Lithospheric thermal structure in the Baltic shield

Geophys. J . Inr. (1991) 106, 611-620

Lithospheric thermal structure in the Baltic shield

V. Pasquale, M. Verdoya and P. Chiozzi Dipart. di Scienze della Terra, Univ. di Genova, 1-16132 Genova, Italy

Accepted 1991 March 26. Received 1991 March 15; in original form 1990 September 21

SUMMARY Steady-state heat conduction modelling was carried out to calculate the crustal temperature field and thermal lithosphere thickness in the Baltic shield. The radiogenic sources at the surface were fixed depending on the age of the crust. Below the uppermost 10 kilometres, a relationship between heat generation and P-wave velocity was applied and seismic structure was used to define individual crustal blocks of specific thermal parameters. Small-scale surface heat-flow density anomalies are interpreted as lateral variations of heat generation within the upper part of the crust, whereas the large anomaly in the southern part of the shield is attributed to an anomalously high mantle heat-flow density. The results are shown on maps outlining mantle heat-flow density, Moho temperature and thermal lithosphere thickness. A relationship between subcrustal temperature and Moho depth was found, except for southern Sweden. The lithospheric thickness is found to exceed 200 km in the Bothnian Gulf-northern-central Finland and part of the Kola peninsula, where temperatures at the Moho are less than 500°C; the temperature gradient at the top of the mantle has, on the average, a value of 7.5 mK m-' and the mantle heat-flow density varies from 19 to 25mWm-2. Towards the south, the lithospheric thickness decreases until it attains a value lower than 100km under southern Sweden. In this latter area, the Moho temperature and mantle heat-flow density are remarkably high, 700"-900 "C and 30-45 mW m-2, respectively, as well as the temperature gradient, which amounts to 10 mK m-'.

Key words: Baltic shield, crustal radiogenic heat generation, lithospheric thickness, mantle heat-flow density, thermal structure.

1 INTRODUCTION

The lithosphere may be specified in a number of ways depending on the physical or chemical characteristics being used to define it. Thermally, it is probably best defined as that portion of the earth (crust and uppermost mantle) with conductive heat transfer overlying an asthenophere with predominantly convective heat transfer and much lower viscosity. For most of the oceanic regions, lithosphere can be defined in terms of a thermal boundary layer at least in a first approximation. For continental areas, the situation is less clear. The evolution of the lithosphere involves both heating and cooling processes of varying intensity and duration, and in tectonically active regions non-conductive heat transfer may dominate (see e.g. Pollack & Sass 1988; Uyeda 1988).

For the stable continental lithosphere, as in the Baltic shield, the situation is clearer. Steady-state heat conduction is predominant as the heat from a transient thermal perturbation associated with tectonogenesis is negligible

(Vitorello & Pollack 1980). Local anomalies of surface heat-flow density can be explained by lateral variations of thermal parameters within the upper crust. The mantle heat-flow density (surface heat-flow density minus crustal radiogenic component) may be dependent on the deepest tectonic processes and its knowledge is therefore important for interpreting the geothermal field on the regional scale (Pasquale, Cabella & Verdoya 1988). Moreover, the conductive geotherms can be extended to the depth at which conditions allow significant solid-state creep, that is to the lower boundary of the thermal lithosphere (Pollack & Chapman 1977; Pasquale 1987).

The topic of this paper is to determine the thermal regime and thermal lithosphere thickness of the Baltic shield for 36 sites whose locations are shown in the map of Fig. 1. Along sites 1-25, which are 60km apart from each other, 2-D thermal modelling is carried out, whereas for the remaining sites the thermal structure is determined through 1-D models. The results provide new information on the thermal state of the crust and uppermost mantle, improving the

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612 V. Pasquale, M . Verdoya and P. Chiozzi

O0 E 1 Z 0 2 k 3 6' I I I I I 1 I I

North cape

Figure 1. Simplified tectonic setting of the Baltic shield (GaB1 et al. 1989) together with the position of the profile along which the 2-D temperature distribution and lithospheric thickness were calculated (solid line). Sites at which crustal P-wave velocity-depth functions were converted into heat generation are indicated triangles. In sites 26-36 lithosphere 1-D thermal models were calculated. 1 = Belomorian province (Archean); 2 = Karelian province (Late Archean-Early Proterozoic); 3 = Lapland Supergroup (Late Archean-Early Proterozoic); 4 = Lapland granulite belt (Early Proterozoic); 5 = Pechenga-Varzuga belt (Early Proterozoic); 6 = Svecofennian province (Early Proterozoic); 7 = Transscandinavian granite porphyry belt (Middle Proterozoic); 8 = Southwest domain (Middle-Late Proterozoic); E.E.P. = East European Platform; K = Kola superdeep borehole.

results previously achieved by us (Pasquale, Verdoya & Chiozzi 1990a; Pasquale, Cabella & Verdoya 1990b) and other researchers (CermBk 1982, 1989; Balling 1990; Baumann 1990). The computed lithospheric thickness is compared to that derived from other geophysical methods.

2 HEAT-FLOW DENSITY FIELD A N D COMPONENTS

We can assume with reasonable approximation that in the shield region under investigation steady-state heat conduction is the only heat transport mechanism in the crust and uppermost mantle. The surface heat-flow density qo is given by the heat originating from the radioactive elements of the crust g c and by a background mantle heat-flow density qm:

where A ( z ) is the heat generation rate down to the depth of the Moho discontinuity 2,.

2.1 Surface heat-flow density

The Baltic shield data set consists of measurements reported by several authors (Parasnis 1975, 1982; Cermik 1979;

Eriksson & Malmqvist 1979; Malmqvist et al. 1983; Kukkonen 1988, 1989).

The geothermal field derived from these data (Table 1) is higher than that reported in the previous heat-flow density map by CermBk & Hurtig (1979). This difference is partly due to the new heat-flow density data covering Finland (Kukkonen 1988, 1989). Moreover, all the data considered in drawing the heat-flow density pattern are corrected for palaeoclimatic effects. The measurements reported by Eriksson & Malmqvist (1979) were corrected according to the authors' comments on the glacial climatic effect. The data reported by Cermik (1979), which pertain to the extreme northern and eastern part of the shield, were corrected on the basis of the palaeoclimatic model recently formulated by Balling (1990).

The western sector of the shield (sites 1-25) is characterized by a generally low regional surface heat-flow density (about 40 mW m-2) in northern and central Sweden, and significantly higher heat-flow density (50-70 mW m-*) in southern Sweden, which is a consequence of a general increase of the thermal regime towards the younger adjacent tectonic provinces of central Europe.

The heat-flow density of the eastern part of the shield (sites 26-36) is of the order of 40n1Wm-~ in the Kola peninsula and increases up to 55mWrr1-~ in southern


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Lithospheric structure in the Baltic shield 613

Table 1. Mean values of crustal radiogenic heat contribution qc in m W m-’ of the sites of Fig. 1 with characteristic range of surface heat-flow density. qr calculated by means of Cermik’s (1989) technique. Standard deviation in brackets. 51 te 90 9c qc

31-36 40-45 18.2 ( 2 . 8 ) 18.0 (2.01

10-20,23-25,30 45-50 20.3 ( 1 . 8 ) 19.5 (1.41

21,22,26-29 50-55 26.8 ( 3 . 4 ) 25.8 (2.71

1-9 > 55 28.0 ( 6 . 2 ) 29.2 ( 5 . 4 1

Finland, following the distribution trend of the surface radiogenic heat generation (Kukkonen 1989).

2.2 Crustal radiogenic component

As one deduces from equation (l), A(z) has to be known in order to determine the crustal radiogenic component qc. The natural radioelement distribution (U, Th and K) in the uppermost crust results from a complex superposition of redistribution processes operating on a different depth and time-scale (CermBk & Rybach 1989). Several approaches of calculation of the radiogenic heat generation in this crustal zone were formulated in literature. Within the surface layer of thickness 10 km, we adopt a technique based on the age dependence of the surface heat generation rate A, (Pasquale et al. 1990b):

A,(pW IT-^) = 3.2 exp (-0.31 t ) , (2)

where t is the geological age in Gyr. A, is assumed to decrease with depth z according to (Lachenbruch 1970):

A(z) =A,exp ( - z / c ) , (3) in which a value of 10 km is assigned to c, the logarithmic decrement of heat generation with depth (Pasquale 1987). From recent determinations, the geological age of the shield ranges from 0.8 to 2.7Gyr and the most parts of the sites have ages over 1.6 Gyr (Pesonen et al. 1989).

In order to treat the geothermal field fine structure which superimposes the regional part, three blocks 10km thick and having A, of 3.5-4pWmP3 were considered. This latter parameter was estimated on the basis of laboratory determinations of acidic rocks of northern Sweden (Lindqvist 1984). Block I, comprising sites 3 and 4, and block 11, sites 7 and 8, are in the southern part of Sweden. The first corresponds to the Transscandinavian granite- porphyry belt, and the second to the Svecofennian domain where there are superficial low-velocity layers constituted by granitoides, metasediments and gneiss rich in mica (Guggisberg 1986). Block 111 (sites 20-22) lies in the northern sector of the profile, where formations are mainly constituted by granitic rocks (Parasnis 1975, 1982). Moreover, according to surface heat production measurements (Kukkonen 1989), in southern Finland (sites 26-30) A, was considered to decrease towards the north from 4 to 3.5 pW mP3. In all these enriched sites, the heat generation distribution versus depth follows equation (3).

Beneath the uppermost crustal layer, A(z) is estimated applying a geophysical-petrological approach, based on the experimental relation between seismic velocity up and radiogenic heat generation rate by Rybach & Buntebarth

(1984). Temperature and pressure effects on up, corresponding to in situ conditions, and composition and geological age of the crust are taken into account as described by Pasquale et al. (1990b). Within individual crustal layers heat production was assumed to decrease exponentially according to (3); for the ith layer

ci = - zi)/ln ((A,/Ai+d, where z is the depth between the upper boundary of the layer zi and the lower one zi+l with heat generation Ai and Ai+l, respectively. In the low-velocity layers heat generation was obtained by linear interpolation between the radiogenic heat production at the top and the bottom (Rybach et al. 1977).

The seismic structure of the crust beneath sites 1-25 was deduced from results along the Fennoscandian seismic refraction profile of 1979 (Guggisberg 1986; Galson & Mueller 1986; Guggisberg & Berthelsen 1987). Crustal thickness varies from about 35 to 55 km and the lower crust has greater thickness where the Moho is deeper. In the remaining sites of the eastern part of the shield seismic data, derived from several deep seismic soundings, are sum- marized by CermAk (1989). The Moho depth is on average about 38 km and it is higher (43 km) in the southwestern part of Finland. Crustal thickness decreases to the north to 35 km and attains a constant value of about 40 km towards the northern and eastern edge of the shield.

Table 1 summarizes the calculated crustal radiogenic heat contribution of the 36 sites divided into four groups according to their typical surface heat-flow density. On the average, qc is about 23mWm-’ and the maximum difference between the groups of sites is around lOmWm-’. For sites with heat-flow density lower than 50 mW m-’ the difference is smaller, as they have nearly the same surface heat production and, in any case, variations of the lower crustal thickness, which has a heat generation lower than 0.1 pW m-’, have negligible influence. For comparison, the crustal heat contribution values deduced by CermAk’s (1989) technique are reported in the same table. According to this author, combining the linear relation between average surface heat-flow density and near-surface heat production (Roy, Blackwell & Birch 1968) with the relation between average surface heat-flow density and reduced heat-flow density (Pollack & Chapman 1977), A, at the top of uppermost 10 km layer is given by 0.04 4,. The mean crustal heat contribution estimated with this technique is slightly lower than that obtained with the method we have proposed. It turns out that for the sites with surface heat-flow lower than 45 mW m-2 the two methods give the same results. A certain discrepancy of about 4 per cent is observed for sites with heat-flow density greater than 50 mW m-’ and, in particular, for southern Sweden, having heat-flow density anomalously higher than that expected on the basis of the geological age.

2.3 Heat flowing out from the mantle

The mantle heat-flow density qm was calculated by applying relation (1). This component of the surface heat-flow density varies from 19 to 45mWm-’ and the mean is about 28mWm-’. Even if the areal distribution of data is inhomogeneous, a well-defined regional pattern is apparent


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Figure 2. A contour

in the contour map of qm (Fig. 2). A wide elliptically shaped zone of low values (<25mWm-’) occurs in the central-northern part of the shield. The minimum of about 20mWm-’ is reached in central Finland. In southern Sweden the mantle heat-flow density varies from 30mWm-’ in the north to 45mWmP2 in the south, in agreement with the regional pattern of the surface heat-flow density.

3 THERMAL STRUCTURE

As lateral inhomogeneities of the thermal parameters exist in the crust and uppermost mantle, the multidimensional approach, when possible, is prefered to determine the deep thermal field of the shield. The temperature-depth computation requires the solution of the Poisson’s equation. The multidimensional problem was numerically treated by applying a finite-element program (Lee, Rudman & Sjoreen 1980).

3.1 Crustal temperature

Figure 3 shows an application of the 2-D method to sites 1-25 beneath which the seismic structure is known. The model, whose horizontal dimension is 1500 km and vertical dimension 60 km, was divided up by columns and rows into a grid of elements defined by nodes of known coordinates. The elements have a width of 30 km and a height increasing with depth z , to provide better resolution of lateral

variations of thermal parameters near the upper part of the crust. The z coordinate of the nodes of the jth row is

zj = { ( j - 1)[2(Nz + 1) -jlIN,(N, + l))h, (4) where h is the thickness of the model, and N, = 30 is the number of rows of the rectangular elements. A Dirichlet- type boundary condition T(x, z = 0) = To@) is applied to the surface, where the horizontal axis of the x coordinate is directed along the strike of the profile and the surface temperature To is assumed to be constant, equal to 5 “C. A Neumann-type condition is used for the vertical boundaries; it was assumed that the sides of the model are so remote from any lateral variations in the heat generation or thermal conductivity that the heat flowing across these boundaries is zero. At the base of the model the heat-flow density corresponds to qm calculated as described in the preceding section.

The variation of thermal conductivity with depth was described by the following relation:

k = ko/(l + bT) , (5)

where ko is the thermal conductivity at surface conditions and b is the temperature coefficient. k is recomputed several times, until in each single element it assumes values in agreement with the final temperatures. The adopted values, corrected for pressure effects as discussed by Pasquale (1987), are listed in Table 2, together with the range of P-wave velocity and the technique used to calculate the heat generation in each crustal material. In the upper mantle the


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I

301 I 1 I I I 1 I 1 I

10 I I 1 I I I I 1 1 I

0 600 1200 krn

F p r e 3. 2-D thermal model and crustal structure of the profile whose layout is shown in Fig. 1. I, 11, I11 = blocks having surface radiogenic heat generation, A,, of 4.0, 3.5 and 3.5 pW mP3, respectively; heat source at depth z (km), A,, is described by A, exp (-2/10), Geothermal flow field qo and heat-flow density data (above): 1 = Parasnis (1975); 2 = Ericksson & Malmqvist (1979); 3 = Parasnis (1982); 4 = Malmqvist er al. (1983). Filled squares indicate data with estimated thermal conductivity. Bars represent standard deviation (for data 1,3,4) and the range of values (for data 2). Hatched line; heat-flow density fleld predicted. Heat-flow density from the mantle qm (below). For details, see text.

heat generation rate was considered to be constant with depth and equal to 0.01 pW m-' (CermBk & Bodri 1986).

In the eastern part, gaps in the available deep seismic profiles did not allow us to perform a multidimensional analysis of the problem. Therefore, in sites 26-36, 1-D thermal models were made to compute the temperature- depth distribution. The thermal parameters are the same as used for the 2-D survey but, in this case, the upper boundary conditions are the mean annual temperature and the heat-flow density at the surface. This technique is described and applied to several areas along the European Geotraverse (EGT) by Pasquale et al. (1990b).

The main results derived from the 2-D profile are reported in Fig. 3. The heat-flow density at the Moho decreases towards the north where it reaches values ranging

Table 2. P-wave velocity range (up), radiogenic heat generation (A) , thermal conductivity ( k ) and temperature coefficient of thermal conductivity (6) used in the thermal models; A,(t), surface heat generation versus geological age; A(up), heat generation versus velocity '(see text).

"P S t r u c t u r e A K b ( k m < - ' I (pW m-'I ( W rn- 'K- l ) (*C-'I

from 23 to 27mWm-*. At great depth isotherms are controlled by the regional anomaly of the heat-flow density field and tend to have a subhorizontal trend toward the oldest part of the shield. The lowest Moho temperature of 500 "C occurs beneath sites 14 and 15 where the thickness of the lower crust is reduced. The highest horizontal geothermal gradients are observed in southern Sweden, where temperature at the Moho reaches its maximum value of 900 "C. The heat-flow density field predicted by the model is in good agreement with the measured one. The maximum discrepancy is of the order of 2 mW m-'.

In the eastern shield crustal temperatures derived from 1-D dimensional thermal modelling are remarkably low. At the Moho the temperature is, on average, 440°C and the lowest value of 320 "C occurs beneath the site 32 where the crust is about 35 km thick.

Figure 4 shows the temperature at the base of the crust. The highest temperatures are found in sites 1-10 having younger age and beneath which the temperature gradient G, at the top of the mantle ranges from 9 to 11 mK m-'. In the other sites the temperature is lower than 600 "C and G, is 6-8 mK m-'.

0 . 00060

o.ooo60

Even if in general a linear increase of the Moho temperature TM with crustal thickness zM is found, the correlation appears to be rather weak (correl. coeff. 0.52).

the sites corresponding to the regional heat-flow density anomaly southwest of the shield. In fact, when sites 1-10 Lower crust 7.0-7.8 A ( V P 1 2.6

Upper mantle > 7.8 0.01 2.6 -0.00025 are excluded from the data analysis one obtains the

First 10 k m < 6.2 A , ( t l 3.0

Upper c r u s t 6.2-6.4 3 .0

Middle crust 6.4-7.0 A l v p 1 2.2 0 The scatter of the data is high because of the influence of

-0.00025


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616 V . Pasquale, M . Verdoya and P. Chiozzi

50

E Y

N

30

Figure 4. A contour

-

-

map of the Moho temperature. The contour interval is 100°C.

following regression line:

zM(km) = 12 + TM/16, 300 5 TM 5 600 "C, (6) with a higher correlation coefficient of 0.72 (Fig. 5).

3.2 Thermal lithospheric thickness

It was assumed that the lower border of the lithosphere, where conductive heat transfer ceases to be dominant, lies

20 c -4 300 400 TM"C 500 600

Figure 5. Moho temperature TM against crustal thickness zM. Linear regressions for the data of sites 11-36 and the 95 per cent confidence limits (hatched lines). Number next to each point is the site code number of Fig. 1.

64'

56O

at the depth at which a sufficiently low viscosity is attained to decouple the lithosphere from the underlying asthenosphere. In order to characterize the rheological behaviour of the upper mantle material we used the ratio between the computed temperature T ( z ) and the mantle melting temperature Tm(z). Variations in this ratio directly correspond to variations in the viscosity of the mantle (Weertman 1970).

The adopted mantle solidus is that deduced by Pollack & Chapman (1977) from numerous available experimental and theoretical mantle melting curves. According to these authors, the value of the temperature expected for the base of the lithosphere in shields must be a fraction of Tm(z) which is lower than 0.85 and perhaps close to 0.75.

Figure 6 shows the 2-D lithospheric thermal model, along the sites 1-25. From the crustal model in Fig. 3, the lower boundary was shifted to the base of the lithosphere which was reasonably fixed to a certain depth and then iteratively recalculated so that the measured and modelled heat-flow density fields have a discrepancy lower than 2 m W m-' and the temperature at the lower boundary is of the order of 0.8Tm(r). In the sites 26-36, where 1-D modelling was performed, the base of the lithosphere was given by the depth at which the conductive geotherms in the upper mantle reach the temperature of 0.8 Tm(z).

Figure 7 displays a map of the thermal lithosphere thickness. The similarity to the mantle heat-flow density map is evident. The lithosphere is thinnest where the mantle heat-flow density qm is high and thickest in shield areas with qm lower than 25 mW m-'.


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Figure 7. A contour map of the lithosphere-asthenosphere boundary as predicted by the thermal models. The contour interval is 50 km.

4 DISCUSSION

In order to evaluate the reliability of these results several tests have been carried out. From a sensitivity analysis of the models it turns out that the computed temperature field largely depends on the assumed parameters such as surface heat-flow density, mantle heat-flow density and upper temperature gradient at the top of the mantle.

mantle melting temperature. If an uncertainty of 10 per cent in the surface heat-flow density is assumed, it follows that the error on the mantle heat-flow density is about 15-20 per cent and on the lithospheric thickness is 15-45 km. The relative error is, on the average, 14per cent on the Moho temperature and of the order of 12 per cent on the


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618 V. Pasquale, M . Verdoya and P . Chiozzi

The different lower boundary conditions assumed for the crustal (Fig. 3) and lithospheric (Fig. 6) models involve temperature differences at the Moho lower than 35 “C. Uncertainties in the criterion for defining the lithospheric boundary [i.e. the range of 0.75-0.85 T,(z)] may produce differences of 5-7 per cent in the calculated surface heat-flow density field. Models are less sensitive to variations in Moho depth and thermal conductivity values, as the heat production in the lower crust is very small and a conductivity variation of 10 per cent should give a mean error of only 6 per cent on the Moho temperature.

The combination of the age dependence of surface heat generation with the petrological-geophysical technique A(up) provides a good basis to find the contribution of crustal radiogenic heat. The application of relation (2) instead of the A ( v p ) technique in the uppermost 10 km layer eludes the problem of the influence of some factors (porosity, fractures, etc.) on P-wave velocity, but does not take into account the redistribution of the radioactive elements mainly due to underground water circulation in microcrack systems (Cermfik & Rybach 1989). Anomalous distributions of heat sources are observed in the northeastern part of the shield by the Kola deep borehole (see Fig. l ) , drilled to the depth of about 12 km (Kremenetsky, Milanovsky & Ovchinnikov 1989). Table 3 shows the radiogenic heat source contribution to heat-flow density calculated from measured concentration of radioactive isotopes in core samples and estimated by the method adopted in this paper. The 7 per cent observed discrepancy between the two models is of the same order of magnitude as the standard error on the heat generation determination from laboratory measurements (Ryback 1988) and as the uncertainty on (2). A significant difference appears in the vertical distribution of the radioactive elements with depletion of the Pechenga sequence and enrichment of the underlaing Archean basement. However, this gives an error of only about 15°C on the temperature at the depth of 11.5 km, as deducible from a 1-D thermal model.

At present the local distribution of heat generation determinations based on petrological study does not allow us to use a direct approach. The choice of the A(up) technique for greater depth was prefered as it allows us to estimate the vertical distribution of heat generation from measured physical properties. The reliability of the crustal radiogenic heat contributions which we calculated is testified by the good agreement with results obtained by applying Cermfik’s (1989) technique (see Table 1).

The whole shield region is characterized by a mean mantle heat-flow density of about 281nWm-~. In the

Table 3. Heat generation models in the Kola peninsula superdeep borehole: (a) estimated on the basis of the method adopted in this paper; (b) from radioactivity measurements of rocks directly sampled (Kremenetsky et al. 1989).

Depth range beo 1 cg 1 c: a 1 5 t r uc t u i - e R a o i o g c n i c heat f l o w

( i ! ” l ( m u 2) ( 2 1 I b l

0 - 6 . 0 4 Pechenga sequence

b.84 - 11.50 O r i h e a n c o m p l e x

n - 11.50

b . 2 2 . 8

4 . 2 6 . 9

1 ” . 4 9 .7

western part of the shield we obtain a mean value of about 23mWm-’ which is slightly lower than that noticed by Cermfik (1989) in the East European platform (26 mW m-’). About 72 per cent of the values falls within the range 21-33mWm-’ found by Sclater, Parsons & Jaupart (1981) as best estimate of mantle heat-flow density of a craton. Values higher than 33 mW m-*, at about 20 per cent of sites, occur in the youngest areas of the shield where elevated geothermal field is observed. Sites having a lower mantle heat-flow density are found in the central part of the shield where surface heat-flow density attains lower values.

The most significant feature of the crustal thermal field is a relatively low Moho temperature obtained beneath the major part of the shield. The lowest Moho temperatures (<4OO0C) occur in the southern part of the Karelian province and part of the adjacent Svecofennides, where mantle heat-flow density and temperature gradient at the top of the mantle reach minimum values (19 mW m-’ and 6 mK m-l, respectively). North of this area of minimum, in the Kola peninsula, and southeast, in the East European platform, a slight increase of the temperature field is observed. In the western part of the shield the temperature increases from north to south with the decrease of the geological age of the crust. Values of temperature at the Moho of 450”-500 “C in the Lapland Supergroup change up to 500”-600 “C in the Svecofennian province of central Sweden.

A distinctively different thermal regime is observed in southern Sweden where intense lateral variation of temperature occurs. Temperatures at the Moho range from 700” to 900°C with the maximum values corresponding to the sub-vertical transition zone proposed by Clowes et al. (1987) which is associated with the juxtaposition of the Transscandinavian granite porphyry belt and the Svecofenn- ian province. South of this zone, the crust is about 12 km less thick than in the northern sector and shows a low-velocity zone within the whole lower crust.

Relation (6) indicates that, except in southern Sweden, the shield has Moho temperatures related to the crustal thickness. In fact, the temperature at the Moho is about 600°C beneath central Sweden, which has a great crustal thickness of about 55 km, whereas it is 320°C beneath site 32 where the crustal thickness is reduced to only 35 km.

The anomalous behaviour of southern Sweden does not follow this type of correlation since this part of the shield is a transition area south of which younger central Europe tectonic units occur. Here, thermal structure and the heat-flow density field observed are in agreement with those previously indicated by us (Pasquale et al. 1988, 1990b) and with those by Balling (1990) and Baumann (1990), who determined slightly lower and higher temperatures at the Moho, respectively.

No correlation between the regional surface geothermal field and crustal thickness of the shield was already pointed out by Meissner, Wever & Fluh (1987). Connections were not even observed with other geophysical features. Post-glacial uplift and large-wavelength free-air gravity anomalies do not correlate well with crustal thickness, since the lithosphere and not the crust is deemed the major structural block in most geotectonic processes (Bungum, Pirhonen & Husebye 1980).

The deep thermal analysis indicates that the shield area is


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characterized by strong horizontal variations of thickness of the lithosphere. The lower boundary of lithosphere, defined on the basis of the ratio between modelled temperature and mantle melting temperature, marks a significative lowering in the viscosity such that lithosphere is decoupled from the underlaying mantle. Although calculation uncertainty has to be considered, the modelling results clearly show a thickening of the thermal lithosphere from the southweast to the northeast. The highest lithospheric thickness (greater than 200 km) is found in the Bothnian Gulf-northern- central Finland area and in part of the Kola peninsula.

A lithospheric thickness greater than 200 km beneath the central-northern shield was also deduced by Pollack & Chapman (1977) from global heat-flow density studies. Sacks, Snoke & Husebye (1979) identified from body wave studies the lithosphere-asthenosphere boundary at a depth of about 250 km. By analysing the surface wave dispersion, Calcagnile (1982) found a thickness exceeding 170 km in the Bothian Gulf-central Finland area. The existence of a thick lithosphere in the northern shield is also supported by geomagnetic induction studies (Jones, Olafsdottir & Tiikkainen 19831, by deep electromagnetic soundings (Hjelt 1984) and, more recently, by seismic refraction surveys of the FENNOLORA profile (Guggisberg 1986).

A marked change in the lithospheric structure of the shield is found beneath southern Sweden, where the thermal lithospheric thickness decreases from 150 to less than 100 km, in agreement with the trend inferred from surface wave data (Calcagnile 1982), and the highest temperature gradient at the top of the mantle and heat-flow density at the Moho are observed.

ACKNOWLEDGMENTS

The Genoa Thermal Geophysics Group is supported by the ‘Consiglio Nazionale delle Ricerche’ and by the ‘Minister0 dell’Universit8 e della Ricerca Scientifica e Tecnologica’.

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Baumann, M., 1990. 1-D and 2-D temperature modelling along the FENNOLORA and EUGENO-S profiles, in Proceedings of the Sixth Workshop on the European Geotraverse ( E G T ) Project: Data Compilations and Synoptic Interpretation, 1989, Einsielden, pp. 401-403, eds Freeman, R . & Mueller, St., European Science Foundation, Strasbourg.

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