The thermal state and strength of the lithosphere in the Spanish Central System and Tajo Basin from crustal heat production and thermal isostasy Alberta Jimenez-Diaz a, b,., Javier Ruiz a, Carlas Villaseca b,c, Rasa Tejero a,b, Raman Capate" • Departamento de Geodindmica, Facultad de Ciencias GeoI6gicas. Universidad Complutense de Madrid.Jose Antonio Novais 2,28040 Madrid, Spain b Instituto de Geociencias, CSIC-UCM.Jose Antonio Novais 2, 28040 Madrid, Spain C Departamento de Petrologi a y Geoquimica. Facultad de Ciencias Geologicas. Universidad Complutense de Madrid.Jose Antonio Novais 2, 28040 Madrid, Spain ABSTRACT Keywords: Heat flow Thermal structure Strength envelopes, Rheology Continentallithosphere Iberian Peninsula In this work we have modeled the thermal structure of the lithosphere of the Spanish Central System and the Tajo Basin, and their implications for lithospheric strength. For this, we have used refined heat­ producing elements (HPE) values to obtain new estimates of heat production rates in the Spanish Central System and Tajo Basin areas, which have been used joined to the relation between topography and thermal structure ofthe lithosphere to calculate the best-fit surface heat flows in the study area. Moreover, we have implemented a temperature-dependent thermal conductivity (appropriate for olivine) for the lithospheric mantle to improve the calculations of temperature profiles in the mantle. The geotherms so obtained, together with the implementation of a new rheological law for the upper lithospheric mantle, have been used to calculate refined estimations of the strength and effective elastic thickness of the lithosphere. We have obtained surface heat flow values of84mWm-2 and �82 mWm-2 for the Spanish Central System and the Tajo Basin, respectively. The thermal state of the lithosphere affects mantle temperatures, and hence may be playing an important role in the uplift and maintenance of the Spanish Central System. 1. Introduction The thermal state and the rheological behavior of the continen­ tal lithosphere depend on many factors (e.g., Afonso and Ranalli, 2004; Chapman and Furlong, 1992; Kohlstedt et aI., 1995; Ranalli, 1997; Ranalli and Murphy, 1987). Due to the relationship between thermal and mechanical structure of the lithosphere, it is neces­ sary to have an adequate knowledge of the thermal parameters, local heat flow and thermal structure, to reduce the uncertainty in strength estimates. For example, the amount and distribution of lithospheric heat-producing elements (HPE) and the values of the thermal conductivities of crust and mantle may affect the results substantially. Thus, the mechanical behavior and rheo­ logical stratification of the lithosphere in continental areas are largely a consequence of local conditions (e.g., Afonso and Ranalli, 2004; Furlong and Chapman, 1987; Ruiz et al., 2006; Watts and Burov, 2003). On the other hand, recent laboratory experiments conducted under controlled microstructural and chemical condi­ tions have shown a significant effect of important parameters on " Corresponding author at: Departamento de Geodinamica, Facultad de Cien­ cias Geol6gicas, Universidad Complutense de Madrid, Jose Antonio Novais 2, 28040 Madrid, Spain. E-mail address: ajimenezdiaz@geo.ucm.es (A. Jimenez-Dfaz). the rheological properties of major silicate rocks (BOrgmann and Dresen, 2008), and have yielded new rheological laws describing the first-order mechanical behavior of the lithospheric materials (e.g., Katayama and Karato, 2008; Keefner et al., 2011 ; Mei et al., 2010). The aim of this work is to model the thermal structure and their implications for lithospheric strength of the Spanish Cen­ tral System (SCS) and the Tajo Basin (TB). The SCS constitutes the most prominent tOJX)graphic elevation in the interior of the Iberian Peninsula separating the Duero and Tajo watersheds. It is flanked by two Cenozoic intracontinental sedimentary basins, the Duero Basin to the north and the TB to the south (Fig. 1). The SCS is a thick-skinned double-vergence (pop-up) intraplate range built as a result of JX)lyphase Alpine tectonic evolution (De Vicente and Vegas, 2009; De Vicente et al., 2004, 2007, 2009; Fernandez-Lozano et al., 2011; Martin-Velazquez et al., 2009), in which the defor­ mation partitioning of the basement in the intraplate convergence setting of Iberia has had a profound influence on the development of topography. The surface heat flow map of the Iberian Peninsula performed by Fernandez et al. (1998) provides some values at the SCS-TB bound­ aryand the north and south of the TB. In contrast, the SCS is not well characterized due to the unavailability of heat flow measurement, and it is necessary to approach the study of the thermal state of the range through other methodologies. In this sense, Tejero and Ruiz 5'W Duero Basin • Madrid Tajo Basin \ c=J Cenozoic basins -- Faults -- Thrusts Fig. 1. Geographical andgeoJogicaJ settings of the study area showing the two mountains ranges (Spanish Central System and Toledo Mountains) separated by the Tajo Basin. Map background is from the Neotectonic Map of Spain (IGME and ENRESA, 1998). (2002) modeled the thermal structure of the lithosphere of this area by using thermal isostasy to improve the calculated geotherms, considering surface heat flow values of 70 and 6S-70mWm-2 for the SCS and the TB, respectively. On the other hand, several works have focused on characterizing the lithospheric strength from estimating the effective elastic thick­ ness of the lithosphere through flexure modeling (Van Wees et al., 1996), the coherence between tOJX)graphy and Bouguer anomaly (G6mez-Ortiz et al., 200sa; Perez-Gussinye and Watts, 2005) or from rheological models (Martin-VeLhquez et al., 2008; Ruiz et al., 2006; Tejero and Ruiz, 2002; Te5auro et ai., 2007, 2009). In the present work, we have used the relation between topog­ raphy and thermal structure to calculate the best-fit surface heat flows. We included a temperature-dependent thermal conductiv­ ity (appropriate for olivine) for the lithospheric mantle to improve the calculations. Moreover, we have used refined HPE values based on bulk rock composition of main lithological formations of the SCS and the Toledo Mountains (e.g. Villaseca et al., 1998, 1999, 2005; this study), and these values have been used to obtaining estimates of heat production rates. Finally, we have used our result for the thermal structure in order to analyze the strength of the lithosphere in the study area. To make this, we have implemented a new rheo­ logical law for the upper lithospheric mantle, largely controlled by low-temperature plasticity of olivine-rich rocks (Mei et al., 2010). All of this provides an opportunity to refine existing thermal and rheological models and litho spheric strength determinations of the study area. 2. Temperature profiles The thermal structure of the lithosphere depends on heat flow, heat sources distribution and thermal conductivity of lithospheric rocks. The temperature profile within the lithosphere has been calculated assuming steady-state conditions and radioactive heat sources homogeneously distributed in three crustal layers and in the lithospheric mantle. The temperature at depth z in each crust layer is Fsz Hz2 Tz = Ts+ T - 2[' (1) where Ts and Fs are the temperature and heat flow at the layer top, k is the thermal conductivity, and H is the volumetric heat produc­ tion rate. The calculations assume k=2.s, 2.5 and 2.1 Wm-1 K-1 for upper, middle and lower crust, respectively, and the surface temperature was taken as 288 K. The thermal conductivity of olivine (the main mineral in the mantle) is strongly temperature-dependent; therefore tempera­ ture profiles in the mantle lithosphere are calculated from (see Ruiz et ai., 2011) dT Feb - PmHm(z - be) dz km(T) (2) where Feb =F - PeHebe is the heat flow at the base of the crust, Pm and Hm are, respectively, the density and heat production rate per mass unity of the mantle lithosphere, be is the base of the crust, and km is the thermal conductivity of the mantle lithosphere. For km we use the thermal conductivity of olivine, which is a function of temperature according to the expression (McKenzie et al., 2005) 3 km� 1+C(:- 273) + LdjTi, i=O (3) where a�5.3, c�0.0015, do�1.753xlO-2, d,�-1.0364xlO-4, d2 = 2.2451 xl 0-7 and d3 = -3.4071 x 10-11. Results obtained from Eq. (3) are similar to those of Hofmeister (1999) for forsterite olivine. For solving Eqs. (2) and (3), we used the Newton iterative method. Moreover, the use of concept of thermal isostasy is useful in order to constrain continental temperature profiles (e.g., Fernandez et al., 1998; Hasterok and Chapman, 2007, 2011; Lachenbruch and Morgan, 1990; Tejero and Ruiz, 2002), by providing a link between the thermal structure of the lithosphere and the elevation of the surface. Elevation above sea level (e) can be expressed by (4) where he and hm are the individual contributions of crust and mantle comJXJnents to the buoyancy of the lithosphere. ho is the buoyant height of sea level above the free asthenosphere surface (hoR::J2.4 km; Lachenbruch and Morgan, 1990). Crust contribution is estimated from 1 he = -(Pa - Pe)be, Pa (5) where be is crust thickness, Pe is mean crust density and Pa is asthenosphere density (3200 kg m-3). Mantle contribution is related to the thermal state of the lithosphere mantle by (6) where IX is the thermal volumetric expansion coefficient (3.5 x 10-5 K-1), bm is the thickness of the lithospheric mantle until the asthenosphere temperature Ta, assumed to be the isotherm of 1350 oc, and Tm is the mean lithosphere mantle temperature given by 1 lbm Tm � b T(z)dz, m 0 (7) which we use here in order to determine the depth of the lithosphere-asthenosphere boundary (lAB). Here we use crustal structure and composition derived from seismic data (Banda et al., 1981; ILIHA DSS Group, 1993; Surifiach and Vegas, 1988), and crustal density derived from gravity data analysis (G6mez-Ortiz et al., 2005b). Mantle heat flow was estimated by subtracting crustal contribution from surface heat flow. We applied the thermal isostasy model by iterative calculation to fit the calculated eleva­ tion to the observed mean elevation (�1250 m and �650 m for SCS and TB, respectively). Table 1 summarizes the parameters used in the calculations. 3. Crustal heat production Uranium, Thorium and Potassium (collectively termed as heat­ producing elements, HPE) abundance determines heat production rates of crustal rocks. Thus, heat production is calculated from HPE abundance by the addition of the contribution of each element as follows (Rybach, 1988) H (I"W m-3) � 10-5 p(9.52Cu + 2.56CTh + 3.48CKl, (8) where Cv and CTh are in ppm and CK in percent, and P is the density (in kg m-3). This method has been used to estimate heat produc­ tion rates of the crust of Central Iberia (Spain), where 196 samples of metamorphic and igneous rocks of Variscan basement were collected and their content on HPE determined. They represent main outcropping lithologies mostly orthogeneiss and granites. Furthermore, granulite xenoliths carried by Upper Permian alka­ line lamprophyres have been interpreted as samples of the lower crust below the SCS (Villaseca et al., 1999). Table 2 summarizes HPE abundances collected from 196 sites covering the Spanish Central System and the Toledo Mountains. HPE abundances and heat production values for the SCS com­ plex comprises outcropping metamorphic rocks mostly of two types: metasedimentary sequences and felsic orthogneissic rocks. These Cambrian-Lower Ordovician metaigneous rocks are the dom­ inant country rocks, and they show higher heat production rates than metasedimentary types because they are enriched in U and K (Table 2). Metabasic rocks have been also described in the SCS, "r---------------------------------, '" c:: 14 12 10 4 12 o 10 � GI '" .c o '0 4 o Z 2 7 , 5 4 3 SCS upper crust (n = 111) H= 2.45IlWm� Montes de Toledo upper crust (n = 58) H=2.36IlWm-3 SCS lower crust (granulit. x."oliths,!! (n= 27) H= 0.96 "Wm' Fig.2. Estimates of heat production rates of rocks from the Spanish Central System upper crust, Toledo Mountains upper crust and Spanish Central System lower crustal granulite xenoliths. but defining a very minor surface; regional geologic maps sug­ gest that felsic orthogneisses constitute approximately the 80% of the metamorphic rock eXJXJsures at least in his eastern half sec­ tor. Otherwise, most of the SCS is occupied by a huge granitic batolith. Proportions determined by mapped lithologies suggest that granites might be 75% of the SCS. The SCS Variscan granites are characterized by an averaged heat production of 2.49 j.1Wm-3 (pondered by granite type and area, Table 2), higher values than those of orthogneissic wall-rocks. This approach yields an aver­ aged heat production value of 2.45 j.1Wm-3 for the outcropping SCS rocks (Fig. 2). In the south, Toledo Mountains is also comprised by metamor­ phic rocks intruded by Variscan granite plutons. Country rocks are dominated by Neoproterozoic-Low Palaeozoic metasedimentary sequences (the Schist-Greywacke Complex), most of low-grade metamorphism (San Jose et al., 1990). The estimated average heat production rate of 2.36 j.1Wm-3 (Table 2 and Fig. 2) for the whole Toledo Mountains area is lower to that obtained for the whole granite-high-grade metamorphic complex of the SCS, mostly due to the lower abundance of granites in the Toledo Mountains area (Table 2). HPE abundances for Tajo sedimentary rocks are not available, although representative heat production rates can be estimated from the surrounding orogenic areas as their Table 1 Parameters used to construct thegeotherms and strength envelopes. Crustal structure and composition derived from seismic data (Banda et ai., 1981; lLlHA DSS Group, 1993; Surifiach and Vegas, 1988), crustal density derived from gravity data analysis (G6mez-Ortiz et ai., 2005b) and rheological parameters from Ranalli (1997). Thickness (km) Thermal conductivity (Wm-1 K-l) Spanish Central System Upper crust (dry granite) 11 2.5 Upper crust (wet granite) Middle crust (quarzdiorite) 14 2.5 Lower crust (felsic granulite) 9 2.1 Tajo Basin Sediments layer 2/F 2.5 Upper crust (dry quarzite) 12/13' 2.5 Upper crust (wet quarzite) Middle crust (quarzdiorite) 9 2.5 Lower crust (felsic granulite) 8 2.1 , Thickness for north/south Tajo Basin, respectively. sedimentary source regions (averaged heat production of 2.40 !"Wm-3 ; Table 1). The averaged heat production rate of the lower crust is esti­ mated to be 0.96 j.1Wm-3; value slightly lower than preliminary estimates (Villaseca et al., 2005), but clearly higher than values usu­ ally considered for the lower crust (Furlong and Chapman, 1987; Hasterok and Chapman, 2011; Rudnick and Gao, 2003; Vila et al., 2010). This is consequence of the markedly felsic comJX)sition of the SCS lower crust, dominated by felsic meta-igneous (95vol%) and pelitic (5 vol%) granulites (Villaseca et al., 1999). The felsic nature of the SCS lower crust is best shown in comparison with other lower-crustal xenoliths suites which, on average, are more mafic than granulite terranes (Villaseca et al., 1999 and references therein). Heat production of the middle crust is not well characterized due to the unavailability of direct measurement, and thus, there is considerable uncertainty with regard to this parameter. The mid-crustal layer is not pervasive globally, and where it exists, tends to have radiogenic heat generation more similar to lower rather than to upper crust (Hasterok and Chapman, 2011 and ref­ erences therein). On the other hand, the middle crust of the area is made of intrusive felsic materials (Villaseca et al., 1999) similar to those forming the upper crust in many continental areas, and seis­ mic velocities findings JX)int to a granodioritic comJX)sition of the middle crust (Banda et al., 1981). Unknown of the depth distribu­ tion of rocks compelled us to consider layer homogeneity, and we have assumed an intermediate (density weighted) heat production rate between upper and lower crust. Finally, for the lithospheric mantle we use a standard heat production rate of 0.02 j.1W m-3 (e.g., Chapman and Furlong, 1992; Hasterok and Chapman, 2011 ). Table 2 Heat production Density (kgm-3) A (MPa-n S-l ) Q (kJmol-1) n (�Wm-3) 2.45 2670 1.8 x 10-9 123 3.2 2.0 x 10-4 137 1.9 1.75 28DO 1.3 x 10-3 219 2.4 0.96 29DO 8.0 x 10-3 243 3.1 2.40 24DO 6.7 x 10-6 156 2.4 2.36 2780 6.7 x 10-6 156 2.4 3.2 x 10-4 154 2.3 1.65 28DO 1.3 x 10-3 219 2.4 0.96 29DO 8.0 x 10-3 243 3.1 4. Strength of the lithosphere The concept of strength envelopes is useful to illustrate a first­ order approximation of the rheological properties of lithosphere (e.g., Brace and Kohlstedt, 1980; Kohlstedt et ai., 1995; Ranalli, 1997; Ranalli and Murphy, 1987). Thus, the strength of the litho­ sphere at any depth is the minimum between the strengths for brittle and ductile deformation. Assuming a prefractured medium with fractures ideally oriented, the brittle strength is calculated according to the expression (e.g., Ranalli, 1997; Ranalli and Murphy, 1987) (a, - a3)b � ,6pg(l - )..)2, (9) where j3 is a coefficient depending on the stress regime (0.75 for tension and 3 for compression), p is the density,g is the acceleration due to the gravity (9.8 m s-2), A is the JX)re fluid factor defined as the ratio of pore fluid pressure to lithostatic pressure (Sibson, 1974), and z the depth. The density of the brittle crust, adequate for rocks in the upper crust, is taken as 2670 kg m-3 fortheSCS and 2780 kg m-3 for the TB (G6mez-Ortiz et al., 2005b). In addition, a sedimentary layer is considered in the TB (Table 1). We use the same hydrostatic pore fluid factor (A = 0.37) for the whole lithosphere. The ductile strength does not depend on the stress regime but it is strongly strain rate- and temperature-dependent (Burov and Diament, 1995; Ranalli and Murphy, 1987; StOwe, 2002), and can be described by a thermally activated power law, (t) '/n ( Q ) (0' 1 - a3)d = A exp nRT ' (10) where e is the strain rate, A, Q, and n are laboratory-determined constants, R is the gas constant (8.31447 jmol-1 K-1), and Tis the Estimates of heat production of rocks from the Spanish Central System and Toledo Mountains. Area (%) Numbers of samples U (ppm) Th (ppm) K(%) H (�Wm-3) Spanish Central System Metamorphic rocks Metabasites <0.1 6 1.37 5.70 0.54 0.86 Metapelites 20 8 2.65 14.82 2.76 2.00 Metagranites 80 41 4.61 12.25 3.69 2.42 Granitic rocks Monzogranites 85 25 3.07 15.63 3.72 2.22 Leucogranites 15 31 8.61 21.23 3.83 4.04 LC Granulite Xenoliths Chamockites <1 4 1.18 2.61 1.87 0.69 Pelites 5 6 0.78 9.14 2.32 1.21 Metaigneous 95 17 0.70 6.54 2.55 0.95 Toledo Mountains Metasediments 55 5 3.66 10.97 2.63 1.98 Granites MTB 35 42 6.38 14.76 3.55 2.99 ACT Migmatites 10 11 3.39 12.88 4.97 2.27 absolute temperature. Strength envelopes are calculated for a strain rate ofl0-15 s-1. The ductile strength of the upper crust is calculated using flow laws for wet/dry granite and wet/dry quartzite for the SCS and TB, respectively. The lower crust of the central Iberian Peninsula is of a felsic granulite nature (Villaseca et al., 1999), and bearing in mind its flow law it should not appreciably contribute to the strength of the lithosphere. It is therefore not taken into account in the present work (see Tejero and Ruiz, 2002; Ruiz et al., 2006). In turn, the middle crust of the area is made of intrusive felsic materials (Villaseca et al., 1999) similar to those forming the upper crust in many continental areas; its mechanical behavior is therefore likely to be similar (Ruiz et al., 2006). For the estimation of the strength of the litho spheric mantle we use dry and wet olivine rheologies, which give upper and lower limits, respectively. The behavior of the upper lithospheric man­ tle is in turn largely controlled by low-temperature plasticity of olivine-rich rocks (Mei et al., 2010), resulting in a rheology signif­ icantly weaker than that usually used for the lithosphere mantle. Under anhydrous conditions, Mei et al. (2010) define a flow law for a quasi steady state deformation of olivine under low-temperature and high-stress, which can be written in terms of differential stress as ( ) _ (e ) '/2 [£k(O) ( 1 J(a1 -a3))] 0' 1 - 0' 3 - exp -- -- Ap 2RT ap ' (11) where Ap = 1.4 x 10-7 s-l MPa-2, Ek( 0) is the zero-stress activation energy (320±SOkjmol-1), and ap is Peierls stress (S.9±0.2 GPa). Thus, for dry olivine we use the minimum strength obtained from Eq. (11) and from the high temperature flow law obtained for artificially dried dunites: A=28,840MPa-ns-1, n=3.6 and Q= 535 kJ mol-1 (Chopra and Paterson, 1984). For wet olivine, we use the flow law of the Anita Bay dunite: A=9550MPa-ns-1, n�3.3S and Q�444 kj mol-1 (Chopra and Paterson, 1984). This flow law places a lower limit on the strength of wet olivine due to its rel­ ative weakness (compared with other wet dunites, such as Aheim dunite). Finally, the total lithospheric strength (Ranalli, 1997) can be defined as s� lbL(a1-a3)(Z)dZ, (12) where (0' 1 -0' 3) is the minor, at z depth, between the brittle and ductile strength, and bI is the mechanical thickness of the litho­ sphere. The base of the mechanical lithosphere is here defined as the depth at which the ductile strength reaches a low value of 10 MPa (McNutt, 1984; Ranalli, 1994), and below which there are no further significant increases in strength, although the exact value selected does not produce significant changes in the calculations due to the eXJXJnential dependence of ductile strength on temper­ ature. Table 1 summarizes the rheological model parameters. 5. Results 5.1. Thermal modeling Fig. 3 shows the geotherms obtained by our thermal model. For the SCS, we have obtained a value of Fs=84mWm-2, with elevation adjustment of ± 1 m (Table 3 and Fig. 3). For the north and south TB, we have obtained values of Fs=81 mWm-2 and Fs = 83 mWm-2 respectively, with elevation adjustment of ±2 m in both cases (Table 3 and Fig. 3). We have calculated the surface heat flow from the thermal isostasy model obtaining through iterative calculation the surface heat flow best-fitting the observed mean elevation. In this sense, a heat flow uncertainty of ±0.1 mWm-2 Temperature ('C) o 250 500 750 1000 1250 1500 1750 O �--�----�---L----�--�----�--� 10 � 70 90 '7' -------------�-.,..-I.::;:.-� ------------. ��c-Pi;':;"'""'5' 30 " , SO " .. 60 � 70 80 90 100 110 120 �---- -750 ·500 -250 ---- ------- - - ---- - --- - 250 500 750 -750 -500 -250 250 500 750 -750 -500 ·250 00 (Mpa) 00 (Mpa) lITHQSPHERIC MANnE 250 500 750 " SO , " SO .. 70 � 80 90 100 110 120 Fig,4, Results of thermal models and strength envelopes, calculated for a strain rate of10-15 S-I, plotted along a NW-SE transverse section of the area. Outer black line binds differential stress estimated for dry rock composition. Inner dashed line denotes differential stress for wet rock composition of the upper crust (quartzite or granite) and lithospheric mantle (peridotite). F,: surface heat flow. Te: effective elastic thickness for wet/dry rheology. Srn(\>:: maximum total lithospheric strength. Moho: crust-mantle boundary. LAB: lithosphere-asthenosphere boundary. Fig, S shows total lithospheric strength for compressional and tensional stresses and, in each case, for dry and wet rheologies, Total strength ranges from �8,2 x 1012 to � 1,2 x 1012 Nm-1, In general, higher total lithospheric strengths are associated with the north TB, while minimum values corresponded to the SCS, These values are consistent with mean integrated strength values estimated under compressional conditions by Tesauro et al. (2009) for the continen­ tal lithosphere in Iberia, In the same way, the contribution of the lithospheric mantle to the total lithospheric strength ranges from � 7 x 1012 to �2 x 1011 Nm-1, These values are consistent with the wavelengths «250 km) of the lithospheric folds, which suggests low mean mantle strength values «1013 Nm-1; Sokoutis et al., 2005), proposed by Mufioz-Martin et al. (2010) from the spectral - 1.0.1013 E is 8.2·10" 6.5·10" £ C> 4.8·10" c � iii 1.0.1012 NTB STB Fig,S, Total lithospheric strength values for dry and wet rocks in compression and tension plotted for the different tectonic units. Higher strength values were obtained for dry rheology and compressive differential stress. Minimum strength values correspond to wet rheology and tensional differential stress. analysis of the gravity and elevation for continental lithosphere at the Africa-Eurasia boundary, Our results can also be interpreted in term of the effective elastic thickness of the lithosphere (re), a measure of the total strength of the lithosphere which integrates the contributions from brittle and ductile layers and from elastic cores of the lithosphere (for a review see Watts and Burov, 2003), We have calculated re from the strength envelopes constructed for the SCS and the TB, Following Burov and Diament (1995), the total effective elastic thickness of an unflexed plate constituted by n detached layers is ( n ) 1/3 Te= ;�l (13) where tei is the mechanical thickness of the layer j, We take the base of each mechanical layer as the depth in which the strength goes down to a value ofl0 MPa (see above), If strength levels at the base layer are higher than 10 MPa, the layer is considered welded to the layer below, The calculations were performed for both wet and dry rheology, For the SCS, the results are 16km for wet rhe­ ology and 28 km for dry rheology, For the north and south TB, the obtained values are 17-20 and 30-34 km for wet and dry rheology, respectively, The lower values for the SCS are greatly resulting of the weaker upper mantle in this zone, which has a more limited contribution to the total strength of the lithosphere, The effective elastic thickness depends on the thermal state of the lithosphere, which determines the thickness and contribution Van Wee, et al. (1996) . I Flexurc modcling ��� ����������j��������� ��� I Bougu" ooh","" Gomez-Ortiz et al. (2005a) Perez-Gussinye and Watts (2005) Rui. et al. (2006) ______________________________ I Free-air admittance - - ------------------------------ Tc,auro ct al. (2007) Martin-Vehizquez et al. (2008) • - Upper limits Rheological modeling ----------------------- - ------ This study Upper limits o 10 20 30 40 50 Effccfi\'C elastic thickness o Central Iberian Peninsula _ Spanish Cenlral System _ Tajo Basin Fig. 6. Compilation of effective elastic thickness values obtained for the study area. '" " 30 '. (Km) 20 " 10·" Spanish Central System " 22 " --"'lI: ___ �_ 10·" Strain Rate (s·') F.=84 mWm-l! --- 10·" '" lajo Basin (North) " • F,=81 mWm-l 30 " � '. (Km) 20 - -