The effect of plate stresses and shallow mantle temperatures on tectonics of northwestern Europe

Global and Planetary Change 27 Ž2000. 23–38 www.elsevier.comrlocatergloplacha

The effect of plate stresses and shallow mantle temperatures on tectonics of northwestern Europe S. Goes ) , J.J.P. Loohuis, M.J.R. Wortel, R. Govers Vening Meinesz Research School of Geodynamics, Utrecht UniÕersity, P.O. Box 80.021, 3508 TA Utrecht, Netherlands Received in revised form 5 April 2000

Abstract Northwestern Europe is tectonically more active, in terms of seismicity, vertical motions and volcanism, than would be expected from its location far from any plate boundaries. In the context of the Netherlands Earth System Dynamics Initiative, we investigated the implications of two recent modeling efforts, of Eurasian plate forces and European mantle structure, for our understanding of the dynamics of these intraplate tectonics. We find that: Ž1. a simple balance between ridge push and collision forces along the southern European boundary does not seem sufficient to explain the observed direction of maximum horizontal compression in northwestern Europe. Our stress model, which imposes dynamical equilibrium on the full Eurasian plate, predicts that collision forces along the African–European boundary are relatively weak and have only a minor effect on the stress field in northwestern Europe; Ž2. seismic velocity anomalies in the shallow mantle imply 100–3008C variations in temperature under northwestern Europe. This regional mantle structure probably plays a significant role in the high level of intraplate tectonic activity and the regional variations in stress and tectonic style. For most tectonically active areas in Europe, observed surface heat flow anomalies coincide with anomalies in mantle velocity. Low velocity anomalies under northwestern Europe coincide with areas of recent volcanism and uplift, but are offset from the regions of maximum surface heat flow. This suggests that the thermal regime of the central European lithosphere is not in a steady state, probably due to changing mantle conditions. The effect of strong variations in lithospheric strength, predicted from the modeled thermal gradients in the shallow mantle, and of dynamic stresses induced by proposed active mantle upwellings may account for Žsome of. the differences between the observed and modeled stress field and will be investigated in future stress models. q 2000 Elsevier Science B.V. All rights reserved. Keywords: intraplate; continental tectonics; plate boundaries; lithosphere; upper mantle; geodynamics

1. Introduction As part of the Netherlands Earth System Dynamics Initiative ŽNEESDI., our project has been con)

Corresponding author. Present address: Institut fur ¨ Geophysik, ETH Honggerberg, 8093 Zurich, Switzerland. Tel.: q41-1¨ 6332907; fax: q41-1-6331065. E-mail address: [email protected] ŽS. Goes..

cerned with understanding the effects of mantledriven processes on tectonics of the Netherlands. Although not located near any active plate boundary, where most tectonic activity is concentrated, tectonic processes have influenced and still influence the evolution of the Netherlands. Of special importance for the Netherlands, that has elevations around sea level, is an understanding of the role tectonic processes play in vertical motions. Kooi et al. Ž1998.

0921-8181r00r$ - see front matter q 2000 Elsevier Science B.V. All rights reserved. PII: S 0 9 2 1 - 8 1 8 1 Ž 0 1 . 0 0 0 5 7 - 1

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estimated the effect of compaction and glacial rebound on vertical movement in the Netherlands for different time intervals, starting in the Miocene and ending with the last 100 years. They find that there is a significant unexplained component of vertical motion that is probably caused by tectonic, mantledriven processes. Tectonic processes further cause fault activity such as in the Roer Graben. Although the Netherlands is not very active seismically, fault structures especially in the south of the country do pose a hazard as illustrated by the M w 5.4 1992 Roermond earthquake Že.g., van Eck and Davenport, 1994.. Because mantle-driven processes operate on the scale of lithospheric plates, they have to be investigated on a European scale. In this paper, we address the effect of plate forces, which has been modeled on the scale of the Eurasian plate ŽLoohuis et al., 1999, 2001., and the effect of thermal shallow mantle structure, which has been inferred from seismic velocities under Europe ŽGoes et al., 2000.. We discuss the results obtained from these separate lines of research with a focus on the Netherlands and northwestern Europe. The term northwestern Europe will be used to refer to the area north and west of the Alps, with the Rhenish Massif and Rhine–Roer Graben system as its central tectonic features ŽFig. 1.. We aim to set the stage for future progress in developing a first order tectonic model for northwestern Europe which, we will argue, hinges on a successful integration of these two lines of work. The Netherlands is cut by a system of roughly northwest–southeast trending faults, starting in the Roer Graben in the southern Netherlands and continuing offshore into the North Sea basin. The grabens in the North Sea were formed in the Mesozoic as part of the Arctic–North Atlantic rifting system. Faults in the southern Netherlands are PermoCarboniferous structures that were first reactivated during the Mesozoic rifting phase that opened the North Atlantic, and caused subsidence in the North Sea Basin. After a phase of transpressional motion in the late Cretaceous to early Tertiary, renewed extension in the Roer Valley Graben started around 36 million years and appears to be continuing today ŽZiegler, 1990, 1992.. The Roer Valley Graben is part of a large-scale system of Cenozoic rifts which cuts through Europe. This rift system ŽFig. 1. includes the Rhine and Leine Grabens in Germany, the

Fig. 1. Map of northwestern Europe showing topography and major tectonic structures: RG—Roer Valley Graben, LR—Lower Rhine Embayment, LeG—Leine Graben, RM—Rhenish Massif, EG—Eger Graben, BM—Bohemian Massif, UR—Upper Rhine Graben, V—Vosges, B—Black Forest, LiG—Limagne Graben, BG—Bresse Graben, MC—Massif Central, TTZ—Tornquist— Teisseyre zone.

Eger Graben in the Bohemian Massif and the Saone, ˆ Limagne and Bresse Grabens in France ŽZiegler, 1992.. Extensional activity along this rift system started in the Eocene in the Limagne and Bresse Grabens in France, and the Rhine Graben in Germany. In the Oligocene, rifting propagated northward into the Leine Graben and the Lower Rhine Embayment, and southward into the Valencia Trough. The Eger Graben was also activated in this time period ŽZiegler, 1992.. The rifting has been accompanied by volcanism, concentrated in the Massif Central, the Rhenish Massif, the Eger Graben and the Valencia Trough, and by uplift, which started in the early Miocene in the Rhenish Massif and the Massif Central, in the mid-Miocene in the Vosges and Black Forest on the flanks of the southern Rhine Graben, and in the Plio-Pleistocene in the Bohemian Massif. The location of both volcanism and faulting appears to be largely controlled by preexisting Variscan and older structures ŽZiegler, 1992.. The synchronous occurrence of uplift and volcanism points to a tectonic origin for the uplift ŽMalzer et ¨ al., 1983; Ziegler, 1992.. Only the parts of the rift system furthest removed from the Alpine front show

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recent or ongoing activity. The Rhenish Massif may have present uplift rates of up to 1.0 mmryear ŽMalzer et al., 1983.. In the Massif Central ŽGranet ¨ et al., 1995. and the northern parts of the Bohemian Massif ŽZiegler, 1992., uplift also appears to be continuing today, while present-day vertical movements in the Upper Rhine Graben are small ŽMalzer, ¨ 1986.. The most recent volcanic activity in the system Ža few thousand to 1-million-year-old. also occurred in the Massif Central, the Rhenish Massif and the Bohemian Massif ŽLippolt, 1983; Downes, 1987; Ulrych and Pivec, 1997.. The stress field in northwestern Europe defined by earthquakes and other stress measurements has a quite consistent N1458E Ž"308. orientation of maximum horizontal compression ŽAhorner et al., 1983; Grunthal and Stromeyer, 1992; Muller et al., 1992.. ¨ ¨ There is, however, a regional variation in the style of deformation ŽMuller et al., 1997a,b, 1992.. Present¨ day seismicity defines zones of active extension in the Rhenish Massif, in the lower Rhine Embayment ŽAhorner et al., 1983. and in the Massif Central ŽMuller et al., 1992.. Eocene–Pliocene extension in ¨ the Rhine Graben has been replaced by dextral strike-slip faulting ŽAhorner et al., 1983.. In Belgium, some Žpredominantly strike-slip. seismic activity occurs along what appear to be reactivated Hercynian structures ŽAhorner et al., 1983.. The Cenozoic European rifting has been attributed to: Ž1. forces at the plate boundaries and a resulting passive upwelling of mantle material ŽIllies, 1975; Sengor, ¨ 1976; Ziegler, 1992.; or to Ž2. the presence of one large or several small active mantle upwellings ŽGranet et al., 1995; Hoernle et al., 1995; Zeyen et al., 1997; Goes et al., 1999.. Tomographic studies show that significant velocity variations exist under northwestern Europe ŽBijwaard et al., 1998; Marquering and Snieder, 1996., indicating that regional upper mantle structure may indeed be a factor in determining surface tectonics. The average orientation of maximum compression has been explained by the combined effect of ridge push forces along the Atlantic ridge system and collision forces along the southern European boundary ŽGrunthal and ¨ Stromeyer, 1992; Richardson, 1992; Golke and ¨ Coblentz, 1996.. The mid-plate occurrence of earthquakes in Europe has been attributed to the existence of old zones of weakness, which are reactivated

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under the current stress field ŽZiegler, 1992.. Regional variations in tectonic style and in the direction of maximum stress have been linked to: Ž1. variations in the force along the southern European boundary, e.g., due to irregular shape of the colliding continental margins ŽIllies, 1975; Sengor, ¨ 1976; Regenauer-Lieb and Petit, 1997.; Ž2. variations in lithospheric strength ŽGrunthal and Stromeyer, 1992.; Ž3. ¨ preexisting structures and a weak lower crust which allow crustal fragments to move independently of each other and of the underlying lithospheric mantle ŽMuller ¨ et al., 1997b.. Lithospheric strength will also be affected by variations in temperature under northwestern Europe. Thus, both plate boundary forces and anomalous mantle structure under central Europe appear to play a role in the tectonic activity of northwestern Europe. Recent stress modeling which treats the Eurasian plate as a whole ŽLoohuis et al., 1999, 2001., and the shallow mantle thermal structure under Europe inferred from seismic velocities ŽGoes et al., 2000., provide some new insights into the dynamics that drive intraplate tectonics in northwestern Europe. In Section 2, we shortly discuss our modeling work. The full EuropeanrEurasian models and the detailed justifications of the modeling assumptions arerwill be discussed elsewhere ŽGoes et al., 2000; Loohuis et al., 2001.. Section 3 concentrates on the implications of these models for our understanding of northwestern Europe tectonics. An integration of the inferred thermal structure and the stress modeling is the subject of ongoing work.

2. Modeling 2.1. Stress model Previous European stress models ŽGrunthal and ¨ Stromeyer, 1992; Golke and Coblentz, 1996. have ¨ considered only the European part of the Eurasian plate with a fixed boundary within, or on the edge of, the stable eastern European platform. These models were able to reproduce the observed large-scale stress pattern. Loohuis et al. Ž1999, 2001. recently developed a stress model for the full Eurasian plate. Considering the whole plate allows for imposing a torque balance on the plate, which provides extra

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constraints on the magnitude of forces that cannot be calculated a priori. Major lithospheric discontinuities, like the Tornquist–Teisseyre Zone or the Ural Mountains, are not part of our first order model. Previous studies that modeled northwestern European stresses opted to have a model plate ending at one of the lithospheric discontinuities. However, shear stresses across these discontinuities may Žpartially. vanish, but normal stresses do not. As a consequence, the boundary conditions across an internal boundary are difficult to assess a priori. Our approach is to start with a mechanically uniform model that includes the whole Eurasian plate to get a sense of the main contributions to the intraplate stress field. Additional complexity like lithospheric discontinuities may be introduced later if a mismatch between observations and model predictions requires it. Loohuis et al. Ž1999, 2001. treat the plate as a thin elastic spherical shell. Forces applied to the plate are: Ž1. ridge push, the gravitational effect of the change in density of the oceanic lithosphere due to cooling with age; Ž2. boundary forces in the direction of relative plate motion between the Eurasian and adjacent plates based on NUVEL1A ŽdeMets et al., 1990.; Ž3. basal drag in the direction of absolute plate motion; Ž4. continental margin forces that account for the difference in gravitational potential between continental and oceanic lithosphere. The ridge push can be calculated directly from age data of the oceanic lithosphere ŽRoyer et al., 1992., and reasonable values for the continental margin force have been previously estimated ŽCoblentz et al., 1994.. The boundary forces can be split in transform resistance forces, which act along the transform segments of the ridges, collision forces along regions of continental collision, and an upper plate resistance for active subduction zones. The direction of the boundary forces and basal drag is taken from the relative and absolute plate motions, respectively. The assumption that the whole plate is not currently undergoing any acceleration, i.e., no net torque, gives constraints for the magnitude of the boundary forces and the basal drag. The no net torque condition only applies when the full plate is modeled, and this provides the motivation for using a stress model that extends far beyond the region of interest in this paper. A finite element method is

used to calculate intraplate stresses that result from the total set of forces, assuming mechanical equilibrium. A constant Young’s modulus of 70 GPa and a Poisson’s ratio of 0.25 are used throughout the model. This approach is similar to the one developed by Richardson et al. Ž1979. and Wortel and Cloetingh Ž1981.. The no net torque condition allows for determining the magnitude of three out of the four unknown forces, i.e., the magnitude of one needs to be assumed. Furthermore, for the Eurasian plate, there is a trade-off between ridge push ŽRP., collision forces ŽCC., transform resistance ŽTF. and the continental margin force ŽCM., because of the close locations of the torque poles associated with these forces. This results in a range of possible models. Additional reasonable assumptions that limit the range of acceptable models are: Ž1. CC and TF should have a positive sign, i.e., they should be resistive forces and Ž2. CC is larger than TF, because of the larger contact area in a continental collision zone than along transform faults. The various allowable sets of forces do not produce very different stress orientations, but do produce different stress magnitudes, and in some places, predict different tectonic regimes. The stress field for an average model, where the magnitude of CC is twice the magnitude of TF and CM is taken equal to 1 P 10 12 Nrm is shown in Fig. 2. In this model, collision forces were not applied along the Italian–Aegean section of the southern boundary, where trench roll-back and back-arc extension are taking place. The absolute plate motion model used for the basal drag is HS2-NUVEL1A ŽGripp and Gordon, 1990.. The stress model shown in Fig. 2 is chosen as a first order reference model and provides a reasonable overall match to the stress observations in Eurasia ŽMuller ¨ et al., 1997a; Zoback, 1992; Zoback et al., 1989.. The full model will be presented, together with an analysis of the sensitivity to various assumptions for the boundary and distributed forces, by Loohuis et al. Ž2001.. In Section 3.1, we discuss the results that are relevant for the western European stress field. 2.2. Thermal structure inferred from seismic Õelocities No direct observations of the structure of the deep lithosphere and mantle can be made, but surface heat

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Fig. 2. European stresses of the reference model of Loohuis et al. Ž1999, 2001.. For this figure, the model is shown up to 708N and 608E, but it covers the whole Eurasian plate. The various types of forces are indicated with different lines: bold dash: transform resistance, solid black: continental collision, solid gray: continental margin, thin line: free boundary. A distributed ridge push and basal drag force are also applied. The maximum and minimum horizontal stress directions are represented by arrows, where outward pointing arrows denote extension and inward pointing arrows denote compression.

flow, images of seismic velocity at depth and gravity measurements all yield information on the thermal andror compositional structure at depth. In Section 2.2.1, we discuss what is known about mantle structure under our region of interest from the most recent European-scale P and S tomographic models ŽBijwaard et al., 1998; Marquering and Snieder, 1996.. In Section 2.2.2, we summarize the constraints provided by the seismic velocities on thermal structure of the lithospheric and sublithospheric mantle down to 200-km depth ŽGoes et al., 2000.. 2.2.1. Seismic Õelocities Local tomographic models, for example for the Rhenish Massif ŽRaikes and Bonjer, 1983. and for the Massif Central ŽGranet et al., 1995., provide detailed information on seismic velocity structure of the lithosphere and sublithospheric mantle, but they do not cover a large enough region to study mantle

processes. The most detailed regional models that do cover the whole European mantle resolve structures on the scale of 100 km and larger. Recent velocity models that have good resolution for our region of interest are the European S velocity model from Marquering and Snieder Ž1996., and the global P velocity model of Bijwaard et al. Ž1998.. The two tomographic models show similar large scale structure ŽFig. 3., especially bearing in mind the very different data sets and inversion techniques that were used to obtain the models. The P velocity model ŽBijwaard et al., 1998. is based on the P, pP and pwP travel time data set of Engdahl et al. Ž1998. and uses an irregular grid in order to minimize the difference in hit count per grid cell. The P velocity model has a lateral resolution of 0.6–1.28 under most of Europe, and shows some smearing along the Žin the upper mantle mostly vertical. ray paths. The shear velocity model

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Fig. 3. P velocity anomalies from the global travel time model of Bijwaard et al. Ž1998. and S velocity anomalies of the European waveform model of Marquering and Snieder Ž1996., under northwestern Europe at a depth of 100 km. Note the different scale used for DVP and DVS to facilitate comparison of P and S velocity anomalies, by accounting for the fact that DVS rDVP is close to 2.0 at this depth.

ŽMarquering and Snieder, 1996. is obtained from fitting waveforms in a time window starting at the S wave arrival and ending after the fundamental mode Rayleigh wave. Using a partitioned waveform inversion which includes the effect of mode-coupling ŽMarquering et al., 1996., a three-dimensional velocity model for the upper mantle is obtained. In central Europe, where the ray density is the highest, struc-

tures on the scale of 0.58 can be recovered ŽMarquering and Snieder, 1996.. In the region shown in Fig. 3, some smearing occurs close to horizontal paths, grossly in north–south direction. Unfortunately, the spatial resolution of the velocity models, as well as the recovery of anomaly amplitudes, decreases from very good, in central Europe, to poorer, as one moves off of continental Europe. Resolution is not good enough to assess uppermost mantle structure under the North Sea with much confidence. Also, neither of these tomographic models provides a very reliable crustal structure. Because of this, and because other information from seismic reflection and refraction experiments Že.g., Prodehl et al., 1992; Remmelts and Duin, 1990. is available on crustal structure, we concentrate on using the tomographic models to study the structure of the deeper lithospheric and sublithospheric mantle, from 50- to 200-km depth. The images of seismic velocity under northwestern Europe exhibit quite a bit of structure ŽFig. 3.. At shallow mantle depths, the most conspicuous anomaly is a low velocity anomaly more or less under the Rhenish Massif. In the shear velocity image, the anomaly is offset somewhat to the north of the Rhenish Massif. This is probably due to smearing along the mainly north–south paths in this region. A second low velocity anomaly of lesser amplitude is seen under the Bohemian Massif. In the depth slice shown in Fig. 3, the anomaly under the Bohemian Massif is clearest in VP , but it is also clear in VS at somewhat larger depths. Other major anomalies that show up in both P and S velocities lie more on the border of our region of interest: a low velocity anomaly under the Massif Central, a high velocity anomaly associated with subduction under the Alps, a low velocity anomaly under the Pannonian basin in the southeast corner of Fig. 3, and a sharp contrast across the Tornquist–Teisseyre zone, which separates Variscan Europe from the Paleozoic shield areas in the east Žclearest in VS in Fig. 3.. Previous work revealed a shallow low velocity anomaly only under the western part of the Rhenish Massif ŽRaikes and Bonjer, 1983.. The larger scale tomographic models used here cannot resolve exactly where the anomalies are located, relative to the surface expression of the Rhenish Massif. However, the low velocity anomaly appears to be a larger scale

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structure that possibly connects to low velocities under the Bohemian Massif, and maybe also to low velocities under southeastern France. Both P and S velocity models find the upper mantle under a good part of Europe, north and west of the Alps, to be relatively slow. The strongest anomalies, under the Rhenish Massif and the Massif Central, continue down to depths of at least 400 km ŽBijwaard et al., 1998; Marquering and Snieder, 1996.. The low velocity anomalies under the two massifs may connect with deeper low velocity anomalies in the upper mantle ŽGranet and Trampert, 1989; Hoernle et al., 1995. and possibly the lower mantle ŽGoes et al., 1999.. Seismic velocities under most of the Netherlands and western France are fast compared to the velocities under Central Europe. 2.2.2. Thermal structure. Previous work has shown that seismic velocities in the uppermost mantle are much more sensitive to temperature than to variations in mantle composition, as long as no partial melt is present ŽGoes et al., 2000; Jackson and Rigden, 1998; Jordan, 1979; Sobolev et al., 1996.. Under the assumption that seismic velocity anomalies could be solely attributed to variations in temperature, Goes et al. Ž2000. inverted the shallow upper mantle P and S wave velocities under Europe for temperature. Independent temperature estimates from P and S velocities were found to be consistent with each other, as well as with temperatures estimated from surface heat flow. It is important in the conversion from velocities to temperature that the effect of anelasticity Žwhich causes damping and dispersion of the seismic waves. is taken into account ŽGoes et al., 2000; Karato, 1993; Sobolev et al., 1996.. Anelasticity depends exponentially on temperature and especially affects how velocity changes with temperature at temperatures close to the mantle adiabat. For the results used here, an anelasticity model which we consider to be an average model Žmodel Q 1 Goes et al., 2000; Sobolev et al., 1996. was assumed. Alternate anelasticity models may yield somewhat lower temperatures in regions where temperatures are close to those of an adiabatic mantle. The temperatures were derived assuming a garnet lherzolite composition ŽJordan, 1979. for the mantle throughout the region. A different composition Že.g., a spinel lherzolite, or a

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more depleted mantle composition. will only have a minor effect on the temperature estimates. Fig. 4 shows several regional geotherms under northwestern Europe estimated from the seismic velocities at depth. Geotherms were averaged over circles with a radius of 150 km, to account for the scale on which anomalies are resolved in the tomographic models. The tomographic images yield estimates of mantle temperature with an uncertainty of "1008C ŽGoes et al., 2000.. Temperatures derived from P and S velocities agree at depths larger than 80 km. At shallower depths, the temperatures inferred from S velocities appear to be influenced by inadequately modeled crustal structure, i.e., topography of the compositional Moho discontinuity ŽGoes et al., 2000.. In the subsequent analyses, we therefore concentrate on the model derived from P velocities, as this model has the better lateral resolution and gives the better estimates of the temperatures at subcrustal depths. The geotherms reflect the pattern already seen in the seismic velocities ŽFig. 3.. Relatively low temperatures are found under the Netherlands and western France, where the geotherms at 200-km depth are slightly below or just reach adiabatic mantle temperatures. The discrepancy between temperatures inferred from P and S wave velocities in these regions is probably due to the decreasing resolution of the velocity models at the edge of the continent. The temperatures under most of Germany and the Bohemian Massif are estimated to be about 2008C higher than the temperatures further west, and reach the mantle adiabat at 150–200-km depth. P and S velocities do yield comparable estimates of temperature for these regions. Mantle structure under the Rhine Graben is warmer than further west, but not as warm as under the Rhenish Massif. Average temperatures under the Rhenish Massif are close to the mantle adiabat up to depths as shallow as 50 km and are 100–3008C higher than the surrounding areas. Very little melt is required to be present under the Rhenish Massif by the velocity models. P and S velocities would not yield similar temperature estimates if a significant amount Žmore than about 1%. of partial melt was present, as the presence of melt affects S waves much stronger than P waves. The melt that we expected to be present because of the recent surface volcanism is probably contained in

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Fig. 4. Average regional geotherms inferred from seismic velocities for northwestern Europe. Averages were taken over the 150-km radius circles shown in the central map. The abbreviations stand for: NL—the Netherlands, NG—Northern Germany, RM—Rhenish Massif, BM —Bohemian Massif, UR—Upper Rhine Graben, WF—Western France. Bold solid lines represent mantle temperatures inferred from P velocity, bold dashed lines represent mantle temperatures inferred from S velocity. Dotted lines give an estimate of the uncertainty in the temperatures, resulting from the uncertainty in the experimental parameters used to determine temperature from seismic velocity ŽGoes et al., 2000.. The thin, larger dashed lines are an extrapolation of P wave velocity-derived temperatures along a steady state conductive geotherm to the surface. The thin, small dashed lines represent a mantle adiabat for a potential temperature of 12808C. The area between the wet and dry mantle solidi ŽThompson, 1992. is shaded.

small pockets Žin some places, temperatures do approach the dry solidus..

3. Implications for the dynamics of northwestern European tectonics 3.1. Northwestern European stress field The modeled stress field in western Europe ŽFig. . 2 has an orientation close to that expected from ridge push alone. Part of the ridge push is balanced

by the CM force as can be seen from the reduction in stress magnitude across the continental margin. The orientation of maximum compression in western Europe, predicted by the reference stress model, is WNW–ESE. Although the modeled stress directions agree with part of the individual observations for northwestern Europe ŽMuller ¨ et al., 1997a, 1992., the average observed stress direction is 208 to 308, more north than the direction predicted from the model. This result is different from previous results ŽGolke ¨ and Coblentz, 1996; Grunthal and Stromeyer, 1992. ¨ obtained with models with a fixed boundary in east-

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ern Europe. The main reason for this difference is the whole plate torque balance, which provides constraints for the resistive forces and does not allow for a strong collision force along the southern boundary. Other sets of forces that satisfy the torque balance, e.g., using a different absolute plate motion model to determine the basal drag, or using different relative magnitudes for the forces, do not yield very different stress orientations in northwestern Europe. Assuming different forces along the southern European boundary only has a small effect on the direction of maximum compression in western Europe, due to the relatively small magnitude of these forces. The ridge push that results from spreading along the MidAtlantic ridge is balanced by forces that are located further east than the fixed boundary assumed in the previous stress models ŽGolke ¨ and Coblentz, 1996; Grunthal and Stromeyer, 1992., instead of being ¨ balanced by forces along the southern European boundary. The observed stress pattern in northwestern Europe ŽMuller et al., 1997a, 1992. has a long wave¨ length character, as expected for a stress field produced by plate forces. Thus, plate boundary forces are probably the main factor controlling the stress orientation in northwestern Europe. However, to explain the direction of the observed present-day stress field, additional complexity, beyond what is incorporated in the reference stress model, seems to be necessary. A different parametrization of the southern boundary of the Eurasian plate, which not only in the Mediterranean but also in Asia is significantly more complicated than modeled so far, may be necessary. Another possibility is that strong gradients in lithospheric strength within the plate cause rotations of the stress field. Such variations in lithospheric structure, for example between the stable eastern Europe platform and western Europe, are known to exist, and significant variations in lithospheric strength are also expected from the variations in mantle temperatures as is discussed in Section 3.3. The modeled stresses in northwestern Europe are predominantly uni-axial ŽFig. 2.. As a result, small changes in the minimum horizontal stress can lead to a variety of deformation styles under the same general maximum horizontal stress. In this respect, changes due to modifications in the forces along the Mediterranean boundary could play a role. Models

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we ran with different forces along the southern European boundary show that this can indeed change the style of faulting in northwestern Europe, but it is unlikely that this sets up a stress field that allows for normal faulting. Previous models of this kind that included the effect of topography ŽMeijer et al., 1997. produce normal faulting at high elevations, but the topography in northwestern Europe is too low for this effect to be significant. Local extension, as observed in the Rhenish Massif ŽBaumann and Illies, 1983. and the Massif Central ŽMuller et al., 1992., ¨ may instead be the result of stresses induced by the relatively hot and thus buoyant mantle under the two Massifs. Variations in lithospheric strength Ždiscussed in Section 3.3. may also play a role in the observed variation in tectonic style in western Europe ŽGrunthal and Stromeyer, 1992; Muller et al., ¨ ¨ 1997b.. 3.2. Upper lithospheric temperatures and surface heat flow To obtain an estimate of temperatures in the upper part of the lithosphere, we extrapolate the mantle temperatures inferred from P wave velocities upward along a one-dimensional steady state conductive continental geotherm ŽChapman, 1986. Žsee Fig. 4.. Crustal properties are not varied laterally, and constant values are used for the depth of the Moho Ž32 km. and the boundary between the upper and lower crust Ž16 km.. The shallow geotherms are not strongly sensitive to the depths of these boundaries ŽChapman, 1986.. Furthermore, the depth of the Moho does not vary strongly under this part of Europe on the scale of the tomographic velocity anomalies. Taking into account some of the stronger small-scale variations in Moho depth Že.g., Prodehl et al., 1992; Remmelts and Duin, 1990. is not warranted by the spatial resolution of the tomographic models. The surface heat flow predicted from the extrapolation of the mantle temperature model is generally consistent with observed European heat flow in both amplitude and location of the anomalies ŽGoes et al., 2000.. In northwestern Europe, however, there are discrepancies between modeled and observed surface heat flow. The range of heat flow values predicted by the shallow mantle temperatures Ž45–90 mWrm2 .

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ŽFig. 5. is similar to the range observed in northwestern Europe Ž45–100 mWrm2 . ŽCermak ´ and Hurtig, 1979; Cermak ´ and Rybach, 1979; Pollack et al., 1993.. Overall, observed high average heat flow values in central Europe are consistent with a relatively warm mantle, and lower heat flow values in the Netherlands and western France are in agreement with a cooler underlying mantle. However, the predicted locations of maximum surface heat flow are over the Rhenish Massifrnorthern Rhine Graben and the Massif Central, while the observed maximum heat flow values are measured in the southern Rhine Graben and north of the Massif Central in the Paris Basin. Surface heat flow reflects both heat generation in the crust, which can be very heterogeneous due to variations in crustal composition, and a contribution from heat flow from the mantle. In areas of active volcanism or shallow fluid flow, the assumption of conductive heat transport may not be justified even at shallow depths and, thus, bias our heat flow estimates. Shallow fluid flow in the thick layers of sediment may be responsible for locally enhancing surface heat flow, for example, in the Rhine Graben, the Paris Basin and graben structures in the Netherlands.

To calculate surface heat flow values from mantle temperatures, we assumed a steady-state thermal structure for the conductive lithosphere. The heat flow values observed for most of the Rhenish Massif are 60–70 mWrm2 , which is not nearly as high as might be expected from the anomalous mantle velocities. Although some of the volcanism in this region dates back as far as the Eocene ŽLippolt, 1983., the assumption of a steady state thermal structure may not be valid ŽR. Van Balen, personal communication, 1999.. Present-day uplift rates, inferred from leveling surveys ŽMalzer et al., 1983., are an order of ¨ magnitude higher than long-term uplift estimates based on dating the Rhine terraces Že.g., Brunnacker and Boenigk, 1983; Meyer et al., 1983.. Furthermore, the most recent phase of wide spread volcanism started only 0.6–0.7 million years ago ŽLippolt, 1983.. Simple 1-D thermal models show that after an increase of 200–3008C in mantle temperature at 50-km depth, it takes around 20 million years to reach a steady state thermal structure throughout the crust again. The sites of activity within the European Cenozoic rift system have moved through time. If the activity is due to active mantle upwellings ŽGranet et al., 1995; Hoernle et al., 1995; Zeyen et al., 1997., a movement of the upwellings as the Alpine front advanced ŽGoes et al., 1999. may be responsible for the migration of activity. Surface heat flow would show a delayed response and therefore reflect previous rather than present upper mantle temperatures. 3.3. Lithospheric strength

Fig. 5. Modeled surface heat flow based on mantle P wave velocities Žblack bars. and observed surface heat flow from the global heat flow data base ŽPollack et al., 1993. Žgray bars.. For the Netherlands, where the global heat flow data base contains only one value, the range in observed heat flow is based on the heat flow maps of Cermak ´ and Hurtig Ž1979. and Ramaekers Ž1992.. The values shown are averages over the circles shown in Fig. 4. The variation is a standard deviation in the case of the heat flow observations, and based on the error estimates of the geotherms ŽFig. 4. in the case of the model heat flow values.

The rheology of lithosphere is strongly sensitive to temperature. The effect of temperature on the lithosphere can be characterized by calculating the thermal and mechanical thickness of the lithosphere. The thermal thickness is a measure for the thickness of the thermal boundary layer, and is usually defined by the depth of an isotherm. The mechanical thickness of the lithosphere is a proxy for the strength of the lithosphere, and is usually defined as the depth at which the strength drops below a reference value, or as the integrated strength divided by a reference strength value. The mechanical thickness is thus a measure of average rheology and depends not only on temperature, but also on the assumed composition and water content.

S. Goes et al.r Global and Planetary Change 27 (2000) 23–38

The thermal thickness, defined by the 13008C isotherm, is quite consistent ŽGoes et al., 2000. with the seismic lithospheric thickness inferred from P waves ŽBabuska and Plomerova, ´ 1992. and displays similar patterns, but is slightly larger than the seismic lithospheric thickness based on surface wave modeling ŽPanza, 1985.. The thermal lithosphere thus defined reaches a minimum thickness of 50–60 km under the Rhenish Massif, compared to a more average central European value of 100–150 km ŽFig. 6.. Under the western Netherlands and western France, the thermal thickness estimated from the tomographic velocities ranges from 150 to larger

33

than 200 km. These variations in thermal thickness are consistent with local tomographic studies, which found significant shallowing of the lithosphere– asthenosphere boundary under the Rhenish Massif ŽRaikes and Bonjer, 1983., while no strong thinning of the lithosphere under the southern Rhine Graben was found ŽGlahn and Granet, 1992.. The mechanical thickness estimates shown in Fig. 6, defined as the depth where mantle strength crosses 1 MPa ŽRanalli, 1994., are based on the extrapolated temperatures inferred from P velocities ŽFig. 4., and used a range of upper crustal, lower crustal and mantle rheologies. References for the rheological

Fig. 6. Estimates of thermal Žblack bars. and mechanical Žgray bars. lithospheric thickness based on the temperatures inferred from mantle P wave velocities. Values are averages for the circles shown in Fig. 4. Thermal thickness is defined as the depth of the 13008C isotherm. Upper and lower estimates are for upper and lower estimates of TP ŽFig. 4.. Mechanical thickness is defined as the depth of the 1 MPa strength contour. An average TP geotherm is used, and the upper and lower estimates are for a range of rheological parameters. The complete strength profiles are shown in solid Žstrong rheology. and dashed Žweak rheology. lines. Negative strength values are tensile, positive strength values are compressional. The rheological parameters used for the upper strength estimate are for dry quartzite in the upper crust ŽJaoul et al., 1984., microgabbro for the lower crust ŽWilks and Carter, 1990., and dry olivine for the mantle ŽKirby, 1983.. The parameters used for the lower strength estimates are for wet quartzite in the upper crust ŽJaoul et al., 1984., Adirondack granulite for the lower crust ŽWilks and Carter, 1990., and wet olivine for the mantle ŽRutter and Brodie, 1988.. The depth of the upper crust–lower crust boundary Ž16 km. and Moho Ž32 km. are the same as used in the modeling of surface heat flow. Note the large uncertainties in strength estimates that result from ill-constrained composition and water content. Additional uncertainty is due to the "1008C uncertainty in temperature.

34

S. Goes et al.r Global and Planetary Change 27 (2000) 23–38

parameters used are given in the caption of Fig. 6. Uncertainties in an appropriate choice of rheological parameters Žspecifically whether to use wet or dry rheologies. and the uncertainties in thermal structure result in a large range of lithospheric strength estimates ŽFig. 6.. Note that thermal thickness can be estimated directly from the thermal mantle model. The estimates of mechanical thickness, however, depend on the shallow thermal structure, and are there for only as good as the assumption of extrapolation along a steady state conductive geotherm. For example, in the Rhenish Massif, the thickness of the thermal lithosphere is probably a reasonable one, but the crustal strength may be underestimated if a steady state geotherm has not been attained, as indicated by the surface heat flow and the recent accelerations of uplift and volcanic activity. The strong thermal gradient between western and eastern Europe along the Tornquist line ŽGoes et al., 2000. should result in a strong gradient in lithospheric strength. Such a gradient may result in less efficient transfer of stresses within the weak parts of the plate and can rotate stress orientations ŽGrunthal ¨ and Stromeyer, 1992.. Previous stress models ŽGolke ¨ and Coblentz, 1996; Grunthal and Stromeyer, 1992. ¨ have used the relative strength and stability of the eastern European platform as the justification for assuming a fixed boundary within it. Future models of the full Eurasian plate, with a reasonable strength difference between western and eastern Europe, will test whether the effect can be large enough to explain the discrepancy between modeled and observed stress orientations in northwestern Europe ŽSection 3.1.. The thermal mantle model further predicts a pattern of relatively low strengths under the Rhenish Massif and relatively high strengths under the western part of the European continent ŽFig. 6.. In previous stress models ŽGrunthal and Stromeyer, 1992., ¨ stresses rotated to a direction more or less parallel to the boundaries of a weak region are introduced in the model domain and, thus, variations in strength may explain some of the scatter in stress orientations in northwestern Europe. In relatively warm regions, the lower crust becomes sufficiently weak to result in a mechanical decoupling of crust and lithospheric mantle, allowing crustal blocks to move independently, as proposed by Muller et al. Ž1997b.. Along the boundaries of ¨

such crustal blocks, various styles of faulting would be expected. Note that defining mechanical thickness by the depth at which the reference strength is crossed is appropriate only if no decoupling upper and lower crust, or lower crust and mantle, occur. Our calculations show that decoupling is certainly possible under the central regions of the area shown in Figs. 3 and 5. Under western France and the Netherlands, temperatures may be too low to establish a decoupling of crust and mantle ŽFig. 6.. Under the Rhenish Massif, our inferred mantle temperatures are so high that the lithospheric mantle has a very low effective strength. 3.4. Topography Topography of Rhenish Massif and Massif Central grossly correlates with the strongest low velocity anomalies in the uppermost mantle, and uplift in these areas continues today ŽGranet et al., 1995; Malzer et al., 1983.. In some parts of the Bohemian ¨ Massif, active uplift is also occurring ŽZiegler, 1992., and this area too seems to be underlain by relatively warm mantle. The southern Rhine GrabenrVosgesr Black Forest region, on the other hand, does not appear to be presently underlain by strongly anomalous mantle, and shows little present-day vertical movement ŽMalzer, 1986.. Since uplift in these re¨ gions has been accompanied by volcanic activity, a thermal origin for the uplift appears likely ŽZiegler, 1992.. The mantle density structure consistent with the temperatures derived from seismic velocities can be used to provide a very first order estimate of the relative isostatically supported topography in the region. Our analyses provide no information on time dependence, i.e., uplift rates or amount of uplift accomplished within a certain time interval. Assuming Airy isostasy, the contribution of the mantle structure between 50- and 200-km depth would give an uplift of around 1.5 km above the Rhenish Massif anomaly relative to the western edge of Europe. Taking into account an average lithospheric flexural strength, this estimate could be reduced by a factor of two to three ŽR. van Balen, personal communication, 1999., and be consistent with the observed topography of up to 850 m above sea level. Again, a steady state assumption may not be representative

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for the Rhenish Massif, and an increase in the strength of the upwelling may be responsible for the recent acceleration in uplift ŽBrunnacker and Boenigk, 1983; Malzer et al., 1983.. The 1–2-km topography of the ¨ Vosges–Black Forest and of the Bohemian Massif can only be partially supported by our inferred present-day thermal structure of the mantle. Larger thermal mantle anomalies in the past and compression associated with the collision in the Alps may have contributed to the uplift there. The thermal anomaly under the Massif Central is large, but here too, the topography can probably not solely be attributed to present-day mantle structure. In many areas, the effect of variations in crustal thickness on topography is much stronger than that of the thermal mantle structure. For example, the Alps and east European Platform are underlain by relatively cold mantle, but taking into account the large crustal thicknesses here would predict relative isostatic uplift rather than subsidence. The effect of relatively cold upper mantle under western France is also partially offset by the thicker crust here, while under the Netherlands, the net effect of crust and mantle predicts isostatic subsidence relative to Germany. Under most of Germany, Moho depths do not vary strongly ŽMeissner et al., 1987., and relative uplift due to anomalous mantle structure under the Rhenish Massif is expected. These calculations should only be taken as a first order indication of the possible effect of mantle structure on topography, since flexural strength, local variations in crustal thickness and deflections of major density interfaces in the mantle play important roles.

4. Conclusions For a mid-plate environment, northwestern Europe is tectonically quite active, with active seismicity and recent volcanism and uplift. The activity is probably attributable to Ž1. stresses associated with the forces acting at the plate boundaries, which can cause earthquakes along preexisting weaknesses, and Ž2. effects of regional mantle processes underneath northwestern Europe. In this paper, we made a first order quantitative assessment of the effects of both types of mantle-driven processes. We discussed the

35

results from two recent plate-scale models that investigated plate forces and mantle temperatures separately. Further modeling that integrates the two types of forcing is clearly necessary for developing a better understanding of the dynamics of northwestern intraplate tectonics and is the subject of ongoing work. A first order stress model for the whole Eurasian plate ŽLoohuis et al., 1999, 2001., based on a set of plate boundary forces that satisfies the no net torque condition, constrains the magnitude for collision forces along the southern European boundary to be relatively small. The modeled orientation of maximum horizontal compression in northwestern Europe is WNW–ESE, which is rotated by 20–308 from the average observed direction of maximum compression ŽMuller ¨ et al., 1997a, 1992.. The modeled stress orientation differs from previous results of partial plate European stress models, which assumed a fixed boundary in eastern Europe ŽGolke and Coblentz, ¨ 1996; Grunthal and Stromeyer, 1992.. These previ¨ ous models used stronger collision forces along the southern boundary, which is not consistent with a whole plate dynamic equilibrium for a homogeneous plate. Thus, additional factors have to play a role, for example, incorrectly modeled complications along the southern boundary of the Eurasian plate, and, perhaps more likely, strong gradients in lithospheric strength induced by the thermal structure of the underlying mantle. In terms of its mainly compressive plate stresses, northwestern Europe does not appear to be an unusual intraplate region. In terms of mantle structure, it is. While the seismic mantle structure under, for example, the eastern US, is relatively fast and uniform, the upper mantle under west and central Europe exhibits significant structure ŽBijwaard et al., 1998; Marquering and Snieder, 1996. with relatively low velocities dominating much of central Europe. Down to at least 200-km depth, these velocity anomalies can probably be largely attributed to thermal structure ŽGoes et al., 2000.. The maximum shallow mantle temperatures under northwestern Europe inferred from P and S wave velocities are located under the Rhenish Massif. The temperatures here are close to that of an adiabatic mantle up to depths as shallow as 50–60 km ŽGoes et al., 2000. and correlate with the occurrence of recent volcanism and accelerated uplift, but not with the observed

36

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surface heat flow pattern. Instead, surface heat flow is high under the southern Rhine Graben ŽCermak ´ and Hurtig, 1979; Cermak ´ and Rybach, 1979., where mantle temperatures do not seem strongly anomalous. These observations may indicate that the thermal structure of the mantle underlying northwestern Europe has changed in the last 10–20 million years, as proposed by Goes et al. Ž1999., and surface heat flow does not reflect a steady state. The stress model predicts the minimum horizontal stress to be close to zero in northwestern Europe. Therefore, even small additional regional stresses may, in combination with existing lithospheric weaknesses, determine the style of faulting, although plate-scale forces are probably responsible for the relatively constant orientation of the maximum horizontal compression. The tectonic regime in northwestern Europe is observed to be rather variable, and includes normal and strike slip faulting in spite of the predominantly compressive stress regime ŽMuller ¨ et al., 1997b, 1992.. Sources of regional stress may be related to complexities along the southern European boundary and to upwelling mantle material under the Rhenish Massif and Massif Central. The orientation of stress and style of faulting may also be affected by lateral and vertical variations in lithospheric strength ŽGrunthal and Stromeyer, 1992; ¨ Muller et al., 1997b.. We predict significant varia¨ tions in lithospheric rheology from the 1008C to 3008C variations in shallow mantle temperature inferred from seismic velocity anomalies by Goes et al. Ž2000.. In the hottest regions under the Rhenish Massif, the lithospheric mantle may have very little strength. Temperatures under most of central Europe may introduce a decoupling of the crust and mantle, but the 100–3008C lower temperatures under the Netherlands and western France could be too low for this to occur. Uncertainties in the calculations of strength are large unfortunately, due to the unknown water content and the uncertainties in the temperature estimates. The topography of the Rhenish Massif may to a first order be consistent with a thermal origin. Topography of the Vosges–Black Forest and of the Bohemian Massif, however, can only be partially explained by inferred present-day thermal structure. Assuming that the active present-day uplift in the Rhenish Massif is the result of the upwelling of

warm mantle material under the region, the same process may contribute to the tilting of the Netherlands ŽKooi et al., 1998. and cause uplift of the Žsouth.eastern Netherlands relative to the western Netherlands. Overall, it seems that plate boundary forces define the long wavelength background for northwestern European tectonics, but many of the regional features probably have to be attributed to regional lithospheric and upper mantle structure and upwelling of mantle material under central Europe.

Acknowledgements We thank Harmen Bijwaard, Wim Spakman, and Roel Snieder for the use of their tomographic models and Ronald van Balen, Paul Meijer and Jan-Diederik van Wees for discussions. We also thank K. Fuchs, B. Muller and an anonymous reviewer for their ¨ comments, which helped us improve the manuscript. All figures were made using GMT3.3 ŽWessel and Smith, 1995.. This is a contribution Žcomponent a1. to the Netherlands Environmental Earth System Dynamics Initiative ŽNEESDI. program, partly funded by the Netherlands Organization for Scientific Research ŽNWO Grant 750-29-601.. This work is part of the research program of the Vening-Meinesz Research School of Geodynamics.

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