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6 2D thermo-mecanical finite element models of accretion in the Central Cyprus margin∗ Contents 6.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . 136 6.1.1 Morphotectonic features in accretionary margins . . . 136 6.2 From the Cyprus trench to the Central Anatolian Plateau . . . 138 6.3 Modeling accretion in the Central Cyprus margin . . . . . . . 139 6.3.1 Model design and parameterization . . . . . . . . . . . 142 6.3.2 Model strategy and representation of results . . . . . . 143 6.4 Model 6.4.1 6.4.2 6.4.3 6.5 Discussion . . . . . . 6.5.1 A new view at 6.5.2 A new view at 6.5.3 A new view at 6.6 Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 158 Results . . . . . . . . . . . . . . . . . . . Standard model . . . . . . . . . . . . . . Suite one - sedimentation rate variations Suite two - viscosity parameters . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 146 . 146 . 149 . 152 . . . . . . . . . . . . . . . . . . . . uplift in South Turkey . . . . . . . forearc highs . . . . . . . . . . . . vertical motions in forearc regions . . . . . . . . . 155 . 155 . 156 . 157 ∗ To be modified and partially submitted to Nature Geoscience as: D. Fernández-Blanco, C. Fuller, M. Utsav, T. Cassola, G. Bertotti, and S. Willett, Explaining the Uplift in South Turkey 133 “Essentially, all models are wrong, but some are useful.” George E. P. Box 135 CAP Model Abstract Thermo-rheological and sedimentological changes acting during subduction in the overriding crust influence the landward formation of uplifted plateau-like terrains. Here, we analyze the impact of these factors on the mechanics of uplift of the southern margin of the Central Anatolian Plateau. We compiled tectonosedimentary and geophysical studies from the Cyprus trench to Central Turkey in an N-S crustal transect, which is used as a constraint to our 2D thermo-mechanical numerical models. Models show that the growth by sediment accretion of the Cyprian forearc basin system, let to thermal weakening, lower crustal viscous deformation and subsequent surface uplift of the modern Taurus Mountains. The viscous rheology of the overriding plate and the amount of sediments in the forearc basin control the type and timing of the surface uplift. We further discuss this thermo-viscous mechanism in the context of other plateau-like areas. 136 6.1 Chapter 6 Introduction Subduction margins over the globe can be categorised as erosional or accretionary depending on the direction of interplate sediment transfer; net crustal loss of sediments and trenchward migration occur in erosional margins, whereas net accretion of sediments and landward migration take place in accretionary margins [e.g. Clift and Vannucchi, 2004]. Accretionary margins develop complex trench-parallel structures, such as an accretionary wedge, a forearc basin system, and a volcanic arc [Gutscher et al., 1996, 1998; Kukowski et al., 2002] (Fig. 6.1). Whereas the structural development of topography in the seaward areas of the system is related to brittle deformation caused by the friction exerted by the subducting slab and the landward accretion of material, the mechanisms of uplift for forearc highs, plateau-like areas with high elevations that develop inland at some accretionary margins, are open to debate [e.g. Fuller, 1996; McNeill et al., 2000]. The modern Central Taurus Mountains (CTM) in southern Turkey, an uplifted area of some 350 km E-W and ∼200 km N-S, can be perceived as the southern flank of the Central Anatolian Plateau (SCAP). Surface uplift in this region has been accurately dated as latest Miocene or younger (post–8 Ma) by recent studies, which attribute the upward motion to break-off of the Cyprus slab [e.g. Cipollari et al., 2012; Schildgen et al., 2012b; Cosentino et al., 2012; Cipollari et al., 2013; Schildgen et al., 2014a]. However, these studies fail to explain (i) why the maximum surface uplift occurs above an area where the Cyprus slab is still presently attached, at least partially [Bakırcı et al., 2012], and miss to address (ii) the coeval subsidence taking place in South Turkey offshore [Walsh-Kennedy et al., 2014] or (iii) the presence of regional contractional structures in this region (chapter 5). In this contribution, we test the hypothesis that the SCAP formed as the forearc high of the Cyprian subduction margin, effectively fragmenting the preceding Miocene forearc basin and differentiating the offshore Cilicia and onshore Mut basins (chapter 5 and Fig. 6.1). We use a series of 2D thermo-mechanical finite element models (FEM) to investigate the mechanics of the surface uplift of the SCAP in relation to the Cyprian subduction margin. Simulations are constrained by the present-day thicknesses and geometries, geological observations, and kinematic constraints of an N-S transect in which we integrated tectonosedimentary and geophysical studies of the Cyprian accretionary margin (Fig. 6.2). Models include, but are not limited to, visco-plastic and critical wedge mechanics, isostatic compensation, and the effects of variations in sediment input, rheological behavior, and thermal conductivity. We look for an overall agreement on the evolution of the major morphotectonic units (between the model and nature), focusing particular attention on the uplifted CTM and its bounding domains to the south and north. We propose applying this approach to other accretionary margins. 6.1.1 Morphotectonic features in accretionary margins The terms assigned to the different morphotectonic features observed in arc-trench regions are defined in the following according to how they are used in this contribution. In general, terms to describe natural examples are used after their original definition as in Karig and Sharman [1975]; Dickinson and Seely [1979], while simpler, more visual names are used instead for the description of features in the models [e.g. Fuller et al., 2006a]. 137 CAP Model Trench Trench-slope Trench fill s.l. basin Trench-slope break Accretionary Forearc-high Forearc basins Residual ? Subductin area of mechanical accretion Not to scale Messaoria Basin Troodos Ophiolite Levantine basin s.l. ? Overriding plate Kyrenia Range Cilicia Basin N Central Taurus Mountains No volcanic arc Mut Basin ? Cyprus S la area of mechanical accretion Not to scale 2D Models Proside Oceanic ? (CVP) Anatolian plate Structural Highs Wedge-top basin lithosphe Not to scale ? unknown mechanism b Proside basin ? unknown mechanism g plate Cyprus Arc S Volcanic arc Intramassif re Forearc-high Forearc probasin Forearc retrobasin Retroside Generic forearc Accretionary wedge Mechanical domain Figure 6.1: Terminology of the features of a conceptual accretionary margin with forearc high and their correlation to the main features seen in our study area and those in our models. Generic terminology is used in the same manner to Karig and Sharman [1975]; Dickinson and Seely [1979]. The terms of “trench-fill basin” and “accretionary forearc basin” are referred to in the model description as “pro-side” and “wedge-top” basins, respectively, and the terms “residual” and “intramassif forearc basins” of Dickinson and Seely [1979] are called, during the modeling description, as “landward” and “seaward” forearc basins, respectively. From the subducting slab landward, accretionary systems are composed of: (i) the trench-fill basin, which is a sedimentary infill at the front of the trench, on top of the contact point between the subducting and the overriding plates; (ii) the trench itself, an area grown by active deformation that includes the trench-slope, shaped by imbricate seaward-verging thrust faults; and (iii) the trench-slope break, in which the slope changes to landward-dipping in response to the appearance of landward-verging thrusts. The trench region might contain accretionary forearc basins inDickinson and Seely [1979] sense that are often carried on top of thrust sheets as piggy-back basins. Inland of the trench-slope break, (iv) forearc basins 138 Chapter 6 might appear as residual forearc basins located between the trench and the arc massif, as intramassif forearc basins with sediments that unconformably lie on the arc massif, or as a constructed basin, a large single forearc basin (not shown in Fig. 6.1) enclosing the two previously described types [Dickinson and Seely, 1979]. Before reaching the volcanic arc, which delimits the forearc from the backarc regions, (v) a forearc high might be present, separating the residual and the intramassif forearc basins or enclosing them within the intramassif forearc basin (Fig. 6.1). In the models, simpler nomenclature is used. Convex-up features that develop in relation to shear zones are termed “structural highs”. Basins commonly appear enclosed between adjacent structural highs and are named based on their location with respect to the pro- and retro-sides of the models (sea- and landward, respectively) and the area of active deformation. Developing at the pro-side of the model, (i) proside basins and (ii) wedge-top basins [DeCelles and Giles, 1996] are differentiated on the basis of their distinctive wedge-top position for the latter, or lack thereof for the former. Heading to the retro-side of the model, a well-developed (iii) structural high separates the pro-side area of active deformation (shear zones) from a stable area, where the (iv) forearc basin develops. If present, the (v) forearc structural high divides the forearc basin in pro- and retro-side forearc basins at the front and back of the forearc high, respectively. The term “negative-alpha basin” is used for model and natural cases alike, in the Fuller et al. [2006a] and Willett and Schlunegger [2010] sense; this is, a basin with a surface slope reverse to that of the wedge on top of which the basin remains stable with no internal deformation, and passively sliding above the subduction thrust, as long as it is restricted and steadily infilled by its bounding highs. 6.2 From the Cyprus trench to the Central Anatolian Plateau A regional transect that runs north from the East Mediterranean to the Central Anatolian Plateau interior is shown in Fig. 6.2. This crustal-scale transect encloses relevant tectonosedimentary and geophysical data from on- and offshore studies. This integrative effort, unique for the area, allows the determination of four key features; (i) the shape of the subducting slab and the geometry of its contact with the continental crust, (ii) the overall distribution of crustal thicknesses, (iii) the relative age and cutting relationships of the main faults, and (iv) the position, geometries and continuity of the Miocene rocks. The transect shows three first-order tectonic features (Fig. 6.2). To the south, the subducting African plate is represented by a continental fragment, the Erastothenes Seamount, and farther north by oceanic crust, starting at >15 km depths from around 35°and northward. The Troodos Ophiolite in the center of the transect is a sliver of oceanic crust in between plates. Toward the north of the transect, the Anatolian overriding plate consists of thickened continental crust. Crustal thicknesses along the transect range from a minimum of ∼25 km to a maximum of approx. 45 km. In the southern sectors of the transect, thickness changes are well detected by the gravimetric signal of Ergün et al. [2005] (Fig. 6.2B) and the Moho models of Koulakov and Sobolev [2006] (Fig. 6.2-C). In the African plate, average thicknesses of ∼28 km are observed below the Erastothenes Seamount south of Cyprus. More to the north, between 34°30’ and 37°N (Fig. 6.2-C), the geophysical models concur on a significant increase in Moho depth, from some 28 km CAP Model 139 to more than 40 km, which we correlate with the steepening (up to 40°) of the subducting slab (Fig. 6.2-E). The thinnest oceanic crust (∼25 km) is seen below the trench area. Northward, thickening occurs in relation to the Troodos Ophiolite, the detachment depth of which is uncertain. This crustal thickening is probably the result of the Troodos emplacement, with probable thrust doubling. This is also the location underneath which the locked underthrust of the Erastothenes Seamount is found. The extent of the continental crust underneath the Troodos Ophiolite and the position of its transition to oceanic crust more to the north remain enigmatic. In the southern regions of the overriding Anatolian plate, a relevant crustal thickening takes place below the CTM. Farther north, the Anatolian crust thicknesses decrease gently from circa 45 km to values in excess of 35 km. Here, Pn tomography was used [Mutlu and Karabulut, 2011] instead of gravity data [Özeren and Holt, 2010] (Fig. 6.2-C), which points to crustal thickness values up to 10 km thicker. South-verging thrusts led to fragmentation of the area into present-day sedimentary basins and structural highs (Fig. 6.2-D). These thrusts have been linked either directly (in the south) [e.g. Stephenson et al., 2004; Calon et al., 2005a, b], or indirectly (in the north) (Chapter 5), with the Cyprian subduction thrust and become older when progressing northward along the section (Fig. 6.2); they are as young as present-day at the trench [e.g. Stephenson et al., 2004], Pliocene or younger in north Cyprus [e.g. McCay, 2010], mid-Pliocene in the Cilicia Basin (Chapter 5), and pre-Miocene in the Mut Basin [e.g. Çiner et al., 2008]. Pliocene contractional structures are seen in the Cilicia Basin, but no evidence of any Miocene or younger thrust system is known for the Mut Basin or for the transition between both basins. This transition is marked instead by a long-wavelength monocline [Çiner et al., 2008] that accommodates ∼4 km of relative vertical displacement between initially laterally equivalent Miocene rocks (Chapter 5). From south to north, Miocene rocks outcrop in four different localities: the Messaoria Basin in Cyprus, the offshore Cilicia Basin, the Mut Basin on the CTM, and the Tuz Gölü Basin, in the plateau interior. Originally deposited on top of a peneplaned surface, these rocks allow the determination of the relative end-result vertical motions since the Miocene. The Miocene basins seen along the transect show dissimilar geometries. Rocks above pre-Miocene substratum thin southward in the asymmetrical Messaoria Basin [Calon et al., 2005a, b; McCay, 2010]. Messinian and younger deposits of the Cilicia Basin also thin southward, but the geometry remains speculative for pre-Messinian Miocene, since its bottom is not seen in the seismic profiles of the area [e.g. Aksu et al., 2005a]. This asymmetry is not seen for the Mut Basin nor farther north. Basin terminations are erosive in all four Miocene basins, with the exception of the southern margin of the Messaoria Basin. 6.3 Modeling accretion in the Central Cyprus margin Our modeling analysis is suitable to describe the central segment of the Cyprus arc, which is normal to the convergence direction and is where contractional structures develop. Our models simulate a transect running roughly S−N, which is parallel to the convergence and the main transport direction of material coming from the Cyprus arc (Fig. 6.3). However, other relevant sources of material are the mountainous areas to the north and northeast. This and other important three dimensional effects that might occur in the area [e.g. Wortel et al., 2010] are not taken into account. B C mGal Turkey 0 -100 0 200 Erastosthenes East Mediterranean Cyprus CVP Turkey Central Anatolia 0 -100 30 20 Figure 6.2: Central Cyprus geologic and geophysical transect (I) (a) Map view, Bouguer gravity analysis and cross-sectional representation of geophysical data available for the study area. ~110 km Vertical exaggeration ~ 3,5 50 Crustal thickness via Pn tomography, as in Mutlu & Karabulut, 2011 Moho depth via real data inversion, as in Koulakov & Sobolev, 2006 Moho depth via seismics + gravity, as in Özeren & Holt, 2010 Moho depth CRUST 2.0, as in Özeren & Holt, 2010 0 200 10 Moho depth CRUST 2.0, after modification by Koulakov & Sobolev, 2006 Reflaction seismics, as in Mart & Ryan, 2002 Bouger gravity analysis, as in Ergun et al, 2005 Aeromagnetic survey MTA - Interpreted by Ates et al., 1999 Taurus Range 50 Cilicia Basin 40 Ergun et al., 2005 Trench Northeast Mediterranean 40 30 20 10 Km mGal Km A 140 Chapter 6 50 Volcanics Oceanic crust Troodos Ophiolite Miocene Paleogene Plio-Q Messinian Mantle lithosphere Continental crust Figure 6.2: Crustal-scale transect running N-S (650 km at around 33°30’ E). The transect stretches from the Levantine Basin to the Tuz Gölü Basin, passing though the Cyprus arc trench, Cyprus island, the Cilicia Basin, and the Taurus Mountains, finally reaching the Central Anatolian interior. Panel A represents a map view of the area. Panel B shows the two major gravimetric studies in the area [Ates et al., 1999; Ergün et al., 2005]. Panel C shows the interpretation of the offshore section c of Ergün et al. [2005], the interpretation of the seismic study performed by Mart and Ryan [2002], and several cross-sections constructed from depth map models of the Moho [Koulakov and Sobolev, 2006; Özeren and Holt, 2010] and Pn tomography [Mutlu and Karabulut, 2011]. The circles are focal epicenters with Mw > 5 recorded in a longitudinal area from 32°30’ E to 34°E in red and 31°30’ E to 34°30’ E in green. Panel D is the upper sector of the transect (up to 10 km depth), shown with ∼ 7 km of vertical exaggeration for the sake of visualization, and it integrates this Thesis and upper crustal studies [Robertson, 1998b; Stephenson et al., 2004; Calon et al., 2005a, b; Çiner et al., 2008; McCay, 2010; Fernández-Blanco et al., 2013]. Panel E is the crustal cross-section for the area, integrating all the aforementioned inputs. (b) Continuation from the previous page. Upper crust and whole crust sections of the study area. Eratosthenes SeaMt Levantine Basin 40 ~110km Vertical exaggeration ~ 2 50 39º 40 38º 30 30 37º 20 20 36º 10 10 35º 0 0 34º ~110km Vertical exaggeration ~ 7 10 10 Turkey 5 5 Km Km East Mediterranean B 0 0 Km Km A CAP Model 141 142 6.3.1 Chapter 6 Model design and parameterization Accretion mechanisms at accretionary margins are often described using the critical wedge theory [Davis et al., 1983; Dahlen, 1984; Larroque et al., 1995; Malavieille, 2010], which defines the geometry of the orogenic wedge as a function of the mechanical properties of the accreting wedge and that of its surface of accretion. However, the critical wedge excludes the ductile properties of these systems, such as the absence of brittle fracturing in their lower ends, which cause relevant deviations from the predicted brittle morphologies. Recent research inclusive of the visco-plastic attributes show the influence of thermal or rheological variations or that of sediment load and/or competence in the strain distribution and deformation patterns within the accretionary wedge [e.g. Fuller et al., 2006a; Simpson, 2010; Willett and Schlunegger, 2010; Fillon et al., 2013]. Our 2D kinematic-dynamic models consist of two coupled domains (Fig. 6.3). The domain where mechanical laws apply represents the crust of a deforming subduction zone. Accretion of incoming sediments is driven by the tangential velocities at the base of the mechanical domain. These velocities decrease toward, and become zero at, the “S” point, which represents the point of contact of the subducting slab and the continental Moho. The thermal domain covers the whole model, including the mechanical domain. For a detailed description of the model formulation and boundary conditions, the reader is referred to Chapter 2, sections 2.4.1 and 2.4.2. The model is consistent with time frame and cross-sectional lengths of interest and simulates 25 Ma of subduction in a transect of 550 km (Fig. ??). Parameters such as amount of material incoming at the trench and convergence velocity are set constant during the simulations. We adopted values of 3 km and 35 mm/y, respectively, which are considerably lower than present for the incoming material at the trench and higher than present for the convergence velocities. Our choice is based on the needed extrapolation in time (25 Ma) of present-day values. The presentday sedimentary thicknesses in the East Mediterranean Sea range from 10 km to 15 km [e.g. Makris and Stobbe, 1984]; these values are surely the largest along the time frame of interest, considering factors such as the narrower confinement of the present-day Mediterranean or the presence of the Nile. Similarly, the underthrust of the Erastothenes Seamount below south Cyprus presently decelerates subduction motion to 9.3 ± 0.3 mm/y [Reilinger et al., 2006], thus a larger, more common, convergence velocity value is given for the time considered for our simulations. The subducting lithosphere is 50 My old at the left side of the model and its 70 km thickness remains constant during the running time. Since thicknesses in the mechanical domain, which represents the crust, change as accretion takes place, an initial thickness of 30 km is chosen on the basis of the similarity between thicknesses in nature and in models after the run is completed, i.e., maximum values of 45 km near the “S” point. The rest of the overriding lithosphere is 80 km thick. Flexural rigidity is set at 2.4 x 1023 N · m for both plates (after Fuller [1996]). Variations of this value of up to four orders magnitude did not produce substantial changes [e.g. Forsyth, 1985, for a discussion on flexural rigidity values]. Densities are commonly accepted values: 2.8 g/cc for the lithosphere (which includes the sedimentary cover), 3.3 g/cc for the mantle and 1.03 g/cc for the overlaying layer of water. Cohesion and internal friction angles control the mechanical strengths in our model. Cohesion, c, is set to 1000 Pa, a value higher than expected for the crust, 143 CAP Model da = Incoming thickness A Vx = Vc 2 Km Crust - Mechanical domain Vc = 10 km/Ma Ds = 2,4 x 1023 N·m Vtg 0 S Fully Coupled Thermo-mechanical model B C = 1000 Pa n = 3,0 Ao = X x 10 -19 Pa-3 . Ma-1 Q = X x 105 J . mol-1 φ = 24o φo = 8o Vtg = Vc Vtg = 0 Do = 2,4 x 10 N·m Full Model - Thermal domain Crust V = Vc (kinematic) Oceanic 3 Lithosphere H = 0 mW/m o T = 0 oC V = Vmech H = 1,25 mW/m3 K = 2,0 W/(m oK) C = 1000 J/(Kg oK) p S K = 2,0 W/(m K) Cp = 1000 J/(Kg oK) dT = 0 dx Vx,y = 0 23 Asthenosphere V=0 H = 0 mW/m3 K = 50 W/(m oK) Cp = 1000 J/(Kg oK) T = 1400 oC Continental Lithosphere V=0 H = 0 mW/m3 K = 2,0 W/(m oK) Cp = 1000 J/(Kg oK) dT = 0 dx Asthenosphere T = 1400 oC Figure 6.3: Model setup, with indication of thermal and mechanical parameters. to maintain model stability. Lower values do not affect the outcomes [Fuller, 1996]. The internal friction angle of the crustal material, φ, is set to 27◦ and the friction angle between the subducting and the overriding plates, φb , to 8◦ . Friction values are set low to include the effect of fluid pressures not explicitly taken into account and imply fluid pressure ratios within the range of those at accretionary wedges [Fuller, 1996, and the references therein]. To let the thermal structure equilibrate, the thermal model runs for 20 My before the crustal model onsets. Surface temperature is 5◦ C, an average between subaereal and subacuatic temperatures, and a value of 1400◦ C is given for the asthenosphere at the base of the model. We use values for thermal conductivity of 2 and 50 W/(m · ◦ K) for the lithosphere and the asthenosphere. Asthenospheric conductivity values are readably high to represent isothermal conditions. Heat production has a value of 0.85 µW/m3 [Jaupart and Mareschal, 2005], occurring only in the mechanical domain. Specific heat, c, is 1200 J/kg ◦ K for both model domains. 6.3.2 Model strategy and representation of results The constraints presented above are used to study the mechanics of uplift of the southern margin of the Central Anatolian Plateau. A standard model, which best represents the known geologic evolution for the central Cyprus arc, is chosen attending to the overall agreement of large-scale structure evolution and vertical motions through time, as well as final geometry. Based on our preferred model, the values of sedimentation rate and viscosity for the parameter space studied are obtained as outcomes. Then we estimate the sensitivity of the model to the influence of each parameter using two suites: (i) Suite One for the sedimentation rate, and (ii) Suite Two for the viscosity, where we study the activation energy, Q, and the power-law viscosity coefficient and exponent, Aµ and nµ . 144 Chapter 6 Representation of the morphotectonic features We intend to quantify the rates of the vertical motions and the geometrical variations within and around the forearc high before and during its development. Changes in the rate of the motions of the different components of the forearc basin system and its resulting evolution can be traced over time following the position of the upper limit of the basement, i.e., the surface representing the top of the mechanical domain at the moment of model initiation. We track the position over time of the lowest and highest points of the top of the basement for the area in and around which the forearc high develops to track the development of basins and forearc highs and their change through time (Fig. 6.4). In this schematic representation, a sedimentary basin is confined between two highs and a structural high is enclosed between two lows. The width of the basins is the distance, measured on a straight line, between its highest bounding points (B0 W for time step 2, or BP W and BR W for time step 3 in fig. 6.4-AB). Similarly, the width of the structural highs is defined as the distance on a straight line between its bounding lowest points (F W for time step 3 in fig. 6.4-AB). The depth/height of each feature is defined differently depending on the morphotectonic features present at every time step. Before the development of the forearc high (time step 2 in fig. 6.4-AB), the depth of the forearc basin is the vertical distance between the deepest point of the basin and its position at the moment of model initiation (B0 D for time step 2 in fig. 6.4). After the onset of the forearc high growth (time step 3 in fig. 6.4-AB), and in order to better represent the relative vertical motions of the forearc high and that of its frontal and rear basins, a new base level change takes place at every time step. Restitution of the base level occurs with respect to an arbitrary point, which coincides with the landward boundary of the forearc basin system at every time step. This point is defined with a deliberate increase of initial model topography in the retro-side of the model, and continuously subsides at a similar pace to that of the area where the forearc high develops. The height of the forearc high (F H), and the depth of the pro- and retrobasins (BP D and BR D, respectively), are measured with respect to the new base level at every time step. The base level change removes (at least part of) the primary subsidence signal and allows the quantification of the relative motions of the forearc high and its bounding basins. It also allows the distinction between the growth of (i) a subsiding high, a feature that subsides less than its surroundings but remains at deeper absolute levels than the reference point, and that of (ii) an uplifting high, an uplifted area that overcomes the primary subsidence and is then treated as the forearc high. Only the case of the uplifting high (forearc high) is represented in the plots. We anticipate that the first-order variations observed in our models would will occur in any accretionary margin (Fig. 6.4). Therefore, the measured widths and depths/heights of each morphotectonic element have been converted into percentages in relation to the maximum basin depth and width of the standard model, and are depicted as the base and altitude of the triangles, respectively. The maximum basin depth accounts for the depth of the initial forearc basin and the depth of the forearc pro-basin, whereas the maximum basin width corresponds with the width of the initial forearc basin (see Fig. 6.4). This representation allows for easy comparison between different time steps of the same model and among different models. Time-step cross-sections of top of the basement Forearc high heigth (FH) Time-step 2 Base level before onset forearc high formation B PW B0W BRD Forearc retrobasin width (BRW) Forearc high width (FW) Forearc initial basin width (B0W) Forearc probasin width (BPW) BRD FW BRW l ch a Base level after onset forearc high formation Time-step 3 ve FH 10 km 20% max. width 3 2 Time-steps Base le D C Maximum basin depth of standard model 5 km 0 Ma 20% 5 Ma 10 Ma 15 Ma 20% 15 Ma 10 Ma 40% 5 Ma 0 Ma 60% 80% 20 Ma Height/depth of the highest and deepest points of the basement at 5 Ma, with respect to the reference point (after forearch high formation) Depth of the deepest point of the basement at 20 Ma (before forearch high formation) Position of basement at 25 Ma (considered zero) Reference point (changed to new position evey time -step after forearc high formation) Maximum basin width of standard model 50 km 20 Ma Reference surface (used every time-step before forearc high formation) 25 Ma Standard Model Figure 6.4: Example of a simple representation of the forearc basin system evolution (left) and simple representation of the evolution of the standard model (right). The upper panels represent the evolution of the basement top surface for a hypothetical case (left) and our standard model (right). The lower panels show the proposed representation of the upper cases. Basins are shown as normal triangles, isosceles are used for the initial forearc basin, a right triangle looking left is used for the forearc pro-basin, and a right triangle facing right is used for the forearc retro-basin. The structural highs are represented as inverted isosceles triangles only if they grow pass the chosen base level. For the right panels, the different time steps are shown as different colors; 20 Ma in orange, 15 Ma in green, 10 Ma in blue, 5 Ma in brown, and present in red. Simple representation of the width and depth evolution of the morphotectonic elements of interest B0D B 2 3 Forearc initial basin depth (B0D) Reference Example 20% 20% 40% Time-steps Forearc probasin depth (BPD) e ng Forearc retrobasin depth (BRD) 20% max. depth 60% km 80% A CAP Model 145 146 6.4 Chapter 6 Model Results In the following, we describe our standard model and analyze the overall geometrical and evolutionary changes between the standard and the rest of the models. 6.4.1 Standard model A sedimentation rate of 0.5 mm/yr and viscous parameters values of Q = 1.7 J/mol, Aµ = 2.05 Pa−n /Ma, and nµ = 2.65 reproduce our preferred model. The complete list of parameter values used in the standard model is shown in Table 6.1. Corresponding to the crust, the mechanical domain of the standard model is shown (mostly in green) in fig. 6.5. At the moment of model initiation (25 Ma) we can observe the geometry of the mechanical domain. A roughly flat topography dominates the area, with the only exception of the topography defined at the retro-side of the model, at the right hand of figure 6.5. At the left-hand of this figure, the curved area below the mechanical domain represents the down-going slab. As seen at 15 Ma, topography has started to develop in the left-hand side of the model (pro-side). This topography is seen in relation with two shear zones, marked in warmer colors than their surroundings. A small asymmetric wedge-top basin developed between the shear zones, which are seen as white horizontal isochrones. This basin is translated between the two shear zones as they migrate retroward (toward the right). At the same time, a large forearc (negative-alpha) basin develops in the center and retro-sides of the model. The subducting slab becomes steeper as the base of the crust becomes more horizontal. Both factors influence the extent of strain around the “S” point at the base of the crust. By 10 Ma, the wedge-top basin migrates retroward and the shear zones become wider and more pronounced. Further deepening and thickening in the forearc basin leads to increasing temperatures and strain accumulation at the base of the crust. As the sediments accumulate in this forearc basin, the initial formation of a frontal branch of a shear zone occurs at the “S” point and basement-basin contact begins to dip toward the left (proward). At 5 Ma, two small basins have developed in the pro-side. Toward the centralright area of the model, two shear zones branch out from the “S” point and produce an uplifted broad bulge. These shear zones clearly delimit the forearc high uplifted area, which dips more markedly toward the left-hand side of the model. The basement further tilts proward (left), and the formation of a monocline is clearly seen, with horizontal deposits present on top of the bulge and dipping away from it to its sides. Onlap relationships take place on both sides of the bulge that are more angular in the pro-side. This bulge has divided the forearc basin into a pro-side negative-alpha basin to the left and a retro-side landward basin to the right; it continues leading the development of onlap relationships to its sides. The deflection of the slab proceeds toward higher angles (∼40°) at the “S” point and the base of the crust evolves to smaller ones. Since 5 Ma, the system evolves in a self-similar manner. The basement of the pro-forearc basin deepens while further uplift takes place in the forearc high, in which an area with very low values of strain is now seen in the uppermost ∼5 km. 147 CAP Model 10 km 20 km 10 km 20 km 10 km 20 km 10 km 20 km 10 km 20 km Turkey East Mediterranean Cyprus Cilicia Basin Taurus Mountains CAP Interior Figure 6.5: Mechanical model evolution in time steps, with zoom-ins into the area where the forearc high develops. Times represented are, from top to bottom:25 Ma, 15 Ma, 10 Ma, 5 Ma, and present. The individual lines on top of the basement represent isochrones and should reflect the overall geometric relationships expected for strata. 148 Table 6.1: Principal parameter values of the standard model Parameter Description Value Mechanical model ̺c ̺m ̺w φ φb c D Sα νc h Sedr Lithospheric density Mantel density Water density Internal friction angle of crustal material Friction angle between subducting and overriding plate Cohesion Flexural rigidity (both plates) Slab angle at S point Convergence velocity Sediment thickness in the subducting plate Sedimentation rate 2800 kg/m3 3300 kg/m3 1030 kg/m3 27◦ 8◦ 1000 Pa 2.4x1023 N · m 40◦ 35 mm/yr 3 km 0.5 mm/yr Thermal model Thickness of the subducting lithosphere Thickness of the overriding lithosphere Age of the subducting lithosphere Thermal model runup Thermal conductivity (lithosphere, asthenosphere) Surface temperature Asthenosphere temperature Heat production Specific Heat Power-law viscosity cohefficient Power-law viscosity exponent Activation energy 70 km 80 km 50 My 20 My 2.0, 50.0 W/(m ·◦ K) 5◦ C 1400◦ C 0.85 µW/m3 1200 J/kg ◦ K 2.05 Pa−n /Ma 2.65 1.7 J/mol Chapter 6 Hs l Ho l M al ith tr unup Kl , a Ts Ta A c Aµ nµ Q 149 CAP Model 6.4.2 Suite one - sedimentation rate variations Recent thermo-mechanical modeling studies have focused on the impacts of sedimentation rate on the structural evolution of orogenic wedges and have successfully shown that sedimentation rate strongly affects the stability of the wedge and the distribution of strain within these systems [Fillon et al., 2013; Cassola, 2013]. Building upon this concept, we analyze how sedimentation rate influences the thermal distribution underneath and within orogenic wedges, and thus affects their brittle or ductile behavior, resulting in differential strain mode and strain localization. In our models, sedimentation is “fill-to-spill” in that we assume that sufficient sediment is available and that areas capable of accommodating sediments are filled to capacity, i.e., areas between structural highs are completely filled. In this section, we study the effects that variations in sedimentation rate produce on the development of the different morphotectonic elements of the margin. Table 6.2: Sedimentation rate parameterization Model Sedimentation rate (mm/yr) Standard Sr1-to-Sr14 0.5 0 to 1.4 (∆0.1) Variation in Sedimentation Rate 0.1 mm/y 15 Ma 0.9 mm/y 10 km 5 Ma 10 km 10 km 0 Ma 10 km 10 km 10 km Figure 6.6: Forearc high development with changes in sedimentation rates. The figure shows 15 Ma, 5 Ma, and 0 Ma snapshots for the models with 0.1 mm/y (left column) and 0.9 mm/y (right column) 150 Chapter 6 As sedimentation rate increases, strain becomes greater in and around the “S” point and is distributed on both sides away from it. An increase in sedimentation rate also promotes the development of a second bulge in the retroward end of the model. The subducting lithosphere becomes steeper retroward (to the right) and the deflection of the base of the crust increases toward the “S” point. That is, the “S” point increasingly subsides with increasing sedimentation, thus the surface uplift per se (with global reference frame) decreases. However, with respect to the “S” point only, or any fixed surface, such as the sea level, a more unexpected behavior is observed. Starting from the model with no sedimentation, we notice that larger sedimentation rates produce larger surface uplift. If we continue to increase accumulation rates, surface uplift eventually decreases until a threshold is reached, after which surface uplift becomes larger again (Fig. 6.7). With low sediment accumulation rates, the initial forearc basin is thin, which contributes to a surface uplift that occurs earlier with respect to model onset (i.e., older), and both upward and downward vertical motions take place over shorter periods. In general, with lower accumulation rates the margin is more susceptible to the influence of the other parameters; for example, the incoming sediment at the trench highly dominates the structural evolution of the margin, developing sharper shear zones that migrate retroward faster and structural highs and lows that are larger in amplitude. As sedimentation rates increase, the overall relative vertical displacement between subsided and uplifted areas is consistently larger and occurs in narrower (horizontally shorter) distances, i.e., there are larger amplitudes in the main subsidence-uplift pair (forearc high–pro-basin). Another effect from increasing the accumulation rates is the development of thicker forearc basins that stabilize the margin; these force deformation outside the basin into its margins and farther away. The surface uplift becomes smaller (Figs. 6.6 and 6.8) and eventually stops happening. The uplifted and subsided areas become broader, and the points of depocenter and maximum uplift move landward. Similarly, as the sedimentation rate increases, sedimentary angular discordances develop more commonly instead of onlap relationships, i.e., the vertical displacements occur more gradually. Variations in sedimentation rates influence virtually every aspect of the growth of the margin. When sediment input is small, the development of structural highs along the margin takes place at high rates and at early stages of the simulations. A broad forearc high develops with very small flanking basins (Fig. 6.8, left-hand). As sedimentation increases, a long wavelength subsidence takes place at model initial stages and the onset of the forearc high formation is delayed (Fig. 6.8, right hand). Eventually, with higher sedimentation values, the forearc high formation stops and only the long wavelength subsidence is observed. We can therefore conclude that the forearc high is a common feature developing in accretionary margins, but its development is gradually halted by the long wavelength subsidence imposed by the sediments. 151 CAP Model Top of the basement Sr14 = 1.3 mm/y Sr11 = 1.0 mm/y Sr8 = 0.7 mm/y Sr6 = 0.5 mm/y Sr4 = 0.3 mm/y Sr1 = 0.0 mm/y Bottom of the crust Figure 6.7: Line trace of the bottom of the crust and top of the basement in models with changes in sedimentation rates after a complete run. All models share a fixed “S” point for easy comparison among them. Sed. rate = 0.1 mm/y 25 Ma 10 Ma Sed. rate = 0.9 mm/y Maximum basin width of standard model 60% 80% 10 Ma 5 Ma 20% 40% 60% 80% 40% 15 Ma 60% 20% 20% 40% 60% 20 Ma 20% 20% 20% Maximum basin depth of standard model 40% 80% 20% Maximum basin depth of standard model 20% 20% 40% 60% 80% Standard Model Maximum basin width of standard model 20% 20% Maximum basin depth of standard model 15 Ma 0 Ma 80% Maximum basin width of standard model 20 Ma 5 Ma 20% 40% 60% 0 Ma Figure 6.8: Forearc evolution with changes in sedimentation rate 80% 152 6.4.3 Chapter 6 Suite two - viscosity parameters The evolution of topography in geologic systems is clearly influenced by the viscosity, and hence the ductile behavior, of the crust. In accretionary margins this might lead to the development of backward thrusting sequences, viscous flow and the formation of forearc ridges [e.g. Pavlis and Bruhn, 1983; Smit et al., 2003; Fuller, 1996]. Here, we want to expand upon these studies to explore the effects that variations of specific viscous parameters (Aµ , nµ , and Q) have in the development of accretionary margins. The relationship between stress and the rate of deformation is modeled by the power-law viscosity equation: −Q nµ , (6.1) Dij = Aµ σij exp RT where Aµ and nµ are constants dependent on the material, Q is the activation energy, R is the molar gas constant, and T is the temperature. Table 6.3 shows the explored range of values for the aforementioned parameters. Table 6.3: Viscosity parameterization: Aµ , nµ , and Q. For every model set (left column), the parameter being modified is marked with italic bold characters. Model Standard A1-to-A7 n1-to-n7 Q1-to-Q7 Power-law viscosity Aµ (Pa−n /M a) nµ 2.05 0.05-4.15 (∆0.7) 2.05 2.05 2.65 2.65 2.35-2.95 (∆0.1) 2.65 Activation Energy Q (J/mol) 1.7 1.7 1.7 1.4-2.0 (∆0.1) The pre-exponential viscosity parameter, Aµ , affects the amount of relative differential motion between the subsided and the uplifted areas, and the age of uplift. Within the parameter range explored, lower values produce a deeper initial forearc basin and slow down the formation of a subsiding high that develops at younger ages than the standard. This occurs only weakly and is unable to compensate for the primary subsidence signal (Fig. 6.9). Subsidence in the initial forearc basin decreases with time until formation of the structural high, which leads to accelerating subsidence in the pro-basin and decelerating in the retro-basin. Higher values of Aµ led to a more pronounced forearc high uplift, which develops at earlier stages than the standard and more widespread stresses, i.e., broader shear zones. The formation of the forearc high outpaces the primary subsidence signal, and surface uplift in this area is eventually larger in absolute value than the subsidence undergone by the initial forearc basin (Fig. 6.9). The rate of the motions after forearc high development accelerates with time for both the forearc high and the pro-basin, leading to differential motions that become more relevant with time; whereas the rate of subsidence in the retro-basin remains roughly constant (Fig. 6.9). Similar to Aµ , nµ exerts control on the relative differential motions and their age and on the diffusion of deformation. Lower values of nµ within the parameter spectrum studied, led to the formation and protracted development of the initial forearc basin, which thickens while decelerating its subsidence rate very slowly (Fig. 6.9). 153 CAP Model The continuous development of the initial forearc basin makes the formation of the forearc high difficult since it cannot outpace the overall subsidence. 25 Ma 10 Ma Viscosity - A = 0.75 Pa-n/Ma Viscosity - A = 3.45 Pa-n/Ma Maximum basin width of standard model Maximum basin width of standard model 20% 20% 80% 5 Ma 20% 40% 60% 80% 20% 20% 40% 60% 20% 20% 40% 60% 10 Ma 80% 60% 15 Ma Maximum basin depth of standard model 40% 20 Ma 80% 20% Maximum basin depth of standard model 40% 20% 20% 20% 60% 80% 15 Ma 0 Ma Standard Model Maximum basin width of standard model Maximum basin depth of standard model 20 Ma 5 Ma 20% 40% 60% 80% 0 Ma (a) Variations in viscosity parameter Aµ Figure 6.9: Forearc evolution with changes in viscosity parameter Aµ . Viscosity - n = 2.85 25 Ma 10 Ma Maximum basin width of standard model Maximum basin width of standard model 60% 80% 15 Ma 10 Ma 5 Ma 20% 40% 60% 80% 20% 20% 40% 20 Ma 60% 60% 40% 20% 20% 20% Maximum basin depth of standard model 40% 80% 20% Maximum basin depth of standard model 60% 40% 20% 20% 20% 80% 20% 80% 15 Ma 0 Ma Standard Model Viscosity - n = 2.45 Maximum basin width of standard model Maximum basin depth of standard model 20 Ma 5 Ma 20% 40% 60% 80% 0 Ma (b) Variations in viscosity parameter nµ Figure 6.9: Forearc evolution with changes in viscosity parameter nµ . 154 Chapter 6 Higher values produce an early uplift that is older than that of the standard. Soon after initiation of the model the forearc high forms, fragmenting the initial forearc basin, which had barely started to form. Uplift rates accelerate with time in the forearc high and its width remains roughly constant. The subsidence rates in the pro-side basin accelerate first but slow down at more advance stages (younger times); whereas the opposite trend in subsidence rates is recorded in the retro-basin, with older deceleration and younger acceleration (Fig. 6.9). Within the parameter spectrum checked (Table 6.3), the activation energy, Q, controls the occurrence and shape of uplift. Its effects can broadly be considered reverse to those seen for nµ . Relatively low values produce an early-staged, older, and larger (than standard) uplift of a convex up area, blocking the development of the large initial forearc basin, while inducing a second uplift in the retro-side. Surface uplift in the forearc high outpaces the primary subsidence and uplift rates accelerate with time. Subsidence rates also accelerate with time for both the proand retro-basins. The width of the pro-basin decreases with time while that of the retro-basin increases until stabilization (Fig. 6.9). High values accelerate the subsidence rates in the initial forearc basin, as the basin becomes slightly narrower (Fig. 6.9). High values lessen the strain distribution and thus increase the stability of the margin, which counteracts the upward motion of a box-shaped area that takes place in later stages of the simulations. Viscosity - Q = 1.6 J/mol 25 Ma 10 Ma 20 Ma 5 Ma 15 Ma 0 Ma Standard Model 80% 10 Ma 5 Ma 20% 40% 60% 80% 20% 20% 40% 60% 20% 20% 40% 60% 15 Ma 80% 60% 20 Ma Maximum basin depth of standard model 40% 80% 20% Maximum basin depth of standard model 20% 20% 40% 60% 80% 20% 20% 20% Maximum basin depth of standard model Viscosity - Q = 1.8 J/mol Maximum basin width of standard model Maximum basin width of standard model Maximum basin width of standard model 20% 40% 60% 80% 0 Ma (c) Variations in viscosity parameter Q Figure 6.9: Forearc evolution with changes in viscosity parameter Q. CAP Model 6.5 155 Discussion Accretionary margins have subduction wedges markedly compartmentalized, along the convergence direction as a result of strain distribution (Fig. 6.1). The overriding plate in subduction wedges undergo brittle deformation in their seaward regions, i.e., the trench, the trench-slope, and the trench-slope break areas, and remain mostly undeformed in their landward forearc basin region. Forearc highs, uplifted flat terrains developed farther inland, are classically thought to form in margins where large amounts of accreted material encounter a “backstop” barrier during their inland migration [e.g. Byrne et al., 1993]. Here we focus on the mechanics of uplift of one of these areas, the southern margin of the Central Anatolian Plateau, and on the relative impact of the possible parameters influencing its development. Our simulations demonstrate that the accretion and sedimentary thickening of the Cyprian forearc basin system promote thermal weakening and viscous deformation in the base of the crust that drive the subsequent surface uplift of the modern Central Taurus Mountains. Models clearly manifest that the viscous rheology of the overriding plate and the amount of sediments in the forearc basin control the type and timing of uplift in the forearc high without the need for a backstop. 6.5.1 A new view at uplift in South Turkey Post-8 Ma surface uplift of the southern margin of the Central Anatolia Plateau in south Turkey is attributed to slab break-off [e.g. Cipollari et al., 2012; Cosentino et al., 2012; Schildgen et al., 2014a], although tomography shows the intact Cyprus slab underlying the area with more pronounced uplift [Bakırcı et al., 2012], an area where regional-scale contractional structures developed during the time-frame of interest (chapter 5). The analysis of the structures observed along the regional 2D geologic transect of the central Cyprus arc demonstrate clear signs of contraction. The type, distribution, relative age, and geometry of these contractional structures are compatible with their development in relation to a wide accretionary subduction system [e.g. Karig and Sharman, 1975; Dickinson and Seely, 1979] (Fig. 6.1). Thus, in this contribution, we attribute the surface uplift of the southern margin of the Central Anatolia Plateau to shortening and the accretion of material from the Cyprus subduction system, a mechanism that is validated by our models (Fig. 6.10). Simulations can reproduce the overall geometry and the main contact relationships as well as the main vertical tectonic events of the south Anatolian upper plate, including the early continuity and later disruption of the Miocene basin(s) in the presence of regional-scale accommodating structures in the seaward side, and in their absence in the landward side of the system. A conceptual representation of this surface motion mechanism in the context of the tectonic evolution of the central Cyprus arc is shown in Fig. 6.10. Mechanical accretion is responsible for the structural highs and lows in the seaward side of the margin, where the development of young contractional structures, leading to ridges and related piggy-back basins, is clearly linked with the subduction system [Aksu et al., 2005a, b; Calon et al., 2005a, b; Hall et al., 2005a, b]. The trench in south Cyprus, the Troodos Ophiolite, and the Messaoria Basin belong to this trench of slope domain, a wedge-like region undergoing active accretion in which seaward imbricate thrusts systems develop. The Kyrenia Range is the trench-slope break, a 156 Chapter 6 100km V.e.=2 Miocene basin formation Middle Miocene Horizontral length ~ 650km Mechanical accretion Miocene basin deepening Upper Miocene Differential vertical motions Present-day Cyprus Cilicia Thermo-viscous accretion Taurus basin CAVP Central Anatolia ~45km Lower Miocene Domains fragmentation S Basement Miocene Accretion ~40km ~35km ~30km Maximum crustal thickness structural high shaped by sea- and landward verging thrusts, dividing the active seaward areas from the landward “stable” forearc basin. The direct effects of accretion, and the vertical motions attached to it, are less readable in the forearc basin area. Protracted accretion of material is nevertheless responsible for the thermal expansion and ductile deformation at the base of the crust, which results in surface uplift in the absence of ground-reaching faults (Fig. 6.10). Viscous deformation N Figure 6.10: Conceptual evolution of the Central Cyprus Arc This process led not only to the development of the uplift of the forearc high and growth of the modern Taurus Mountains, but also to the subsidence occurring in the intervening seaward area. The south Turkey surface uplift coupled with offshore subsidence divided the forearc basin into a landward uplifted forearc basin, the Mut Basin, and a seaward subsided forearc basin, the Cilicia Basin, while developing a Miocene monocline (chapter 5) at ∼8 Ma and later times [e.g. Cosentino et al., 2012]. 6.5.2 A new view at forearc highs Although accretionary subduction margins are very diverse (plate convergence rate, age of subducting oceanic crust, incoming sediment at the trench, dip of the subduction plate, etc.) even within the same system [e.g. Vannucchi et al., 2013], the upper-plate morphology all over these margins is remarkably similar. Forearc highs are often, but not always, seen in the internal parts of these margins. A common point of view is that areas of relatively larger strength in the interior of the forearc regions, backstops, induce a strength contrast that controls the structure and limits the growth of the wedge. Under this view, forearc high formation is driven CAP Model 157 by forced mechanical accretion against these larger-strength areas, which results in upward terrain growth. The presence of a strength contrast in the inland areas seems plausible given the variations in deformation history or metamorphic grade between the pre-existing rocks and those formed during the development of the forearc basin system. However, the apparent absence of backstop in some margins and the fact that the uplift of the forearc high takes place within and on the stronger region are indications that other mechanisms may also influence the development of forearc highs and that of the subduction wedge [Fuller, 1996]. Our models prove that the forearc high develops above the area where the friction exerted by the subducting plate does not affect the overriding plate. This singularity might indeed be induced by stronger material in the retro-side of the model, but also, and more simply, could be induced by the geometrical characteristics of the wedge itself, i.e., the shape of the subducting slab, which controls the point where the slab loses contact with the overriding plate. This was already postulated by Fuller et al. [2006a] as “trench-slope breaks and forearc basins can form as a natural consequence of the landward increase in slab dip in a margin with no lithologic strength contrasts.” Besides, in our simulations, the uplift of the forearc high is seen for common values of thermal conductivity, viscosity, and sedimentation rate, and becomes more probable as the system ages. Variation in the values of these parameters favors or complicates forearc high development without the need of a material-strength backstop. Our models show that even in the less favorable lowviscous rheologies forearc highs eventually arise. Therefore, forearc highs should no be considered an isolated feature conditioned to backstop presence, but rather an integral part that develops in accretionary wedges at earlier stages of evolution in the forearc basin systems if the conditions are favorable, and at advanced stages if they are not. Also, we understand that forearc highs grow by protracted accretion and do so in combination with feedback between the amount of sediments in the upperplate crust and the thermal blanketing of these forearc basin sediments. These factors lead to a thermally activated viscosity drop at the base of the crust and subsequent surface uplift. 6.5.3 A new view at vertical motions in forearc regions Our models of accretionary margins show that if the amount of incoming sediments at the trench remains constant, variations in the sedimentation rate control the age and occurrence of surface uplift in the forearc high regions. Assuming an overfilled forearc basin in which the sediments available from the surrounding source areas outpace the accommodation space available, the amount of sediments that could potentially enter the system depends solely on subsidence. Therefore, regional subsidence in the forearc basin could cause an uplift in the forearc high. We observe that two counteracting effects compete during the growth of the forearc basin system: the thermal restrictions caused by the sediments in the forearc basin system and their sedimentary load. Sediments have a lower thermal conductivity than basement rocks, and therefore lower heat transfer rates, which ultimately lead to changes in the thermal gradient underneath the forearc basin. As the sedimentary infill thickens, the temperature gradient becomes more abrupt, with increasingly larger temperatures at the base of the crust. Upward circulation is re- 158 Chapter 6 strained by the thermal resistivity or “blanketing” effect, caused by the forearc basin sediments. Increasing temperatures at the base of the crust lead to lower viscosities, and thus lower resistance to deformation, which in turn favors surface uplift. On the contrary, larger sedimentary loads lead to larger isostatic subsidence. Since both effects act simultaneously, the surface uplift induced by rheological changes at the crustal base, and the subsequent sharper differences in the thermal temperature column, might be partially or completely masked due to regional subsidence. These opposite controlling factors cause the region of the forearc basin to undergo a sequenced downward-upward motion in which surface uplift of the forearc high takes place while subsidence in the surrounding areas persist, at least in a relative frame. 6.6 Conclusions The models successfully reproduce upper plate deformation structures and modern sedimentary architectures of the central Cyprus arc and prove that thermally activated viscous deformation at base of the Anatolian crust, yielded by the accretion of incoming sediments of the Cyprus arc, is a plausible mechanism driving the surface uplift that raised the modern Taurus Mountains. Frictional-mechanical deformation occurring in the retro-side of the model controls the time sequence of structures and topographies/bathymetries seen in the trench, trench-of-slope and seaward areas of the forearc system. Viscous-thermal deformation dominates in the landward areas of the forearc system and controls the development of the forearc high and its shape. Continued accretion thickens the forearc basin, which results in increasing temperature at the base of the crust and subsequent weakened viscous flow of the deforming material. The time delay between the onset of the accretionary system and the moment of uplift change in relation to the velocity of accretion and the sediment input, as well as the rheological and thermal state of the wedge. Higher viscosity produces an older, more pronounced uplift. Decreasing values modify the uplift from an older rounded-shaped uplift, to a box-shaped younger uplift, to the absence of uplift. The margin stability increases when thicker basins develop in relation to either larger incoming sediment input at the trench or larger accumulation rates. This thickening, which shifts deformation toward the forearc basin margins, leads to younger uplift times and can force the avoidance of its occurrence, i.e., the subduction system needs longer times to develop the forearc high. We conclude that the topography and vertical tectonic motions seen in the Cyprian subduction margin developed as a consequence of mechanical accretion in Cyprus and related deep-seated deformation in south Turkey.