The European Alps, located in south-central Europe, record the closure of several ocean basins in the Mediterranean region during convergence between the African and European Plates. This contribution gives a short overview of the overall architecture of the western and central Alps of Europe and their forelands (Po plain and northern foreland), based mostly on three recent geophysical-geological transects, the locations of which are given in Figure 1. The evolution of the Alpine system is discussed in time slices, starting with Cretaceous orogeny and ending with evidence for very recent movements in the area of the Rhine Graben. Some aspects of neotectonics and earthquake hazard will be addressed as well.

Figure 1. Sketch map of the Alps, indicating locations of the three geophysicalgeological transects depicted in Figure 3.

The Major Tectonic Units of the European Alps

The simplified sketch map of the Alps in Figure 1 highlights the transition from the central to the western Alps, which will be discussed along three major transects. The Insubric line marks the northern and western boundary of the southern Alps. The southern Alps are characterized by a dominantly south-verging fold-and-thrust belt whose southern tip is stratigraphically sealed by the Messinian (7 Ma) unconformity below the Po plain. At the base of this very young (Miocene) foreland prism we find the Adriatic middle and lower crust, including the Adriatic mantle, from which 10to 15-km thick slices, consisting of basement and its Mesozoic sedimentary cover have been detached (see Figure 21.1.3c). This style of deformation points to the availability of a potential décollement (detachment) horizon within the granitic upper crust at a depth interval of 10–15 km, corresponding to temperatures in the range of 250°–375°C (assuming a gradient of 25°C/km). This depth probably corresponds to the brittle–plastic transition within granitic crust. The transition is due to the onset of crystal plasticity in quartz (at about 270°C), and/or reaction-enhanced ductility due to breakdown reactions of feldspar at about 250°C. Before this post-collisional Miocene shortening, during Paleogene plate convergence and collision, the lithosphere of the present-day southern Alps (the Adriatic Plate) formed the upper plate, under which the Penninic Valais and Piemont-Liguria Oceans, an intervening microcontinent (Briançonnais), and the European continental margin were subducted (see Figure 5a–c). Figure 1 also depicts the outlines of the Ivrea geophysical body, which represents the western edge of the Adriatic Plate. The Ivrea zone, a belt of south Alpine lower-crustal rocks, is the surface expression of the Ivrea geophysical body. Because this lower crust has been exhumed to moderate depth (corresponding to less than 300°C) during Mesozoic rifting, it represents a particularly rigid part of the south Alpine basement at the west-northwest front of the Adriatic indenter.

Most of the roughly 100-km Oligo-Miocene dextral strike-slip along the east-west striking eastern branch of the Insubric line (the Tonale line) has been taken up by dextral strike-slip movements along the Simplon ductile shear zone and the Rhone-Simplon line. Hence, from Oligocene to probably recent times, the western Alps are kinematically part of the west-northwest moving Adriatic indenter, causing west-northwest directed thrusting along the Penninic front of the western Alps and within the Dauphinois foreland. The Rhone-Simplon line continues to act as a major discontinuity up to the present day, both in terms of seismic activity and in the character of the stress regime. During the latest stages of orogeny, this west-northwest directed indenting by the Adriatic Plate possibly migrated further into the foreland, also affecting the western Molasse Basin and causing arcuate folding in the Jura Mountains.

The central Alps are characterized by ongoing north-south shortening that occurred during the Oligocene-Miocene, coeval with west-northwest to east-southeast shortening in the western Alps. These diverging transport directions necessitate an orogen-parallel extension, the effect of which is best documented by the Simplon normal fault and the exhumed amphibolite-grade Lepontine dome in its footwall. Oligo-Miocene exhumation of the Lepontine dome is the result of the combined effect of orogen-parallel extension, backthrusting along the Insubric line, and fast erosion.

The units north of the Insubric line consist of the Austroalpine nappes outcropping in eastern Switzerland and extending into Austria. Although paleogeographic provenance of these units is similar to that of the southern Alps, they consist of completely rootless slivers of basement and cover that have been detached (or delaminated) from their lithosphere during Cretaceous orogeny. These nappes have been stacked towards the west-northwest and their former (Cretaceous) tectonic front runs almost perpendicular to the present-day Alps in eastern Switzerland (Grisons). The Sesia-Dent Blanche unit of the western Alps underwent an Alpine tectono-metamorphic history that is different from that of the Austroalpine nappes and the southern Alps (subducted near the Cretaceous-Tertiary boundary). However, its pre-Alpine basement exhibits close similarities to that of the southern Alps.

The Penninic units are of extremely heterogeneous paleogeographic provenance, containing remnants of oceanic lithosphere, a continental fragment referred to as Briançonnais, as well as basement of the European margin. Deformation is penetrative and polyphase, and most of the Penninic units are overprinted by metamorphism, except for the Préalpes Romandes that have been detached and transported towards the northern foreland during the Eocene.

The Helvetic nappes have been detached from their former crystalline basement, which must be looked for in the lowermost Penninic nappes. The units still attached to the European lithosphere consist of the external massifs and their cover, slightly detached from the lower crust during the Miocene, when deformation started to migrate into the foreland, eventually displacing the western Molasse Basin and the Jura Mountains by up to 30 km from the Serravallian (12 Ma) onwards. The southern Rhine Graben represents an Eocene-Oligocene continental rift, kinematically linked to the Bresse Graben situated west of the Jura Mountains and ultimately to the opening of the western Mediterranean Basin (but not to the Alps). The geometry of Oligocene extensional faulting exerts a profound influence on Miocene to recent movements in the Jura Mountains and their northern margin in the southern Rhine Graben.

Figure 2. Paleogeographic map of the Alps, indicating the present-day position of the major paleogeographic units of the Alps.

The Major Paleogeographic Units of the Alps

The major paleogeographic units of the Alps are shown in Figure 2. Many of the units are only preserved as extremely thin slivers that were detached from the subducting lithosphere and that accreted as slices (so-called nappes) to the upper plate (Austroalpine and south Alpine units). A description of the main units follows.

European margin

This consists of external massifs and their cover (extending northward underneath the Molasse Basin) and Helvetic cover nappes, whose crystalline substratum lies within the deepest part of the Lepontine dome (lowermost “Penninic” nappes). Note that the European-derived basement can be traced southward almost to the Insubric line. This demonstrates very substantial exhumation of formerly subducted and newly accreted European lithosphere during the formation of the Lepontine dome. Its later exhumation is due to a combined effect of retro-flow (backfolding and backthrusting along the Insubric line) and unroofing by orogen-parallel extension and erosion during the post-collisional stages of orogeny. Some of these units also underwent Tertiary eclogitization.

Valais Ocean

The remnants of this ocean predominantly consist of Cretaceous Bündnerschiefer, grading into Tertiary flysch and at least partly deposited onto oceanic lithosphere. Eclogitic mafic rocks are preserved in the Versoyen of the western Alps, while blueschists and other low temperature-high pressure rocks are preserved in the Engadine window. The Valais Ocean opened near the Jurassic-Cretaceous boundary; its remnants presently define a northern Alpine suture zone that closed during the Late Eocene.

Briançonnais microcontinent

This microcontinent was attached to Iberia (Spain) and formed the northern passive continental margin of the Jurassic Piemont-Liguria Ocean, before it broke off the European margin in conjunction with the opening of the Valais Ocean. The Mesozoic cover of the Briançonnais microcontinent largely consists of platform sediments with frequent stratigraphic gaps (“mid-Penninic swell”). Its basement is preserved in the form of the Tambo-Suretta, Maggia, and Bernhard–M. Rosa nappes in the eastern, central, and western Penninic realm, respectively. Detached cover nappes form a substantial part of the Préalpes Romandes.

Piemont-Liguria Basin

This consists of oceanic lithosphere formed during the Middle Jurassic to Early Cretaceous and is characterized by a classic Alpine ophiolite suite. Seafloor spreading was followed by the deposition of radiolarites and aptycha limestones. During the Cretaceous, the deposition of trench deposits (Avers Bündnerschiefer of Eastern Switzerland and schistes lustrées of western Switzerland) indicates that the southern (Apulian) margin of this basin had been converted into an active margin. In eastern Switzerland, the PiemontLiguria units (Arosa and Platta unit) were involved in Cretaceous orogenic activity. However, the Piemont-Liguria Ocean did not completely close before the onset of orogeny in the Tertiary.

Margna-Sesia fragment

A small fragment of the Apulian margin, rifted off Apulia during the opening of the Piemont-Liguria Ocean and was later incorporated into the accretionary wedge along the active northern margin of Apulia.

Apulian margin

North of the Insubric line, this southern margin is only preserved in the form of rootless basement and cover slices (Austroalpine nappes). South of the Insubric line it corresponds to the southern Alps and their lithospheric substratum, the Adriatic Plate (which is part of the larger “Apulian” Plate).

Neotethys Ocean

This is a third oceanic domain, associated with the so-called Meliata-Hallstatt ocean, formed during the Triassic. In the Alps, only the distal passive margin is preserved (Hallstatt), while ophiolitic remnants (Meliata) are found in the Dinarides and the western Carpathians. This ocean is of significance mostly for the role it plays in understanding Cretaceous (or Eoalpine) orogeny.

Three Alpine Transects and Their Deep Structure

The major features common to all three transects, schematically sketched in Figure 3, are

(1) ESE– S-directed subduction of the European lithosphere,

(2) a gap between European and Adriatic Moho, and

(3) the presence of wedge-shaped bodies of lower crust, largely decoupled from the piling up and refolding of thin flakes of upper-crustal material (the Alpine nappes).

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Figure 3. Three schematic geophysical-geological cross sections through the western and central Alps (profile traces indicated in Figure 1). Superimposed circles mark earthquake foci for the 1980–95 time period. These foci are projected into the sections from within a 30-km traverse.

However, there are substantial differences in the geometry and kinematic evolution of the eastern transect (Figure 3c) as compared to the western and central transects (Figure 21.1.3a and b, respectively). Based on a study of these sections, the following observations are made:

• In the eastern (NRP-20 East) transect (Figure 3c) the Adriatic Moho descends northward and toward its contact with the lower (European) crust, while in the central and western transects (Figure 3a and b) this same Adriatic Moho rises toward the surface when approaching the contact zone with the European lithosphere. This contrast is also expressed in the surface geology. In the eastern transect, the southern Alps form an impressive south-vergent foreland fold-and-thrust belt (“retro-wedge”) riding above the Adriatic lower crust, while this same Adriatic lower crust is exposed in the Ivrea zone, situated at the Southeast end of the central and western transects. The Ivrea zone and Ivrea geophysical body wedge out eastward and do not extend into the area covered by the eastern transect.

• In the eastern (NRP-20 East) transect, a wedge of Adriatic lower crust is found above European lower crust and below European upper crust at its northern end. This slice of Adriatic lower crust was wedged into the European lithosphere during the Miocene, splitting apart along the boundary between the upper and lower crust. For the western (ECORSCROP) and central (NRP-20 West) transects a geometrically different process of wedging is inferred. In these latter two cases, however, the lower crustal wedge is interpreted to be derived from the European lithosphere. Hence the wedges of lower crust seen in the eastern (Figure 3c), and the western and central transects (Figure 3a and b), respectively, are of different origin and, thus, cannot be laterally connected. This implies that the Adriatic lower crust descends below the Penninic nappe stack in the eastern profile, while it rises to the surface in the western transect, independently supports the conclusion that the two lower crustal wedges are not laterally connected.

• In the area immediately north of the Insubric line, the eastern transect (NRP-20 East) exhibits substantial backthrusting and backfolding of all the Penninic nappes, including the Valais suture zone. This was associated with exhumation of the amphibolite grade Lepontine dome and the deep-seated Bergell Massif. In the western transect (ECORS-CROP), however, backthrusting does not affect the Valais suture zone and appears to be restricted to the units above this suture (within the Briançonnais upper crust). Note that Barrovian-type amphibolite-grade rocks have not been exhumed to the surface in the western transect, where the Insubric line only exhibits minor vertical offset.

• The orogenic lid of the Austroalpine nappes, under which Penninic and Helvetic nappes were accreted in the eastern transect, is absent in the western and central transects.

• Accurately located earthquakes, when orthogonally projected over a maximum distance of 30-km onto the transects of Figure 3 (see Schmid and Kissling, 2000), show significant differences in the depth of seismogenic regions. The maximum depth of earthquakes is situated near the Moho in the northern and southern forelands along the eastern traverse (Figure 3c). Beneath the Penninic units (i.e., within the Lepontine metamorphic dome), they are restricted to the thickened upper crust. Also note that the Adriatic lower crustal wedge is quiescent. Coincidence of the lower limit of seismicity with predicted isotherms based on thermal modeling suggests that the 500°C isotherm controls the cataclastic–plastic transition. Quiescence within the Adriatic lower crustal wedge further suggests that, in the case of the central Alps, stress transmission between the European and Adriatic lithosphere is largely restricted to upper crustal levels. In contrast, the earthquake distribution along the western and, to a lesser degree, along the central transects exhibits a wide, east-dipping corridor of foci affecting the entire transect, including the allochthonous European lower crust (Figure 3a and b). Thus, mechanical coupling and stress transmission between the Adriatic microplate and the European lithosphere occur along a deep-reaching seismogenic zone. This indicates a contrasting (with respect to the eastern transect) thermal regime, primarily because of the following substantial differences in kinematic evolution. First, oblique convergence and collision in the western Alps before 35 Ma must have led to a significantly smaller volume of accreted radiogenic upper crustal rocks, as compared to the central Alps, the latter being characterized by head-on convergence and collision. Secondly, double-verging displacements of the central and western Alps after 35 Ma allowed for orogen-parallel extension in the central Alps (Lepontine dome), associated with updoming of the isotherms. The Penninic realms of the central and western Alps differ significantly not only in deeper crustal architecture but also in the thickness of the seismogenic zone. Some of the earthquake foci (down to a depth of about 10–15 km are known to be associated with normal faulting within the axial zone of the Alps, while strike-slip and/or thrusting modes prevail in the northern and southern forelands of the Alps. The cause of this normal faulting within the central parts of the Alps remains uncertain (gravitational collapse and/or buoyant rise of the lithospheric root). However, compression in both forelands suggests ongoing compressional coupling between the Adriatic and European Plates, although no focal solutions are available for the deep (~15 km) foci.

In summary, the three transects reveal major differences between the central Alps (Figure 3c) and western Alps (Figure 3a and 3b). As shown in Figure 1, the boundary between these two different segments of the Alpine chain coincides with the Rhone-Simplon line, which to this day continues to be seismically active and separates different modern stress domains.

Inferences Concerning Rheologic Behavior

The maximum depth of the seismogenic zone, which is assumed to coincide with the maximum depth of cataclasis (i.e., friction-controlled and dilatant deformation) is a widely disputed topic. Some partly speculative inferences can be made based on field observations and deductions from the geometry of the present-day deep structure of the Alps. Observations made by structural geologists focusing on the study of deformation microstructures indicate the onset of crystal-plasticity at different temperatures for different minerals under natural strain rates. Anhydrite (associated with the décollement horizon of Jura-type folding) may deform crystal-plastically above about 70°C; calcite exhibits significant non-cataclastic deformation above about 180°C; quartz starts to deform by crystal plasticity above 270°C. Feldspar does not start to deform by crystal-plasticity below 450°C–500°C, but breakdown reactions in feldspar may promote reaction-enhanced ductility at lower temperatures, provided that water is available. Minerals such as hornblende and pyroxenes are more flow resistant than feldspar, and olivine does not start to deform crystal-plastically below 700°C.

At first, these data suggest a fairly shallow base for the seismogenic zone in the quartz-rich upper crust. However, elevated pore pressures are able to displace the brittle–plastic transition to greater depth (i.e., higher temperatures). On the other hand, deformation of more mafic lower crustal rocks is predicted to be controlled by cataclastic deformation at temperatures less than 450°C–500°C (i.e., down to Moho depth for an undisturbed geotherm within the foreland), assuming that the strength of these rocks is due to feldspar in the absence of significant amounts of quartz. Contrary to predictions based on the extrapolation of experimentally determined flow laws, therefore, lower crustal rocks may be flow resistant and deform by cataclastic processes. Hence, it is not surprising to find deep foci within lower crustal rocks, as shown in Figure 21.1.3 for the case of the Alpine forelands. However, lower crustal rocks may become weak within overthickened crustal roots of mountain belts, and/or if heat flow is elevated.

The geometry of the deep structure along the transects given in Figure 3 independently suggests that lower crustal rocks are flow resistant. This contrasts with a common assumption among geologists that the lower crust is generally “weak.” Lower crustal wedging demands the lower crust to remain little deformed (or undeformed) and calls for décollement horizons at the top, as well as at the base of the lower crust. Since quartz already starts to deform by crystal plasticity above 270°C, the upper crust may easily detach from the lower crust. Given the high strength contrast between feldspar and olivine, detachment at the base of the lower crust may also occur, provided that temperatures of around 450°–500°C (onset of crystal-plasticity in feldspar) are reached within the lowermost continental crust.

It is interesting to note that the Adriatic lower crustal wedge in the NRP-20 East profile (Figure 3c) is presently aseismic, in contrast to the European lower crustal wedges in Figure a, b. This points to differences in the thermal regime between the central and western Alps. In case of the NRP-20 East profile the lower limit of seismicity roughly coincides with the 500°C isotherm, as predicted by thermal modeling along this transect. This independently supports the inference that the base of the seismogenic zone in the lower crust coincides with the cataclastic–plastic transition for feldspar near the 500°C isotherm.

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Figure 4. Correlation table showing an attempt to date deformation phases and metamorphism along the eastern transect.

Evolution of the Alpine System and Its Forelands in Time Slices

Our brief discussion of the history of the Alpine system in central and western Europe focuses on the evolution along the eastern (NFP-20 East) transect of Figure 3c, where the timing of events is best constrained. Figure 4 gives a timetable of orogenic activity, while Figure 5 depicts cross sections along this eastern transect for different time slices.

Figure 5. Scaled and area-balanced sketches of the kinematic evolution of the eastern central Alps from early Tertiary convergence (a and b), to collision (c), and post-collisional shortening (d–g).

Cretaceous Orogeny

Cretaceous (or Eo-Alpine) orogeny in the eastern Alps is regarded as independent and unrelated to Tertiary orogeny because of its different kinematic scenario (top to WNW, hence almost orogen-parallel thrusting) and because it is separated from Tertiary convergence by an extensional event during the Late Cretaceous (“Ducan extension” in Figure 4). Apart from the Austro-Alpine nappes, it only affects the Piemont-Liguria units of eastern Switzerland (Arosa Platta), while the rest of the Penninic units remain largely unaffected by this orogeny, which did not propagate further to the west beyond eastern Switzerland, nor downsection into the Briançonnais units.

The attribution of a pre-Adamello phase in the southern Alps (the main deformation is of Miocene age) to Cretaceous orogenic activity is uncertain, but a precursor of the Insubric line must have been active because of the separation between the detached crustal flakes of the Austro-Alpine nappe system and the Adriatic lithosphere, which remained intact. However, the southern margin of the Piemont-Liguria margin represented an active margin, as documented by the accretionary wedge of the schistes lustrées and by the eclogitization of the Sesia unit at around the Cretaceous-Tertiary boundary.

During the various stages of Tertiary orogeny, the pre-structured Austro-Alpine nappe system, together with the Arosa-Platta ophiolites, formed a rigid upper plate (referred to as “orogenic lid” in Figure 5a–d), of which the southern Alps formed part (not depicted in Figure 5a–c, but present at the southern margin of these figures).

Early Tertiary Convergence and Subduction

(65–50 MA) During the Paleocene, the Briançonnais terrane enters the subduction zone, thereby closing the last remnants of the Piemont-Liguria Ocean in eastern Switzerland, the youngest sedimentary cover of which now forms an accretionary wedge consisting of the Avers Bündnerschiefer (Figure 5a). This southern ocean likely remained open for a longer period of time in the western Alps. After about 200 km of north-south convergence (13 mm/y) the distal margin of Europe (the future Adula nappe) enters the subduction zone at around 50 Ma, then closing the Valais Ocean (Figure 5b). Penetrative deformation during this time interval is largely restricted to the southernmost Penninic units, that is, the Briançonnais terrane (Tambo-, Surettaand Schams nappes, see Figure 4) and the Avers Bündnerschiefer of the Piemont-Liguria Ocean (Figure 5b).

Tertiary Collision (50–35 Ma)

During the middle and late Eocene (between Figure 5b and d), an additional 200-km N–S plate convergence (corresponding to 15 mm/y) was taken up by the incorporation of the Valais Ocean and the distal European margin into a growing accretionary wedge below the orogenic lid formed by the Austro-Alpine nappes. Figure 4 illustrates the migration of deformation and metamorphic events towards the northern foreland, reaching the area of the future Helvetic nappes by the end of the Eocene. Note that a total of some 400 km of north-south convergence across the central Alps involves substantial sinistral strike-slip movement across the future western Alps. Hence, the western Alps formed under a sinistrally transpressive scenario during Early Tertiary convergence and collision, with west-directed movements postdating Tertiary collision (see post-collisional stage 1).

Since the Alpine nappes in Figure 5 exclusively consist of thin slices of upper crustal basement and/or its cover, detached from their lower crustal and mantle substratum, all European (and Valaisan) lower crust (including parts of the upper crust) must have been subducted together with the mantle lithosphere (Figure 5c). Hence, north-vergent nappe stacking during this collisional stage took place within an accretionary wedge that starts to grow as more non-subductable upper crustal granitic material of the European margin enters the subduction zone. Radiogenic heat production within this granitic basement, perhaps in combination with slab break-off (depicted in Figure 5e) leads to a change in the thermal regime and to Barrovian-type (i.e., Lepontine) metamorphism.

Figure 5 ctnd. Scaled and area-balanced sketches of the kinematic evolution of the eastern central Alps from early Tertiary convergence (a and b), to collision (c), and post-collisional shortening (d–g).

Postcollisional Stage 1 (35–20 Ma)

Further growth of the accretionary wedge leads to retro-thrusting of part of the material entering the subduction zone above the steeply north-dipping Insubric line (Figure 5e and f). A “singularity point” develops within the lower part of the upper crust, separating the subducting part of the European crust from that part of the wedge which is backthrusted, sheared, and exhumed by erosion (this singularity point is near the bent arrow depicted in Figure 5f).

As can be seen from Figure 4, forward thrusting in the Helvetic nappes (i.e., the Glarus Thrust) is contemporaneous with retro- or backward thrusting along the Insubric line. The Alps evolve into a bivergent orogen, with a southern and northern foreland. Interestingly, the transition into bivergent thrusting coincides with increased rates of erosion due to the pop-up of the central Alps between fore- and retrothrusts, resulting in the transition from flysch-type to molasse-type sedimentation in the northern foreland.

North-south-directed plate convergence during this first post-collisional period amounts to about 60 km, slowing down to about 4.5 mm/y. In map view, this time interval coincides with the west-northwest-directed movement of the Adriatic Plate, now decoupled from the central Alps via dextral strike-slip movement along the Tonale line (about 100 km). Kinematically, the western Alps are now part of the west-northwest-moving Adriatic Plate and are separated from the central Alps along the Simplon ductile shear zone and later on by the Rhone-Simplon line (see Figure 21.1.1). Note that continental rifting in the Rhine and Bresse Grabens falls into this same time interval. However, this rifting is kinematically unrelated to shortening across the Alpine system, which remains in compression throughout.

Postcollisional Stage 2 (20–7? Ma)

Continued crustal overthickening within the central part of the Alpine Orogen by bivergent (retro- and pro-wedge) thrusting eventually led to rapid propagation of the deformation front from the Insubric line towards the Po plain (southern Alps), as well as towards the northern foreland (thrusting at the base of the Aar Massif and within the southern Molasse Basin) at around 20 Ma. This is depicted in Figure 5e; the timing constraints are given in Figure 4. Regarding the southern Alps, deformation stopped at around 7 Ma (Messinian unconformity).

In the northern foreland, however, the situation is more complex. During the late Serravallian (12 Ma), deformation suddenly stepped further into the foreland, incorporating the western part of the Molasse Basin and the Jura Mountains into the orogenic wedge. While décollement along Triassic evaporites is recognized by most authors as being responsible for this forward stepping of the deformation front onto the northernmost Jura Mountains up to the southern Rhine and the Bresse Grabens, two questions remain unanswered:

• Did thin-skinned deformation stop at around 7 Ma in the Jura Mountains, that is, contemporaneous with foreland deformation in the southern Alps?

• How exactly did the arc of the Jura Mountains form? By clockwise rotation of the western part of the Molasse Basin and the northern Alps? Or by west to northwest directed indentation of the western part of the central Alps?

In regard to the first question we can argue that present-day deformation is thick-skinned, hence it is likely that Jura-folding was a short-lived event (12–7 Ma). Regarding the second question we favor an indentation model, as there is evidence for counterclockwise rather than clockwise rotation of the Adriatic Plate during the Miocene. Assuming that relatively fast plate convergence across the Alpine system of Switzerland stopped at around 7 Ma, the 60-km plate convergence over the duration of this second post-collisional episode amounts to about 0.5 cm/y. Thus, plate convergence remained practically unchanged between 35 and 7 Ma. It is interesting to compare this figure of 5 mm/y to present-day shortening estimates across the Alpine system, as we do in the following section.

Recent Movements in the Upper Rhine Graben

Recent results from work in progress in the framework of EUCOR-UGENT program concerning the area of the Upper Rhine Graben in the Sundgau area west of Basel are shown in Figure 6. The Sundgauschotter have been deposited during a very short time interval from 3.2 to 2.6 Ma. Presently they outcrop within a 20-km wide corridor between the Vosges and the frontal Jura Mountains (see outlines of the base of the Sundgauschotter in Figure 6). The base of the Sundgauschotter forms an excellent reference horizon for inferring relative vertical movements during the last 3 m.y. or so (their basal part does not need to have been deposited everywhere 3.2 Ma, but they certainly must have been deposited before 2.6 Ma), provided that this basal contact is assumed to be near-horizontal at the time of deposition. The fact that these gravels were deposited in a braided river environment indicates that bases were nearly planar, with very minor slope from east to west.

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Figure 6. Contour map of the base of the Pliocene (3.2–2.6 Ma) Sundgau gravel deposits west of Basel. Assuming an originally planar base of these braided river deposits, the contours indicate folding with an amplitude of up to 200 m. This indicates substantial tectonic movements during the last 3 Ma, postdating thin-skinned Jura-type tectonics.

The contour map of the bases of the Sundgauschotter (Figure 6) suggests substantial relative vertical uplift of the southernmost part of the depositional corridor with respect to the northernmost occurrences (in the order of 250 m). Moreover, two very pronounced en-echelon anticlines, gently folding the bases of the Sundgauschotter are inferred north of the Vendlincourt fold in the Rechésy area, these gentle folds being directly observable within Upper Jurassic limestones and Oligocene deposits below the bases of the Sundgauschotter. Note also that the bases of the Sundgauschotter are affected by at least part of the folding to be observed in the Ferrette fold.

The geometry of these folds, particularly in the area immediately east of Montbéliard, suggests thickskinned reactivation of basement faults formed during Oligocene rifting. Hence, we propose a thickskinned origin for the approximately north-northwest to south-southeast directed ongoing shortening, as indicated by the northernmost (post 3 Ma) folds in the Basel area affecting the Sundgauschotter, and suggest that this shortening is very probably going on at present. This pattern suggests that thin-skinned Jura folding may indeed have stopped around 7 Ma. Such a postulate is compatible with:

(1) the present-day stress field in the Jura Mountains, as determined by in situ stress measurements, indicating that the Jura belt is no longer an active thin-skinned foldand-thrust belt,

(2) the historical Basel earthquake, which reactivated a deep-seated basement fault, and

(3) the occurrence of intracrustal earthquakes within the Molasse Basin (Figure 3c).

Closing Remarks

To some extent, the considerable complexities in structure and evolution of the European Alps, outlined in this essay, reflect the fact that our knowledge of their surface geology and deep structure is exceptionally detailed when compared with that of other orogens. Also, the Alps are particularly well exposed. Furthermore, the existence of along-strike culminations, such as the Lepontine dome (Figure 1), allows for along-strike projections of deep-seated structures to laterally adjacent areas. Such lateral (or axial) projections may cover a depth interval of up to 30 or 40 km (e.g., Figure 3c). Such projections have been a classical tool of Alpine geology since the pioneering work of Argand in the beginning of the twentieth century.

Our knowledge concerning the deep structure is relatively recent and results from a series of geophysical transects carried out cooperatively by an international group of French, Italian, Swiss, German, and Austrian scientists (ECORS-CROP, NFP-20, and very recently TRANSALP in the eastern Alps). The combination of geologic and geophysical data offers the unique opportunity to unravel the crustal structure of this orogen from its former (now eroded) surface 10 km above sea level, down to the crustal roots at more than 50 km depth.

However, in spite of the “model” character the Alps have always had, there are also good reasons to believe that they are special and that many of their features should not be extrapolated to other parts of the world. Features such as, for example, the very considerable along-strike variations in the overall architecture of the orogen, the superposition of two orogenies (Cretaceous and Tertiary) in one mountain belt, and, the former existence of three oceanic domains (Valais, Piemont-Liguria, and Neotethys Oceans) may be rather special. We will not know for certain how special the Alps are before we have data in comparable detail from other orogens. In the meantime, the Alps will continue to serve as a case study concerning the integration and interpretation of a very large set of geologic and geophysical data.

Additional Reading

Beaumont, C., Fullsack, P. H., and Hamilton, J., 1994. Styles of crustal deformation in compressional orogens caused by subduction of the underlying lithosphere. Tectonophysics, 232, 119–132.

Becker, A., 1999. In situ stress data from the Jura mountains—new results and interpretation. Terra Nova, 11, 9–15.

Burkhard, M., 1990. Ductile deformation mechanisms in micritic limestones naturally deformed at low temperatures (150–350°C). In Knipe, R. J., et al., eds., Deformation Mechanisms, Rheology, and Tectonics, Geological Society Special Publication, 54, 241–257.

Burkhard, M., and Sommaruga, A., 1998. Evolution of the western Swiss Molasse basin: Structural relations with the Alps and the Jura belt. In Mascle, A., et al., eds., Cenozoic Foreland Basins of Western Europe, Geological Society Special Publication, 134, 279–298.

Froitzheim, N., Schmid, S. M., and Frey, M., 1996. Mesozoic paleogeography and the timing of eclogite-facies metamorphism in the Alps: A working hypothesis. Eclogae Geologicae Helveticae, 89, 81–110.

Giamboni, M., Ustaszewski, K., Schmid, S. M., Schumacher, M., and Wetzel, A. (submitted). PlioPleistocene transpressional reactivation of Paleozoic and Paleogene structures in the Rhine-Bresse transform zone (Northern Switzerland and Eastern France). International Journal of Earth Sciences.

Meyer, B., Lacassin, R., Brulhet, J., and Mouroux, B., 1994. The Basel 1356 earthquake: which fault produced it? Terra Nova, 6, 54–63.

Schmid, S. M., Pfiffner, O. A., Froitzheim, N., Schönborn, G., and Kissling, E., 1996. Geophysicalgeological transect and tectonic evolution of the Swiss-Italian Alps. Tectonics, 15, 1036–1064.

Schmid, S. M., Pfiffner, O. A., and Schreurs, G., 1997. Rifting and collision in the Penninic zone of eastern Switzerland. In Pfiffner, O. A., et al., eds., Deep Structure of the Alps, Results from NFP 20, Birkhauser Verlag.

Schmid, S. M., and Kissling, E., 2000. The arc of the western Alps in the light of geophysical data on deep crustal structure. Tectonics, 19, 62–85.