Mantle flow uplift of western Anatolia and the ... - Wiley Online Library

JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 117, B11412, doi:10.1029/2012JB009306, 2012

Mantle flow uplift of western Anatolia and the Aegean: Interpretations from geophysical analyses and geodynamic modeling Tolga Komut,1 Robert Gray,2 Russell Pysklywec,2 and Oğuz H. Göğüş3,4 Received 12 March 2012; revised 31 July 2012; accepted 26 September 2012; published 29 November 2012.

[1] The Western Anatolian and Aegean region demonstrates a complex geologic history of horizontal and vertical tectonics. Active normal faulting and exhumation zones indicate that Western Anatolia has experienced significant extension since the Oligocene-Early Miocene (30 Ma). Our geophysical analyses demonstrate that the region is also uplifted relative to an elevation that would be expected given an isostatic response to the lithosphere structure. Namely, topography “residuals” indicate a residual uplift of about 1500 m over 200 km sections of Western Anatolia and the Aegean. Admittance functions between free-air gravity and topography indicate that the regional topography is isostatically uncompensated and as it approaches 50 mGal/km at the longest wavelengths, the uncompensated topography is likely owing to an underlying mantle flow component. Using forward geodynamic modelling we consider an idealized section of Western Anatolian lithosphere based on tomographic inversions and examine the magnitude and pattern of surface topography to reconcile with the geophysical observables. The models consistently show a plateau-type uplift (and horizontal extension) through Western Anatolia with an amplitude and wavelength consistent with the residual topography calculations. Together, the geophysical analyses and modelling provide independent quantitative evidence that the thin Anatolian-Aegean lithosphere is being buoyed upwards by underlying mantle flow. The mantle flow may be associated with active lithosphere delamination beneath the region; a process that would also explain the ongoing crustal extension. Citation: Komut, T., R. Gray, R. Pysklywec, and O. H. Göğüş (2012), Mantle flow uplift of western Anatolia and the Aegean: Interpretations from geophysical analyses and geodynamic modeling, J. Geophys. Res., 117, B11412, doi:10.1029/2012JB009306.

1. Introduction [2] The Aegean-Anatolian region is an active region of rather complex horizontal and vertical tectonics. We distinguish these such that horizontal tectonics involve the lateral tectonic motion of the lithosphere (e.g., resulting in contraction/extension of the lithosphere), whereas vertical tectonics relates to the large-scale uplift/subsidence of the lithosphere. To first order, the Aegean region is a back-arc

1 Department of Geophysics, Faculty of Engineering, Çanakkale Onsekiz Mart University, Çanakkale, Turkey. 2 Department of Earth Sciences, University of Toronto, Toronto, Ontario, Canada. 3 Institute of Geophysics and Tectonics, School of Earth and Environment, University of Leeds, Leeds, UK. 4 Also at Department of Geology, Çanakkale Onsekiz Mart University, Çanakkale, Turkey.

Corresponding author: T. Komut, Department of Geophysics, Faculty of Engineering, Çanakkale Onsekiz Mart University, Terzioglu Campus, 17100 Çanakkale, Turkey. ([email protected]) ©2012. American Geophysical Union. All Rights Reserved. 0148-0227/12/2012JB009306

area, characterized by extensional tectonics associated with the subducting slab along the south Aegean tectonic boundary between the African plate and Anatolia (Figure 1). This extensional motion is accommodated by active seismicity and normal faulting throughout the Aegean. The extension also affects Western Anatolia where NE to E-W trending basins bounded by normal faults dominate the neotectonics (Figure 1). Velocity vectors derived from GPS surveys show westerly directed motion in central Turkey becoming progressively more W-WSW-directed to the west, with respect to the Eurasian plate [Oral et al., 1995; Reilinger et al., 2006, 2010]. According to geodetic evidence [Oral et al., 1995], extension in Western Anatolia grabens decreases from west to east and increases from north to south. [3] Active normal faulting and exhumation zones indicate that Western Anatolia has experienced significant extension since the late Oligocene-Early Miocene (30 Ma) [e.g., Jolivet and Faccenna, 2000; Okay and Satır, 2000; Bozkurt and Sozbilir, 2004; Işık, 2004]. The dynamics of the extensional tectonics in the Aegean region are not well understood. McKenzie [1978] suggested that three main forces drive the extension: (1) forces on the boundaries created by

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KOMUT ET AL.: UPLIFT OF WESTERN ANATOLIA

Figure 1. Regional map showing primary tectonic features. WA- Western Anatolia; SM- Marmara Sea; LB- Levantine Basin; Aeg- Aegean; Pe- Peleponesia; NAF- North Anatolian Fault; SF-Dead Sea Fault; EAF- East Anatolian Fault.

the collision (escape tectonics) of the African and Arabian plates with the Eurasian plate [Dewey and Şengör, 1979; Şengör and Yılmaz, 1981]; (2) regional gravitational forcing, maintained by high topography in eastern Anatolia or postcollisional crustal over thickening in Anatolia [e.g., Dewey, 1988; Seyitoğlu and Scott, 1991]; and (3) the forces at the base of the lithosphere resulting from the peel-back (or rollback) of the subducted African slab [e.g., Le Pichon and Angelier, 1979; Meulenkamp et al., 1988; Royden, 1993]. With respect to Eurasia the accelerated motion of the Aegean region, west of Anatolia, towards the Ionian and Levantine basins [Oral et al., 1995; Le Pichon et al., 1995; McClusky et al., 2000; Reilinger et al., 2010] cannot be explained by escape tectonics. On the other hand Özelçi [1973] proposes that upwelling mantle may be the cause of the observed extension in the Aegean region. Özelçi [1973] suggests that Anatolia is undercompensated according to Bouguer gravity anomalies where, for example, there are high positive (up to 100 mGal) Bouguer gravity anomalies on shore in Western Anatolia including the Marmara region. There is also geomorphologic and geological evidence for regional uplift: incised meanders since at least the Pliocene [Arpat and Bingöl, 1969; Ardos, 1995; Özelçi, 1973] and raised marine and alluvial terraces in Northwestern Anatolia [Komut and Kapan-Yeşilyurt, 2010]. Westaway et al. [2004] investigated the surface uplift along the Gediz graben river terraces and suggested that 400 meters (m) of uplift occurred in the last 3 Myrs in Western Anatolia. The evidence is not restricted to local areas. It is very common throughout Western Anatolia including the Biga peninsula and central Anatolia, suggesting long-wavelength uplift of Western Anatolia and the surrounding region.

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[4] There are various independent data sets pointing to the existence of hot upper mantle in the study area. Substantial thermal activity is revealed by numerous hot springs, fumaroles, hydrothermal alterations and recent mineralizations [Göktürkler et al., 2003]. In addition, surface heat flow in Western Anatolia is high with respect to the surrounding area [Tezcan, 1995]: the mean surface heat flow is 107 45 mWm2 [Ilkişik, 1995] and the Curie Point depths in the area are very shallow (about 10 km) [Bilim, 2007]. A heat flow map [Akın et al., 2006] derived from magnetic data also shows high values in Western Anatolia. Young basaltic volcanism of Plio-Quaternary age is extensive over Anatolia [Çoban, 2007], this is interpreted to mean that sub-continental mantle temperatures are much higher (1300–1450 C) than normal. Several geochemical studies suggest that the apparent lack of mantle lithosphere beneath Western Anatolia is responsible for widespread volcanism in the region [Aldanmaz et al., 2000; Altunkaynak, 2007; Dilek and Altunkaynak, 2009; Ersoy et al., 2010]. Helium-carbon relationships in geothermal fluids also suggest hot upper mantle underlying the crust [Mutlu et al., 2008] in the study area. In agreement with the body of surface evidence for elevated temperatures beneath Western Anatolia and the possible presence of underlying hot mantle, P-wave tomography data [van Hinsbergen et al., 2010; Amaru, 2007; Biryol et al., 2011; Kömeç Mutlu and Karabulut, 2011] shows a significant slow velocity anomaly beneath Western Anatolia that lies between the (interpreted) subducted African slab and Western Anatolian crust. [5] Stretching lineations in core complexes of the overriding crust of Western Anatolia are oriented approximately NNE-SSW [Bozkurt and Oberhänsli, 2001; Jolivet et al., 2004; Işık, 2004]. SKS splitting directions are also directed approximately NNE-SSW in the area [Polat et al., 2006; Jolivet et al., 2009a, 2009b] and are parallel to the directions of active extension derived from differential GPS velocities [Kreemer et al., 2004]. These data are consistent with the possibility that mantle upwelling and divergence may be in part responsible for the extensional tectonics in the Western Anatolian region. We propose that the vertical forcing associated with rising mantle is an important factor of the observed uplift working in tandem with horizontal tectonic advection. Namely, actively rising mantle may be helping support anomalous (non-isostatic) uplift of the surface as “dynamic topography”. [6] We explore a range of geological and geophysical observables to determine an anomalous component of surface topography in Western Anatolia. This includes admittance functions of free-air gravity anomalies and topography [D’Agostino and McKenzie, 1999] to demonstrate whether there is a component of uncompensated topography at long wavelengths. In addition, various “residual” topography determinations are made using isostatic rules on several available datasets of topography and crustal structure. The observables are interpreted with forward numerical geodynamic modeling where lithospheric tectonics is driven by the mantle structure imaged from seismic tomography. Analyses of these independent data components will provide insight into the potential role of the mantle in driving the vertical tectonics of the region. In the larger context of the geodynamics of the eastern Mediterranean, the work will

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Figure 2. (a) Topography of Western Anatolia and surrounding region. (b) Crustal thickness of Western Anatolia and surrounding region derived mostly from available receiver function analysis [Tezel et al., 2010; Zhu et al., 2006; van der Meijde et al., 2003; Sodoudi et al., 2006; Zor et al., 2006]. For the Marmara area [Bécel et al., 2009] and Southern part of Aegean Sea [Makris, 1976], deep seismic data provided additional information. The Crust 2.0 global model data was used for the rest. (c) Residual topography map of Western Anatolia and surrounding region. Interpolated crustal thickness values re-sampled and topography calculated using simple isostasy rules. The map shows the difference between observed and this calculated topography. Solid grey line shows the location of the residual topography plots in Figure 2d, the P-wave tomography section (Figure 6a) and the section along which the results of the forward numerical geodynamic models apply. (d) Residual topography plots along the solid grey line in Figure 2c. demonstrate the potential significance (or insignificance) of mantle forcing on regional surface tectonics.

2. Uplift of Western Anatolia 2.1. Residual Topography [7] Residual topography is the difference between the observed topography and the expected topography resulting from isostatic compensation of the lithosphere - i.e., the uncompensated component of the topography where positive and negative deviations indicate undercompensation and overcompensation conditions, respectively. We calculated the residual topography of Western Anatolia and the surrounding region from two separate sources: (1) the CRUST2.0 database; and (2) available receiver function data and controlled source deep seismic data. Receiver function data across Anatolia and the Aegean were used to estimate

crustal thickness across most of the region [Tezel et al., 2010; Zhu et al., 2006; van der Meijde et al., 2003; Sodoudi et al., 2006; Zor et al., 2006]. For the Marmara area [Bécel et al., 2009] and the Southern part of the Aegean Sea [Makris, 1976], deep seismic data provided additional information. The CRUST2.0 global model data [Bassin et al., 2000; G. Laske et al., CRUST 2.0: A new global crustal model at 2 2 degrees, 2001, http://igppweb.ucsd. edu/gabi/rem.html], a compilation of crustal thickness derived from surface wave data, free-oscillation data, body wave travel times and a detailed compilation of ice and sediment thickness data, provided broad but lower resolution estimates of crustal thickness. [8] Figure 2b shows an interpolated crustal thickness map derived from the seismic data and the CRUST2.0 global model data. The receiver function estimates were used where available and the deep seismic data and CRUST2.0 data

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Figure 3. Residual topography of Western Anatolia and surrounding region calculated assuming oceanic crustal density of 2900 kg/m3 for oceanic crust in the Mediterranean and the Western Black Sea and continental crustal density of 2800 kg/m3 for continental landmass of Western Anatolia. filled in data-absent areas (viz., for the latter, North of Thrace and Mediterranean Sea). Western Anatolia shows thin crust (down to 24 km thick) in the westernmost areas with thickening to 35 km to the east. The Aegean Sea is characterized by continental-type crust according to deep seismic soundings of Makris [1976] and the average crustal thickness of the Aegean Sea including coastal Western Anatolia is 27 km (Figure 2b). The crustal thickness gradually increases to the north, reaching up to 45 km at 46 N latitude on the Eurasian plate. [9] Using the calculated crustal thickness values (Figure 2b) and regional topography data from the ETOPO2 Global Relief Database (Figure 2a), residual topography was calculated and mapped (Figure 2c). For these calculations, an estimate of crustal density is required. We assume a crustal density of 2800 kg/m3 and a mantle density of 3280 kg/m3 for the entire map area to calculate the residual topography. Our assumption is a reasonable representative for the majority of the study area—particularly Western Anatolia and the Aegean—and the simplified density allows for external repeatability of our calculations and transparent interpretation of the residual calculations. We tested the sensitivity of the calculated residual topography to spatial variations of crustal composition/density (e.g., oceanic crust for the Mediterranean in the south and continental crust in the continental landmass of Western Anatolia). Our sensitivity analysis shows that incorporating spatial variations of

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density in the residual topography calculations does not affect the strong positive residual topography signal for Western Anatolia and the strong negative residual topography signal for the more southern regions and the northern portion of the map area (Figure 3). This supports our assumption of an average crustal density of 2800 kg/m3 for the entire map area. The calculated residual topography data was scaled to be a deviation from the mean of the dataset (i.e., the mean of the residual topography dataset was subtracted from the data). Essentially, the data show all of Western Anatolia as being anomalously high. The positive topography residuals reach up to +2 km and the uplifted zone extends over 600 km in width. Surrounding areas to the north and south of the uplifted Aegean/Anatolian region are characterized by low residual topography, indicating isostatic compensation of observed topography. The deep negative in the Eastern Mediterranean may be related to a subduction-induced low; this feature wraps around the entire Aegean subduction system and up towards the Adriatic-Apennines [Shaw and Pysklywec, 2007]. The deep negative residual topography in the northern boundary of the map area (1500 m) may be due to the deep passive margin along the north-western shore of the Black Sea. [10] Profiles of residual topography along a north-south line through Western Anatolia (Figure 2d) are extracted from the data sets. The receiver function profile highlights the anomalous (isostatically undercompensated) plateau-like uplift across Western Anatolia. This uplift is more muted in the CRUST2.0 data. We attribute this to the larger crustal thickness values of Western Anatolia in the CRUST2.0 database (Figure 4) and the somewhat coarse 2 degree 2 degree resolution of the CRUST2.0 data in a regional study like this one. With this low resolution, moderate wavelength variations in crustal thickness–e.g., near this boundary between African ocean plate and Anatolian continental plate in the eastern Mediterranean–can “blur” features of dynamic topography. In both profiles there are adjacent residual topography lows to the north and south of Western Anatolia. 2.2. Gravitational Admittance [11] Spectral analysis of the relationship between free-air gravity anomalies and topography can be used to place constraints on the mechanism(s) of support of topographic loads [e.g., McKenzie and Fairhead, 1997; McKenzie, 1994; D’Agostino and McKenzie, 1999; Simons et al., 2000]. These studies suggest that short-wavelength topography (150 km) and free-air gravity anomalies may be supported by isostasy, assuming a lithosphere with no elastic strength or zero thickness, or convective circulation within the mantle. To distinguish between the roles of isostasy and mantle circulation in supporting the long-wavelength topography characteristic of Western Anatolia, we calculate the admittance between free-air anomalies and topography. The free-air admittance is defined as Z(k) = g/t, where k is the magnitude of the wavenumber, g is the two-dimensional Fourier transform of the free-air anomalies and t is the Fourier transform of topography [Dorman and Lewis, 1970, Forsyth, 1985]. In order to minimize geologic noise we calculate Z(k) by employing the multitaper method [Thomson,

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Figure 4. (a) Crustal thickness of Western Anatolia and surrounding region derived mostly from available receiver function analysis [Tezel et al., 2010; Zhu et al., 2006; van der Meijde et al., 2003; Sodoudi et al., 2006; Zor et al., 2006]. For the Marmara area [Bécel et al., 2009] and Southern part of Aegean Sea [Makris, 1976], deep seismic data provided additional information. The Crust 2.0 global model data was used for the rest. (b) Crustal thickness of Western Anatolia and surrounding region derived exclusively from the CRUST2.0 database.

1982]. Using this method, Z(k) = ShDg(k)/Shh(k), where ShDg (k) is the estimated cross-spectral density function of the free-air anomalies and topography and Shh(k) is the estimated power-spectral density function of topography. The multitaper technique provides an optimal spectrum estimate, S, by reducing the estimation variance of the spectrum. This is achieved by calculating S as a weighted average over a number k of independent direct spectral estimates Sl with weights ll [Simons et al., 2000]. For a more detailed explanation of the multitaper method and its application in calculating the free-air gravitational admittance the reader is referred to Simons et al. [2000] and McKenzie and Fairhead [1997]. [12] To study the long-wavelength gravity field of Western Anatolia we obtained the free-air gravity data from the Earth Gravitational Model EGM2008 [Pavlis et al., 2008]. This gravitational model is complete to spherical harmonic degree and order 2159. Topography data were taken from the 30 arc sec (1 km) digital elevation model SRTM30 [Becker et al., 2009]. Both datasets were projected onto a Cartesian grid and re-sampled at the same rate. In order to detect the gravitational contribution from largescale mantle processes no background gravity field was subtracted from the EGM2008 gravitational model. The free-air gravity and topography data were transformed using the multitaper method with NW = 4. Conventionally, only

the first 2NW – 1 eigentapers are used [Simons et al., 2000]. We use the same approach in this contribution. [13] Figure 5a shows the two-dimensional free-air admittance calculated for the red box shown in Figure 1. For wavenumbers between 0.04 and 0.12 2p/km, Z(k) for Western Anatolia increases steadily from 50 to 95 mGal/km. This suggests that for wavelengths smaller than 150 km topography is isostatically uncompensated (i.e., supported by lithosphere flexure), given that the expected value for topographic loads supported by isostatic compensation is zero. Figure 5b illustrates that the coherence, equivalent to correlation in the spectral domain, for wavenumbers greater than 0.04 2p/km is between 0.7 and 0.8, showing that most of the short-wavelength free-air gravity signal is correlated with topography. At shorter wavenumbers the admittance does not approach zero (i.e., the expected value for topographic loads supported by isostatic compensation). Rather, for wavenumbers between 0.04 and 0.01 2p/km the admittance approaches 50 mGal/km. Long-wavelength free-air gravity anomalies supported by isostasy produce a Z(k) = 0 mGal/km, whereas those supported by convective circulation in the mantle produce Z(k) =50 mGal/km [McKenzie and Fairhead, 1997]. Consequently, the admittance observed at long wavelengths is not consistent with an admittance predicted for isostatic support. As a result, the free-air admittance values for Western Anatolia at long wavelengths in Figure 5a suggest

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Figure 5. (a) Gravitational free-air admittance and (b) coherence versus wavenumber for the region shown in Figure 1. that the long-wavelength topography characteristic of the region is supported by mantle convection.

3. Geodynamic Models for Western Anatolia [14] Using forward numerical geodynamic modeling we explore whether the observed anomalous uplift of Western Anatolia is consistent with predicted mantle flow beneath the region. For the highest resolution mantle structure available, we use P-wave teleseismic tomography data from Amaru [2007] and van Hinsbergen et al. [2010]. Their P-wave velocity anomaly cross-sections cross north-south across Western Anatolia. The locations of the residual topography profiles (Figure 2d) were chosen to coincide with these seismic sections. Figure 6a shows the P wave velocity anomalies digitized from the Amaru [2007] and van

Hinsbergen et al. [2010] tomography inversions. Taking a node spacing of 200 100 from the digitized data, we converted the P-wave velocity anomalies to density variations using the velocity/density scaling function from Karato [1993]. Specifically, the (∂ln r/∂ln vp) vs. depth scaling relation of Karato’s [1993] Figure 2 that can be solved for r with given vp in the depth-dependent relation. Thermal anomalies were derived from the density variations using thermal expansivity: r (T) = r0(1 a(DT)), where ro is the reference density, a is the coefficient of thermal expansivity, and DT is temperature variation (material physical parameters used in the numerical models are detailed in Table 1). The DT data were superimposed on a background temperature field of 1623 K (chosen to result in temperatures at the base of the lithosphere of 1600 K) to produce the full temperature field (Figure 6b). This represents an approximation

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Figure 6. (a) Seismic tomography profile, after van Hinsbergen et al. [2010] and Amaru [2007]. The position of the P-wave tomography section is along the solid grey line of Figure 2c. (b) Set-up of reference experiment A1 and temperature variations derived from the P-wave tomography section (Figure 2). In the forward numerical geodynamic experiments the reference density of the 30 km thick wet quartzite crust (rc) is 2840 kg/m3, 120 km thick wet olivine mantle lithosphere (rm) is 3300 kg/m3 and that of the 1250 km thick sub-lithospheric mantle (asthenosphere) (ra) is 3300 kg/m3. The initial temperature conditions of: 1) the surface of the computational domain is 25 C; 2) the Moho is 550 C; and 3) the base of the computational domain is 1350 C. Velocity boundary conditions are free-slip on the side-walls and no slip on the bottom. The top of the computational domain is a free surface to allow for the development of topography. No surface erosion/deposition is imposed at the surface, although these may have the potential for influencing the longer term tectonic evolution of the lithosphere [Pysklywec, 2006; Gray and Pysklywec, 2012b]. The computational domain is 1400 km x 1400 km with a numerical resolution of 201 nodes (horizontal) 101 nodes (vertical). to what may be considered “realistic” temperatures in all parts of the lower mantle (i.e., yielding zones of relatively low temperature in certain areas of the lower mantle). However, our specific test runs indicate that changes in the absolute value of temperature in the lower mantle has little influence on the resulting surface topography. Rather it is the

thermal anomalies and temperatures in the upper mantle, for which the temperatures in the model have been specifically “tuned” that govern the dynamic topography. Furthermore, by applying a uniform background temperature, the model results are easily reproduced.

Table 1. Physical Parameters for Reference Experiment A1a Mechanical Parameters

Crust

Mantle Lithosphere

Sub-lithospheric Mantle

Ref. Density r0 (kg/m ) Ref. Density T0 (K) f1 (degrees) f2 (degrees) (I′2)1/2 (1) (I′2)1/2 (2) Cohesion C0 (MPa) Flow Law A (Pa-ns1) N Q (kJ/mol) Heat capacity cp (J/kg/K) Thermal conductivity k (W/m/K) Thermal expansivity a (K1) Radioactive heat production H (W/m3)

2800 293 15 15 1.5 0.5 1 wet quartzite (1) 5.04 1028 4 223 793 2.25 2.0 105 0

3300 293 15 12 1.5 0.5 1 wet olivine (2) 4.89 1015 3.5 515 793 2.25 2.0 105 0

3300 293 15 12 1.5 0.5 1 wet olivine (2) 4.89 1015 3.5 515 793 2.25 2.0 105 0

3

a

Flow laws: (1) Gleason and Tullis [1995]; (2) Hirth and Kohlstedt [1996].

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[15] The most prominent interpreted feature on the temperature profile includes the presence of a hot region directly beneath Western Anatolia. This is consistent with other seismic evidence for absent mantle lithosphere or at least partial replacement of mantle lithosphere by hot mantle [Al-Lazki et al., 2004; Genç and Yürür, 2010]. Below the hot region is a cold body which may represent a fragment of the subducting slab associated with the Africa-Eurasia collision [van Hinsbergen et al., 2010]. [16] The calculated thermal anomalies beneath Western Anatolia were used as input for a forward numerical geodynamic model (i.e., seismically-derived data for the structure of the mantle beneath Western Anatolia). Though seismic velocity variations depend on composition and temperature, Schmid et al. [2006] showed that compressional wave velocity anomalies beneath the Mediterranean are dominated by shear modulus heterogeneity and that the ratio between bulk sound and shear wave velocity heterogeneity is indicative of a thermal origin. Consequently, though our simple conversion from P-wave velocity anomaly to temperature ignores compositional effects, the conversion is consistent with regional interpretation. [17] The forward numerical geodynamic models use the numerical code SOPALE. SOPALE solves the coupled conservation equations of mass, momentum and internal energy: r • ðruÞ ¼ 0;

ð1Þ

r • sij þ rg ¼ 0;

ð2Þ

rcp ð∂T=∂t þ u • rTÞ ¼ kr2 T þ rH;

ð3Þ

that govern the behavior of incompressible, plane-strain viscous-plastic media using arbitrary Lagrangian-Eulerian finite element techniques [Fullsack, 1995]. This system of equations is completed by an associated linearized equation of state: r ¼ ro ½1 a ðT To Þ:

ð4Þ

[18] In equations (1)–(4), r, u and T represent the density, velocity and temperature fields, respectively. The variables g, cp, a, k, H and t are the acceleration due to gravity, specific heat capacity, thermal expansivity, thermal conductivity, rate of internal heat production per unit mass and time, respectively. In the experiments presented in this study, flow in the model is driven by the internal thermal/ density heterogeneities only (e.g., there is no imposed plate convergence). In the numerical experiments the velocity and temperature fields are allowed to evolve self-consistently with time for 63,000 years (100 time-steps 630 years/ time-step). The temperature data were superimposed on an idealized model of the crust-mantle system where the top 30 km of the model was defined as having crust-like rheological behavior and a thermal gradient defined by a simple linear geotherm from 293 K at the surface to 773 K at the Moho; the rest of the model was designated as mantle material with temperatures defined by the seismicallydetermined thermal anomalies on the background of 1623 K (Figure 6b). In the model, viscous deformation is governed

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by non-Newtonian power law creep where the effective viscosity heff is a function of the second invariant of the deviatoric strain rate tensor, İ′2, and the temperature T: ð1=n–1Þ ðQ=nRTÞ heff ðİ′2 ; TÞ ¼ A1=n İ2′ e

[19] We used A, n, and Q (viscosity parameter, power exponent, and activation energy) based on the wet-quartzite of Gleason and Tullis [1995] and the wet-olivine of Hirth and Kohlstedt [1996] for the crust and mantle, respectively (Table 1). A two orders of magnitude viscosity increase is imposed at 660 km depth in accordance with Mitrovica and Forte [1997]. In addition to the viscous rheology, the model regions can yield according to a Mohr-Coulomb criterion: sy ¼ psinФ þ Co;

ð5Þ

where, sy, p, F and Co are the yield stress, dynamic pressure, angle of internal friction and cohesion, respectively. In the viscoplastic numerical model the deviatoric stress is determined at each computational node as the lesser value of either the yield stress sy or viscous stress sv, where sv = 2heffİ′2 . For the crust, the angle of internal friction is F = 15 and the cohesion Co = 1 MPa. For the mantle, F decreases linearly from 15 (F1) to 12 (F2) over a strain interval of 0.5 to 1.5 and Co = 1 MPa (see Pysklywec et al. [2000, 2002] for a more detailed description of the numerical and/or rheological formulation of SOPALE). [20] Figure 7 shows the mantle dynamics of Experiment A1. The viscosity plot (Figure 7c) illustrates the result of the rheological formulation: a strong crust and upper mantle lithosphere (that thins beneath the centre of the section) underlain by a weak upper mantle and stronger lower mantle. The velocity field (Figure 7a) indicates upwelling flow in the uppermost mantle beneath the centre of the profile. Beneath this region there is downwelling associated with the fast/cold anomaly at depth. The strain rate field (Figure 7b) shows active upper mantle flow with circulation cells of wavelengths 400–600 km. There is no localized deformation and low strain rate in the crust and uppermost mantle owing to its high strength. [21] The plotted dynamic topography for the different models represents a snapshot of the dynamic surface deflections associated with the computed underlying mantle flow (Figure 7d). Experiment A1 shows a long wavelength uplift across the centre of the model with amplitudes reaching +1.5 km. This reflects the response of the lithosphere to underlying upwelling flow; since the hot rising mantle anomaly is near the surface of the model, the surface of the model is affected more by the hot rising mantle anomaly than by any downwards forcing associated with the deeper cold mantle anomaly. Dynamic topography lows develop on either side of the uplift region as downgoing flow influences topography in these regions. Figure 7d plots the scaled horizontal component of surface strain rate. A positive value (+1 in this scaled calculation) indicates extension whereas a negative value (1) represents contraction. We chose to scale the surface strain rate to +1 (extension) and 1 (contraction) in order to simplify the surface strain rate data calculated in the numerical experiments and the identification of zones of extension or contraction at the surface of the

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Figure 7. For experiment A1: (a) temperature field and flow velocity vectors; (b) second invariant of the deviatoric strain rate tensor; (c) viscosity; and (d) surface topography and surface strain rate (İxx). numerical experiments. In experiment A1, a band of extension develops in association with the upwelling mantle material and surface uplift. This is bounded by zones of contraction to the sides of the uplift/extension zone. [22] We conducted a series of alternate experiments (Table 2) to consider how some of the primary modeling parameters control the resulting topography. Experiment A2 (Figure 8a) shows the result of decreasing the viscous rheology of the quartzite crust by a factor of 10. A similar pattern of central rising flow and adjacent downwelling develops and this supports a topography of similar form and magnitude to experiment A1 where broad uplift of 1.5 km is surrounded by topography lows. Again the uplift is characterized by surface extension and bounding contraction. In experiment A3 the (upper and lower) mantle viscosity is decreased by a factor 100 in comparison to A1. Again, a similar flow pattern arises, inducing a broad plateau-like uplift with adjacent low topography and extension across the plateau (Figure 8b). [23] In experiment A4, a “break” is introduced into the lithosphere to model the location of the Africa-Eurasia plate boundary (Figure 8c). The break is simulated by a weak zone 300 km away from the right margin of the box, 30 km

wide, located between the surface and 150 km depth at an angle of 60 from the horizontal and with a viscosity of 5 1019 Pa s. While this is a very crude approximation, the weak zone is effective in structurally decoupling the plates as an end-member simplification of a tectonic plate boundary. The lithospheric break manifests in the resulting topography as a local “trench”, but otherwise the signal is similar to previous models with broad uplift over the central hot mantle region. The surface strain rate plot (Figure 8c) shows very localized contraction right at the model trench; Table 2. List of Numerical Experiments Experiment

Description: Variation From Reference Model

A1 A2

reference model viscous rheology of wet quartzite crust decreased by factor 10 viscous rheology of wet olivine upper and lower mantle decreased by factor 100 introduction of weak zone (h = 5.0 1019 Pa s) to represent Africa-Eurasia boundary extension of seismically-derived thermal anomalies to surface of model

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4. Conclusions [25] The independent data sets and forward numerical geodynamic modeling are consistent with the interpretation that Western Anatolia is anomalously uplifted and that this uplift may be a consequence of underlying rising hot mantle material. Residual topography calculations from separate data indicate that the region is anomalously uplifted by over a kilometer. The crustal thickness is not sufficient to isostatically support the surface topography, indicating the area is undercompensated. Perhaps most significantly, the gravitational free-air admittance for Western Anatolia indicates clearly that there is an important component of uncompensated topography at wavelengths greater than 500 km. Specifically, the long-wavelength gravitational free-air admittance approaches 50 mGal/km, a magnitude that is } consistent with topographic support by convective circulation in the mantle [McKenzie and Fairhead, 1997]. When mantle structure inferred from seismic tomography data beneath Western Anatolia is used to drive neotectonics, there is plateau-type uplift across the region and topographic depressions adjacent to the sides of the uplift. The modeled mantle flow also drives surface extension across the uplift zone. [26] Figure 9 shows a direct comparison of the calculated residual topography profiles with the geodynamic prediction A1 on a map of Western Anatolia. The receiver function residual topography data and numerical model agree with anomalous uplift across Western Anatolia with an amplitude of 1.5 km. The CRUST2.0 data does not agree with this uplift, but the data set may not be suitable for drawing a discrete section of topography owing to its possible overestimation of crustal thickness in the study area and its low latitudinal and longitudinal resolution. The admittance Figure 8. Surface topography and surface strain rate (İxx) function calculation is isolated to data from the Aegean and from experiments: (a) A2, (b) A3, (c) A4, (d) A5. Western Anatolia, and it is in agreement with the receiver function-derived residuals and geodynamic model that rising mantle flow is actively supporting high topography in this this interrupts the broad uplift/extension caused by the region. Interestingly, all of the profiles agree with the upwelling mantle material. residual topography lows to the north and south of Anatolia. [24] In experiment A5 the crustal material is still present Although these are not the focus regions of this study (and but the seismically-derived thermal anomalies are imposed all the way to the surface of the model, rather than super- are outside of the admittance function calculation window), it seems that the subduction forcing to the south and some imposing a 30 km thick crustal geotherm (Figure 8d). This mantle feature to the north may be important in supporting results in unrealistically high crustal temperatures and anomalies [Tezcan, 1995]; for example, temperatures of some anomalous subsidence there. Some very small magnitude, shorter wavelength features appear on the observed 1000 C at a depth of 10 km. Nevertheless, topography is still residual topography sections, but do not manifest on the defined by a broad plateau-type uplift in the centre of the observed profile. These minor features may be attributed to a model and extension across this zone. In experiment A5, host of factors that are not contained in the numerical models after 63,000 yrs the topography starts to deviate from this such as elastic flexural effects of the lithosphere and comconsistent pattern with essentially negligible values across positional/density heterogeneities in the crust and mantle. the centre of the region and significant uplift (to 4 km) at the far left of the box. These variations are caused by the small very near-surface seismic anomalies that are removed by the 5. Discussion imposed crustal geotherm in experiments A1–A4 (Figure 6). [27] In Figure 10 we plot elevation versus crustal thickA simple conversion of these very shallow velocity anoma- ness for various sites (based on seismic receiver function lies directly to temperature is problematic; other factors such locations) across Western Anatolia. The plot also shows as compositional variations may be responsible for such similar data for the North American Cordillera by Hyndman surface, short wavelength velocity anomalies. However, the and Currie [2011]. In their study, they interpret the data as model is a useful illustration that generally the nearest- separating out into Cordilleran (hot back-arc mountain belts) surface thermal flow features dominate in the resultant and stable areas [Hyndman and Currie, 2011] with 1600 m surface topography. elevation difference for similar crustal thickness. The 10 of 14

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Figure 9. Residual topography plots for Western Anatolia using receiver function data and CRUST2.0 and the dynamic topography for numerical experiment A1. elevation difference between Western Anatolia and surrounding “stable” areas are less than the difference between the North American Cordilleran and stable areas. That is, the region does not seem to exhibit the same contrast between regional lithospheric structure. Further, we interpret that the high magnitude residual and dynamic topography (1500 m) is influencing significantly Western Anatolia and surrounding areas broadly. [28] Our study suggests active mantle support for surface topography; one may consider the nature of the mantle flow in the region. We speculate that the active upwelling mantle flow interpreted from the study may be related to flow induced by delamination of mantle lithosphere from crust beneath Western Anatolia. Plateau-type uplift of Eastern Anatolia has been linked with mantle lithosphere delamination [Göğüş and Pysklywec, 2008a, 2008b]. A similar series of geophysical observables exist for Western Anatolia to those that support the interpretation of delamination in Eastern Anatolia. Broad-scale (500 km) zones of low (2000 m) between Western Anatiolia and “cold stable areas”. (b) Geographical location of black and green dots in the graph. [32] Aside from active mantle dynamics, interpretations have been made suggesting that various aspects of the surface tectonics can be explained by crustal dynamics. The crustal thickness of Western Anatolia (Figure 2b) may suggest an uneven crust-mantle boundary geometry. The high temperature of the crust presents the possibility of gravitydriven lateral flow of middle/lower crustal domains and development of dome exhumation. Tirel et al. [2008] mention the importance of the geometry of the crust-mantle boundary to explain the dynamics and structural development of metamorphic core complexes. This approach enables lateral flow and crustal thinning independent of mantle processes (although they may be working in tandem). [33] Acknowledgments. The collaboration between the authors was facilitated by a grant from the Council of Higher Education of Turkey. In addition, R.G. was funded by an Alexander Graham Bell CGS NSERC scholarship, and R.P. was funded by Natural Sciences and Engineering Research Council of Canada (NSERC) Discovery Grant. O.H.G. is funded by (NERC) U.K. Natural Environment Research Council. We appreciate careful and thoughtful reviews by Diane Arcay and an anonymous referee.

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