ORIGINAL PAPER

Michael Schwarz Æ Andreas Henk

Evolution and structure of the Upper Rhine Graben:insights from three-dimensional thermomechanical modelling

Received: 7 October 2003 / Accepted: 19 September 2004 / Published online: 21 December 2004� Springer-Verlag 2004

Abstract The evolution and geometry of the TertiaryUpper Rhine Graben were controlled by a continuallychanging stress field and the reactivation of pre-existingcrustal discontinuities. A period of WNW-ESE exten-sion in the late Eocene and Oligocene was followed bylateral translation from the early Miocene onwards. Thisstudy utilizes 3D finite element techniques to simulateextension and lateral translation on a lithospheric scale.Brittle and creep behaviour of lithospheric rocks arerepresented by elastoplasticity and thermally activatedpower-law viscoplasticity, respectively. Contact elementsallocated with cohesion and frictional coefficients areused to describe pre-existing zones of weakness in theelastic-brittle field. Our results suggest that (1) extensionis accommodated along listric border faults to midcru-stal depth of 15–16 km. Beneath, pure shear stretchingoccurs without a need for localized shear zones in lowercrust and upper mantle. (2) Ductile flow at midcrustaldepth across the graben accounts for the pronouncedhalfgraben morphology. Thereby, the shape of the bor-der faults, their frictional coefficients, and sedimentaryloads have profound effects on the rate of ductile flowacross the graben. (3) Horizontal extension of 8–8.5 kmand sinistral displacement across the rift of 3–4 kmare needed to accommodate the observed sedimentthickness.

Keywords Upper Rhine Graben Æ Basin evolution Æ3D finite element modelling Æ Contact elements ÆDuctile flow

Introduction

This study uses a 3D thermomechanical finite elementapproach to investigate the structure and evolution ofthe Cenozoic Upper Rhine Graben (URG) on alithospheric scale.

The Upper Rhine Graben forms a major segment ofthe Cenozoic Rift system of Western and CentralEurope (Prodehl et al. 1992; Ziegler 1992; Fig. 1a).Limited to the north by the Rhenish Massif and theSwiss Jura to the south, the NNE trending some 40 kmwide graben runs over 300 km from Basel to Frankfurtcrossing boundaries of various tectonostratigraphicunits of the Variscan internides (Fig. 1b). The develop-ment of the URG was controlled by a continuallychanging ambient stress field and pre-existing crustaldiscontinuities (Illies 1975; Schumacher 2002). Periodsof extension alternated with phases of lateral translation,superimposed by lithospheric bulging in the southernURG area which may be related to isostatic andcollisional events in the Alpine domain (Pfiffner et al.1998; Laubscher 2001; Schumacher 2002).

In recent years, the crustal structure in the URGarea was the target of many seismic and geologicalinvestigations (Meier and Eisbacher 1991; Brun et al.1992a). Although they provided detailed informationon the geometry of the rift structure, the style oflithospheric extension below the inferred faults, thedepth to detachment, and the amounts of horizontalextension and lateral displacement are still beingdebated (Meier and Eisbacher 1991; Ziegler 1992;Groshong 1996).

In this paper we are interested in the consequencesof Tertiary extension and strike slip events on the URGstructure and evolution in terms of basin geometry,sediment layer thicknesses, Moho elevation, andshoulder uplift on a lithospheric scale. Our numericalapproach is 3D in order to include the lateral crustalheterogeneities in the URG area and the varyingambient stress field. We regard this study as a first step

M. Schwarz (&) Æ A. HenkGeologisches Institut,University of Freiburg, GermanyE-mail: Michael.Schwarz@geologie.uni-freiburg.de

Int J Earth Sci (Geol Rundsch) (2005) 94: 732–750DOI 10.1007/s00531-004-0451-2

to understand the complex interaction between exten-sion and strike slip in a spatially and temporallyvarying stress field with a pre-existing structuralinventory.

Basin evolution

The geographical position of the URG was determinedby pre-existing crustal discontinuities which are relatedto late Palaeozoic wrench tectonics (Edel and Weber1995). Major shear zones delineating the URG are evi-denced by geophysical data (Edel and Fluck 1989), re-gional kinematic analyses (Krohe 1992), and widelydistributed NNE-SSW trending dyke swarms of Viseanage (Floettmann and Oncken 1992). A further importantstructural trend is revealed by NE to ENE trending faultzones which form the boundaries of Palaeozoic troughsand highs as well as between various tectonostrati-graphic units of the Variscan internides (Chorowiczand Defontaines 1993). Most prominent is the Hun-sruck-Taunus (HT) fault, which represents the boundarybetween the Rhenohercynian and Saxothuringian do-mains, and the Lalaye-Lubine-Baden-Baden fault(LLBB) separating the Saxothuringian from theMoldanubian domain in the south (Fig. 1 b).

In late Priabonian times, a N-S to NNE-SSW com-pression led to a reactivation of NNE- and ENE-ori-ented dislocation zones in a transtensional strike slipmode (Berger 2002; Schumacher 2002; Fig. 2). Atreleasing bends isolated depocentres came into existencewhich were superimposed on fault-bounded late Palae-ozoic troughs (Chorowicz and Defontaines 1993). Be-fore the end of the Priabonian these individual subbasinscoalesced to the NNE trending graben (Schumacher2002).

From latest Priabonian to the beginning of theMiocene, the evolving URG was subjected to exten-sion with r1=vertical and r3=WNW-ESE to W-E(Larroque and Laurent 1988; Fig. 2). Subsidence andsedimentation affected the entire graben and led tosporadic marine connections with the North Sea and theMolasse basin (Sissingh 1998).

From Miocene to the present, the URG area hasbeen governed by a homogeneous NW-SE directedorientation of r1 on a regional scale (Bergerat 1985).The rift was put into a strike slip setting, activated in aleft-lateral sense (Illies 1975). Subsidence and sedimen-tation were restricted to the northern graben segment(Doebl 1967, 1970). The synrift deposits in the centraland southern segments were uplifted and partly erodeddue to northward migration of the Alpine forebulgeinto the southern URG (Laubscher 1992) and trans-pressional reactivation of the central graben segmentaround Karlsruhe (Illies and Greiner 1979). Subsidenceresumed in the southern parts of the URG in Quater-nary and probably already Upper Pliocene timeswhereas the rift shoulders were being uplifted (Bartz1974; Demoulin et al. 1998). Releasing bends in anotherwise compressive regime are accounted forthe recent vertical motions in the southern URG(Laubscher 2001; Schumacher 2002).

Fig. 2 Diagram showing the lithostratigraphy of the URG, basedon Sissingh (1998) and Berger (2002); kinematic phases and stressorientations (black r1, grey r2, white r3) modified afterSchumacher (2002)

Fig. 1 a Tectonic setting of the URG in Central Europe; greyCenozoic graben fill, dark Cenozoic volcanic rocks. b Simplifiedtectonic map of the URG with seismic reflection lines (HTHunsruck-Taunus fault, LLBB Lalaye-Lubine-Baden-Baden zone)

733

Structural and thermal constraints for the numericalmodel

Overall Graben geometry and style of lithosphericextension

In seismic reflection lines DEKORP9S and DEKORP9Nacross the URG prism-shaped Tertiary sediments and anasymmetrical Moho uplift reveal a pronounced halfgra-ben geometry of the rift zone (Wenzel et al. 1991; Brunet al. 1992b; Figs. 3a, b; for location of the profiles seeFig. 1b). In the south, the graben-bounding main fault islocated at the western flank (Fig. 3a). The polarity of thegraben switches to the eastern boundary fault in thecentral segment near Strasbourg (Fig. 3b). Smoothanticlinal structures in the graben interior and listricborder faults are found in the south (Brun et al. 1992b)whereas intricate fracture fields, slight antithetic flexur-ing, and steep dipping midcrustal reflectors are supposedto indicate more planar border faults for the northernpart of the URG (Durst 1991; Meier and Eisbacher 1991;Wenzel et al. 1991).

There are no indications for localized shear zonescutting through the entire crust and upper mantle whichcould support a whole lithospheric simple shear mech-anism of extension in the URG rift zone (Brun et al.1992a). Probably rift-related offsets and faults, respec-tively, can be traced only to depths of 14–18 km near thebottom of the upper crust. Beneath, the lower crustshows a ductile behaviour (Wenzel et al. 1991; Brunet al. 1992b). In contrast to the southern URG, the linedrawing of the seismic profile DEKORP9N exhibits aMoho offset along westward dipping reflectors and ahigher crustal thickness west of the URG (Fig. 3b). Alithospheric shear zone fading in the lower crust wassuggested for this region (Meier and Eisbacher 1991;Wenzel et al. 1991). However, seismic refraction datashow no such crustal thickening beneath the western riftshoulder (Zeiss et al. 1990). This contradiction is likelyto be attributed to a velocity pull down effect of the upto 6 km thick Permocarboniferous sedimentary se-quences in the west compared with surface-exposedgranitoides, gabbros, and gneisses on the easternshoulder. The seismic reflection line DEKORP1Ccrossing the DEKORP9N profile west of the URGdisplays no apparent crustal thickening after taking intoaccount the different seismic velocities of basement rocks

and the Permocarboniferous sedimentary infill (Henk1993).

The Moho map of Zeiss et al. (1990) shows areduction of crustal thicknesses from 30 km elsewheretowards the URG rift zone. The depth of the crust-mantle boundary increases gradually from 24 km in thesouthern rift segment to 26.5 km in the central partsaround Karlsruhe and, finally, to 28 km at the northerntermination of the URG at the HT. The higher Mohouplift in the southern URG may be attributed to aninteraction of rift-related thinning and lithosphericbulging in the northern foreland of the Alps (Gutscher1995).

The amounts of horizontal extension and sinistraldisplacement are only loosely constrained for theURG. Cross sectional balancing in WNW-ESE direc-tion yielded a horizontal extension of 2.5–7 km infer-ring detachment levels in the middle and lower crust aswell as at the Moho (Doebl and Teichmuller 1979;Meier 1989; Groshong 1996). The amount of sinistraldisplacement along the NNE trending faults of theURG rift system from the early Miocene onward is notknown actually. The match of discontinuous NE-SWelongated geomagnetic anomalies crossing the borderfaults in the northern URG (Bosum and Ullrich 1970)might indicate that the horizontal displacement alongone border fault is less than the resolution of 2–3 kmfor the geomagnetic methods. A 1–3 km sinistral dis-placement of Cenozoic age along the eastern borderfault is reported from local kinematic analyses in thesouthern URG (Laubscher 2001) as well. The left-lat-eral displacement of the LLBB of more than 15 kmalong NNE trending faults of the central URG(Fig. 1b) can be attributed primarily to Palaeozoicwrench tectonics (Maass 1988; Edel and Weber 1995)and thus cannot provide a reliable basis for estimatinghorizontal displacements in the Cenozoic.

The surface heat flow of the URG shows values be-tween 60 mW/m2 and 140 mW/m2 (Haenel and Staroste1988). The heat flow variations are caused by a domi-nant convective heat transport system inside the Meso-zoic and Tertiary sediments which was probablyinitiated 2–3 Ma ago (Doebl and Teichmuller 1979;Teichmuller and Teichmuller 1979; Clauser and Villinger1990). The present-day conductive heat flow at theboundary between synrift and prerift sediments is re-ported to be 70–80 mW/m2 (Clauser and Villinger 1990;

Fig. 3 Line drawings of seismic reflection profiles. a Southern profile DEKORP9S, modified after Brun et al. (1992b); grey Cenozoicgraben fill. b Northern profile DEKORP9N, modified after Wenzel et al. (1991); open circles Permocarboniferous sediments, darkPermocarbonian volcanic rocks, grey Cenozoic graben fill

734

Lysak 1992) with temperatures at the Moho rangingfrom more than 600–700�C (Banda and Cloetingh 1992).

Tectonic framework of the northern URG

In the northern URG, the area of Cenozoic subsidenceand sedimentation is considerably widened and com-prises the Mainz basin and the Hanau basin as well(Fig. 4). This fanning out from a single graben to severalfault-controlled sub-basins at the northern terminationof the URG can be ascribed to reactivation of the pre-existing structural grain in this region. Thereby, theeastern border fault of the URG separates two areas ofdifferent structural evolution during the Cenozoic.

In the Mainz basin and the URG north of Karlsruhe,ESE to SE trending fault zones are added to the afore-mentioned NE- and NNE-oriented zones of weakness.This additional structural inventory can be related to thelate orogenic evolution of the Variscides west of theRhenish shear zones (Henk 1993; Edel and Weber 1995).For the Oligocenian rift deposits, the gradual shift of thedepocentre axis towards the western border fault be-tween Karlsruhe and Darmstadt (Doebl 1967; Derer2003) may be caused by extension along both SE- and

NNE-oriented dislocation zones during the Oligocene.In a farther northward direction, the SE trending faultsin combination with NE-oriented Variscan weak zonesresulted in a step by step reduction of the Oligoceniangraben fill and separation of the URG into small-scalegrabens (Anderle 1968, 1974; Plaumann 1991). TheMiocene to recent extension in the northern URG isprobably concentrated along these SE-oriented normalfaults (Anderle 1974; Mayer et al. 1997).

To the east of the URG, Cenozoic extension was alsotaken up by N trending fault zones bordering the Hanaubasin. These faults terminate against a NE trendingtranscurrent fault in the south (Fig. 4). They came intoexistence in the late Permian (Marell 1989) and wereinvolved in the evolution of the URG between the lateChattian and early Aquitanian (Golwer 1968; Diederich1987). The extension broadened to the N trending nor-mal faults of the eastern rift shoulder. Thus, the wid-ening of extension in the northern URG and the erosionof Chattian sediments in the southern URG resulted in amain depocentre in the central graben segment for thisperiod only mimicing a releasing bend under NE-SWdirected compression as proposed by Bergerat (1985)and Schumacher (2002). Therefore, we hold to the evo-lution of the URG with a WNW-ESE to W-E directedextension in the Oligocene followed by a strike slip eventfrom early Miocene to the present.

The widening of extension in the northern grabensegment reduced the amount of horizontal WNW-ESEdirected extension for the graben proper to about 4 km.The structural grain of the Mainz basin and the Hanaubasinmaybe takenup to anadditional 2 km (Meier 1989).

As mentioned before, the shape of the border faults,the depth to detachment, and the friction on faults arenot well known along the strike of the rift. In combi-nation with the rheology of crustal rocks and the ther-mal boundary conditions in the URG area, thesedisputes entail a strong impact on the evolution of sed-imentary sequences and uplift of Moho and shoulders,respectively. Hence, before simulating the evolution ofthe URG and quantifying the amount of horizontalextension and lateral translation, we switch in aparameter study in order to put some limits on the shapeand depth of the border faults, its frictional coefficients,the crustal rheology, and the thermal conditions at theprerift stage.

Stratigraphic constraints for numerical model

The thickness of the graben fill and the uplift of the riftflanks are estimated as a comparative data set to eval-uate the modelling results. Moreover, we use a Mohouplift of some 3.5–4 km of the central graben segmentaround Karlsruhe as a further standard value forassessing the numerical results because this region hasbeen less affected by lithospheric bulging processes in thenorthern Alpine foreland and the widening of extensionto the rift flanks of the northernmost URG.

Fig. 4 Tectonic framework of the northern URG with principalnormal faults; shaded Precenozoic rocks, white Cenozoic rocks,HG Hattersheimer Graben, NG Nidda Graben, dashed line maindepocentre axis during the extension interval; modified after Schad(1962), Straub (1962), Anderle (1968), and Meier (1989)

735

Sediment thickness data

Sediment thicknesses and compaction parameters aretaken from decompaction analyses of Heling (1969) andRoll (1979) as well as from isopach maps of Doebl(1967, 1970; Table 1). We compare the distribution andthickness of the rift sediments calculated with thenumerical model simulations with isopach maps cover-ing the extension interval from latest Priabonian to earlyAquitanian as well as with isopach maps from present-day (Figs. 5a, b). The sedimentation of the middleAquitanian Corbicula beds can be attributed to com-paction of older, extension-related units alone (Roll1979). Therefore, this unit is added to the isopach mapof deposits which are connected to the extension fromlatest Priabonian to early Aquitanian (Fig. 5a). More-over, as we neglect the rift initiation with formation ofisolated depocentres in this study, the thicknesses of theearliest rift deposits and their compaction effects aresubstracted in the isopach maps.

A problem arises from uplift and erosion in thesouthern graben segment from the late Oligocene on-

wards. In this region, decompaction of shaly units sug-gests a maximum overburden of 500 m above the presentoutcrop level and thus an overall sediment thickness of2,100 m for the Oligocenian series (Roll 1979). In thecentral graben segment, the thickness of sedimentaryunitsrelated to the extension interval amounts to over 2,800 m(Fig. 5a). The central and southern graben segmentsshowed nearly the same subsidence before bulging pro-cesses in the southern URG likely initiated in the lateOligocene (Sissingh 1998; Schumacher 2002). Therefore,the erosion of 500 m and non-deposition of about 700 msedimentary rocks may be ascribed rather to a northwardshift of the Alpine forebulge than to a slow-down ofextension rates in the southern URG. As the impacts ofthe Alpine forebulge are neglected in this paper we keepthe value of 2,800 m along the western main fault for thesouthern graben segment as well.

The course of the isopachs (Fig. 5a, b) highlights therepeated shift of the depocentre axis across the graben.In the central graben segment near Strasbourg, the steepgradients of the sediment thickness indicate a fault-controlled change in rift polarity by a northwest-dippingweak zone. This consideration is supported by reducedsedimentary thicknesses just south of Strasbourg due toan uplifted footwall position in this region (Fig. 5a, b).The reversals in rift polarity in the northern and centralgraben segment are already visible in the subsidencepattern of the earliest synrift deposits (Schad 1962;Doebl 1967, 1970) which clearly reflects the influence ofpre-existing structural grain on the structural asymmetryalong the strike of the URG.

Uplift of rift flanks

Erosion of the rift flanks affected predominantly theMesozoic platform strata. The total uplift of the graben

Table 1 Sediment properties

Units/System Compactionparameter(10�3/m)

Initialporosity

Quaternary 0.27 0.45Upper Tertiary I/II 0.35 0.45Hydrobia beds 0.55 0.63Cerithium beds/Corbicula beds 0.48 0.59Niederrodener beds/Freshwater beds 0.27 0.42Cyrena marls 0.60 0.48Rupel Clay 0.53 0.57Pechelbronn beds 0.45 0.56Lymnea marls 0.55 0.55

Fig. 5 Isopachs with 100 mintervals for graben deposits.a Isopachs of series related toextension from latestPriabonian to early Aquitanian;south of the erosion line onlyisopachs of stratigraphical unitscompletely preserved are shown(uppermost Lymnea marls toCyrena marls). b Isopachs ofseries related to extension andstrike slip from latestPriabonian to present-day

736

shoulders can be resolved by backstacking the inferredamount of eroded material onto the present-daytopography and subtraction of the prerift peneplain of200 m above present sea level (Fig. 6). The involvementof provenance analyses and Miocenian pediments spo-radically preserved allow to reconstruct roughly theuplift history. We have to restrict the reconstruction tothe eastern rift flank because Cenozoic tectonic move-ments and erosion of basement at some places on thewestern graben shoulder make an estimate of the erodedrock column highly doubtful for this region. To assessthe eroded sediment column, the thickness of theMesozoic sediment cover in the nearby graben availablefrom Roll (1979 and references herein) and Groschopfet al. (1996) are used. Locations and elevations of mostlymiddle Miocenian pediments are taken from Fischer(1978 and references herein). The values in terms of thealtitude of the prerift peneplain range from less than100 m to more than 300 m above present sea level (Roll1979; Sissingh 1998). Therefore, assuming an averagealtitude of 200 m for this surface, the amount of totaluplift in Fig. 6 shows an uncertainty of approximately±150 m. As rifting was taking place when bulgingprocesses began to affect the southern URG in the lateOligocene the vertical motions till early Miocene can beregarded as an upper bound for uplift related to theextension process.

Uplift related to the extension interval is highest with1,300 m at the latitude of Strasbourg because the mainfault switches to the eastern rift flank in this region. Tothe south, provenance analyses indicate an uplift of theeastern rift flank due to extension in the Oligocene of800–900 m which equals the thickness of the Mesozoicseries in this area (Illies 1962; Duringer 1988). At thewestern border of the southern graben, the abundance ofbasement rocks in the proximal rift deposits suggestsOligocene uplift to be at least a few hundred metersmore than 800–900 m. This estimate is confirmed by thelocation of the main fault at the western rift flank in thesouthern URG.

Modelling approach

A finite element technique is used to simulate the ther-momechanical evolution of the URG during the Ceno-zoic. The numerical equations are solved by using theANSYS (ANSYS Inc., Houston, Texas) softwarepackage. A comprehensive description of the capabilitiesof ANSYS regarding the numerical simulation of geo-dynamic processes is given by Henk (1998). Thenumerical set up consists of a mechanical and a thermalmodel which are coupled via transfer of displacementsand temperatures.

Modelling concept and model geometry

Numerical modelling of the URG evolution is done intwo steps. It covers the extension interval from latestPriabonian (34 Ma ago) to the early Aquitanian (23 Maago) and the following left-lateral strike slip period tothe present-day. The earlier rift initiation with formationof isolated depocentres is neglected in this study. Themechanical boundary conditions for the strike slip eventare restricted to times with sedimentological record.Thereby, the resumption of sedimentation in Quaternarytimes with thicknesses mostly less than 200 m is added tounits which were deposited between 23 Ma and 14 Ma.Consequently, for simulating the strike slip reactivation,the model is loaded with mechanical boundary condi-tions between 23 Ma and 11 Ma.

The 3D numerical model comprises horizontally a250 km long and over 200 km wide cut-out of the URGrift zone between Freiburg and Frankfurt (Fig. 7a).Vertically, the model always extends to the depth wherethe equivalent stress falls below 10 MPa in the mantle, atleast to a depth of 40 km including a 30-km-thick-crust(Fig. 7b). The finite element grid is subdivided bothvertically and horizontally in domains with differentthermal and mechanical properties (Table 2). The Sax-othuringian and Rhenohercynian domains are regardedas a whole. Pronounced gravity highs are additionallyimplemented in the model.

The pre-existing weak zones aremodeled using contactelements which allow discrete displacements betweenindividual fault-bounded blocks. Both border faults, theHT, and aSE trending fault are included in themodel. Thedip, maximum depth, and shape of the border faults areset as variables for the parameter study whereas the HT isimplementedwith a listric shape down to a depth of 16 kmand an average dip of 35� according to seismic profiles andbalancing approaches (Meissner and Bortfeld 1990; Henk1993). The separation of the URG into small-scale gra-bens towards theHT is simulated by a smaller depth of theborder faults of up to 10 kmnorth of the SEorientedweakzone (Fig. 7c). We use the present fault traces as a tem-plate for the orientation of the weak zones in the numer-ical model. To incorporate both border faults into thenumerical model, the distance between them is reduced by

Fig. 6 Along-rift-border section of the eastern rift shoulderindicating total uplift and vertical motions related to the extensioninterval

737

the assumed amount of horizontal extension, varyingalong the strike of the URG between 4 km and 7 km. Theborder faults of theMainz basin and the Hanau basin arenot included in the modelling study because there is nodetailed comparative data set in terms of sedimentthickness and shoulder uplift available for verifying thenumerical results for this region. Furthermore, the dipangle and themaximumdepth aswell as the displacementsalong these faults are only loosely known.

Thermal and mechanical boundary conditions

Extension and strike slip in the URG area are modelledby velocity boundary conditions acting on the left(western) and right (eastern) side of the mechanicalmodel. At the base, the front and the back side of themodel lithostatic pressures (Fig. 7d) are assumed. AWNW-ESE direction of extension perpendicular to theborder faults is imposed, the strike slip is simulated by

displacements which are approximately parallel to theborder faults. The amount and orientation of thevelocity boundary conditions are assumed to be constantboth horizontally and vertically. Isostatic rebound for-ces are simulated by adding linear elastic springs at thebase of the model which counteract the lithostaticpressures. Furthermore, buoyancy forces arising fromthermal dilatation are incorporated.

Temperature in the thermal model is constrained byexponentially decreasing radiogenic heat production inthe lithosphere, surface temperature, and a constantheat flow imposed at the base of the model. No heatflux across the lateral model boundaries is allowed.For several model runs, we change the basal heat flowand radiogenic heat production to allow for the pos-sible range of Moho temperatures of more than 600–700�C in unperturbed central European lithosphere(Banda and Cloetingh 1992; Lysak 1992) with initialsurface heat flows varying between 60 mW/m2 and80 mW/m2.

Fig. 7 Setup for finite element model. a Oblique view of the initial element grid (32,000 nodes, 27,000 elements) showing trend of the faultsand domains with different thermal and mechanical properties; thick lines main faults. b Oblique view with contact elements; upper crustdown to fault tip removed, border faults sketched with planar geometry. c Lateral view of the northern part of the model from east;smaller depth of the border faults between the discrete SE trending fault and the HT. d Plan view of initial element grid with mechanicalboundary conditions; filled arrows direction of extension, open arrows direction of strike slip movements

738

Material laws

In the finite element model, deformation in the brittledomain is described by an elastic-perfectly plasticmaterial law. A power-law is used for steady-state vis-cous behaviour at higher temperatures and pressures.Brittle strength for upper crustal levels is linearly pro-portional to pressure (Byerlee 1978). Recasting thisequation according to Sibson (1974) results in a strengthgradient of 12 MPa/km for extension and 21 MPa/kmfor the strike slip regime in all Variscan terranesassuming a hydrostatic fluid pressure and a frictionalcoefficient of 0.8 for Central Europe. The creep param-eters used for describing the viscous rheology are sum-marized in Table 2. As a reference rheology, parametersare taken from experiments on felsic granulite and

anorthosite for upper and lower crust, respectively(Shelton and Tullis 1981; Wilks and Carter 1990). Creepparameters of wet dunite represent the mantle rheology(Chopra and Paterson 1981). The depth of transitionbetween the localized frictional-plastic behaviour andthe domain of ductile-nonlinear viscous creep is updatedat each time step depending on the temperature, equiv-alent strain rate, and equivalent stress.

The shifts of the depocentre axis across the grabenare taken into account by pre-implementation of ele-ments with an apparent frictional coefficient reduced to0.25. These elements delineate additional pre-existing SEtrending weak zones in the northern URG and the fault-controlled accommodation zone in the central part(Fig. 7a). Moreover, we introduce a strain softeningmechanism for the frictional-plastic field in order to take

Table 2 Thermal and mechanical material properties

Thermal propertiesSpecific heat (W/kg/K) 1300Graben fillThermal conductivity (W/m/K) 3.0–2.6a Seipold (1998)Radiogenic heat production (W/m3) 1.00·10�6Upper crust (incl. Mesozoic cover)Thermal conductivity (W/m/K) 3.0–1.6a Seipold (1998)Radiogenic heat production (W/m3) 4.00·10�6b Royer and Danis (1988)Lower crustThermal conductivity (W/m/K) 1.8–1.6a Seipold (1998)Radiogenic heat production (W/m3) 0.47·10�6 (constant) Henk (1997)Upper mantleThermal conductivity (W/m/K) 2.9–3.3a Seipold (1998)Radiogenic heat production (W/m3) 0.02·10�6 (constant) Henk (1997)

Mechanical propertiesVolumetric thermal coeff. (1/K) 3·10�5Strength gradient for normal faulting (MPa/km) 12Strength gradient for transcurrent faulting (MPa/km) 21Upper crust (incl. Mesozoic cover)Poisson’s ratio 0.24 Holbrook et al. (1988)Young’s modulus (Pa) 0.5·1011 Henk (1997)Density at 273 K (kg/m3)Graben fill (matrix) 2650Moldanubian 2700Saxothuringian/Rhenohercynian 2750 Grosse et al. (1992)Gravity high of Kraichgau 2810 Plaumann et al. (1986)Gravity high of Saverne 2840 Rousset et al. (1992)Creep parameters Exponent Activ. const (K) Strain rate

coeff. (Pa�n/s)Granulite, felsic 3.1 29226 2.01·10�21 Wilks and Carter (1990)

Lower crustPoisson’s ratio 0.27 Holbrook et al. (1988)Young’s modulus (Pa) 0.8·1011 Henk (1997)Density at 273 K (kg/m3) 3000 Gutscher (1995)Creep parameters Exponent Activ. const (K) Strain rate

coeff. (Pa�n/s)Anorthosite 3.2 28745 2.06·10�23 Shelton and Tullis (1981)

Upper mantlePoisson’s ratio 0.30 Holbrook et al. (1988)Young’s modulus (Pa) 1.0·1011 Henk (1997)Density at 273 K (kg/m3) 3400Creep parameters Exponent Activ. const (K) Strain rate

coeff. (Pa�n/s)Olivine, wet 4.7 59899 5.25·10�28 Chopra and Paterson (1981)

Whenever possible, properties specific for Saxothuringian and Moldanubian crust are used; only creep parameters for reference rheologylistedaTemperature-dependentbExponentially decreasing

739

into account the subsidence on secondary faults newlycreated by extension in the graben interior. As, in reality,there is no shift of extension to regions outside thegraben except to pre-existing N trending faults border-ing the Hanau basin, the onset and rate of strain soft-ening can be adjusted in such a way that no migration ofextension into the footwall blocks of the model takesplace. Taking this into account, a more distributedsubsidence in the graben is achieved by a linear reduc-tion of the apparent coefficient from 0.7 to 0.25 with thesecond invariant of the deviatoric strain from 0.2 to 0.6for a identical mesh resolution in all models.

The frictional strength on the embedded faults isdescribed by an apparent frictional coefficient between0.1 and 0.6 without cohesion. We treat the apparentcoefficient of friction as an independent variable whichcan take on a range of values derived from laboratoryexperiments and natural examples (Dieterich and Con-rad 1984; Byerlee 1990; Hill 1992). As there is no indi-cation of a different depth range of the border faultsalong the strike of the URG, an asymmetric extension inthe URG area is seeded by a minimum difference of 0.1in the frictional coefficients for opposite border faults.

Surface processes

The thermal and mechanical effects of erosion and sed-imentation are incorporated by activation and deacti-vation of predefined elements. This approach allows forimplementation of the physical properties of the addedor eroded material.

Sedimentation always kept up with subsidence duringthe evolution of the URG. Water depths presumablynever exceeded 200 m (Derer 2003 and references here-in). Consequently, the accommodation space createddue to extension in the numerical model is completelyfilled with sediments which are compacted to an expo-nential porosity-depth relationship (Athy 1930) at thenext time step. The compaction parameter and initialporosities are listed in Table 1. The sedimentation levelvaries over time and along the length of the rift zone andranges between 100 m and 200 m above present sea levelincreasing from north to south. Sea level variations (Haqet al. 1988) for periods of marine conditions, a preriftpeneplain with an altitude of 200 m, and a maximumwater depth of 200 m are included into the calculationsof the sediment accommodation space.

The erosion in the URG area is modelled as beingproportional to regional elevation. The Oligocene riftphase with a uplift of 800–900 m on the eastern grabenshoulder of the southern URG led to a concomitanterosion of primarily Jurassic rocks to the tune of300±100 m whereas younger rift sediments containalmost exclusively older sedimentary units and basementrocks (Illies 1962; Fischer 1978; Duringer 1988;Groschopf et al. 1996). In order to reach this erodedcolumn for the Oligocene rift phase, we simulate erosionin the URG area by removing a fraction of five percent

of crustal rocks uplifted above the erosion level of 200 mat each time step and apply this functional relationshipto the present day. The proportionality between eleva-tion and erosion rate implies no erosional retreat of therift flanks. In fact, the regression of water divides ismarginal in the northern URG area which is probablyunaffected by the aforementioned lithospheric bulgingfrom Miocene onwards. Thus, we understand the ero-sion being a function of elevation as a sufficientapproximation of the nonlinear diffusion process of rockdenudation in the URG area.

Results of the parameter study

Herein, we outline the first order-effects of the shape anddepth of the border faults, its coefficients of friction, therheology of crustal rocks, and the thermal boundaryconditions on the evolution of the URG. The velocityboundary conditions are applied in such a way that thehorizontal extension reaches the reported maximum va-lue of 7 km at the end of the extension phase and 3 km atmost along the border fault for the strike slip interval. Werestrict the presentation and discussion of the results tocriterions used here to approximate the parameters for asubsequent best fit model. These criterions are a maxi-mum subsidence of about 2,800 m on the main faults forthe extension interval, an uplift of some 800–900 mon theopposite graben shoulders, and an upwelling of the crust-mantle boundary of 3.5–4 km. In order to keep within thelimits, the results of the parameter study are primarilyreferred to a vertical cross-section running throughFreiburg at the southern edge of the 3Dmodel. There, themaster fault is found at the western graben flank.

At first, the main faults are assigned with a frictionalcoefficient of 0.3. The friction on their stronger coun-terparts is 0.4. A basal heat flow of 33 mW/m2 leads toan initial surface heat flow of 73 mW/m2, a temperatureof some 660�C at the Moho, and to a maximum depth of15 km for the border faults.

Shape and depth of border faults

A model with planar faults is thought to stand for thenorthern URG with its intricate fracture fields and steepdipping reflectors visible in the seismic line of Durst(1991). The planar faults dip 60�. The listric faultgeometry is taken from the line drawings of the seismicprofiles DEKORP9N and DEKORP9S (Wenzel et al.1991; Brun et al. 1992b). The curved faults are imple-mented with the same dip angle of 60� near the surfacebut flatten out to 30� towards the depth of 15 km.

For both fault geometries, the extension producesboth footwall and hangingwall tilting and an asymmetrywith respect to subsidence, shoulder uplift, and Mohoelevation (Fig. 8a). The use of listric faults amplifiesconsiderably the structural asymmetry of the rift incontrast to the ‘‘planar case’’. At the end of the exten-

740

sion phase, the overall subsidence of 2,300 m along thewestern listric border fault is up to 500 m higher than fora planar fault whereas a complementary reduction of thesediment thicknesses of a few hundred meters is pre-vailed near the eastern border fault. But on a whole, theaccommodation space created for sediments is higher forthe ‘‘listric case’’. Mantle uplift is more localized in thelistric-fault model and the Moho is elevated by about3.3 km, 500 m higher than in the ‘‘planar case’’.

The higher accommodation space for the ‘‘listriccase’’ is due to a stronger ductile flow in the viscousregime at depths of 15–18 km. The lower average dip oflistric faults provokes a lateral displacement between thelocation of the maximum subsidence on the hanging wallof the western master fault and the highest uplift of theMoho being more easterly beneath the downthrownblock. This results in a higher maximum subsidence forthe ‘‘listric case’’. Intensified by sedimentary loading, astrong ductile flow between the maximum depth of theborder faults at 15 km and the top of the lower crust at18 km is initiated which is oriented across the grabentowards the eastern graben flank. Consequently, a listricwestern border fault canalises the ductile flow in thiszone of high mobility in an eastward direction and leadsto stronger anticlockwise rotation of the hangingwallblock (Fig. 8b). In contrast, the upward directed iso-static restoring forces dampen the subsidence in thehanging wall of the planar western fault on account ofthe close spatial relationship between basin and Mohouplift. These observations likewise hold in a laterally

reversed manner for a profile across the central grabensegment around Karlsruhe where the master fault islocated in the east. The higher and narrower uplift of theMoho for the ‘‘listric case’’ is a consequence of astronger interference of upward directed isostaticrestoring forces.

Listric faults cause a lower internal deformation in thedownthrown block (Fig. 8c). This results in a later wid-ening of extension to the graben interior than for planarborder faults. The weakening starts to develop after 3–4 Ma for the ‘‘planar case’’ and up to 1 Ma later formodels with listric faults. For both fault geometries, thestrain softening provokes most of the strain to beaccommodated in the graben proper. A pronouncedviscous hardening below both faults comes into exis-tence. The creep regime is soon replaced by a frictional-plastic one down to 15–16 km. This change in rockbehaviour had to be allowed for by a pre-implementationof contact elements down to a depth of at least 15 km.

A model with faults being listric in the elastic-brittlefield of the upper crust comes closer to reality in respect ofthe rift geometry, the amount of subsidence and Mohouplift along the entire URG. In contrast, the ‘‘planarcase’’ results in a underestimation of sediment thicknessesand Moho uplift along the length of the rift zone. Mod-elling with planar faults being steeper or deeper reachingwould lead to higher thicknesses of the synrift deposits butconcomitantly to a even lower Moho uplift. Conse-quently, we regard only listric faults along the entire strikeof the URG for the remainder of the parameter study.

Fig. 8 Cross sections at thesouthern front of the numericalmodel showing a comparison ofplanar and listric faults; thickfault trace main fault; allfigures with no verticalexaggeration. a Verticaldisplacements with 100 mintervals at the end of theextension phase; shaded area ofsubsidence in Precenozoic units;graben fill not shown.b Displacement vectors at theend of the extension phase, lightgrey graben fill, dark grey areaof subsidence in Precenozoicunits. c Total equivalent strainin percentage at the onset ofstrain softening in the grabeninterior for planar faults, i. e.about 4 Ma after beginning ofextension

741

Friction on border faults

For several model runs, we changed the frictional coef-ficients on both border faults but kept a constant dif-ference of 0.1 over time for each model run. As expected,lower friction forces lead generally to a narrower, butlower Moho uplift, and to rugged rift topography withhigher subsidence on both faults (Fig. 9). But if thefriction coefficient is assumed to be less than 0.4 on themain faults, the rate of ductile flow across the grabenincreases significantly and modifies this general correla-tion. In consequence, the subsidence on the main faultsis higher whereas both graben and shoulder are upliftedat the opposite rift flank. If frictional coefficients on bothfaults are lower than 0.4, the total horizontal pressuregradient driving material towards the stronger faultdiminishes again. Thus, the difference in subsidence be-tween the border faults reaches an asymptotic value forlow friction coefficients, given a constant difference be-tween the coefficients over time.

Several models were run with friction coefficients onthe stronger faults up to 0.6 whereas the friction coeffi-cient on the main faults were held constant at 0.3. Thehigher differences in the coefficients enforce slightly thesubsidence and shoulder uplift on the main faults butdiminish exceedingly the same items on the strongercounterparts to values of only 200–300 m. In contrast,for a combination of 0.3/0.4 or lower, the uplift of 650–700 m on the graben shoulders lying opposite to themain faults comes closer to the value of 800–900 mestimated before for these regions. Therefore, we regardthe combination of 0.3/0.4 as an upper bound for fric-tion coefficients in order to provoke both a considerableshoulder uplift and the accumulation of high sedimentthicknesses in combination with a pronounced halfgra-ben morphology.

Lower crustal rheology and thermal conditions

The thermal conditions at prerift stage and creepparameters for the lower crust are varied in order tocover the range of compositions, surface heat flows, andMoho temperatures proposed for the lithosphere ofCentral Europe. On the basis of geochemical observa-tions, the lower crust is supposed to be of granuliticcomposition with a high content of quartz (Mengel et al.1991). This would favour the application of a lowercrustal rheology to be even weaker than used for thereference model. Otherwise, some geophysical and geo-dynamical considerations imply that the lower crust maybe as stiff as the upper mantle (Holbrook 1988; Seyferthand Henk 2000).

Compared to the reference rheology, we adopt aweaker lower crust using a flow law for felsic granulite(Wilks and Carter 1990). Creep parameters for dryColumbia diabase (Mackwell et al. 1998) are the otherend member representing a very stiff lower crust.Moreover, the flow law data for a another dry diabase

(Strehlau and Meissner 1987) is used as an intermediaterheology which is still considerably stiffer than the ref-erence rheology for a given temperature and strain rate.Background heat flow and radiogenic heat productionare adjusted in order to vary the initial surface heat flowfrom 60 mW/m2 to 80 mW/m2 leading to Moho tem-peratures from 600�C to 700�C and to a maximum depthof the upper crustal faults between 12 km and 18 km.

For all variations of creep parameters and initialthermal conditions discussed here, extension and strikeslip in the URG area leave the heat flow at the boundarybetween synrift and prerift sediments and the tempera-ture at the uplifted Moho nearly unchanged (Fig. 10).On a regional scale, the thinning of radiogenic layersbelow the upper crustal faults leads to a negligiblecooling of 5–15�C from midcrustal depth down to thebase of the model. The insulating effects of the graben

Fig. 9 Subsidence of the graben and uplift of the Moho as afunction of the friction coefficient on the border faults; difference incoefficients held constant over time for each model run, shaded arearange of uplift of the Moho in the central URG assuming an initialcrustal thickness of 30 km, dotted line maximum sedimentthicknesses related to extension

742

fill keep the temperatures in the upper crust of the modelnearly constant.

Numerical modelling with the end member rheologiesfails to reproduce the thermomechanical evolution of theURG. The weak felsic granulite is squeezed out rapidlyfrom underneath the main fault towards the grabeninterior causing a breakdown of the model runs after4–5 km of horizontal extension. This observation of arheology being too weak for the URG area also holdsfor a combination with a low initial surface heat flow of60 mW/m2, 600�C at the Moho, and frictional coeffi-cients of more than 0.4 (Fig. 11a). In contrast, the stiffColumbia diabase in the lower crust after Mackwell(1998) is considerably hardened beneath the basin, evenfor a high initial surface heat flow of 80 mW/m2 and700�C at the Moho. A frictional-plastic regime comesinto existence in the lower crust (Fig. 11b) but NNEtrending shear zones in the upper levels of the lowercrust are not observed in the URG area.

The main features of the intermediate rheology arecompared with the model in which the weaker referencerheology was used for assuming the initial surface heatflow of 73 mW/m2. We note that the thermomechanicalevolution of the URG rift zone is almost identical. Asfor the reference rheology, the region at depths of 15–18 km below the graben block shows significant viscoushardening (Fig. 11c) due to an outward directed ductile

flow. This lateral flow in combination with low frictioncoefficients still controls the subsidence of the grabenblock and the structural asymmetry of the rift system.The higher stiffness of the lower crust only results in amarginal reduction of subsidence on the main faults, anincrease of subsidence on the stronger counterparts, andto a higher uplift of the Moho of 200 m at most.Modelling the evolution of the URG only allowsexcluding very weak and very stiff lower crustal rheologyin Central Europe as well as an initial surface heat flowbeing lower than the present-day’s 70–80 mW/m2. Be-tween these end members, the rheology of the lowercrust is of secondary importance for the thermome-chanical evolution of the URG.

To summarize the parameter study, listric borderfaults down to a depth of at least 15 km and a combi-nation of around 0.3/0.4 for the frictional coefficients arethe structural limits for the best fit model. Furthermore,we keep up the creep parameters of anorthosite to rep-resent the lower crust for an initial surface heat flow of75 mW/m2 corresponding to a Moho temperature ofabout 680�C. Taking into account additionally therotation of the border faults to lower dip angles duringextension, the contact elements feature a dip of at least70� near the surface and flatten out to about 30� at thedepth of 15 km.

Numerical results of the structure and evolutionof the URG

The numerical model is designed to simulate the evolu-tion of sedimentary sequences in the URG rift systemincluding deposition and compaction. As the maininterest of the study lies in the quantification of theamounts of horizontal extension and strike slip dis-placements we concentrate on the presentation ofnumerical calculations which refer to the comparative

Fig. 10 Modelled present-day temperature field for a basal heat

flow of 33 mW/m2 in the URG area at the southern front of thenumerical model

Fig. 11 Strength profiles presented as equivalent stress for different creep parameters in lower crust at the southern front of the numericalmodel; path for profile along the main fault in the west, then in vertical direction down to the base of the model; dashed line strengthenvelope of reference rheology (anorthosite and surface heat flow of 73 mW/m2, 660�C at the Moho); in all figures friction coefficient of0.3 on the main fault. a Strength envelope of weaker felsic granulite with surface heat flow of 60 mW/m2, 600�C at the Moho after 4 km ofhorizontal extension. b Strength envelope of very stiff dry diabase (Mackwell et al. 1998) with surface heat flow of 80 mW/m2, 700�C at theMoho at the end of the extension interval. c Strength envelope of ‘intermediate’ dry diabase (Strehlau and Meisner 1987) with surface heatflow of 73 mW/m2, 660�C at the Moho at the end of the extension interval; dotted line strength envelope for reference rheology at theonset of extension

743

data set at the end of the extension period and at pres-ent-day. For modelling the extension period, the left andright boundary of the model are moved horizontallywith a constant velocity of 1.1 mm/a from 34 Ma to29 Ma (uppermost Lymnea marls to Cyrena marls). Thiscorresponds to an average strain rate of 1.5·10�16/s. Thedeposition of Chattian units (Bunte Niederrodener Bedsand lower Cerithium Beds between 29 Ma and 26 Ma) ismodelled by moving the lateral sides of the model with aconstant velocity of 0.7 mm/a resulting in an averagestrain rate of 1.2·10�16/s. The remaining 3 Ma of theextension period (Upper Cerithium Beds and CorbiculaBeds) are simulated again by a velocity boundary con-dition of 1.0 mm/a. We keep up this value for the fol-lowing strike slip event between 23 Ma and 11 Mawhich also includes the Quaternary resumption of sedi-mentation in the URG area.

Displacements at the end of the extension interval (latestPriabonian—early Aquitanian)

In the central graben segment around Karlsruhe, a hor-izontal extension in WNW-ESE direction between 8 kmand 8.5 km is required to accommodate the observedsediment thickness (Fig. 12a). There, a short-wavelengthuplift of the Moho has developed with a maximumamount of some 4 km (Fig. 12b). To the south, the de-pocentre axis and the line of maximum Moho upwellingshift at the latitude of Strasbourg to the western rift flankof the southern URG where the computed sedimentthicknesses are a mirror image of the central grabensegment for the same amount of horizontal extension(Fig. 12c). The numerical model correctly simulates thegradual shift of the depocentre axis between both riftflanks in the northern URG as well. However, it does notreconstruct the sediment thickness in the URG north ofKarlsruhe. The computed thicknesses are mostly lowerthan observed sediment thicknesses, in some places morethan 600 m. In contrast overestimation of the sedimentthickness of up to 600 m is computed in the directhangingwall of the SE trending fault where the subsi-dence is enhanced at the junctions between the borderfaults and the SE trending weak zone. The horizontalextension and Moho uplift considerably fade in thenorthernmost URG towards the HT to some 4 km and2 km, respectively (Fig. 12b). By a shallowing of theborder faults to the north and their termination at theHT, the northern part of the model experiences morepure shear stretching and a broad but small uplift ofthe crust mantle-boundary comes into existence in thisregion.

Despite a orthogonal extension in WNW-ESE direc-tion, sinistral and dextral displacements of up to 1,500 malong each border fault take place (Fig. 12d). In thecentral and southern graben segment, the strike slipmovements are concentrated on fault sections where thefrictional coefficient along the strike of a border faultchanges due to a switch of the rift polarity. There, inside

the graben block, the lateral displacements are alwaysdirected towards the weak sector of the border fault withthe highest thickness of the graben series. In return, aductile flow at depths of 15–18 km parallel to the borderfault is directed out of the zones of high sedimentthickness. This lateral flow amplifies the structuralasymmetry along the strike of the URG for the thirddimension as it does two dimensionally across the gra-ben for opposite border faults with different frictionalcoefficients. The horizontal displacements just south ofthe SE trending weak zone of the northernmost URGare caused by extension both on the border faults andthe discrete SE trending fault.

Displacements during the strike slip phase (middleAquitanian to present-day)

For a horizontal extension of 8–8.5 km, a sinistral dis-placement across the entire graben of 3.5–4 km is suffi-cient to create the additional accommodation space forthe maximum sediment thickness of 3,400 m in thenorthern URG (Fig. 13a). The left-lateral motions reacha maximum value of nearly 3.5 km on the eastern borderfault south of Frankfurt. Including the few hundreds ofmetres of sinistral displacement on this fault segmentduring the extension phase, the total maximum amountof sinistral translation along the eastern border fault canbe stated to be around 4 km. The left-lateral movementsalong the western border fault come up to 1.5 km at thesame latitude. This value can be cleared with the sameamount of dextral displacement resulting from theextension phase (Fig. 12d). Thus, no horizontal dis-placement along the western border fault is recorded atthe latitude where the sinistral displacement along theeastern fault is highest.

The modelled subsidence is primarily concentrated tothe hangingwall of the SE trending weak zone and theeastern border fault of the northern URG. It is highestat the junction between both fault orientations where theaddition of 1,600 m thick sediments related to the strikeslip period sums up to the total amount of 3,400 m(Fig. 13a). From this junction to just north of Kar-lsruhe, the eastern border fault is reactivated in atranstensional mode on account of its N-S orientation.In contrast, the western border fault strikes NNE-SSWat these latitudes leading to sediment thicknesses of only700 m for the strike slip period. Thus, the numericalexperiment shows an eastward shift of the depocentreaxis during the strike slip interval which is consistentwith natural observations for the northern graben seg-ment (Schad 1962; Doebl 1967; Derer 2003).

An uplift of about 100 m for the central graben isobserved (Fig. 13a). The central graben segment whichdisplays a more northeasterly orientation acts as arestraining bend in the numerical model. In contrast, theNNE-SSW striking southern URG is reactivated in asinistral strike slip mode with only marginal verticaldisplacements.

744

In the numerical model, the strike slip mode in theURG area leads to an additional Moho uplift of morethan 1,000 m (Fig. 13b). The upwelling trends NW-SEsouth of the normal fault which crosses the URG withthe same orientation. Neither the amount nor the trendis seen on the map of Zeiss et al. (1990).

Shoulder elevation

During the extension interval of the numerical simula-tion, the rift shoulders are elevated to maximum valuesof around 1500 m above the prerift-peneplain near the

western main fault in the Moldanubian domain of thesouth and 1,400 m along the eastern master fault in theSaxothuringian farther north, except some 1,000 m forthe area of the Kraichgau gravity high (Fig. 14). Theuplift of the rift flanks along the secondary border faultsis calculated to be 800–900 m. For the following strikeslip period, only the altitude of the graben flanks in thenorthern parts of the model change noticeably. Itreaches more than 500 m just south of the junction be-tween the SE trending weak zone and the eastern borderfault. The western graben shoulder at the same latitudedisplays a maximum uplift of 300 m for the strike slipinterval.

Fig. 12 Plan view of numericalmodel at the end of theextension interval. a Lines ofsame horizontal displacement inmetre (below zero to WNW;above zero to ESE). b Uplift ofMoho in metre as dashed lines.c Isopachs in 100 m; shadedarea where sediment thicknessis overestimated at 400 m andmore, dotted area wheresediment thickness isundervalued at �400 m andless; sediments related to theextension period only partlypreserved south of the erosionline. d Horizontal displacementin 100 m parallel to the borderfaults

745

A comparison of the results with the calculationdisplayed in Fig. 6 allows scrutinizing the uplift historyof the eastern graben shoulder in more detail (Fig. 15).The modelling results show a good agreement in esti-mated uplifts related to extension from Freiburg to nearKarlsruhe if the Cenozoic fault system north of Freiburgis neglected. Between Karlsruhe and the northern ter-mination of the Kraichgau gravity high, the modelledshoulder uplift approximately matches the total uplift.However, farther north, the total uplift was overesti-mated at about 500 m where significant subsidenceduring the strike slip period occurred. The overestima-tion equals the uplift of the eastern flank in the strike slipinterval from early Miocene to present-day.

Discussion

The numerical experiments have shown that a horizontalextension in WNW-ESE direction of about 8–8.5 kmcan explain the observed sediment thicknesses, shoulderuplift, and Moho upwelling for the central graben seg-ment around Karlsruhe from latest Priabonian to earlyAquitanian. For the southern URG, only the Mohoelevation and uplift of the eastern rift shoulder serve asstandard values for assessing the results of the numericalmodel due to missing sedimentary units for the Chattianonward. A good agreement between these comparativedata and the model output is obtained for the sameamount of horizontal extension as for the central grabensegment. This supports the implicit assumption that thesouthern URG was subjected to extension in the Chat-tian and early Aquitanian as well.

Fig. 14 Shoulder uplift in 100 m; dashed lines for the uplift at theend of the extension phase, stippled lines for the uplift related to thestrike slip period, shaded areas gravity highs

Fig. 13 Plan view of numericalmodel at present-day.a Isopachs in 100 m; shadedarea of uplift in the grabenproper of around 100 m, onlyhorizontal displacement due tostrike slip reactivation isdisplayed. b Uplift of Moho inmetre due to strike slipreactivation as dashed lines

746

Despite a very detailed implementation of the struc-tural grain in the numerical model, the main differencesbetween the modelling results and the comparative dataset occur in the graben segment north of Karlsruhe.There, the thickness of stratigraphic units related to theextension interval are underestimated over a large area.In contrast, the total amounts of Moho elevation andrift shoulder uplift, respectively, are overvalued. Themodelled horizontal extension decreases to values of4–6 km in the northernmost URG. This is in accordancewith estimations from crustal balancing approaches forthe same region (Meier 1989). Consequently, the missingsediment thicknesses in the northern URG can not bereconstructed by assuming higher amounts of horizontalextension in WNW-ESE direction. In addition, thiswould lead to an even higher uplift of Moho and riftshoulders in the northern graben segment for listricborder faults. We suggest that a more detailed numericalmodel, possibly in combination with different mechani-cal boundary conditions, could alleviate the large-scalediscrepancies for the northern URG. The model sim-plified the step by step separation of the URG towardsthe HT by implementation of one discrete SE trendingfault. This creates a narrow but high elevation of theMoho in NW-SE orientation near this weak zone duringthe strike slip period. As a further consequence, thesediment thickness related to the extension phase in thedirect hangingwall of the discrete SE trending fault isoverestimated. In contrast, isopach maps for the exten-sion period (Fig. 5a; Schad 1962; Doebl 1967; Derer2003), gravity modelling data (Plaumann 1991), andseismic images of the crust mantle boundary (Zeis et al.1990) suggest a more distributed deformation on addi-tional SE and ENE trending weak zones. Interestingly,SE oriented and graben-parallel faults often terminateagainst each other in the northern graben segment. Thefault traces in plan view are highly concave both towardsthe footwall and towards the hangingwall (Fig. 4).Therefore, the structural grain of the graben could nothave been formed by a local deviation of the r2 axis atthe dip of evolving faults towards the hangingwall asgenerally suggested by Mandl (2000). It has been created

by a coalescing of these pre-existing faults with differentstrike orientations, likely due to a similarity of r2 and r3in the northern URG area during the extension interval.However, for testing to what extent the implementationof additional structural grain and different mechanicalboundary conditions could remedy deficiencies of thelarge-scale numerical model, there is still a lack of de-tailed information on the strike, depth, geometry, andfriction of faults which lead to a separation of the URGinto small-scale grabens.

A sinistral displacement across the entire rift of 3.5–4 km is calculated in order to accommodate the maxi-mum sediment thickness related to the strike slipinterval. The left-lateral movements along one borderfault reach about 4 km which is a little more than thehighest standard value deduced from kinematic andgeomagnetic indications as mentioned before. Thismodelled amount of sinistral displacement could easilybe reduced by assuming a few degrees higher dip of theSE trending fault than 60�. The rough similarity be-tween modelled and real sinistral displacement makesthe possibility of another stretching direction thanWNW-ESE for the URG area from the latest Priabo-nian to early Aquitanian very doubtful. This observa-tion can be substantiated by additional results of themodelling study. For several model runs, we also used avelocity boundary condition being more E-W orientedas suggested by analog modelling results and 3D bal-ancing techniques (Mart and Dauteuil 2000; Behrmannet al. 2003). However, this approach led to an increaseof the left-lateral displacements beyond those maximumstandard values. The Moho uplift took on a morenorth-westerly trend in the northern URG as alreadyseen for the strike slip interval (Fig. 13b) and the sedi-ment thickness in the ENE trending transfer zone at thelatitude of Strasbourg was significantly reduced.Moreover, in all model runs, the implemented HT doesnot localize the deformation during extension and strikeslip reactivation. The fault reveals a sticking behaviourand is passively deformed even for low friction coeffi-cients of about 0.3. Consequently, the NE trendingfaults bordering the URG in the north and in the southdo not impose kinematic constraints for a WNW-ESEoriented extension, in contrast to the findings of Hart-hill (2002). Thus, the results of the modelling studyconfirm the suggestion of an approximate WNW-ESEdirected extension on the regional scale deduced fromregional kinematic analyses (Bergerat 1985; Larroqueand Laurent 1988; Lacombe et al. 1993; Michon et al.2003). The kinematic evolution of the rift system duringthe extension interval is possibly completed by exten-sion parallel to the rift zone due to a similarity of r2and r3 in the northern URG. It must be stressed thatthe numerical simulations do not provide any con-straints for the structural evolution of the URG in timeintervals without sedimentary record. These includeerosional hiatus at the top of the Hydrobia Beds, be-tween Upper Tertiary I and Upper Tertiary II as well asbetween Serravallian and Quaternary. Therefore, we

Fig. 15 Length section of the eastern rift shoulder showingcalculated and modelled uplift

747

cannot exclude that large fluctuations of the horizontaldisplacement rate might have occurred during Neogeneand Quaternary in the URG area.

The numerical experiments highlight the predominantinfluence of fault friction on the lithospheric evolution ofthe URG rift system. At first view, this seems to becontradictory to some modelling studies on extension incontinental settings (e.g. Boutilier and Keen 1994) whichfound no fundamental dependence of graben subsidenceon frictional coefficients. The ratio of upper crustal tolower crustal thickness of 1.5:1 and high temperatures inthe model lithosphere provoke a zone of a high mobilitybetween the maximum depth of upper crustal faults(15 km) and the top of the lower crust (18 km). In thiszone, low friction coefficients enhance considerably aductile flow towards the opposite rift shoulders which isagain modified by the shape of the border faults and thedifference of friction for conjugate border faults. Thehigh mobility of this region becomes evident if the lateralductile flow is impeded by implementation of contactelements down to the top of the lower crust at the depthof 18 km. As, to the best of our knowledge, there are nolaboratory data available to constrain the strain soften-ing in the semibrittle and viscous regime quantitativelyfor a natural example, the border faults were assignedwith the smallest numerically stable fault friction of5 MPa between 15 km and 18 km. The subsidence atboth border faults was only 100–300 m higher along thestrike of the URG than it was for faults reaching a depthof only 15 km. Thereby, the influence of fault friction onthe amount of subsidence decreased noticeably. Thus, itis very likely that extension in the URG area is accom-modated along border faults which sole out into adetachment level at around 15–16 km.

Numerical modeling of the strike slip period of theURG rift system leads to an inversion of the centralgraben segment to the tune of only 100 m. However, acomparison of the isopachs in Figs. 5a and b suggests anuplift of about 500 m for the central rift segment aroundKarlsruhe in post-Early Aquitanian times. In this region,the base of the graben fill related to rifting was upliftedfrom more than 2,800 m to around 2,300 m if the com-paction of earlier Tertiary sediments is negligible. Sedi-ments of remarkable thickness related to the strike slipinterval are not encountered. Consequently, a uplift of400 m must be caused by other processes involved in theevolution of the central URG than the compressionalreactivation of the rift system. The same amount ofvertical motion can be seen on the eastern rift shoulder asthe difference between the total uplift and uplift relatedto extension both in the numerical models and in thebackstacking approach (Fig. 6). Therefore, it can beconcluded that a process possibly unrelated to the strikeslip reactivation of the URG affected both graben andshoulder. This process seems to terminate in the southernarea of the Kraichgau gravity high where the total upliftof the eastern graben shoulder nearly equals the upliftrelated to extension from latest Priabonian to earlyAquitanian. This observation emphasizes the need for a

numerical model which incorporates the northwardmigration of the Alpine forebulge into the URG areaduring the Miocene.

Conclusions

We have presented a 3D numerical model for the evo-lution of the URG rift zone which accounts for the mainstructural and stratigraphic features. Our model com-prises two tectonic events after the initiation of thegraben in the late Priabonian: an extension in WNW-ESE direction normal to the border faults from latestPriabonian to early Aquitanian, followed by a sinistralstrike slip mode to present-day. We arrive at the fol-lowing conclusions from the modelling study:

1. A horizontal extension of 8–8.5 km is needed toaccommodate the sediments related to the extensioninterval for most parts of the rift. The depth todetachment is 15–16 km. Beneath, pure shear occurswithout a need for NNE trending shear zones inlower crust and upper mantle.

2. The uplift of the Moho related to extension can be ashigh as about 4 km for the central and southerngraben segments. To the north, the lithosphere ismore broadly thinned, the uplift of the Moho grad-ually decreases to values of about 2 km.

3. Extension and strike slip reactivation lead to ashoulder uplift of 1,500 m at most in the footwall ofthe master faults and some 800 m for rift flanks alongsecondary border faults.

4. The border faults exhibit low friction coefficients of0.3–0.4. Supported by sedimentary loading, low faultfriction provokes a ductile flow in the viscous regimeacross the graben away from the weak master faults.This considerably enhances the structural asymmetryof the URG rift zone.

5. A sinistral displacement across the entire rift of 3–4 km is required to accommodate the preserved sed-iments related to strike slip reactivation of the URG.A stretching direction other than approximatelyWNW-ESE on a regional scale leads to an overesti-mation of the amount of left-lateral displacements inthe URG area for the Cenozoic.

Modelling the evolution of the northern URG hasfailed due to a disregard of the complex pre-existingstructural grain in this region. These structures are onlypoorly known today. A simulation of the northern URGmay also require different mechanical boundary condi-tions because highly arcuate fault traces in plan view andthe termination of weak zones of different orientationagainst each other indicate a similarity of r2 and r3

during the extension interval.The results of the simulation let us assume that the

central graben segment was marginally uplifted inconsequence of a transpressive reactivation in Post-Oligocenian times. Other processes have got to be takeninto account to explain the considerable uplift of both

748

graben and shoulders in this region for Mioceneonwards. This hypothesis should be tested by includingthe lithospheric bulging processes in simulations of thestructural evolution in the URG area.

Acknowledgements This paper is a contribution to the EuropeanUpper Rhine Graben Evolution and Neotectonics (EUCOR-UR-GENT) Project. Nina Kukowski and Charles Gumiaux are kindlythanked for their thorough and detailed reviews. The authorsacknowledge financial support by the Deutsche Forschungsgeme-inschaft.

References

Anderle HJ (1968) Die Machtigkeiten der sandig-kiesigen Sedi-mente des Quartars im nordlichen Oberrheingraben und derostlichen Untermainebene. Notizbl hess L-Amt Bodenforsch96:185–196

Anderle HJ (1974) Block tectonic interrelations between NorthernUpper Rhine Graben and Southern Taunus Mountains. In:Illies JH, Fuchs K (eds) Approaches to taphrogenesis: pro-ceedings of an international Rift symposium held in KarlsruheApril 1972. Schweizerbart, Stuttgart, pp 243–253

Athy LF (1930) Density, porosity and compaction of sedimentaryrocks. Bull Am Assoc Pet Geol 14:1–24

Banda E, Cloetingh S (1992) Physical properties of the lithosphere.In: Blundell D, Freeman R, Muller S (eds) A continent re-vealed. Cambridge University Press, pp 71–80

Bartz J (1974) Die Machtigkeit des Quartars im Oberrheingraben.In: Illies JH, Fuchs K (eds) Approaches to taphrogenesis:proceedings of an international Rift symposium held in Kar-lsruhe, April 1972. Schweizerbart, Stuttgart, pp 78–87

Behrmann J, Hermann O, Horstmann M, Tanner D, Bertrand G(2003) Anatomy and kinematics of oblique continental riftingrevealed; a three-dimensional case study of the Southeast UpperRhine Graben (Germany). AAPG Bull 87:1105–1121

Berger JP (2002) Eocene-Pliocene time scale, stratigraphy andterrace dating. Abstract Volume to 2nd EUCOR-URGENTWorkshop, p 5

Bergerat F (1985) Deformations cassantes et champs de contraintetertiaires dans la plate-forme Europeenne. Mem Sci Terre,Universite P et M Curie, pp 1–315

Bosum W, Ullrich HJ (1970) Die Flurmagnetometermessung desOberrheingrabens und ihre Interpretation. Geol Rundsch59:83–106

Boutilier RR, Keen CE (1994) Geodynamic models of fault-con-trolled extension. Tectonics 13:439-454

Brun JP, Gutscher MA, DEKORP-ECORS team (1992a) Deepcrustal structure of the Rhine Graben from DEKORP-ECORSseismic reflection data: a summary. Tectonophysics 208:139–147

Brun JP, Wenzel F, ECORS-DEKORP team (1992b) Crustal-scalestructure of the southern Rhine Graben from ECORS-DE-KORP seismic reflection data. Geology 19:758–762

Byerlee JD (1978) Friction of rocks. Pure Appl Geophys 116:615–626

Byerlee JD (1990) Friction, overpressure, and fault normal com-pression. Geophys Res Lett 17:2109–2112

Chopra PN, Paterson MS (1981) The experimental deformation ofdunite. Tectonophysics 78:453–473

Chorowicz J, Defontaines B (1993) Transfer faults and pull-apartmodel in the Rhine Graben from analysis of multisource data.J Geophys Res 98:14339–14351

Clauser C, Villinger H (1990) Analysis of conductive and convec-tive heat transfer in a sedimentary basin, demonstrated for theRheingraben. Geophys J Int 100:393–414

Demoulin A, Launoy T, Zippelt K (1998) Recent crustal move-ments in the southern Black forest. Geol Rundsch 87:43–52

Derer CE (2003) Tectono-sedimentary evolution of the northernUpper Rhine Graben (Germany), with special regard to theearly syn-rift stage. PhD thesis, University of Bonn, pp 1–103

Diederich G (1987) Bruchtektonik des nordlichen hessischenBuntsandstein-Odenwalds. Geol Jb Hessen 115:305–313

Dieterich JH, Conrad G (1984) Effect of humidity on time- andvelocity-dependent friction in rocks. J Geophys Res 89:4196–4202

Doebl F (1967) The Tertiary and Pleistocene sediments of thenorthern and central part of the Upper Rhine Graben. AbhGeol L-Anst Baden-Wurttemberg 6:48–54

Doebl F (1970) Die tertiaren und quartaren Sedimente des sudli-chen Rheingrabens. In: Illies JH, Muller S (eds) Graben prob-lems: proceedings of an international Rift symposium held inKarlsruhe, October 1968. Schweizerbart, Stuttgart, pp 56–66

Doebl F, Teichmueller R (1979) Zur Geologie und heutigen Geo-thermik im mittleren Oberrheingraben. Fortschr Geol Rheinl uWestf 27:1–17

Duringer P (1988) Les conglomerats des bordures du rift cenozoi-que rhenan. Dynamique sedimentaire et controle climatique.PhD thesis, University of Strasbourg, pp 1–261

Durst H (1991) Aspects of exploration history and structural stylein the Rhine Graben area. In: Spencer AM (ed) Generation,accumulation, and production of Europe’s hydrocarbons. SpecPubl Europ Assoc Petr Geosci 1:247–261

Edel JB, Fluck P (1989) The upper Rhenish Shield basement(Vosges, Upper Rhine Graben and Schwarzwald): main struc-tural features deduced from magnetic, gravimetric and geolog-ical data. Tectonophysics 169:303–316

Edel JB, Weber K (1995) Cadomian terranes, wrench faulting andthrusting in the central Europe Variscides: geophysical andgeological evidence. Geol Rundsch 84:412–432

Fischer F (1978) Zur Frage der alten Landoberflachen in denVogesen und im Oberrheingraben. Annales Universitatis Sa-raviensis 14:42–58

Floettmann T, Oncken O (1992) Constraints on the evolution ofthe Mid German crystalline rise; a study of outcrops west of theRiver Rhine. Geol Rundsch 81:515–543

GolwerA (1968) Palaogeographie desHanauer Beckens imOligozanund Miozan. Notizbl hess L-Amt Bodenforsch 96:157–184

Groschopf R, Kessler G, Leiber J, Maus HJ, Ohmert O, SchreinerA, Wimmenauer W (1996) Erlauterungen zu Blatt Freiburg iBr. Geol Karte von Baden-Wurttemb 1:50000. Stuttgart, p 364

Groshong RH (1996) Construction and validation of extensionalcross sections using lost area and strain, with application to theRhine Graben. In: Buchanan PG, Nieuwland DA (eds) Moderndevelopments in structural interpretation. Geol Soc Spec Publ99:79–87

Grosse S, Behr HJ, Edel JB, Heinrichs T (1992) The gravity fieldalong the central segment of the EGT. Tectonophysics 207:97–121

Gutscher MA (1995) Crustal structure and dynamics in theRhine Graben and the Alpine foreland. Geophys J Int 122:617–636

Haenel R, Staroste E (1988) Atlas of geothermal resources in theEuropean Community, Austria and Switzerland. Schafer,Hannover, pp 1–74

Haq BU, Hardenbol J, Vail R (1988) Mesozoic and Cenozoicchronostratigraphy and eustatic cycles. Soc Econ PalaeontolGeol Soc Spec Pub 42:71–108

Harthill N (2002) The tectonic basis of the geothermal potential ofthe Oberrheingraben. Abstract Volume to 7th GeothermischeFachtagung, pp 143–151

Heling D (1969) Relationships between initial porosity of Tertiaryargillaceous sediments and palaeosalinity in the Rhinegraben(SW-Germany). J Sedim Petrology 39:246–254

Henk A (1993) Subsidenz und Tektonik des Saar-Nahe-Beckens(SW-Deutschland). Geol Rundsch 82:3–19

Henk A (1997) Gravitational orogenic collapse vs plate boundarystresses: a numerical modelling approach to the Permo-carbon-iferous evolution of Central Europe. Geol Rundsch 86:39–55

749

Henk A (1998) Thermomechanische Modellierungen zur postkon-vergenten Krustenequilibrierung in den Varisciden. GeotektForsch 90:1–124

Hill DP (1992) A note on ambient pore pressure, fault-confined porepressure, and apparent friction. Bull Seis Soc Am 83:583–586

Holbrook WS, Gajewski D, Krammer A, Prodehl C (1988) Aninterpretation of wide-angle compressional and shear wave datain Southwest Germany: Poisson’s ratio and petrological impli-cations. J Geophys Res 93:12081–12106

Illies JH (1962) Prinzipien der Entwicklung des Rheingrabens,dargestellt am Grabenabschnitt bei Karlsruhe. Mitt Geol Sta-atsinst Hamburg 31:58–121

Illies JH (1975) Recent and palaeo-intraplate tectonics in stableEurope and the Rhinegraben rift system. Tectonophysics29:251–264

Illies JH, Greiner G (1979) Holocene movements and state of stressin the Rhine graben rift system. Tectonophysics 52:349–359

Krohe A (1992) Structural evolution of intermediate-crustal rocksin a strike slip and extensional setting (Variscan Odenwald, SWGermany): differential upward transport of metamorphiccomplexes and changing deformation mechanisms. Tectono-physics 205:357–386

Lacombe O, Angelier J, Byrne D, Dupin JM (1993) Eocene-Oli-gocene tectonics and kinematics of the Rhine-Saone continentaltransform zone (Eastern France). Tectonics 12:874–888

Larroque JM, Laurent P (1988) Evolution of the stress field in thesouth of the Rhine Graben from the Eocene to the present.Tectonophysics 148:41–58

Laubscher H (1992) Jura kinematics and the Molasse basin.Eclogae geol Helv 85:653–675

Laubscher H (2001) Plate interactions at the southern end of theRhine Graben. Tectonophysics 343:1–19

Lysak SV (1992) Heat flow variations in continental rifts. Tec-tonophysics 208:309–323

Maass R (1988) Die Sudvogesen zu variszischer Zeit. N Jb GeolPalaont Mh 10:611–638

Mackwell SJ, Zimmermann ME, Kohlstedt DL (1998) High-tem-perature deformation of dry diabase with application to tec-tonics on Venus. J Geophys Res 103:975–984

Mandl G (2000) Faulting in brittle rocks: an introduction to themechanics of tectonic faults. Springer, Berlin Heidelberg NewYork, pp 1–434

Marell D (1989) Das Rotliegende zwischen Odenwald und Taunus.Geol Abh Hessen 89:1–128

Mart Y, Dauteuil O (2000) Analogue experiments of propagationof oblique rifts. Tectonophysics 316:121–132

Mayer G, Mai PM, Plenefish T, Echtler H, Luschen E, Wehrle V,Muller B, Bonjer KP, Prodehl C, Fuchs K (1997) The deepcrust of the southern Rhine Graben: reflectivity and seismicityas images of dynamic processes. Tectonophysics 275:15–40

Meier L (1989) Ein Modell fur die Tiefenstruktur und Kinematikim Bereich des nordlichen Rheingrabens. PhD thesis, Universityof Karlsruhe, pp 1–117

Meier L, Eisbacher G (1991) Crustal kinematics and deep structureof the northern Rhine Graben. Tectonics 10:621–630

Meissner R, Bortfeld RK (1990) DEKORP Atlas. Springer, BerlinHeidelberg New York

Mengel K, Sachs PM, Stosch HG, Worner G, Loock G (1991)Crustal xenoliths from the Cenozoic fields of West Germany:implications for structure and composition of the continentalcrust. Tectonophysics 195:271–289

Michon L, Van Balen RT, Merle O, Pagnier H (2003) The Ceno-zoic evolution of the Roer Valley Rift System integrated at aEuropean scale. Tectonophysics 367:101–126

Pfiffner OA, Schlunegger F, Buiter SJH (1998) The Swiss Alps andtheir peripheral foreland basin: stratigraphic response to deepcrustal processes. Tectonics 21(1):10.129/2000TC9000039

Plaumann S (1991) Die Schwerekarte 1:500 000 der BundesrepublikDeutschland (Bouguer-Anomalien), Blatt Mitte. Geol Jb E46:3–16

Plaumann S, Groschopf R, Schadel K (1986) Kompilation einerSchwerekarte und einer geologischen Karte fur den mittlerenund nordlichen Schwarzwald mit einer Interpretation gravi-metrischer Detailvermessungen. Geol Jb E 33:15–30

Prodehl C, Mueller St, Glahn A, Gutscher MA, Haak V (1992)Lithospheric cross sections of the European Cenozoic rift sys-tem. Tectonophysics 208:113–138

Roll A (1979) Versuch einer Volumenbilanz des Oberrheintalgra-bens und seiner Schultern. Geol Jb A 52:3–82

Rousset D, Bayer R, Guillon D, Edel JB (1992) Structure of thesouthern Rhine Graben from gravity and reflection seismicdata (ECORS-DEKORP program). Tectonophysics 221:135–153

Royer JJ, Danis M (1988) Steady state geothermal model of thecrust and the problem of the boundary conditions; applicationto a rift system, the southern Rhine Graben. Tectonophysics156:239–255

Schad A (1962) Voraussetzungen fur die Bildung von Er-dollagerstatten im Rheingraben. Abh Geol L-Anst Baden-Wurttemberg 4:29–40

Schumacher ME (2002) Upper Rhine Graben: role of pre-existingstructures during rift evolution. Tectonics 21:10.1029/2001TC900022

Seipold U (1998) Temperature-dependence of thermal transportproperties of crystalline rocks—a general law. Tectonophysics291:161–171

Seyferth M, Henk A (2000) Deformation, metamorphism andexhumation: quantitative models for a continental collisionzone in the Variscides. In: Franke W, Haak V, Oncken O,Tanner D (eds) Orogenic processes: quantification and model-ling in the variscan belt. Geol Soc Spec Publ 179:217–230

Shelton G, Tullis J (1981) Experimental flow laws for crustal rocks.EOS 62:396

Sibson RH (1974) Frictional constraints on thrust, wrench, andnormal faults. Nature 242:542–544

Sissingh W (1998) Comparative Tertiary stratigraphy of the RhineGraben, Bresse Graben, and Molasse Basin: correlation ofAlpine foreland events. Tectonophysics 300:249–284

Straub EW (1962) Voraussetzungen fur die Bildung von Er-dollagerstatten im Rheingraben. Abh Geol L-Anst Baden-Wurttemberg 4:123–136

Strehlau J, Meissner R (1987) Estimation of crustal viscosities andshear stresses from an extrapolation of experimental steadystate flow data. In: Fuchs K, Froidevaux C (eds) Composition,structure and dynamics of the lithosphere-asthenosphere sys-tem. Geodyn Ser Am Geophys Union 16:69–87

Teichmuller M, Teichmuller R (1979) Zur geothermischen Ges-chichte des Oberrheingrabens. Zusammenfassung und Au-swertung eines Symposiums. Fortschr Geol Rheinld u Westf27:109–120

Wenzel F, Brun JP, ECORS-DEKORP team (1991) A deepreflection seismic line across the northern Rhine Graben. EarthPlanet Sci Lett 104:140–150

Wilks KR, Carter NL (1990) Rheology of some continental lowercrustal rocks. Tectonophysics 182:57–77

Zeiss S, Gajewski D, Prodehl C (1990) Crustal structure ofsouthern Germany from seismic refraction data. Tectonophys-ics 176:59–86

Ziegler PA (1992) European Cenozoic rift system. Tectonophysics208:91–111

750