Journal of Geodynamics 34 (2002) 47–61 www.elsevier.com/locate/jgeodyn
Geothermal modeling along a two-dimensional crustal profile in Southern Portugal§ A. Correiaa,*, J. Sˇafandab a
Department of Physics, University of Evora, Largo Dos Colegiais, 7000 Evora, Portugal Geophysical Institute, Acad. Sci. Czech Rep., Boc˘nı´ II, 141 31 Praha 4, Czech Republic
b
Received 3 April 2001; received in revised form 11 December 2001; accepted 12 December 2001
Abstract Crustal temperatures along a S–N trending profile crossing the main geological units of Southern Portugal, the South Portuguese Zone and the Ossa Morena Zone, were calculated by solving numerically the heat conduction equation in a 2-D geothermal model of the crust. Use of observed surface heat flow and heat production based on the empirical relationship between radiogenic heat sources and observed seismic velocities, yields temperatures at the crust/mantle boundary, which are higher than the upper limit (about 700 C) of crustal temperatures based on magnetotelluric surveys. The most uncertain parameter of the model is the upper crust heat production, whose low value implies a large non-crustal component of surface heat flow, which must be conducted through the whole crust and increases the vertical temperature gradient. To overcome this uncertainty in heat sources, the Moho heat flow was estimated using Vitorello and Pollack’s [J. Geophys. Res. (1980) 983] estimate of the background heat flow of 27 mW m2 arising from below the zone of crustal enrichment (identified in our model with the upper crust) and transient thermal perturbation, represented in Variscan units by 5–10 mW m2. Taking into account the heat production of the middle and lower crust, the mean value of the Moho heat flow should be 28– 33 mW m2. The Moho heat flow was allowed to vary (with variation restricted to 0.2–1.0 mW m2 per 5 km) around 28 or 33 mW m2 along the profile in order to improve the fit between the calculated and measured surface heat flows. Because this approach yields a surface heat flow lower than observed, the residual differences were minimized by increasing the heat production of the upper crust, where the production was allowed to vary along the profile (with variation restricted to 0.02–0.2 mW m3 per 5 km). If Moho heat flow is allowed to vary by a total of about 9–25 mW m2, the upper crustal heat production necessary for the calculated surface heat flow to match the observed one is in the range 2–3 mW/m3, which indicates a mostly granitic upper crust. The resulting Moho temperatures are below 700 C everywhere along the profile in all model configurations considered. # 2002 Elsevier Science Ltd. All rights reserved.
§
Paper presented at the conference ‘‘Geothermics at the Turn of the Century’’, University of Evora, Portugal, April 2000 * Corresponding author. Fax: +351-66-702306. E-mail address:
[email protected] (A. Correia). 0264-3707/02/$ - see front matter # 2002 Elsevier Science Ltd. All rights reserved. PII: S0264-3707(01)00080-1
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1. Introduction The crustal temperatures within the two main geotectonic units of southern Portugal, namely the Ossa Morena Zone (OMZ) and the South Portuguese Zone (SPZ) (Fig. 1) were estimated by Correia and Ramalho (1999). The calculations were based on the steady-state heat conduction equation in one-dimensional geothermal models of the Earth’s crust representative for the two units. The crucial information was provided by the updated heat flow density map of southern Portugal (Fig. 2) (Correia and Ramalho, 1998) and seismic refraction surveys (Mendes-Victor et al., 1993), completed by measured radiogenic heat production values and thermal conductivity data. The paper by Correia and Ramalho (1999) provided characteristic crustal geotherms within the two different geotectonic units. Inferred Moho temperatures were 550–570 C in OMZ and 870–880 C in SPZ. The latter are incompatible with other geophysical and geological data from the area. This paper addresses the problem of thermal coupling between both units and also the question of too high a Moho temperature in SPZ, by solving the steady-state heat conduction equation in a two-dimensional geothermal model crossing both units.
2. Geological setting The study area belongs to the southwestern part of the Iberian Massif, which forms the core of the Iberian Peninsula and is of Variscan age. As mentioned above, it consists of two main units, the Ossa Morena Zone to the northeast and the South Portuguese Zone to the southwest, separated by a major overthrust known as the Ferreira-Ficalho Overthrust (FFO) dipping approximately NE (Fig. 1). Near the FFO, gabbro intrusions are found and synmetamorphic serpentinites change to protomylonites in a zone about 1 km wide (Ribeiro et al., 1979). Precambrian and lower Paleozoic rocks of the OMZ show intense deformation caused by different deformation phases with distinct trends and widespread volcanic synorogenic magmatism. In the SPZ, the older rocks date from the late Paleozoic, namely Upper Devonian formations, and acid volcanism is more common. Plutonism is almost insignificant and metamorphism is low grade. Between OMZ and SPZ there is a conspicuous tectonic and paleogeographic polarity; both the ages of the flysch formations and the main deformation become younger towards the southwest. The E–W zone of high heat flow is close to the FFO and overlaps with the Iberian Pyrite Belt. The belt represents an imbricated SW-trending complex. This area is rich in metalliferous ores, such as copper and zinc, and has complex geological and tectonic patterns.
3. Thermal parameters Because both the contact between OMZ and SPZ and the zone of high surface heat flow trend approximately E–W, the best two-dimensional approximation of the problem is a S–N-trending profile. Its position is shown in Fig. 2. The profile crosses the FFO at 200 km and the section we are interested in lies roughly between 100 and 300 km. Its whole length is 360 km so that the boundary conditions of horizontal symmetry applied at its sides do not influence solution of the
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Fig. 1. Ossa-Morena (I.b) and South Portuguese (I.c) zones in the frame of the structural units of mainland Portugal (modified from Calado, 1991) and the position of the 2-D geothermal profile. The Ferreira–Ficalho Overthrust (FFO) separates the two units. The other units shown are: I.a=Central-Iberian Zone; I.a’=Middle Galicia Subzone; I.c’=Iberian Pyrite Belt; II=Meso-Cenozoic Western and Southern Sedimentary Basins; III=Tagus and Sado Tertiary Basins. The units I.a, I.a’, I.b, I.c and I.c’ belong to the Hercynian Massif.
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Fig. 2. Thermal profile running S–N along the 7 450 W meridian. Isolines are surface heat flow in mW m2.
heat conduction equation in the central zone. The depth of the profile (30 km) coincides in the OMZ with the crustal thickness, while in the SPZ the depth of the Moho boundary is slightly smaller (29 km). From the standpoint of geothermal modeling, the most relevant difference between the OMZ and SPZ is the observed surface heat flow, which is about 60 mW m2 in the OMZ and ranges from 70 to 90 mW m2 in the SPZ. This difference is well evidenced by heat flow measurements. The zone of high heat flow of 90 mW m2 is an interpretation of high, but rather scattered data (Correia and Ramalho, 1998) and its more confident contouring would require additional heat flow data in the area. New interpretations of several seismic refraction surveys (Mendes-Victor et al., 1993) have shown that the OMZ exhibits a three-layer structure with 12-km-thick upper crust with average seismic velocity 6.1 km s1, an 11-km-thick middle crust with velocity of 6.3 km s1, and a 7-kmthick lower crust with velocity of 6.8 km s1. The SPZ also shows a three-layer structure with a 14-km-thick upper crust with velocity of 6.1 km s1, a 7-km-thick middle crust with velocity of 6.2 km s1, and an 8-km-thick lower crust with average seismic velocity of 7.0 km s1. This
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structure of the two geological units was used as the basis of the two-dimensional geothermal model. The heat productions of the upper, middle and lower crust were estimated in the first approximation by the empirical relationship between radiogenic heat production and compressional seismic velocity proposed by Rybach and Buntebarth (1982, 1984): ln A ¼ 12:6 2:17Vp
ð1Þ
where A is the heat production in mW m3 and Vp is the seismic velocity of compressional waves in km s1. The resulting heat production model is shown in Fig. 3. Since seismic velocities in the middle and lower crust differ in the two zones, heat productions differ, too. The lowermost kilometer in the SPZ belongs to the mantle and its production was estimated as 0.01 mW m3. The most conspicuous feature of the crustal production model derived in this way is the low production of the upper crust, about 0.5 mW m3. This can be compared with values measured on rock samples collected in the region (Correia et al., 1993; Correia and Jones, 1997). They show, on the average, heat production values of 2.9 mW m3 for granites, 1.6 mW m3 for diorites, 0.8 mW m3
Fig. 3. The surface heat flow along the profile and heat production and thermal conductivity values in the three-layer crustal model inferred from seismic data. The resulting heat source distribution yields a very low crustal component, about 10 mW m2, of the surface heat flow. The depth of Moho discontinuity is 30 km in the north and 29 km in the south, where the lowermost kilometer of the model belongs to the mantle and has conductivity of 4 W m1 K1 and heat production of 0.01 mW m3.
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for microdiorites and 0.2 mW m3 for gabbros. The average of the uniformly weighted values of these four rock types is higher, about 1.4 mW m3, than results from the relationship (1). But because the relative content of the individual rock types and the depth dependence of the radioactive elements is uncertain, we kept in this first approximation the value obtained by Eq. (1). The thermal conductivity values for both OMZ and SPZ are taken from Correia and Jones (1995) and are 2.7, 2.5 and 2.1 W m1 K1 for the upper, middle and the lower crust, respectively (Fig. 3).
4. 2-D thermal modeling The steady-state heat conduction equation @=@x ðkðx; zÞ@T=@xÞ þ @=@z ðkðx; zÞ@T=@zÞ ¼ Aðx; zÞ
ð2Þ
where T is temperature, k thermal conductivity, A heat sources and x, z are horizontal and vertical coordinates, respectively, was solved by the finite-difference method. The used difference scheme, described by Sˇafanda (1985), enables models to be considered in which the boundaries of discontinuities of thermal conductivity or heat sources need not pass through the grid points. Whereas the solution of (2) is uniquely defined in the one-dimensional case, e.g. by the surface heat flow and surface temperature, and these quantities can be used as the boundary conditions, in a two-dimensional model, this is not the case. It follows from the theory of elliptic differential equations that solution of (2) in a two-dimensional geothermal model need not exist for every combination of the surface temperature and heat flow, which are now functions of x, and that its dependence on these functions can be discontinuous (so-called ill-posed problem). To overcome this difficulty, the temperature or heat flow at the lower boundary has to be known and used as a boundary condition there. Consequently, only one function, surface temperature or surface heat flow can be used as an upper boundary condition. The boundary conditions used in our modeling were the surface temperature of 15 C, the heat flow at the bottom and conditions of horizontal symmetry along the sides of the model. The heat flow along the bottom of the model is unknown, but it can be determined from the non-crustal component of the surface heat flow. The problem is that, in order to calculate the non-crustal surface heat flow component, we must use the model of crustal heat production. Because of our low confidence in estimates of the crustal heat production, we decided to estimate the heat flow at the bottom, which is practically identical with the Moho heat flow (and will be therefore referred to as the Moho heat flow or shortly qM), in a way which is independent on the most unreliable heat production of the upper crust. To do so, we used the ideas of Vitorello and Pollack (1980) about the three components of the surface heat flow. The first component, which stems from the radiogenic heat released in the zone of crustal isotopic enrichment, was identified in our model with the heat generated in the upper crust. The second component, which represents the residual heat of the transient thermal perturbation associated with the tectogenesis, can, according to Vitorello and Pollack (1980), reach 5–10 mW m2 in Variscan massifs. The third component is a background flux, estimated by Vitorello and Pollack (1980) at 27 mW m2, which comes from the crust below the zone of isotopic enrichment and from the mantle. To extract the part coming from the mantle from this third compoment of
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27 mW m2, we used the heat production values obtained by Eq. (1) for the middle and lower crust. They indicate that the heat flow stemming from the middle and lower crust is in the OMZ about 4.6 mW m2 and in the SPZ 3.7 mW m2. If we accept 4 mW m2 as the mean value for the whole model, the mantle part of the third component is then 23 mW m2. This value together with 5–10 mW m2 of the second component yield a Moho heat flow of 28–33 mW m2. In the following calculations we, therefore, considered two values of the average Moho heat flow, 28 and 33 mW m2. As a further step, we allowed the Moho heat flow to vary along the lower boundary of the profile around the mean value of 28 or 33 mW m2 to improve the fit between calculated and observed surface heat flows, when the Moho heat flow is used as a boundary condition at the bottom in solving Eq. (2). We used a regularization algorithm (Stromeyer, 1984; Sˇafanda, 1985), in which the mean-square value of the horizontal variation of qM is limited in order to avoid unrealistic differences in the heat flux from the mantle. The regularization algorithm in fact seeks the qM which minimizes the difference between the observed and calculated surface heat flows on a set of qM-functions with fixed upper limit of the mean-square horizontal variation. The qMfunction found was then shifted without changing its shape to allow its mean value along the profile to be 28 or 33 mW m2. Two examples of qM-function with the mean value of 33 mW m2 determined in this way are shown in Fig. 4. The situation of practically uniform Moho heat flow is examined by Ellsworth and Ranalli (in press). The Moho heat flow with variation 0.2 mW m2 per 5 km (5 km is the mean size of the horizontal grid) (Fig. 4a) decreases smoothly from 37 mW m2 below the SPZ to 28 mW m2 below the OMZ, i.e. by 9 mW m2. The calculated surface heat flow is about 11 mW m2 higher than the qM due to the crustal heat production, but it is by 15–45 mW m2 lower than the observed surface heat flow and without any sign of elevation below the narrow zone of high observed values. The Moho heat flow with variation 1 mW m2 per 5 km (Fig. 4b), which allows for greater horizontal changes of qM than in the previous case, reveals the maximum of 48 mW m2 below the zone of high observed heat flow and its amplitude along the profile is 25 mW m2. The maximum appears also on the curve of the calculated surface heat flow, but in comparison with the observed heat flow it is broader and its relative amplitude is lower. The misfit between the calculated and observed surface heat flows amounts to 14–37 mW m2. Whereas the 9 mW m2 increase of qM with variation 0.2 mW m2 per 5 km along the profile towards the south can be explained by a smaller age of the SPZ and perhaps also by possible differences in composition and therefore heat production of the subcrustal lithosphere (Vitorello and Pollack, 1980), the amplitude of 25 mW m2 of qM with variation of 1mW m2 per 5 km is much harder to explain within the frame of the conductive thermal regime of the lithosphere. However, because of the complex structure of the Iberian Pyrite Belt, where the zone of the high observed heat flow is located, the existence of a high conductivity zone reaching into the subcrustal lithosphere cannot be excluded. We kept therefore the qM with variation 1m W m2 per 5 km as the extreme, but not completely unrealistic case. The attempts to fit the observed heat flow along the studied profile by geophysically acceptable Moho heat flow for the given thermal conductivity and heat production model of the crust have shown that: (a) the zone of high observed surface heat flow cannot be explained only as the effect of the high Moho heat flow below it, and (b) the calculated surface heat flow for the Moho heat flow varying around 28–33 mW m2 is lower than the observed heat flow even outside the zone of its positive anomaly. This means that the crustal heat production in our model is underestimated.
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Fig. 4. Moho heat flow varying around 33 mW m2 with variation (a) 0.2 or (b) 1 mW m2 per 5 km and corresponding surface heat flow curves calculated before and after heat source optimization in the upper crust. The fit between the observed and calculated surface heat flow improves as the horizontal variation of the optimized heat production increases from 0.02 to 0.2 mW m3 per 5 km (panel a). Panel (b) shows only the curves calculated for the highest and lowest variation of the optimized production (full lines). They are compared with the same curves resulting from the model with pressure- and temperature-dependent conductivity (dashed lines). See text for details.
Most probably it is the heat production of the upper crust which must be increased. A mean value of 1.4 mW m3, based on the measurements of the surface samples instead of the adopted value of 0.53 mW m3, would lower the calculated surface heat flow deficit by 10–12 mW m2, having 10 mW m2 unexplained. We therefore decided to improve the fit between the observed and calculated surface heat flows by the procedure of optimization of heat sources (Stromeyer, 1984; Sˇafanda, 1985) in the upper crust. The optimization consists in search for an additional upper crustal heat production distribution f(x), constant in the vertical direction within the upper crust, which varies along the profile in order to minimize differences between the observed and calculated surface heat flow. Consequently, the sought function f depends on the qM considered. Because the character of the problem is similar to that of the Moho heat flow calculations, we used the same regularization algorithm. The algorithm seeks a heat generation function which, when added to the assumed heat production, minimizes the differences between observed and calculated surface heat flows on a set of functions with fixed upper limit of the mean-square horizontal variation (Sˇafanda, 1985). We considered variations in the range 0.02–0.2 mW m3 per 5 km. Examples of resulting upper crust distributions are shown in Fig. 5. These productions were calculated for the mean Moho heat flow 33 mW m2 and its variation 0.2 mW m2 per 5 km (a) or 1 mW m2 per 5 km (b). The improvement of the coincidence between the observed and calculated surface heat flows, when the optimized production is used in solving Eq. (2) is evident. Considering production with the lowest variation of 0.02 mW m3 per 5 km practically boosts the ‘‘pre-optimization’’ curve of the calculated surface heat flow to the levels of the observed one. The
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Fig. 5. Optimized heat production of the upper crust along the profile. Curves are for variations 0.02, 0.04, 0.06, 0.08, 0.1 and 0.2 mW m3 per 5 km and for the Moho heat flow varying around 33 mW m2 with horizontal variation 0.2 (a) and 1 mW m2 per 5 km (b).
Fig. 6. Comparison of isotherms (in C) corresponding to the two models with the mean Moho heat flow 28 mW m2 and its horizontal variation 0.2 mW m2 per 5 km, which differ in upper crust heat production. Heavy-line isotherms are for the heat production with horizontal variation 0.2 mW m3 per 5 km. The dashed-line isotherms refer to production with horizontal variation 0.02 mW m3 per 5 km.
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upper crust production necessary for it is about 22.5 mW m3 for Moho heat flow varying along 33 mW m2. For qM with mean value of 28 mW m2, the curves are very similar, but the values are higher by about 0.4 mW m3. With increasing variation of the optimized production, the improvement is evident also in the zone of the high observed heat flow. Nevertheless, to achieve the near-perfect coincidence in this zone, the values as high as 4–4.8 mW m3 are necessary. This is appreciably more than measured surface values, but because these high values are flanked by below-average production (curve 1 in Fig. 5), one can speculate about the enrichment in radioactive elements of the zone on the expenses of the adjacent areas of the upper crust in this part of the Iberian Pyrite Belt. The ultimate aim of the modelling is to provide the possible range of the crustal temperatures in OMZ and SPZ consistent with the observations. We have shown that for mean Moho heat flow estimates based on general ideas on the crustal thermal regime, the only way to achieve agreement between calculated and observed surface heat flows is to allow for horizontal variations of the heat production in the upper crust, together with moderate variations in Moho heat flow. The degree of this variation influences appreciably the agreement, but has only a limited effect on the crustal temperatures. It can be seen from Fig. 6, where the isotherms for the two models are given, which differ only in the upper crust production, the one with variation 0.02 and the other 0.2 mW m3 per 5 km. As expected, the largest temperature difference occurs at the bottom of the upper crust below the zone of the high surface heat flow and amounts to about 40 C. The temperature differences due to the two other model parameters, namely the mean value of the Moho
Fig. 7. Comparison of isotherms corresponding to the two models with Moho heat flow variation 0.2 mW m2 per 5 km and the upper crust heat production variation 0.2 mW m3 per 5 km, which differ in the mean level of the Moho heat flow. The heavy-line isotherms are for the mean Moho heat flow 33 mW m2, the dashed-line isotherms for the mean Moho heat flow 28 mW m2.
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heat flow (28 or 33 mW m2) and the Moho heat flow variation (0.2 or 1 mW m2 per 5 km) attains a maximum at the bottom of the whole crust, because qM represents the boundary condition at the bottom. The effect of the different mean values of qM is depicted in Fig. 7. The Moho temperature below the zone of the high surface heat flow for qM with mean value 33 mW m2 is higher by 50 C than for the mean value of 28 mW m2. Temperature differences due to different variation of qM, 0.2 or 1 mW m2 per 5 km, amount nearly to 100 C at the base of the crust (Fig. 8). The range of Moho temperatures yielded by the models is 400–500 C in OMZ. The values in the SPZ are generally higher (500–660 C) and reveal greater variability along the profile due to the existence of the surface heat flow positive anomaly. In all the above considered models of the crust we used the same thermal conductivity model (Fig. 3) as in the one-dimensional modelling by Correia and Ramalho (1999). Conductivity in this model was not explicitly temperature-dependent, but the room condition conductivity values of the expected rock types were adjusted a priori to the expected mean temperatures in the three crustal layers. To show the sensitivity of our calculations to this parameter, we used as an alternative the general model of crustal conductivity provided by Chapman and Furlong (1992). The conductivity is explicitly pressure- (depth) and temperature-dependent in this model according the formula kðT; zÞ ¼ k0 ð1 þ czÞ=ð1 þ bTÞ
ð3Þ
Fig. 8. Comparison of isotherms corresponding to the two models with the mean Moho heat flow 33 mW m2 and the upper crust heat production variation 0.2 mW m3 per 5 km, which differ in the horizontal variation of the Moho heat flow. The heavy-line isotherms are for the variation 1.0 mW m2 per 5 km, the dashed-line isotherms for the variation 0.2 mW m2 per 5 km.
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where T is temperature in degrees Celsius, b and c are constants, and k0 is the conductivity measured at 0 C and atmospheric pressure. We used a coefficient b of 0.0015 K1 for the upper crust which is a value intermediate between experimentally determined values for granite and granodiorite. For the middle and lower crust a temperature coefficient b of 0.0001 K1 was used. A pressure coefficient c of 0.0015 km1, is used throughout the crust. Zero depth and zero temperature values ko are 3.0 W m1 K1 for the upper and 2.6 W m1 K1 for the middle and lower crust, respectively. Due to the temperature-dependent thermal conductivity, the system of algebraic equations for temperature to be solved during the finite-difference solution of Eq. (2) is nonlinear and the numerical technique used for the temperature-independent model cannot be applied. Therefore, we used the approach described in Sˇafanda (1988), where the conductivity is adjusted to temperature by an iterative process. Temperature differences between the temperature-dependent and temperature-independent models of the conductivity are shown in Fig. 9. The temperature-dependent conductivity corresponding to the temperature distribution depicted in Fig. 9 is given in Fig. 10. The temperature-dependent conductivity drops below the value 2.7 W m1 K1, used for the upper crust in the temperature-independent model, at the depth of 2–3 km. From this depth downwards, up to the upper/middle crust boundary, the vertical temperature gradient is higher in the temperature-dependent model and therefore the temperature difference is highest, about 50 C at the bottom of the upper crust. The temperature-dependent conductivity of the middle and lower crust stays close to its zero temperature value of 2.6 W m1 K1 since the
Fig. 9. Comparison of isotherms corresponding to the two models with the mean Moho heat flow 33 mW m2 with horizontal variation 1 mW m2 per 5 km and the upper crust heat production with variation 0.2 mW m3 per 5 km, which differ in the conductivity distribution. The heavy-line isotherms are for the pressure- and temperature-dependent conductivity (see Fig. 10), the dashed-line isotherms for the pressure- and temperature-independent conductivity (see Fig. 3).
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Fig. 10. Pressure- and temperature-dependent conductivity (W m1 K1) used as an alternative to the pressure- and temperature-independent conductivity (see Fig. 3), corresponding to the temperature distribution shown in Fig. 9, i.e. for mean Moho heat flow 33 mW m2 with horizontal variation 1 mW m2 per 5 km and upper crust heat production with variation 0.2 mW m3 per 5 km. The conductivity of the middle and lower crust stays close to its zero temperature value of 2.6 W m1K1 since its assumed temperature dependence is 15 times smaller than for the upper crust.
assumed value of coefficient b is 15 times smaller here than in the upper crust. The conductivity of the middle crust is comparable in both models, whereas in the lower crust the temperaturedependent conductivity is higher, which brings about diminishing temperature differences. The highest Moho temperatures in the temperature-dependent conductivity model are only by few degrees higher than those for the temperature-independent model. The conductivity increase due to its pressure (depth) dependence is very small throughout the crust and it lowers the Moho temperatures along the profile by 10–15 C at most. The fit between the observed and calculated surface heat flows corresponding to the temperature-dependent model, whose temperature and conductivity are shown in Figs. 9 and 10 is given in Fig. 4b. The calculated heat flow values are by about 1 mW m2 lower than for the case with temperature-independent conductivity in the zone of the high surface heat flow. This is due to the refraction of the heat flow outside the region of the low conductivity (Fig. 10) in the upper crust of this zone.
5. Conclusions The numerical modeling along a two-dimensional profile of the crust in southern Portugal has provided estimates of crustal temperatures in the the two main geotectonic units of the area, the
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Ossa Morena and the South Portuguese Zone. Contrary to the previous one-dimensional modeling carried out by Correia and Ramalho (1999), the calculated Moho temperatures do not exceed 700 C in any of the model configurations considered and are consistent with the other geophysical and geological data, in particular with the results of a detailed deep magnetotelluric survey (Correia and Jones, 1997). They range in the interval 400–500 C in the Ossa-Morena Zone and 500–670 C in the South Portuguese Zone. The comparison of Moho temperatures resulting from the different model configurations suggests that the uncertainty of the calculations is of the order 50–100 C. The main reason for the lower temperatures obtained by the present models is the lower Moho heat flow estimate, on the average 28–33 mW m2, based on general ideas of the crustal thermal regime (Vitorello and Pollack, 1980) and independent from a great extent on the heat production values based on the empirical relationship between radiogenic heat production and compressional seismic velocity (see also Ellsworth and Ranalli, this volume). The surface heat flow corresponding to such Moho heat flow values and to the heat production based on the heat production— seismic velocity relationship is by about 20 mW m2 lower than the observed heat flow even outside the zone of its positive anomaly south of the Ferreira–Ficalho Overthrust. This deficit of heat flow was attributed to the underestimated heat production, 0.5 mW m3, of the upper crust. Consequently, the increase of this value to 2–3 mW m3 was necessary to match the observed and calculated surface heat flows. According to measurements on surface rock samples collected in the area, the only rock type, which could ensure such a production is granite with a mean value of 2.9 mW m3. The modeling has also shown that the narrow zone of the high heat flow of 90 mW m2 observed south of the Ferreira–Ficalho Overthrust is most probably of the crustal origin. The heat production of the upper crust necessary to explain it, 4–4.8 mW m3, is appreciably higher than the measured surface values. Because of the complex geological structure of the Iberian Pyrite Belt, which partially overlaps with the zone, it is possible that the observed increase of the surface heat flow has other causes. Thermal conductivity heterogeneity or regional groundwater circulation might be among them, but the experimental data are not sufficient yet to resolve the problem. References Calado, C., 1991. Carta de Nascentes Minerais 1:1.000.000. Portugal. Atlas de Ambiente. Ministe´rio do Ambiente e Recursos Naturais, Direcc¸a˜o-Geral de Recursos Naturais. Chapman, D.S., Furlong, K.P., 1992. Thermal state of the continental lower crust. In: Fountain, D.M., Arculus, R., Kay, R.W. (Eds.), Continental Lower Crust. Elsevier, Amsterdam–London–New York, pp. 179–199. Correia, A., Jones, F.W., 1995. A magnetotelluric survey in a reported geothermal area in southern Portugal. Proc. World Geothermal Congr 2, 927–931. Correia, A., Jones, F.W., 1997. On the existence of a geothermal anomaly in Southern Portugal. Tectonophysics 271, 123–143. Correia, A., Jones, F.W., Dawes, G., Hutton, V.R.S., 1993. A magnetotelluric deep crustal study in south-central Portugal. Studia geoph. geod 37, 331–344. Correia, A., Ramalho, E., 1998. New heat flow density data from southern Portugal: a geothermal anomaly revisited. Tectonophysics 291, 55–62. Correia, A., Ramalho, E., 1999. One dimensional thermal models constrained by seismic velocities and surface radiogenic heat production for two main geotectonic units in southern Portugal. Tectonophysics 306, 261–268.
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