaviour of the slab in 3D experiments fits predictions based on the F1 − F2 diagram constructed using the 2D experiments (Fig. 6). Viscous blubs only appear when F2 is larger than about 2.5, and both F1 and F2 must be large (larger than 1.2 and 2.5, respectively) to have break-off. Experiment 1 demonstrates that the plate boundary is likely to absorb several hundreds of kilometres of continent–continent convergence before an indenter-
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Fig. 5. (a) Total N–S shortening between the marker line close to the trench and the boundary of the box, for experiment 2. This value is different from the shortening shown in Figs. 2 and 3 in which we reported the sum of all the shortening that occurred once within the plate, including within the already subducted parts of the plate. (b) Shortening rate vs. E–W position in experiment 2. In the western part of the experiment, the shortening occurs as a consequence of the continental collision. In the eastern part of the experiment, extension occurs above the oceanic subduction zone. Before break-off, shortening varies smoothly from east to west. After break-off, all the western part of the subducted slab is pulling the eastern subduction zone, and provokes the sharp variation of the shortening rate. Arrows highlight the 3.3 cm offset of the shortening domain between periods 30–33 h and 33–36 h. This offset results from the lateral migration of the slab break-off.
like geometry appears. In fact, previously published 2D experiments already evidenced that the oceanic subduction is followed by an episode of continental subduction, that may absorb large amounts of continental plate, particularly when the slab does not break or stretch at depth (Chemenda et al., 1996; Regard et al., 2003). In the 3D configuration, the subduction of the indenting continent is favoured both by the pull of the previously subducted oceanic plate, and by the pull resulting from the lateral active oceanic subduction.
Only during the last phases of this experiment, the different efficiency of the subduction system is accommodated in the surface with internal deformation or via the rigid rotation of the upper continental block. If we allow for a larger subduction system (larger initial ocean width, experiment 2), the slab pull level on the subducting slab can increase up to the point that slab breaks (Regard et al., 2003). The lateral propagation of this process (several cm/year) results in a tremendous concentration of the slab pull effect at the tip of the propagating tear (Yoshioka and Wortel, 1995). This, in turn, may induce a large-scale back-arc extension above the oceanic subducting plate (Wortel and Spakman, 1992; Buiter et al., 2002). This process favours the appearance of an indentation-like geometry between the two lithospheric plates. Then, the entrance of the indenter within the large continent results on the one hand from the shortening accommodated in front of the indenter, but also from the southern motion of the large continent on the sides of the indenter. Most of the experimental sets previously built to study the indentation tectonics focus on the rheology of the upper plate, without taking into account the rheology of the colliding/subducting plate. The indenter and the oceanic plate, however, generally correspond to the same advancing plate, and their motion is not independent. To permit a different behaviour of the large continent, in front of the indenter on the one hand, and above the subduction zone on the other hand, it is necessary to deform both lithospheric plates. This is much easier when a slab break-off narrows the contact between the subducting oceanic plate and the collision zone. In this case (experiment 2), we observe that the tectonic regime within the large continent passes from shortening in front of the indenter, to back-arc extension above the oceanic subduction. Moreover, back-arc extension above the oceanic subduction sharply accelerates as soon as the slab below the collision zone breaks, partly decoupling the motion of the indenter from that of the ocean. Before the onset of break-off, in contrast, the tectonic regime in the large continent remains rather homogeneous (e.g. Fig. 4a). In experiment 1, the slab below the collision zone and below the oceanic subduction preserves its integrity, smoothing the tectonic regime within the large continental plate, and only small amounts of back-arc extension are observed above the subduction zone. This extension is smaller and appears later than in experiment 2, because
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Fig. 6. F1 and F2 dimensionless parameters for experiments. Symbols are filled for 2D experiments (Regard et al., 2003) and empty for 3D experiments (this paper); bold symbols correspond to the two experiments described in this paper.
the deformation of the continental plate above the subduction zone is still largely influenced, 1000 km East of the indenter, and 50 My after the oceanic closure, by the shortening that is developing in front of the indenter. Anyway, a slab break-off below the collision zone only partly decouples the behaviour of the main continent in front of the indenter from that prevailing above the subduction zone. In experiment 2, we observe a roughly 800 km transition zone from the region undergoing pure collision to the region where back-arc extension occurs perpendicularly to the subduction zone. In fact, the transition zone located between collision and back-arc extension has already been observed in analogue experiments, like those of Martinod et al. (2000), which simulated the interaction between Anatolia and Eurasia using a piston indenting collapsing sand–silicone plate. These experiments did not take into account the rheology of the indenter-bearing plate, and the width of the transition zone was only controlled by the rheology of the upper continent. Comparison of experiments 1 and 2 clearly shows that the geometry and rheology of the subducting plate also influence the characteristics of the transition between collision and subduction. Comparison between experimental results and natural systems is limited by the inherent (over)-
simplification of our experiments. However, our results first underline that the way continental lithosphere deforms during indentation depends not only on the upper plate strength or the lateral boundary conditions. A strong and rigid cratonic area should be more prone to escape laterally as a rigid block than a softer, postcollisional lithosphere. The lateral escape mechanism can be very pronounced if a pre-existing weakness zone exists within the upper continent. However, the rheology of the indenter-bearing plate must also be considered. In this sense, our experiments can provide insights into natural situations where the indenting plate shows a transition from ocean to continent, as for the case of the Arabia–Eurasia convergence. 4.1. Comparison with the Arabia–Eurasia collision The Arabia–Eurasia collision zone of Bitlis (Fig. 7) is flanked on both sides by active subduction at the Hellenic and Makran trenches. The complete consumption of the (Neo-) Tethyan ocean and related collision of the Arabian plate with Eurasia occurred only in the Cenozoic in correspondence of the Middle East, with a progressive migration of the collision towards the east from 40 Ma onward (Dercourt et al., 1986). The
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Fig. 7. Map of the present-day tectonics of the Middle East. The main tectonic structures are modified after Dercourt et al. (1986). White arrows indicate local deformation and solid arrows indicate velocities with respect to stable Eurasia (DeMets et al., 1990; McClusky et al., 2000; Vernant et al., 2004). KF: Kephalonia fault; NAF: North-Anatolian fault; EAF: East-Anatolian fault (McClusky et al., 2000); NGF: Nayband and Gowk faults (Walker and Jackson, 2002); NZF: Neh and Zahedan faults (Freund, 1970; Tirrul et al., 1983; Walker and Jackson, 2002); CF: Chaman fault (Lawrence et al., 1992).
African–Arabian platform collided with Eurasia, first in Turkey around 40 Ma (Hempton, 1985) and later in Iran from 20 Ma (Berberian and Berberian, 1981). Beneath Makran and beneath the Hellenic trench, deep earthquakes distribution and arc volcanism show that oceanic subduction has been, and is still active today (e.g. Fytikas et al., 1984; Byrne et al., 1992; Carbon, 1996). Such a similar setting, however, corresponds to quite different tectonic regime on the surface (Fig. 7). The Hellenic arc formed as a result of extension acting in the Aegean Domain from about 25–30 Ma onwards (Jolivet et al., 1994; Gautier et al., 1999; Faccenna et al., 2003). The reason why the Hellenic trench retreats so fast during the Neogene has been the subject of a long-standing debate. Several possible mechanisms have been proposed to explain the origin of extension in the Aegean. Most of the models agree to consider gravitational forces, such as those related to slab pull (Meijer and Wortel, 1997) or related to the collapse of the pre-existing mountain belt (Jolivet et al., 1994; Le Pichon et al., 1995; Gautier et al., 1999), as the dominant factor controlling large-scale extension in the Aegean. Analysing the evolutionary pattern of the eastern Mediterranean region, one can speculate on the way experiment 2 resembles the pattern of defor-
mation observed moving from Anatolia to the Hellenic arc. A simple alternative solution, indeed, would be to consider that the extensional process in the Aegean could be also driven by a rapid increase in the slab pull force due to slab break-off occurring below eastern Anatolia. The traction exerted by the detached portion of the slab may be transferred further to the west, below the Hellenic trench. The absence of actively subducting slab below Bitlis favours this hypothesis. Tomographic images, moreover, show a large-scale high velocity anomaly below Anatolia at transition zone depth (Piromallo and Morelli, 2003). This would confirm that there, the slab detached from the upper plate whereas, below the central portion of the Hellenic trench, the subduction zone is still active. The present-day dynamics of the Aegean is close to what observed in the subduction part of experiment 2, even if length scale is approximate. At the junction between the Hellenic and Cyprus arcs a tip of a break-off may accelerate the southeastward motion of the trench. A similar hypothesis has already been formulated by Wortel and Spakman (1992) and Meijer and Wortel (1996). They propose that the evolution of the Neogene dynamics of Aegea result from slab detachment that would have initiated in the Northwest (under Dinarids) and then
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migrated towards the Aegea. In fact, new published tomographic images do not confirm slab detachment below Yugoslavia, nor below the Hellenic arc (Piromallo and Morelli, 2003). The situation is very different on the eastern side of the Arabia indenter, in Iran. In western Iran, the convergence between the Arabian and Asian continental plates is accommodated by several thrust-belts (mainly the Zagros and Alborz ranges, e.g. Vernant et al., 2004; Alavi, 1994; Berberian, 1995) while to the East, most of the shortening appears to be accommodated in the Makran accretionary prism (Byrne et al., 1992; Kopp et al., 2000, Vernant et al., 2004). Two main N–S strikeslip zones accommodate the transition between continental collision and oceanic subduction (Regard et al., 2004): the Nayband–Gowk domain located North of the Hormuz strait and moving at around 2 mm/year (Walker and Jackson, 2002), and the East Iranian ranges where the Neh and Zahedan faults appear to accommodate, 12 and 2.5 mm/year, respectively (Freund, 1970; Tirrul et al., 1983; Walker and Jackson, 2002). This pattern of deformation resembles more the situation of experiment 1. The strong development and the high elevation of the Makran accretionary prism, although partly explained by the large sediment supply, suggest that the slab-pull force is moderate. Moreover, the flatlying shallow Wadati–Benioff zone corresponding to the Makran oceanic subduction (Byrne et al., 1992) suggests that the old and heavy Neo–Tethys oceanic lithosphere is laterally supported by the continental subductions, and that the behaviour of the collision zone is strongly coupled with that of the subduction zone via the subducting plate. The geology of the collision–subduction transition on both sides of the Arabian plate differs on many points, and probably many parameters explain the different dynamics of the Makran and Aegean subduction zones. For example, the convergence velocity between Africa and Eurasia in Aegea (∼10 mm/year, McClusky et al., 2000) differs from that between Arabia and Eurasia through eastern Iran (∼25 mm/year, Sella et al., 2002; Vernant et al., 2004). Moreover, the India–Asia continental collision, East of the Makran, probably exerts a much stronger influence on the Makran tectonics than the Apulia–Europe collision zone does on the Aegean area. However, the comparison of the dynamics and active tectonics prevailing on both sides of the Arabia–Eurasia collision zone, with the presented ex-
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periments suggests that deep processes affecting the slab may partly explain the different behaviour reported on the two sides of the Arabia indenter. Slab detachment below the western part of the Arabia–Eurasia collision zone could actually increase the slab-pull below the Hellenic arc and favour extension in the Aegean domain.
5. Conclusion We explore, using laboratory experiments, how the deep behaviour of the subducting plates affect the phenomenon of indentation. These simple experiments highlight that the tectonics within the large continent does not only depend on the rheology of this continental plate, as often assumed in analogue or numerical models, but is also strongly dependant on the behaviour of the small indenting continent and of the neighbouring oceanic plate. As a matter of fact, it is both necessary to deform the upper plate, and the subducting/colliding oceanic/continental plate, to obtain the formation of an indenter-like geometry of the plate boundary. In particular, we observe that slab break-off following the continental subduction of the small continent favours the indentation process, because it increases compression in front of the indenter, and extension above the neighbouring oceanic subduction. We check that behaviour of the slab at depth within the mantle is controlled by the F1 and F2 parameters defined in Regard et al. (2003). The larger are F1 and F2 , the easier is the slab deformation, and the sharper is the transition between subduction and collision in the upper plate. Acknowledgments V. Regard’s modelling is part of his Ph.D. thesis collaborative project between the CEREGE (UMRCNRS 6635) and the Laboratory of Experimental Tectonics (University Roma Tre) supported by the French Research Ministry (Action sp´ecifique, Cotutelle de th`ese) and the “Galil´ee” project (Minist`ere des Affaires Etrang`eres). This study was developed in the framework of a wide co-operative French–Iranian project, supervised by D. Hatzfeld and M.G. Ashtiany. A portion of the program was set apart for studying the “modelling of young collision process”, research funding being provided by the program Int´erieur de la
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Terre (INSU-CNRS). Special thanks are due to J.-C. Thomas for helpful discussions and to Chad Hall, Susanne Buiter and Peter Molnar for reviews and criticism which significantly improved this paper.
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