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Global and Planetary Change 27 Ž2000. 23–38
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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
24
S. Goes et al.r Global and Planetary Change 27 (2000) 23–38
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
S. Goes et al.r Global and Planetary Change 27 (2000) 23–38
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
25
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
26
S. Goes et al.r Global and Planetary Change 27 (2000) 23–38
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
S. Goes et al.r Global and Planetary Change 27 (2000) 23–38
27
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
28
S. Goes et al.r Global and Planetary Change 27 (2000) 23–38
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
S. Goes et al.r Global and Planetary Change 27 (2000) 23–38
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
29
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
30
S. Goes et al.r Global and Planetary Change 27 (2000) 23–38
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-
S. Goes et al.r Global and Planetary Change 27 (2000) 23–38
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
31
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 .
32
S. Goes et al.r Global and Planetary Change 27 (2000) 23–38
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
S. Goes et al.r Global and Planetary Change 27 (2000) 23–38
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
S. Goes et al.r Global and Planetary Change 27 (2000) 23–38
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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