.New Frontiers in Integrated Solid Earth Sciences Phần 5 pptx

160 F. Roure et al. 0 10 20 30 40 50 100 150 200 250 Km 10 20 30 40 50 60 –10 0 10 20 30 40 50 60 0 10 20 30 40 50 60 100 150 200 Km 100 150 200 250 Km 100 150 200 250 Km 70 60 50 40 70 60 50 40 30 –10 60 –10 30 0 Electromagnetic data (d) Thermal data (c) –10 Surface waves (b) Body waves (a) Fig. 9 Thickness of the European lithosphere as determined by (a) seismic tomography; (b) surface wave tomography; (c) geothermics; (d) magnetotellurics (after Artemieva et al., 2006) deeper crust and mantle rocks depends on their chem- ical composition, temperature and pressure conditions (combination of burial and regional heat flow) and their water content. Furthermore, the crust and lithospheric mantle may be locally weakened by the occurrence of pre-existing discontinuities related to earlier deforma- tion phases (Ziegler et al., 1998; Ziegler and Cloetingh 2004). Such inherited weakness zones, represented by e.g., crustal scale faults or eclogitized continental crust inserted into the sub-crustal mantle, may be character- ized by considerably reduced strengths as compared to surrounding crustal and mantle domains. Analogue modelling techniques were first devel- oped, and are now routinely used by many lab- oratories to simulate thin-skinned deformation of sedimentary rocks (see Colletta et al., 1991, and references therein), but also of the lithosphere as a whole (Sokoutis et al., 2005, 2007, and references therein), using specific analogue materials for mod- elling either brittle or ductile sediments, the crust or the mantle. Further numerical developments are, however, required when investigating the effects of other parameters during basin evolution, such as pore- fluid pressure, temperature and mineralogical phase transitions. In the following paragraphs, we shall summarize some recent advances achieved in documenting and understanding the rheology and long term behaviour of the European lithosphere, as well as a few dedicated case studies outlining (1) the incidence of deep active décollements on surface topography, (2) the respec- tive effects of coupling and strain partitioning between the foreland and hinterland during the development of selected intramontane basins, as well as the over- all dynamics of (3) intracratonic basins and (4) passive margins. Achievements and Challenges in Sedimentary Basins Dynamics 161 Lithosphere Strength and Deformation Mode The strength of continental lithosphere is controlled by its depth-dependent rheological structure in which the thickness and composition of the crust, the thickness of the lithospheric mantle, the potential temperature of the asthenosphere, and the presence or absence of flu- ids, as well as strain rates play a dominant role. By con- trast, the strength of oceanic lithosphere depends on its thermal regime, which controls its essentially age- dependent thickness (Kuznir and Park, 1987; Cloet- ingh and Burov, 1996; Watts, 2001; see also Burov, 2007). Figure 10 gives synthetic strength envelops for three different types of continental lithosphere and for oceanic lithosphere at a range of geothermal gradients (Ziegler and Cloetingh, 2004). These theoretical rheo- logical models indicate that thermally stabilized conti- nental lithosphere consists of the mechanically strong upper crust, which is separated by a weak lower crustal layer from the strong upper part of the mantle- lithosphere that in turn overlies the weak lower mantle- lithosphere. By contrast, oceanic lithosphere has a more homogeneous composition and is characterized by a much simpler rheological structure. In terms of rheology, thermally stabilized oceanic lithosphere is considerably stronger than all types of continental lithosphere. However, the strength of oceanic litho- sphere can be seriously weakened by transform faults and by the thermal blanketing effect of thick sedimen- tary prisms prograding onto it (e.g., Gulf of Mexico, Niger Delta, Bengal Fan; Ziegler et al., 1998). The strength of continental crust depends largely on its composition, thermal regime and the presence of fluids, and also on the availability of pre-existing crustal discontinuities (see also Burov, 2007). Deep- reaching crustal discontinuities, such as thrust- and wrench-faults, cause significant weakening of the oth- erwise mechanically strong upper parts of the crust. Such discontinuities are apparently characterized by a reduced frictional angle, particularly in the presence of fluids (Van Wees, 1994). These discontinuities are prone to reactivation at stress levels that are well below those required for the development of new faults. Deep reflection-seismic profiles show that the crust of Late Proterozoic and Paleozoic orogenic belts is generally characterized by a monoclinal fabric that extends from upper crustal levels down to the layered lower crust and Moho at which i t either soles out or by which it is truncated (Figs. 2, 3, 4; see Bois, 1992; Ziegler and Cloetingh, 2004). This fabric reflects the presence of deep-reaching lithological inhomogeneities and shear zones. The strength of the continental upper lithospheric mantle depends to a large extent on the thickness of the crust but also on its age and thermal regime (see Jaupart and Mareschal, 2006). Thermally stabilized stretched continental lithosphere with a 20 km thick crust and a lithospheric mantle thickness of 50 km is mechanically stronger than unstretched lithosphere with a 30 km thick crust and a 70 km thick lithospheric mantle (compare Fig. 10b, d). Extension of stabilized continental crustal segments precludes ductile flow of the lower crust and faults will be steep to listric and propagate towards the hanging wall, i.e., towards the basin centre (Bertotti et al., 2000). Under these con- ditions, the lower crust will deform by distributing ductile shear in the brittle-ductile transition domain. This is compatible with the occurrence of earthquakes within the lower crust and even close to the Moho (e.g., southern Rhine Graben: Bonjer, 1997; East African rifts: Shudofsky et al., 1987). On the other hand, in young orogenic belts, which are characterized by crustal thicknesses of up to 60 km and an elevated heat flow, the mechanically strong part of the crust is thin and the lithospheric mantle is also weak (Fig. 10c). Extension of this type of litho- sphere, involving ductile flow of the lower and middle crust along pressure gradients away from areas lack- ing upper crustal extension into zones of major upper crustal extensional unroofing, can cause crustal thin- ning and thickening, respectively. This deformation mode gives rise to the development of core complexes with faults propagating towards the hanging wall (e.g., Basin and Range Province: Wernicke, 1990; Buck, 1991; Bertotti et al., 2000). However, crustal flow will cease after major crustal thinning has been achieved, mainly due to extensional decompression of the lower crust (Bertotti et al., 2000). Generally, the upper mantle of thermally stabilized, old cratonic lithosphere is considerably stronger than the strong part of its upper crust (Fig. 10a) (Moisio et al., 2000). However, the occurrence of upper man- tle reflectors, which generally dip in the same direc- tion as the crustal fabric and probably are related to subducted oceanic and/or continental crustal material, 162 F. Roure et al. Fig. 10 Depth-dependent rheological models for various litho- sphere types and a range of geothermal gradients, assuming a dry quartz/diorite/olivine mineralogy for continental lithosphere (Ziegler, et al., 1995; Ziegler et al., 2001). (a) Unextended, thick- shield-type lithosphere with a crustal thickness of 45 km and a lithospheric mantle thickness of 155 km. (b) Unextended, “nor- mal” cratonic lithosphere with a crustal thickness of 30 km and a lithospheric mantle thickness of 70 km. (c) Unextended, young orogenic lithosphere with a crustal thickness of 60 km and a lithospheric mantle thickness of 140 km. (d) Extended, cratonic lithosphere with a crustal thickness of 20 km and a lithospheric mantle thickness of 50 km. (e) Oceanic lithosphere. Modified from Ziegler et al. (2001) Achievements and Challenges in Sedimentary Basins Dynamics 163 suggests that the continental lithospheric mantle is not necessarily homogenous but can contain lithological discontinuities that enhance its mechanical anisotropy (Vauchez et al., 1998; Ziegler et al., 1998). Such dis- continuities, consisting of eclogitized crustal material, can potentially weaken the strong upper part of the lithospheric mantle. Moreover, even in the face of similar crustal thicknesses, the heat flow of deeply degraded Late Proterozoic and Phanerozoic orogenic belts is still elevated as compared to adjacent old cratons (e.g., Panafrican belts of Africa and Arabia; Janssen, 1996). This is probably due to the younger age of their lithospheric mantle and possibly also to a higher radiogenic heat generation potential of their crust. These factors contribute to weakening of for- mer mobile zones to the end that they present rhe- ologically weak zones within a craton, as evidenced by their preferential reactivation during the break-up of Pangea (Ziegler, 1989; J anssen et al., 1995; Ziegler et al., 2001). Concerning rheology, the thermally destabilized lithosphere of tectonically active rifts, as well as of rifts and passive margins that have undergone only a relatively short post-rift evolution (e.g., 25 Ma), is considerably weaker than that of thermally stabilized rifts and of unstretched lithosphere (Figs. 10 and 11, Ziegler et al., 1998). In this respect, it must be realized that during rifting, progressive mechanical and ther- mal thinning of the lithospheric mantle and its substi- tution by the upwelling asthenosphere is accompanied by a rise in geotherms causing progressive weakening of the extended lithosphere. In addition, its permeation by fluids causes its further weakening (Fig. 11). Upon decay of the rift-induced thermal anomaly, rift zones are rheologically speaking considerably stronger than unstretched lithosphere (Fig. 10). However, accumula- tion of thick syn- and post-rift sedimentary sequences can cause by thermal blanketing a weakening of the strong parts of the upper crust and lithospheric mantle of rifted basins (Stephenson, 1989). More- over, as faults permanently weaken the crust of rifted basins, they are prone to tensional as well as com- pressional reactivation and tectonic inversion (Roure et al., 1994, 1997; Ziegler et al., 1995, 1998, 2001, 2002; Brun and Nalpas, 1996; Roure and Colletta, 1996). In view of its rheological structure, the continen- tal lithosphere can be regarded under certain condi- tions as a two-layered visco-elastic beam (Fig. 12; Reston, 1990, Ter Voorde et al., 1998). The response of such a system to the build-up of extensional and com- pressional stresses depends on the thickness, strength and spacing of the two competent layers, on stress magnitudes and strain rates and the thermal regime (Zeyen et al., 1997; Watts and Burov, 2003). As the structure of continental lithosphere is also regionally heterogeneous, its weakest parts start to yield first once tensional as well as compressional intraplate stress levels equate their strength (Ziegler et al., 2001). The flow properties of mantle rocks control the thickness and strength of the lithospheric plates, the degree of coupling between moving lithospheric plates and the pattern and rate of asthenospheric convection, and the rate of melt extraction at mid-ocean ridges. To be able to understand the dynamic behaviour of the outer parts of the solid Earth, notably the dynamics of lithospheric extension and associated rifting and sed- imentary basin development, a detailed knowledge of the rheology and evolution of the upper mantle (30– 410 km depth) and between the 410 and 670 transi- tion zones is essential. At present, these flow proper- ties are surprisingly poorly known. Experimental work has yielded constitutive equations describing various types of flow in mantle rocks, but it is not clearly established to what extent the experimentally observed flow mechanisms are relevant for natural crust and mantle conditions. A second problem is that trace amounts of water and melt can cause drastic weaken- ing of mantle rocks and may cause the development of upper mantel convective instabilities (Lustrino and Wilson, 2007). Such fluid-related weakening effects are widely recognized as, for example, controlling the strength of trans-lithospheric faults in the substratum of active basins. However, only limited data are avail- able on such effects, and a quantitative, mechanical understanding suitable for extrapolation to nature is lacking. These problems can be addressed by means of experimental studies, scanning and transmission electron microscopy (SEM, TEM) and field stud- ies on exposed upper mantle rocks. Integration of these approaches aims at arriving at quantitative, mechanism-based descriptions of mantle rheologies suitable for use in modelling the dynamics of the upper mantle and transition zone. Field-based studies involv- ing structural geological and EM work on upper mantle rocks deformed in a variety of geological environments 164 F. Roure et al. normal continental lithosphere extended, thermally rejuvenated lithosphere strength (MPa) depth (Km) –1000 1000 1500–500 5000 strength (MPa) –1000 1000 1500–500 5000 0 10 20 30 40 50 60 70 80 90 100 (a) (b) depth (Km) 0 10 20 30 40 50 60 70 80 90 100 UC: granite UM: dunite LC: o-pyroxene Moho UC: granite UM: dunite LC: o-pyroxene Moho wet dry 0 250 500 750 1000 1250 1500 temperature (°C) 0 250 500 750 1000 1250 1500 temperature (°C) tension compression wet dry tension compression Fig. 11 Depth-dependent rheological models for dry and wet, unextended ‘normal’ cratonic lithosphere and stretched, t her- mally attenuated lithosphere, assuming a quartz/diorite/olivine mineralogy. (a) Unextended, cratonic lithosphere with a crustal thickness of 30 km and a lithospheric mantle thickness of 70 km. (b) Extended, thermally destabilized cratonic lithosphere with a crustal thickness of, 20 km and a lithospheric mantle thickness of 38 km. Modified from Ziegler et al. (2001) upper crust lower crust asthenosphere mantle lithosphere MSC MSL strong weak Fig. 12 Kinematic model for extension of rheologically stratified lithosphere. See strength profile on left side of diagram. MSC and MSL indicate the base of the mechanically strong crust and mechanically strong lithosphere, respectively. From Reston (1990) may provide information on flow mechanisms occur- ring in the upper mantle. Therefore, special attention has to be paid to upper mantle rocks showing possible asthenospheric flow structures, which developed when the rocks contained some fluid or partial melts. In addi- tion, attention has to be paid to upper mantle shear zone rocks as such shears probably control the extensional strength of the lithosphere. Lithospheric Folding: An Imp ortant Mode of Intraplate Basin Formation Folding of the lithosphere, involving its positive as well as negative deflection (see Figs. 13 and 14), appears to play a more important role in the large- scale neotectonic deformation of Europe’s intraplate Achievements and Challenges in Sedimentary Basins Dynamics 165 upper crust lower crust upper mantle VV erosion/sedimentation strong strong weak weak strong weak Moho Fig. 13 Schematic diagram illustrating decoupled lithospheric mantle and crustal folding, and consequences of vertical motions and sedimentation at the Earth’s surface. V is horizontal short- ening velocity; upper crust, lower crust, and mantle layers are defined by corresponding rheologies and physical properties. A typical brittle-ductile strength profile (in black) for decoupled crust and upper mantle- lithosphere, adopting a quartz-diorite- olivine rheology, is shown for reference domain than hitherto realized (after Cloetingh et al., 1999). The large wavelength of vertical motions asso- ciated with lithospheric folding necessitates integration of available data from relatively large areas (Elfrink, 2001), often going beyond the scope of regional struc- tural and geophysical studies that target specific struc- tural provinces. Recent studies on the North German Basin have revealed the importance of its neotectonic structural reactivation by lithospheric folding (Marotta et al., 2000). Similarly, the Plio-Pleistocene subsidence acceleration of the North Sea Basin is attributed to stress-induced buckling of its lithosphere (Van Wees and Cloetingh, 1996; Unternehr and van den Driess- che, 2004). Moreover, folding of the Variscan litho- sphere has been documented for Brittany (Bonnet et al., 2000), the adjacent Paris Basin (Lefort and Agar- wal, 1996) and the Vosges-Black Forest arch (Ziegler et al., 2002; Dèzes et al., 2004; Bourgeois et al., 2007; Ziegler and Dèzes, 2007). Lithospheric folding is a very effective mechanism for the propagation of t ec- tonic deformation from active plate boundaries far into intraplate domains (e.g., Stephenson and Cloetingh, 1991; Burov et al., 1993; Ziegler et al., 1995, 1998, 2002). Type-1 model simulating the collision between two different lithospheric blocks Type-2 model representing a cold lithosphere with a strong upper mantle Σ belt 34% 26% Upper Crust Lower Crust Upper Mantle Asthenosphere A B 5 cm Model OCR - SL 17 Model OCR - SL 15 λ 1 λ 2 c) d) a) b) λ 1 A 5 cm B Lower Cr ust Upp er Cru st Upper Mantle Asthenosphere P re-cut suture A AB B 5 cm 5 cm Fig. 14 Analogue tectonic modelling for continental lithosphere folding. Top: uniform lithosphere. Bottom: lithosphere blocks separated by suture zone (after Sokoutis et al., 2005) 166 F. Roure et al. At the scale of a micro-continent that was affected by a succession of collisional events, Iberia provides a well-documented natural laboratory for lithospheric folding and the quantification of the interplay between neotectonics and surface processes (Fig. 15; Cloet- ingh et al., 2002). An important factor in favor of a lithosphere-folding scenario for Iberia is the compati- bility of the thermo-tectonic age of its lithosphere and the wavelength of observed deformations. Well-documented examples of continental litho- spheric folding come also from other cratonic areas. A prominent example of lithospheric folding occurs in the Western Goby area of Central Asia, involving a lithosphere with a thermo-tectonic age of 400 Ma. In this area, mantle and crustal wavelengths are 360 and 50 km, respectively, with a shortening rate of ∼10 mm/year and a total amount of shortening of 200– 250 km during 10–15 Myr (Burov et al., 1993; Burov and Molnar, 1998). Quaternary folding of the Variscan lithosphere in the area of the Armorican Massif (Bonnet et al., 2000) resulted in the development of folds with a wave- length of 250 km, pointing to a mantle-lithospheric control on deformation. As the timing and spatial pat- tern of uplift inferred from river incision studies in Brittany is incompatible with a glacio-eustatic ori- gin, Bonnet et al. (2000) relate the observed verti- cal motions to deflection of the lithosphere under the present-day NW–SE directed compressional intraplate stress field of NW Europe (Fig. 16). Stress-induced uplift of the area appears to control fluvial incision rates and the position of the main drainage divides. The area located at the western margin of the Paris Basin and along the rifted Atlantic margin of France has been subject to thermal rejuvenation during Meso- zoic extension related to North Atlantic rifting (Robin et al., 2003; Ziegler and Dèzes, 2006) and subse- quent compressional intraplate deformation (Ziegler et al., 1995), also affecting the Paris Basin (Lefort and Agarwal, 1996). Levelling studies in this area (Lenotre et al., 1999) also point towards its ongoing deformation. The inferred wavelength of these neotectonic litho- sphere folds is consistent with the general relationship a b Topographic evolution Analysis of modelled topography Increasing shortening Surface topography Moho topography –750 –749 –748 –747 –746 –745 –744 –743 –742 Topography (mm) –750 –749 –748 –747 –746 –745 –744 –743 –742 Topography (mm) –750 –749 –748 –747 –746 –745 –744 –743 –742 Topography (mm) -743 -745 0 50 100 150 200 250 300 360mm Fig. 15 Analogue modelling of intraplate continental litho- sphere folding of Iberia (Fernández-Lozano et al., 2008). Left: incremental shortening and topographic evolution. Top right:2D section of the final stage. Bottom right:3Dviewofthefinal stage. Notice the pop-up structures in the upper crust (layered), the ductile flow of the lower crust (orange), and the folded man- tle lithosphere (light grey) Achievements and Challenges in Sedimentary Basins Dynamics 167 011 22 33 64 72 56 48 40 64 72 56 48 40 Fig. 16 Present-day stress map of Europe showing orientation of maximum horizontal stress axes (SHmax). Different symbols stand for different stress indicators; their length reflects the data quality, “A” being highest. Background shading indicates topo- graphic elevation (brown high, green low). This map was derived from the World Stress Map database (http://www.world-stress- map.org) that was established between the wavelength of litho- spheric folds and the thermo-tectonic age of the litho- sphere on the base of a global inventory of lithospheric folds (Fig. 17; see also Cloetingh and Burov, 1996; Cloetingh et al., 2005). In a number of other areas of continental lithosphere folding, also smaller wave- length crustal folds have been detected, for example in Central Asia (Burov et al., 1993; Nikishin et al., 1993). Thermal thinning of the mantle-lithosphere, often associated with volcanism and doming, enhances litho- spheric folding and appears to control the wavelength of folds. Substantial thermal weakening of the litho- spheric mantle is consistent with higher folding rates in the European foreland as compared to folding in Cen- tral Asia (Nikishin et al., 1993), which is marked by pronounced mantle strength (Cloetingh et al., 1999). Linking the Sedimentary Record to Processes in the Lithosphere Over the last decades basin analysis has been in the forefront of integrating sedimentary and lithosphere components of previously separated fields of geol- ogy and geophysics (Fig. 18). Integrating neotecton- ics, surface processes and lithospheric dynamics in the reconstruction of the paleo-topography of sedimen- tary basins and their flanking areas is a key objective of integrated Solid-Earth science. A fully integrated approach, combining dynamic topography and sedi- mentary basin dynamics, is also important considering the societal importance of these basins on account of their resource potential. At the same time, most of the human population resides on sedimentary basins, often close to coastal zones and deltas that are vulnerable to geological hazards inherent to the active Earth system. One major task of on-going research is to bridge the gap between historic and geological time scales in ana- lyzing lithospheric deformation rates. Major progress has been made in reconstructing the evolution of sed- imentary basins on geological time scales, incorporat- ing faulting and sedimentary phenomena. From this, we have considerably increased our insights into the dynamics of the lithosphere at large time slices (mil- lions of years). On the other hand, knowledge on present-day dynamics is rapidly growing thanks to the high spatial resolution in quantification of earth- quake hypocenters and focal mechanisms, and ver- tical motions of the land surface. Unification, cou- pling and fully 3-D application of different modelling approaches to present-day observations and the geo- logical record will permit to strengthen the recon- structive and predictive capabilities of process quan- tification. Particularly an intrinsically time-integrated approach will permit to assess in greater detail t he importance of the geological memory of lithospheric properties on present-day dynamics. This is one of the key parameters for predicting future vertical motions. Mechanical Controls on Basin Evolution: Europe’s Continental Lithosphere Studies on the mechanical properties of the Euro- pean lithosphere revealed a direct link between its thermo-tectonic age and bulk strength (Cloetingh et al., 2005, Cloetingh and Burov, 1996; Pérez-Gussinyé and Watts, 2005). On the other hand, inferences from P and S wave tomography (Goes et al., 2000a, b; Rit- ter et al., 2000, 2001) and thermo-mechanical mod- elling (Garcia-Castellanos et al., 2000) point to a 168 F. Roure et al. Iberia Iberia Iberia (south) Iberia (north) Brittany Arctic Canada Central Australia Trans Continental Arch of North America Central Asia Central Asia mantle folding whole lithosphere folding upper crustal folding 0 100 200 300 400 500 600 700 800 0 200 400 600 800 1000 1200 1400 Iberia Iberia Iberia (south) Iberia (north) Brittany Arctic Canada Central Australia Trans Continental Arch of North America Central Asia Central Asia mantle folding whole lithosphere folding upper crustal folding 0 100 200 300 400 500 600 700 800 Wavelength [km] 0 200 400 600 800 1000 1200 1400 Time/Age [Ma] Fig. 17 Comparison of observed (solid squares) and modelled (open circles) wavelengths of crustal, lithospheric mantle and whole lithospheric folding in Iberia (Cloetingh et al., 2002c) with theoretical predictions (Cloetingh et al., 1999) and other estimates (open squares) for wavelengths documented from geological and geophysical studies (Stephenson and Cloetingh, 1991; Nikishin et al., 1993; Ziegler et al., 1995; Bonnet et al., 2000). Wavelength is given as a function of the thermo- tectonic age at the time of folding. Thermo-tectonic age corre- sponds to the time elapsed since the last major perturbation of the lithosphere prior to folding. Note that neotectonic folding of Variscan lithosphere has recently also been documented for Brittany (Bonnet et al., 2000). Both Iberia and Central Asia are characterized by separate dominant wavelengths for crust and mantle folds, reflecting decoupled modes of lithosphere folding (Cloetingh et al., 2005). Modified from Cloetingh et al. (2002) Methods for studying uplift and erosion Geomorphology Maximum burial Sedimentology Fission tracks + He-dating Structural geology Fig. 18 Role of constraints from structural geology, geo- chronology, geomorphology and sedimentology in linking the sedimentary record to lithospheric processes (cartoon for coastal Norway by Japsen) pronounced weakening of the lithosphere in the Lower Rhine area owing to high upper mantle temperatures. However, the Late Neogene and Quaternary tecton- ics of the Ardennes-Lower Rhine area appear to form part of a much wider neotectonic deformation sys- tem that overprints the Late Paleozoic and Mesozoic basins of NW Europe. This is supported by geomor- phologic evidence and the results of seismicity studies in Brittany (Bonnet et al., 1998, 2000) and Normandy (Lagarde et al., 2000; Van Vliet-Lanoë et al., 2000), by data from the Ardennes-Eifel region (Meyer and Stets, 1998; Van Balen et al., 2000), the southern parts of the Upper Rhine Graben (Nivière and Winter, 2000), the Bohemian Massif (Ziegler and Dèzes, 2005, 2007) and the North German Basin (Bayer et al., 1999; Littke et al., 2008). Lithosphere-scale folding and buckling, in response to the build up of compressional intraplate stresses, can cause uplift or subsidence of relatively large areas at time scales of a few My and thus can be an impor- tant driving mechanism of neotectonic processes. For instance, the Plio-Pleistocene accelerated subsidence of the North Sea Basin is attributed to down buckling of the lithosphere in response to the build-up of the present day stress field (Van Wees and Cloetingh, 1996; Unternehr and van den Driessche, 2004). Sim- ilarly, the Vosges-Black Forest arch, which at the level of the crust-mantle boundary extends from the Massif Central into the Bohemian Massif, was rapidly uplifted during the Burdigalian (±18 Ma) and since Achievements and Challenges in Sedimentary Basins Dynamics 169 then has been maintained as a major topographic fea- ture (Ziegler and Fraefel, 2009). Uplift of this arch is attributed to lithospheric folding controlled by com- pressional stresses originating at the Alpine collision zone (Ziegler et al., 2002; Dèzes et al., 2004; Ziegler and Dèzes, 2005, 2007; Bourgeois et al., 2007). An understanding of the temporal and spatial strength distribution in the NW European lithosphere may offer quantitative insights into the patterns of its intraplate deformation (basin inversion, up thrusting of basement blocks), and particularly into the pattern of lithosphere-scale folding and buckling. Owing to the large amount of high quality geophys- ical data acquired during the last 20 years in Europe, its crustal configuration is rather well known (Dèzes and Ziegler, 2004; Tesauro et al., 2008) though signifi- cant uncertainties remain in many areas about the seis- mic and thermal thickness of the lithosphere (Babuska and Plomerova, 1992; Artemieva and Mooney, 2001; Artemieva, 2006). Nevertheless, available data helps to constrain the rheology of the European lithosphere, thus enhancing our understanding of its strength. So far, strength envelopes and the effective elastic thickness of the lithosphere have been calculated for a number of locations in Europe (Fig. 19, Cloetingh and Burov, 1996). However, as such calculations were made for scattered points only, or along transects, they provide limited information on lateral strength varia- tions of the lithosphere. Although lithospheric thick- ness and strength maps have already been constructed for the Pannonian Basin (Lankreijer et al., 1999) and the Baltic Shield (Kaikkonen et al., 2000), such maps were until recently not yet available for all of Europe. As evaluation and modelling of the response of the lithosphere to vertical and horizontal loads requires an understanding of its strength distribution, dedicated efforts were made to map the strength of the European foreland lithosphere by implementing 3D strength cal- culations (Cloetingh et al., 2005). Strength calculations of the lithosphere depend pri- marily on its thermal and compositional structure and are particularly sensitive to thermal uncertainties (Ranalli and Murphy, 1987; Vilotte et al., 1993; Ranalli, 1995; Burov and Diament, 1995). For this rea- son, the workflow aimed at the development of a 3D strength model for Europe was two-fold: (1) construc- tion of a 3D compositional model and (2) calculat- ing a 3D thermal cube. The final 3D strength cube was obtained by calculating 1D strength envelopes for each lattice point (x, y) of a regularized raster cov- ering NW-Europe (Fig. 20a). For each lattice-point the appropriate input values were obtained from a 3D compositional and thermal cube. A geological and geo- physical geographic database was used as reference for the construction of the input models. For continental realms, a 3D multi-layer compo- sitional model was constructed, consisting of one mantle-lithosphere layer, 2–3 crustal layers and an overlying sedimentary cover layer, whereas for oceanic areas a one-layer model was adopted. For the depth to the different interfaces several regional or European- scale compilations were available that are based on deep seismic reflection and refraction or surface wave dispersion studies (e.g., Panza, 1983; Calcagnile and Panza, 1987; Suhadolc and Panza, 1989; Blundell et al., 1992; Du et al., 1998; Artemieva et al., 2006). For the base of the lower crust, we strongly relied on the European Moho map of Dèzes and Ziegler (2004) (Fig. 2.1a). Regional compilation maps of the seismo- genic lithosphere thickness were used in subsequent thermal modelling as reference to the base of the ther- mal lithosphere (Babuska and Plomerova, 1993, 2001; Plomerova et al., 2002) (see Fig. 20b). Figure 21a shows the integrated strength under compression of the entire lithosphere of Western and Central Europe, whereas Fig. 21b displays the inte- grated strength of the crustal part of the lithosphere. As evident from Fig. 21, Europe’s lithosphere is char- acterized by major spatial mechanical strength varia- tions, with a pronounced contrast between the strong Proterozoic lithosphere of the East-European Platform to the northeast of the Teisseyre-Tornquist Zone (TTZ) and the relatively weak Phanerozoic lithosphere of Western Europe. A similar strength contrast occurs at the tran- sition from strong Atlantic oceanic lithosphere to the relatively weak continental lithosphere of West- ern Europe. Within the Alpine foreland, pronounced northwest-southeast trending weak zones are recog- nized that coincide with such major geologic struc- tures as the Rhine Graben System and the North Danish-Polish Trough, that are separated by the high- strength North German Basin and the Bohemian Mas- sif. Moreover, a broad zone of weak lithosphere characterizes the Massif Central and surrounding areas. In the area of the Trans-European Suture Zone, which corresponds to a zone of terranes that were [...]... positions sensor positions 1 1 mbar 25 2 2 3 3 In- situ stress curves, measured in experiment M -51 20 15 10 1 5 2 3 0 85 1 2 3 173 261 349 437 52 5 613 701 789 877 9 65 1 053 1141 1229 1317 14 05 seconds 5 top pull-apart basins in a dominantly compressional regime The Vienna Basin is probably the most famous and archetype of this type of basins It developed on top of the Alpine allochthon after emplacement... 40 Type 2 130 60 140 Th(°C) 80 Th(°C) 60 110 150 secondary FI in fluorite 1 C1a sample 1 50 55 Th(°C) 120 C1b Type 1 45 110 a D 0 2 4 F >180 B 0 4 8 0 4 8 -3 Type 2 C1b Type 1 45 -9 .5 50 -2 -1 Tm(°C) 55 60 Th(°C) C1a sample 2 -7.0 Tm(°C) secondary FI in fluorite 2 65 -6.0 0 0.00E+00 5. 00E-02 1.00E-01 1 .50 E-01 2.00E-01 2 .50 E-01 3.00E-01 3 .50 E-01 4.00E-01 4 .50 E-01 0 c A 10 20 Surface oil seeps along the... apparently resulting in a transition from stress-induced endMesozoic basin inversion to Plio-Pleistocene lithospheric folding-induced basin subsidence (Cloetingh et al., 2008) The sub-circular, about 450 km wide Paris Basin, which contains up to 3 km of mainly Mesozoic sediments, is generally regarded as a typical intracratonic basin This basin is superimposed on the deeply eroded internal parts of... 10 400 11 12 13 14 15 geologic units 600 800 1000 Temperature (C°) continental domain 250 300 350 400 450 50 0 55 0 crust mantle Depth (Km) COT oceanic domain 0 50 100 150 200 1 2 3 4 5 Fig 29 Present bathymetry and structural setting at the Norwegian passive continental margin The profile is a cross-section through a 3D model constrained by seismic and well data consistent with gravity and temperature... Sea-RingkøbingFyn chain of highs, the Northern and Southern Permian Basins (Ziegler, 1990) During this phase up to 3 ,50 0 m of Rotliegend clastics and Zechstein evaporites were deposited in the Southern Permian Basin and about 1 ,50 0 in the Northern Permian Basin During the Triassic and Early Jurassic regional thermal subsidence of both basins continued (van Wees et al., 2000), but was overprinted by regional... patterns resulting from erosion and sediment deposition provides a key step in linking the dynamics of hinterland uplift and basin subsidence and the associated mass flux The prospect of increasingly higher resolution in space and time will provide a better understanding of factors controlling the topographic evolution on continents and the subsidence of sedimentary basins along their margins During the last... lithosphere c) Depth thermal lithosphere b) thermal input 60°0'0N a) compositional input 0°0'0E 10°0'0E 20°0'0E 30°0'0E 50 °0'0N b 10°0'0W 10°0'0E base thermal lithosphere (in km) 240 120 80 40 10°0'0W 90 72 60 40 50 40 42 .5 45 30 32 .5 35 37 .5 25 27 .5 6 20 surface heatflow (in mW/m2) 40°0'0N 80 120 10 40 thermal age (in Ma) oceanic lithosp depth to moho (in km) 0°0'0E Fig 20 a) From crustal thickness (top... external forces operating at the surface of the Solid Earth To understand the interrelationship, interdependence and feedback mechanisms between internal and external forcing, lithospheric and surface processes need to be constrained Age dating of detrital minerals in sedimentary basins yields information that can be used to infer detailed spatial and temporal denudation patterns in sediment source areas... allochthon and the underlying African foreland crust Plio-Quaternary inversion of the Chelif depocentre, involving folding and erosion of Pliocene series along basin bounding faults, accounts for renewed transpression along this segment of the Dorsale Calcaire Intracratonic Basins Intracratonic basins developed in the interior of continents, generally far from active plate boundaries In cross-section they... assessed In support of the petroleum industry, numerical tools were developed that couple various surface processes during basin modelling, Achievements and Challenges in Sedimentary Basins Dynamics 1 95 (a) 3My Δα = 30Km 0 1 3 2 4 7 6 5 8 10 9 11 12 13 14 15 13 14 15 16 17 (b) 6My Δα = 60Km 0 1 2 34 5 6 7 9 8 10 12 11 (c) 9My Δα = 90Km 0 1 35 2 7 6 8 9 10 11 13 15 14 16 17 18 19 22 23 26 27 28 29 (e) 15My . al. 0 10 20 30 40 50 100 150 200 250 Km 10 20 30 40 50 60 –10 0 10 20 30 40 50 60 0 10 20 30 40 50 60 100 150 200 Km 100 150 200 250 Km 100 150 200 250 Km 70 60 50 40 70 60 50 40 30 –10 60 –10 30 0 Electromagnetic data (d) Thermal. o-pyroxene Moho UC: granite UM: dunite LC: o-pyroxene Moho wet dry 0 250 50 0 750 1000 1 250 150 0 temperature (°C) 0 250 50 0 750 1000 1 250 150 0 temperature (°C) tension compression wet dry tension compression Fig (mm) -743 -7 45 0 50 100 150 200 250 300 360mm Fig. 15 Analogue modelling of intraplate continental litho- sphere folding of Iberia (Fernández-Lozano et al., 2008). Left: incremental shortening and topographic