1. Introduction
Continental collision leads to formation of a convergent mountain belt and to surface uplift. Erosion of this mountain belt feeds clastic sediments to the adjacent basins (e.g. Burbank, Reference Burbank2002). Tracing the different sources of synorogenic detritus provides information on the nature and extent of lithotectonic units exposed during the successive stages of continental collision. Classic provenance analysis based on sandstone petrography led to identifying several remarkable pebble lithologies, which are of particular importance for discussing the nature and location of potential source areas (e.g. Dickinson & Suczek, Reference Dickinson and Suczek1979; Garzanti, et al. Reference Garzanti, Vezzoli, Ando, Lave, Attal, France-Lanord and Decelles2007). More recently, low temperature (LT) thermochronology of detrital grains has developed into a powerful tool to trace the thermal and exhumation history of orogens (see references in Carter & Bristow, Reference Carter and Bristow2000; Bernet & Spiegel, Reference Bernet, Spiegel, Bernet and Spiegel2004; Bernet & Garver, Reference Bernet, Garver, Reiners and Ehlers2005). In the Alps as in other collision belts, the search for ophiolite clasts in synorogenic sediments is traditionally of particular importance, especially if Cr-spinel is abundant as this heavy mineral is characteristic of an ophiolitic source (e.g. Najman & Garzanti, Reference Najman and Garzanti2000; Garzanti, Ando & Scutella, Reference Garzanti, Ando and Scutella2000). When Cr-spinel are rare or absent, which is the case in the western Alpine synorogenic sediments (Evans & Mange-Rajetzky, Reference Evans, Mange-Rajetzky, Morton, Todd and Haughton1991), serpentinites mainly derived from hydrated peridotites can be useful to depict ophiolitic sources as serpentinites are particularly resistant to fluvial transport (McBride & Picard, Reference McBride and Picard1987). Usually, serpentinite clasts are seldom studied because of difficulties in determining their mineral species, mostly distinguishing between chrysotile, lizardite and antigorite. However, the mineralogy of serpentinite provides direct information on the ophiolite metamorphic history. The use of Raman spectroscopy allows overcoming the difficulties of serpentine characterization (Rinaudo, Gastaldi, Belluso, Reference Rinaudo, Gastaldi and Belluso2003; Auzende et al. Reference Auzende, Daniel, Reynard, Lemaire and Guyot2004; Groppo et al. Reference Groppo, Rinaudo, Cairo, Gastaldi and Compagnoni2006). Applied to different ophiolite-bearing thrust sheets along the Queyras–Monviso transect, this method shows that the nature of serpentine minerals evolves along the subduction wedge of the western Alps, in accordance with the different P–T conditions estimated from petrologic metamorphic studies (Auzende et al. Reference Auzende, Guillot, Devouard and Baronnet2006). On this basis, we propose here to apply the same method to sand grains and pebbles of serpentinite from synorogenic molasse deposits in the foreland basin remnants of the western Alps.
We chose the Barrême basin, a small piggy-back basin with well-preserved lower Oligocene to lower Miocene molasse formations (e.g. Evans & Elliott, Reference Evans and Elliott1999; Callec, Reference Callec2001 and references therein). The sampled syn-collisional molasse deposits are known to bear the very first detrital input from the internal Alpine zone, often termed ‘exotic clasts’ (historical references in Chauveau & Lemoine, Reference Chauveau and Lemoine1961; Morag et al. Reference Morag, Avigad, Harlavan, McWilliams and Michard2008). The early Oligocene is also the time when these internal zones underwent a major orogenic phase related to what may be regarded as the climax of the collision in the western Alps enhancing relief formation (Lardeaux et al. Reference Lardeaux, Schwartz, Tricart, Paul, Guillot, Bethoux and Masson2006; Morag et al. Reference Morag, Avigad, Harlavan, McWilliams and Michard2008; Bernet & Tricart, Reference Bernet and Tricart2011). Considering these constraints allows us to discuss the possible origin of the analysed detrital serpentinites in the context of the Alpine relief formation, from ~ 34 to 23 Ma ago.
2. Geologic setting
The Alps are located on the boundary between the European and African plates. Alpine evolution along the Eurasia–Africa boundary was initially dominated by plate divergence, which induced Mesozoic continental rifting and ocean opening. Since Late Cretaceous times, plate convergence has resulted in subduction and collision (Rosenbaum & Lister, Reference Rosenbaum and Lister2005 and references therein). The structural geometry of the western Alps classically resulted from indentation of the southern margin of Europe by the Adriatic microplate or African promontory after the closure of the intervening Tethyan Ocean (e.g. Coward & Dietrich, Reference Coward, Dietrich, Coward, Dietrich and Parks1989; Platt et al. Reference Platt, Behrmann, Cunningham, Dewey, Helman, Parish, Shepley, Wallis and Weston1989; Schmid & Kissling, Reference Schmid and Kissling2000), associated with strain partitioning in an oblique collision context between Adria and Europe (Malusà, Polino & Zattin, Reference Malusà, Polino and Zattin2009), and possibly with slab retreat (Vignaroli et al. Reference Vignaroli, Faccenna, Jolivet, Piromallo and Rossetti2008). In this study we only consider the southern part of the western Alpine arc (Fig. 1a, b). The internal zone consists of a pile of metamorphic nappes with dominant Palaeogene N- or NW-directed structures (Choukroune et al. Reference Choukroune, Ballèvre, Cobbold, Gautier, Merle and Vuichard1986). The structure of the metamorphic internal zone (Fig. 1b, c) consists of a refolded stack of units derived from the ocean (Piedmont zone) and its European margin (Briançonnais zone) (Tricart et al. Reference Tricart, Lardeaux, Schwartz and Sue2006).
Figure 1. Geologic setting of the study area. (a) Geographic setting of the western Alps showing location of (b). (b) Tectonic sketch map of the western Alps showing the Barrême basin in the external zone. The internal zone is bounded by the Penninic deep crustal thrust to the west with: QS – Queyras Schistes lustrés; V – Monviso. The rectangle locates Figure 7b. (c) Sketch-section of the present-day structure. Presently, metamorphic serpentinite with antigorite but without lizardite (HT-blueschists and eclogites) crops out along both sides of the watershed (Eastern Queyras Schistes lustrés, Monviso).
The Briançonnais zone mainly displays upper Palaeozoic to Mesozoic sediments and pre-Alpine basement rocks. A remarkable nappe stack involves the pre-, syn- and post-rift Tethyan sediments originating from a stretched margin (e.g. Claudel & Dumont, Reference Claudel and Dumont1999). This stack of cover nappes was shortened during collision and, being a mechanically contrasted multilayer, gave rise to regional W- and E-verging folds and associated thrusts. The latter are known as the Briançonnais backfolds and backthrusts and correspond to the present-day Alpine fan-shaped structure (Tricart, Reference Tricart1984) metamorphosed under greenschist-facies conditions (Goffé et al. Reference Goffé, Schwartz, Lardeaux and Bousquet2004 and references therein). The age of backthrusting is debated, and is either ascribed to the Oligocene (Tricart, Reference Tricart1984) or to the Eocene (Malusà, Polino & Martin, Reference Malusà, Polino and Martin2005) period. The Briançonnais basement consists of pre-Alpine magmatic and metamorphic rocks. Its Permo-Carboniferous sedimentary cover is variably re-worked during the Alpine orogeny. The metamorphic evolution of these basement slices, where upper-blueschist- and/or eclogite-facies conditions have been deciphered (Goffé et al. Reference Goffé, Schwartz, Lardeaux and Bousquet2004), contrasts with the evolution of the cover nappe pile. These significant metamorphic gaps are consistent with the existence of severe tectonic decoupling between the Briançonnais units.
The composite Piedmont zone (Fig. 1b, c) corresponds to the present-day juxtaposition of different structural levels of the palaeo-subduction wedge from the blueschist accretionary wedge (Schistes lustrés) to the eclogitized internal crystalline massif. The Schistes lustrés (Fig. 1c) consist of high pressure (HP) units including metamorphic marls, clays and limestones. Marls dominate (calcschists) and enclose hectometric to kilometric tectonic boudins of Triassic dolomites or Jurassic ophiolites sometimes interpreted as olistholiths (Tricart & Lemoine, Reference Tricart and Lemoine1986; Deville et al. Reference Deville, Fudral, Lagabrielle, Marthaler and Sartori1992). Burial under blueschist-facies conditions during Late Cretaceous to early Eocene times built an accretionary wedge that was strongly redeformed when collision relayed subduction in late Eocene times (Agard et al. Reference Agard, Monié, Jolivet and Goffé2002; Lardeaux et al. Reference Lardeaux, Schwartz, Tricart, Paul, Guillot, Bethoux and Masson2006; Tricart & Schwartz, Reference Tricart and Schwartz2006). In detail, the metamorphic conditions evolved from LT-blueschist conditions in the western part to high temperature (HT) blueschist in the eastern part of the complex (Tricart & Schwartz, Reference Tricart and Schwartz2006; Schwartz et al. Reference Schwartz, Tricart, Lardeaux, Guillot and Vidal2009; Bousquet et al. Reference Bousquet, Oberhänsli, Goffé, Wiederkehr, Koller, Schmid, Schuster, Engi, Berger, Martinotti, Siegesmund, Fügenschuh and Froitzheim2008).
The Monviso ophiolitic unit, squeezed between the Schistes lustrés complex and the Dora-Maira internal crystalline massif, contains major remnants of the Tethyan oceanic lithosphere that were strongly deformed and metamorphosed under eclogite-facies conditions (Lombardo et al. Reference Lombardo, Nervo, Compagnoni, Messiga, Kienast, Mével, Fiora, Piccardo and Lanza1978; Schwartz et al. Reference Schwartz, Lardeaux, Guillot and Tricart2000; Rubatto & Hermann, Reference Rubatto and Hermann2001) during Eocene times (Duchêne et al. Reference Duchêne, Blichert-Toft, Luais, Télouk, Lardeaux and Albarède1997). Contrasted eclogitic conditions (e.g. Schwartz et al. Reference Schwartz, Lardeaux, Guillot and Tricart2000) indicate that the Monviso massif is an imbricate of units rapidly exhumed (1 cm yr−1) within the subduction channel during Eocene times and accreted beneath the Schistes lustrés complex under blueschist-facies conditions at 20–35 km depth (Schwartz et al. Reference Schwartz, Lardeaux, Guillot and Tricart2000; Schwartz, Allemand & Guillot, Reference Schwartz, Allemand and Guillot2001). The Monviso eclogites are separated from the Dora-Maira massif by a ductile normal fault (Blake & Jayko, Reference Blake, Jayko, Roure, Heitzmann and Polino1990; Philippot, Reference Philippot1990; Schwartz, Allemand & Guillot, Reference Schwartz, Allemand and Guillot2001).
Located in the lowermost structural position, the Dora-Maira internal crystalline massif corresponds to a stack of more or less deeply subducted continental basement slices involved in a ‘dome-like’ structure (Fig. 1c). Here again, significantly contrasted metamorphic conditions have been inferred (Chopin, Henry & Michard, Reference Chopin, Henry and Michard1991; Compagnoni & Rolfo, Reference Compagnoni and Rolfo2003). Quartz-bearing eclogite-facies rocks crop out at the top of the Dora-Maira dome and overlie a coesite-bearing eclogitic unit. This pile of thin (< 1 km) high-to-ultrahigh pressure metamorphic units overlies the lowermost Pinerolo-Sanfront blueschist unit along a thrust contact. The latter unit is similar, with respect to their lithologies, structural position and metamorphic evolution, to the Briançonnais basement slices.
In contrast to the internal zone, the external zone of the western Alps consists of pre-Alpine crystalline basement (Pelvoux and Argentera massifs) of the proximal European margin and its detached sedimentary cover of the Southern Subalpine Chains (Fig. 1b). The Alpine external zone is less shortened (Gratier, Ménard & Arpin, Reference Gratier, Ménard, Arpin, Coward, Dietrich and Parks1989) and only slightly metamorphosed, ranging up to greenschist-facies conditions (Goffé et al. Reference Goffé, Schwartz, Lardeaux and Bousquet2004). Polystage shortening mainly occurred from Oligocene times onwards (e.g. Ford et al. Reference Ford, Duchene, Gasquet and Vanderhaeghe2006; Dumont et al. Reference Dumont, Champagnac, Crouzet and Rochat2008 and references therein). Non-metamorphic and gentle fold–thrust structures first developed in the Provence area until Eocene times in response to N–S compression. Superimposed Alpine structures sensu stricto developed from Oligocene times onwards when the effects of the Apulian plate indentation, first restricted to the internal arc, spread outwards through the external arc (Ford, Lickorish & Kusznir, Reference Ford, Lickorish and Kusznir1999). From middle Miocene times onwards up to the Present the external zone remains subject to active shortening while the internal zone undergoes extension (Sue & Tricart, Reference Sue and Tricart2003; Tricart et al. Reference Tricart, Schwartz, Sue and Lardeaux2004).
In front of the internal zone, a flexural basin developed since middle Eocene times, and was the locus of deep sea turbidite sedimentation. The main witnesses to this flysch episode are the Grès d'Annot in the south and the Grès du Champsaur in the north (e.g. Sinclair, Reference Sinclair2000; Sissingh, Reference Sissingh2001; Ford & Lickorish, Reference Ford, Lickorish, Joseph and Lomas2004), whose sedimentation started in middle and late Eocene times, respectively. Everywhere, sedimentation ended abruptly in early Oligocene times, owing to emplacement of the Helminthoid flysch nappe (exotic flysch in Fig. 1c; Kerckhove, Reference Kerckhove1969) coming from the internal zone. The structuring of the western Alpine belt in early Oligocene times was accompanied by enhanced mountain building (see discussion in Morag et al. Reference Morag, Avigad, Harlavan, McWilliams and Michard2008; Bernet & Tricart, Reference Bernet and Tricart2011). Rapid erosion of this mountain belt might explain the fast regional cooling of the western Alpine belt during Oligocene times and the sudden supply of coarse metamorphic detritus along the northwestern flank of the western Alps (e.g. Morag et al. Reference Morag, Avigad, Harlavan, McWilliams and Michard2008; Bernet et al. Reference Bernet, Brandon, Garver, Balestieri, Ventura and Zattin2009), and in small basins developed on top of the belt (e.g. Polino, Ruffini & Ricci, Reference Polino, Ruffini and Ricci1991; Carrapa, Wijbrans & Bertotti, Reference Carrapa, Wijbrans and Bertotti2003). Nevertheless, the volume of these clastic sediments is globally modest, and the Adriatic foredeep on the internal side of the western Alps was starved of orogenic detritus (Garzanti & Malusà, Reference Garzanti and Malusà2008). As a matter of fact, the early Oligocene period is also the time of the flysch-to-molasse transition in the European foreland basin when flysch basins closed and were incorporated into the collision wedge. From that time onwards, small molasse basins formed in more and more external locations, as the Alpine deformation front migrated outwards through the external arc (e.g. Ford, Lickorish & Kusznir, Reference Ford, Lickorish and Kusznir1999). This is well illustrated in the Barrême basin (Fig. 1b), where thin Palaeogene marine detrital formations (Fig. 2) represent a shallow water equivalent to the more proximal Grès du Champsaur and Grès d'Annot flysch formations (e.g. Joseph & Lomas, Reference Joseph, Lomas, Joseph and Lomas2004). The overlying molasse sediments in the Barrême basin were deposited as the basin was transported on the Digne thrust sheet (de Graciansky, Reference de Graciansky1972; Artoni & Meckel, Reference Artoni, Meckel, Mascle, Puigdefàbregas, Luterbacher and Fernàndez1998; Evans & Elliot, Reference Evans and Elliott1999). Subsequently, in late Oligocene times the stack of the metamorphic Briançonnais and Piedmont nappes was exhumed in brittle conditions (e.g. Malusà et al. Reference Malusà, Polino, Zattin, Bigazzi, Martin and Piana2005) and transported along the Briançonnais thrust front (Fig. 1). Possible modifications of detrital pathways at this stage will be investigated in this paper.
Figure 2. Simplified stratigraphic log of the Barrême basin molasse series. Depositional ages are from Callec (Reference Callec2001).
3. The sampled material
The Eocene to lower Oligocene so-called ‘Nummulitic’ transgressive sequence covered the European foreland of the outward migrating Alpine belt (Sinclair, Reference Sinclair1997; Ford, Lickorish & Kusznir, Reference Ford, Lickorish and Kusznir1999). The ‘Nummulitic’ series was mainly deposited in a deep and underfilled flexural basin and it generally overlies a continental erosional surface. This sequence always shows the succession of three sedimentary facies of increasing water depth (the so-called ‘Nummulitic trilogy’): (1) nearshore, Nummulites-bearing platform limestones; (2) hemipelagic foraminiferal (Globigerina) marls (Blue Marls Formation), and (3) clastic sediments. In the Barrême, the latter first appeared within the Blue Marls Formation as thin turbiditic sandstones, the Grès de Ville Formation. The latter formation is capped by lower Oligocene marine conglomerates (Clumanc and Saint-Lions conglomerates) interpreted as channelized fan delta slope deposits (Artoni & Meckel, Reference Artoni, Meckel, Mascle, Puigdefàbregas, Luterbacher and Fernàndez1998; Fig. 2).
In the Barrême basin, the Grès de Ville (Fig. 2) consists of silty mudstone and thin-bedded fine-grained sandstone with an erosive base, graded bedding and cross-stratification. They recorded N-directed palaeocurrents thanks to various sedimentary structures (flute casts, grooves, crescent and prod marks) and their clastic material originates from the Provence–Corsica–Sardinia domain (Evans & Elliott, Reference Evans and Elliott1999; Callec, Reference Callec2001). By contrast, the Clumanc and Saint-Lions conglomerates are interpreted as a Gilbert delta with W- to SW-directed transport (Callec, Reference Callec2001). They are dominantly composed of coarse pebbles (generally 5 to 20 cm) with a sandy and fine-grained pebbly matrix (Callec, Reference Callec2001) (Fig. 3). Different pebble populations were recognized early on by geologist Termier (1895 in Chauveau & Lemoine Reference Chauveau and Lemoine1961; de Graciansky, Lemoine & Saliot, Reference de Graciansky, Lemoine and Saliot1971) and especially Bodelle (Reference Bodelle1971), and are summarized in Table 1.
Table 1. Pebble population density of Clumanc and Saint-Lions conglomerates, modified after Bodelle (Reference Bodelle1971)
The lower layers of the lower Oligocene Clumanc conglomerate contain mainly limestone pebbles, and a lower proportion of metamorphic and eruptive rock. The upper layers contain more exotic rocks such as granitoids and various ophiolititic pebbles (gabbros, pillow-lava cortex, dolerites and serpentinites).
Figure 3. Conglomerates and sandstones from the Oligo-Miocene Barrême basin. (a) Outcrop view of the upper layer of the Clumanc Formation showing coarse conglomerate channels with erosive base and graded bedding incised in sandstones. The latter are interpreted as turbiditic deposits (Artoni & Meckel, Reference Artoni, Meckel, Mascle, Puigdefàbregas, Luterbacher and Fernàndez1998; Evans & Elliott, Reference Evans and Elliott1999; Callec, Reference Callec2001). (b) Close-up view of matrix-supported pebbles in coarse sandstones. Centimetric serpentinite pebbles are underlined by black arrows, others are limestones. (c) Investigated serpentinite pebbles: sizes range from 3 cm to 6 cm. (d) Outcrop view of the Grès verts Formation: coarse sandstone-filled channels interbedded in flood-plain fine-grained silty-sandstones are interpreted as fluvial channels and over-bank deposits, respectively (Evans & Elliott, Reference Evans and Elliott1999). (e) Facies of the Grès verts sandstones. (f) Isolated grains of the Grès verts Formation, bulk composition to the left and sorted serpentinite grains to the right.
The Barrême thrust-top W-verging syncline, accommodating W–E shortening, was fed by a deltaic system providing local Cretaceous limestone (lowermost Clumanc conglomerate beds) and further internal Alpine clasts, including metamorphic and eruptive rocks (upper Clumanc beds and Saint-Lions conglomerates) (Fig. 2; Table 1). The metamorphic clasts of internal origin were initially mentioned by Termier (1895 in Chauveau & Lemoine, Reference Chauveau and Lemoine1961). These internal (or Penninic) pebbles mainly consist of ophiolitic clasts (oceanic crust or mantle) and of associated pelagic sediments, in particular the easily recognizable red radiolarian cherts (Fig. 3). The ophiolitic clasts are mainly massive or brecciated basalts but some variolithic fragments of pillow-lava cortex and rare gabbros have also been found. Morag et al. (Reference Morag, Avigad, Harlavan, McWilliams and Michard2008) described garnet-blueschist pebbles. The 40Ar–39Ar geochronology of phengites from these metamorphic pebbles yielded a 34 ± 3 Ma age, suggesting a short time lag between exhumation of these rocks and their deposition in the lower Oligocene (34–28 Ma) basin. The Penninic provenance of the exotic pebbles in the Clumanc and Saint-Lions conglomerates (Fig. 2; Table 1) is consistent with the observed western to southwestern sedimentary transport directions. The latter strongly differs from the previous (Eocene to earliest Oligocene) N-directed transport, which longitudinally fed the Barrême basin from the uplifted Provence–Corsica–Sardinia domain (e.g. de Graciansky, Reference de Graciansky1982; Evans & Mange-Rajetzky, Reference Evans, Mange-Rajetzky, Morton, Todd and Haughton1991; Callec, Reference Callec2001; Joseph & Lomas, Reference Joseph, Lomas, Joseph and Lomas2004). The high amount of serpentinite pebbles we found was unexpected considering the literature (e.g. Chauveau & Lemoine, Reference Chauveau and Lemoine1961).
The Grès verts are dated as Aquitanian (Evans & Mange-Rajetsky, Reference Evans, Mange-Rajetzky, Morton, Todd and Haughton1991), which is younger than the Clumanc and Saint-Lions conglomerates (Fig. 2; Table 1). These fine silty-sandstone and siltstones show crossed-bedding typical of low-energy fluvial channels with ephemeral over-bank deposits. The heavy mineral content of the Grès verts is varied, but includes many serpentinite grains (Fig. 3) with a suspected source corresponding to the Embrunais-Ubaye nappes (Evans & Mange-Rajetsky, Reference Evans, Mange-Rajetzky, Morton, Todd and Haughton1991).
We analysed 38 small serpentinite pebbles, each 1 to 3 cm in diameter, sharing a pronounced flat and blunt shape, sampled in the ‘La Poste’ conglomerate at Clumanc (Table 2; Fig. 3). More precisely, the majority of our serpentinite pebbles come from the lowermost layers, while historically serpentinite pebbles were essentially described in the middle layers of the La Poste conglomerate (e.g. Chauveau & Lemoine, Reference Chauveau and Lemoine1961) (Table 1). Therefore, we sampled the very earliest deposits of exotic clasts in the molasse basin, which is important for the subsequent interpretation. In addition, we sampled the stratigraphically younger continental deposits of the Aquitanian Grès verts in the same Barrême molasse series. This sandstone formation is exceptionally rich in serpentinite grains, and 45 single grains were analysed for this study (Table 2; Fig. 3).
Table 2. Location and petrologic nature of serpentinite pebble samples in the Barrême basin. The proportion of antigorite (Atg) versus lizardite (Lz) present in pebbles is statistically indicated
4. Methodology: Raman spectroscopy
Serpentinite minerals are hydrous phyllosilicates (up to 13 wt% water) formed during the hydration of basic and ultrabasic rocks. Serpentinites, shown here with a simplified structural formula of (Mg, Fe2+)3 Si2O5(OH)4, are made of superposed tetrahedral and octahedral layers. The structures of the three serpentinite species lizardite, chrysotile and antigorite are distinguished by different distortions and spatial arrangements of the tetrahedral and octahedral layers. Layers are flat in lizardite, rolled up in chrysotile and corrugated in antigorite (e.g. Wicks & O'Hanley, Reference Wicks, O'Hanley and Bailey1988). Natural serpentinites sampled in high-grade metamorphic environments show antigorite as the most abundant species (Scambelluri et al. Reference Scambelluri, Muntener, Hermann, Piccardo and Trommsdorff1995, Guillot, Hattori & De Sigoyer, Reference Guillot, Hattori and De Sigoyer2000; Auzende et al. Reference Auzende, Devouard, Guillot, Daniel, Baronnet and Lardeaux2002; Groppo & Compagnoni, Reference Groppo and Compagnoni2007). Experimental studies have also shown that antigorite is the stable variety under high pressure conditions (Ulmer & Trommsdorff, Reference Ulmer and Trommsdorff1995; Wunder & Schreyer, Reference Wunder and Schreyer1997; Auzende et al. Reference Auzende, Daniel, Reynard, Lemaire and Guyot2004). Lizardite and chrysotile are the major varieties of pseudomorphic textures observed in low-grade metamorphic serpentinites from oceanic lithosphere or ophiolites (Andréani et al. Reference Andréani, Mével, Boullier and Escartin2007). In this study, the different serpentinite species were characterized by Raman spectroscopy, coupled with petrographic observations on polished thin-sections of clasts. Raman spectroscopy was performed at the Laboratoire des Sciences de le Terre, Ecole Normale Supérieure de Lyon (ENS Lyon), France, with a Horiba Jobin-Yvon LabRam HR800 apparatus. The light source was an argon-ion laser to achieve 514.5 nm wavelengths. An Olympus™ BX30 open microscope equipped with × 50 and ×80 long-working distance objectives was coupled to a spectrometer to focus the laser beam to a 3 μm spot diameter. The Raman signal was collected in the backscattered direction. Acquisition time was about 60 seconds, distributed over three accumulation cycles, with a laser output power on the sample adjusted to between 10 and 60 mW, depending on the intensity of the signal. Raw or polished serpentinite clast samples were indifferently used. The spectral resolution was 1 cm−1 using 1800 gr mm−1 grating. The spectral regions from 150 to 1150 cm−1 and from 3600 to 3720 cm−1 (Fig. 4) were investigated because they include lattice vibrational modes characteristic of serpentinite species (Rinaudo, Gastaldi & Belluso, Reference Rinaudo, Gastaldi and Belluso2003; Auzende et al. Reference Auzende, Daniel, Reynard, Lemaire and Guyot2004; Groppo et al. Reference Groppo, Rinaudo, Cairo, Gastaldi and Compagnoni2006). The bands detected in these spectral regions are indicative of the molecular and crystalline structure of the sample. For all spectra the assignments of the band position and band width were determined using the Peakfit© software.
Figure 4. Characteristic Raman spectra for the different varieties of Alpine serpentinite with band assignments according to Rinaudo, Gastaldi & Belluso (Reference Rinaudo, Gastaldi and Belluso2003), Groppo et al. (Reference Groppo, Rinaudo, Cairo, Gastaldi and Compagnoni2006) and Auzende et al. (Reference Auzende, Guillot, Devouard and Baronnet2006).
In the low frequency region, four main peaks near 230, 390, 690 and 1100 cm−1 characterize the spectra of lizardite and chrysotile (Rinaudo, Gastaldi & Belluso, Reference Rinaudo, Gastaldi and Belluso2003; Groppo et al. Reference Groppo, Rinaudo, Cairo, Gastaldi and Compagnoni2006). At lower frequencies, intense antigorite peaks occur at 226, 373, 680 and 1043 cm−1. Therefore, the peak positions are different in comparison to other serpentines, and the peaks are much broader (Fig. 4). The peak at 1043 cm−1 allows easy distinguishing of the antigorite variety. Although weak, differences between chrysotile and lizardite can be clearly identified because of the sharpness of the Raman lines. In particular, a single high frequency band at 1100 cm−1 is observed in chrysotile whereas several convoluted bands are observed between 1060 and 1100 cm−1 in lizardite (Rinaudo, Gastaldi & Belluso, Reference Rinaudo, Gastaldi and Belluso2003; Groppo et al. Reference Groppo, Rinaudo, Cairo, Gastaldi and Compagnoni2006). In the high frequency region (Fig. 4), the OH vibrational mode shows a spectrum with an intense single peak at 3697 cm−1 in the case of chrysotile and a bimodal shaped spectra for the lizardite and antigorite (Fig. 4). The spectrum of lizardite is characterized by a fine bimodal shape with a primary band at 3685 cm−1 and a secondary band at 3703 cm−1. The antigorite species presents a spectrum broader than the lizardite with a primary band near 3665 cm−1 and a secondary band centred at 3697 cm−1. One to 20 Raman spectra were acquired on each lower Oligocene pebble and one spectrum was acquired on each single grain coming from the upper Oligocene/lower Miocene sandstones (Table 2).
5. Serpentinite mineralogy
For the 38 serpentinite clasts from the Clumanc conglomerate that we analysed, all the Raman spectra are typical for antigorite (Fig. 5a). No other variety of serpentinite was observed. This observation is confirmed by petrographic thin-section analysis. The initial ultrabasic rocks were fully recrystallized and the antigorite minerals form a typical interlocking texture that replaces olivine. The original magmatic pyroxene was totally replaced by oriented antigorite flake aggregates constituting bastite (Fig. 6).
Figure 5. Typical Raman spectroscopy spectra of serpentinite varieties obtained from detrital ultramafic clasts (LP01) or sand grains (MB153 and MB154) from the Oligocene–Miocene Barrême basin. For each Raman spectrum the signal was analysed on two acquisition spectral windows and the characteristic bands were indicated. (a) and (b) Raman spectra of antigorite obtained, respectively, from the Clumanc conglomerate (sample LP01) and from the coarse to fine-grained sand fraction of the Grès verts (sample MB154). (c) and (d) Raman spectra of lizardite obtained, respectively, from the coarse-grained fraction of sample MB154, and from the homogeneously fine-grained sandstone sample MB153.
Figure 6. (a) Macroscopic view of the serpentinite pebble from the Clumanc conglomerate (LP01 sample). (b) Typical microstructures observed in thin-section under a polarizing optical microscope (crossed polarized). The matrix was full recrystallized into flakes of antigorite (atg) with an interlocking structure. The primary peridotitic pyroxene was replaced by oriented antigorite flake aggregates corresponding to ‘bastite’ antigorite (bst).
In the Aquitanian sandstone ‘Grès verts’, 45 single grains were analysed; the Raman spectroscopy analysis of the single grains (Table 2) shows a mixture of dominant lizardite and subordinate antigorite. More precisely, lizardite is the only serpentine species present (sample MB153, Fig. 5c; Table 2) in the fine-grained, well-sorted sandstone layers. In contrast, the moderately to poorly sorted sandstone layers, with a larger variety of grain sizes in the fine- to coarse-grained range, display 20% of antigorite in the coarser fraction (sample MB154, Fig. 5b; Table 2) but lizardite still dominates (sample MB154, Fig. 5d; Table 2). Nowhere were chrysotile and talc detected.
6. Discussion
6.a. Origin of serpentinite pebbles and sand grains
Despite the difficulty of precisely establishing the P–T stability field of serpentine species (e.g. Evans, Reference Evans2004) (Fig. 7a), an empiric approach has been developed by comparing the worldwide occurrence, including the western Alps, of serpentinite species with the P–T conditions recorded by associated metamorphic rocks in HP–LT metamorphic environments (Guillot, Hattori & De Sigoyer, Reference Guillot, Hattori and De Sigoyer2000; Auzende et al. Reference Auzende, Devouard, Guillot, Daniel, Baronnet and Lardeaux2002, Reference Auzende, Guillot, Devouard and Baronnet2006; Groppo et al. Reference Groppo, Rinaudo, Cairo, Gastaldi and Compagnoni2006; Guillot et al. Reference Guillot, Hattori, Agard, Schwartz, Vidal, Lallemand and Funiciello2009). In summary, chrysotile and lizardite form under greenschist-facies conditions in oceanic crust or ophiolites, while lizardite alone testifies to upper-greenschist and LT-blueschist-facies conditions (Fig. 7a). Antigorite progressively replaces lizardite under intermediate blueschist-facies conditions. Antigorite becomes the exclusive serpentinite variety under HT-blueschist- to eclogite-facies conditions (Fig. 7a).
Figure 7. (a) Phase diagram of antigorite and lizardite (after Evans, Reference Evans2004) with Lz – lizardite, Atg – antigorite, Tlc – talc, Brc – brucite, Fo – forsterite. Up to 380°C the lizardite is entirely destabilized. The blueschist pressure–temperature (P–T) conditions in the Queyras Schistes lustrés are indicated. Metamorphic facies are from Spear (Reference Spear1993): PP – prehnite pumpellyite, GS – greenschist, EA – epidote amphibolite. (b) Simplified tectonic map of the Schistes lustrés of the Piedmont zone. This domain is bounded by the Briançonnais zone to the west and by the internal crystalline massif of Dora-Maira and by the Monviso eclogitic unit to the east. The Schistes lustrés complex corresponds to a calcschist-rich accretionary wedge, enclosing scattered ophiolitic bodies, metamorphosed under HP-facies conditions, from LT-blueschist-facies conditions (lizardite and antigorite-bearing serpentinite) to the west to HT-blueschist-facies conditions (only antigorite-bearing serpentinite) to the east.
Thus, antigorite-bearing serpentinite pebbles from the lower Oligocene Clumanc conglomerate are only derived from HP–LT ophiolitic rocks metamorphosed under HT-blueschist- or eclogite-facies conditions, i.e. the internal zone (Fig. 1, 7b).
As palaeocurrent indicators favour a NE provenance for the detrital grains, the source of these pebbles has to be more precisely searched for in the SE vicinity of the Pelvoux massif and not in more southerly HP–LT metamorphic assemblages such as the Voltri massif, presently exposed to the south of the Tertiary Piedmont basin (Fig. 1b). Pure antigorite serpentinites are an important component of the HT-blueschist Schistes lustrés unit along the drainage divide between France and Italy (Fig. 7b). Pure antigorite serpentinites also crop out further east, in the Monviso eclogitic complex in Italy (Schwartz et al. Reference Schwartz, Lardeaux, Guillot and Tricart2000; Auzende et al. Reference Auzende, Guillot, Devouard and Baronnet2006).
The occurrence of both lizardite and antigorite grains in the upper Oligocene/lower Miocene serpentinite-bearing sandstones (Grès verts) contrast with the sole occurrence of antigorite pebbles in the lower Oligocene conglomerates. This suggests that the coarser antigorite fraction in the upper Oligocene/lower Miocene sandstone could derive from the same source, i.e. the HT-blueschist Schistes lustrés (or the Monviso eclogites) (Fig. 7a, b). Nevertheless, the occurrence of lizardite suggests a different source area, a less metamorphosed ophiolitic source, having escaped subduction and collision processes and only recording oceanic hydrothermal metamorphism. In the present-day Alpine framework, only the upper ophiolitic unit in the Chenaillet massif (Montgenèvre pass, upper Durance valley) escaped the subduction–collision metamorphic imprints, representing a part of the Tethyan oceanic lithosphere obducted onto the subduction wedge, towards the European continental margin (Goffé et al. Reference Goffé, Schwartz, Lardeaux and Bousquet2004; Chalot-Prat, Reference Chalot-Prat, Foulger, Natland, Presnall and Anderson2005; Schwartz et al. Reference Schwartz, Lardeaux, Tricart, Guillot and Labrin2007). An alternative interpretation is that both grains (antigorite and lizardite) were derived from a unique source in which lizardite and antigorite coexisted. This has been observed in the present-day eastern part of the Piedmont Schistes lustrés complex (Fig. 7b), which was metamorphosed only under LT-blueschist-facies conditions (Agard, Vidal & Goffé, Reference Agard, Vidal and Goffé2001; Goffé et al. Reference Goffé, Schwartz, Lardeaux and Bousquet2004; Tricart & Schwartz, Reference Tricart and Schwartz2006). Concerning the antigorite-bearing upper Oligocene/lower Miocene sands, one cannot preclude the possibility that they partly result from recycling of Oligocene detrital sediments.
6.b. The heterogeneity of sources in terms of metamorphic grade
Contrary to the serpentinite clasts, other Penninic pebbles within lower Oligocene conglomerates have been analysed in the lower 70s (e.g. de Graciansky, Lemoine & Saliot, Reference de Graciansky, Lemoine and Saliot1971). The basalts and their derivatives (‘prasinite’, ‘diabases’) essentially bear zeolite or greenschist-facies metamorphic assemblages, more rarely a blueschist-facies imprint. The rare gabbros bear different types of blueschist assemblages; ‘calcschist’ pebbles that likely originate from the Schistes lustrés complex are also observed (Morag et al. Reference Morag, Avigad, Harlavan, McWilliams and Michard2008). The abundant red-purple pebbles of radiolarian cherts only show a very low-grade metamorphic imprint as radiolarites are preserved (Cordey et al. in press).
The upper Oligocene/lower Miocene sandstones contain a mixture of grains of serpentinites, unmetamorphosed radiolarites and LT (lawsonite-bearing) blueschist basalts. As for the lower Oligocene molasse, the upper Oligocene/lower Miocene molasse shows detrital grains were derived from source areas that experienced contrasted metamorphic conditions from those of the LT-blueschist facies to low-grade metamorphic conditions.
This heterogeneity both in the lower Oligocene and upper Oligocene/lower Miocene molasse deposits implies the erosion of a stack of ophiolite nappes that recorded contrasted P–T conditions, having the lithologic and metamorphic signatures of the Palaeogene Alpine accretionary wedge (Agard et al. Reference Agard, Monié, Jolivet and Goffé2002; Lardeaux et al. Reference Lardeaux, Schwartz, Tricart, Paul, Guillot, Bethoux and Masson2006; Yamato et al. Reference Yamato, Agard, Burov, Le Pourhiet, Jolivet and Tiberi2007; Guillot et al. Reference Guillot, Hattori, Agard, Schwartz, Vidal, Lallemand and Funiciello2009).
The provenance of detrital sediments can be discussed. Indeed, it is classically admitted that the source of ophiolitic clasts in the lower Oligocene molasse of the Barrême basin was located in the Embrunais-Ubaye nappes (e.g. Evans & Mange-Rajetzky, Reference Evans, Mange-Rajetzky, Morton, Todd and Haughton1991). In the present-day situation, these nappes represent the closest outcrops of the internal or Penninic units, in an E to NE direction from the Barrême basin, which is the provenance direction of exotic detritus according to sedimentological analyses (Callec, Reference Callec2001). Because these nappes are devoid of any HP–LT metamorphic imprint, they cannot represent the source of HP–LT metamorphic detritus like the antigorite-bearing serpentinite we sampled in the Barrême basin.
6.c. The sudden arrival of mixed metamorphic clasts into the Barrême molasse basin
The sudden erosion in early Oligocene times of the stack of variably metamorphosed nappes inherited from the Palaeogene accretionary wedge fits well with two observations concerning the internal zone, SE of the Pelvoux massif.
First, structural analysis allows identifying a severe early Oligocene shortening. The corresponding post-nappe folds and thrusts are associated with the second and third regional schistosities in the Briançonnais zone. They display, respectively, outward and inward vergence, resulting in the so-called Briançonnais fan-shaped structure (Tricart et al. Reference Tricart, Lardeaux, Schwartz and Sue2006 and reference therein). In the whole internal zone, it has been proposed that this major syncollisional shortening was responsible for the bending of the arc (Thomas et al. Reference Thomas, Claudel, Collombet, Tricart, Chauvin and Dumont1999; Collombet et al. Reference Collombet, Thomas, Chauvin, Tricart, Bouillin and Gratier2002).
Second, detrital thermochronological analysis of Barrême Oligocene molasse recently allowed the estimation of very short lag times from fission-track ages for zircons grains (Bernet & Tricart, Reference Bernet and Tricart2011; Jourdan et al. Reference Jourdan, Bernet, Tricart, Guillot, Schwartz and Dumont2011), consistent with 40Ar–39Ar ages of phengites in pebbles from the same sediments (Morag et al. Reference Morag, Avigad, Harlavan, McWilliams and Michard2008). Comparison to in situ thermochronological data acquired from the internal zone led Bernet & Tricart (Reference Bernet and Tricart2011) and Jourdan et al. (Reference Jourdan, Bernet, Tricart, Guillot, Schwartz and Dumont2011) to propose erosion-driven fast exhumation (1–3 mm y−1) in early Oligocene times. Tectonic exhumation is likely to have been enhanced by erosional processes. Consequently, at that time, these internal zones are supposed to have formed an elevated and narrow mountain range. In summary, converging approaches allow us to propose that the source of mixed metamorphic and very low-grade metamorphic ophiolitic detritus that suddenly fed into the Barrême molasse basin during early Oligocene times was an elevated and possibly narrow mountain range (Fig. 8a). The volume of Oligocene sediments preserved in the proximal Alpine foreland (Barrême basin) is quite low but a much larger amount of sediment is ponded within the Oligocene grabens of the Liguro-Provençal rift, the Rhône valley and the Gulf of Lion (Séranne, Reference Sérrane, Durand and Jolivet1999).
Figure 8. Palinspastic maps and cross-section reconstructions showing the evolution of the relief zones from early to late Oligocene times (based on Schmid & Kissling, Reference Schmid and Kissling2000; Ford et al. Reference Ford, Duchene, Gasquet and Vanderhaeghe2006). (a) During early Oligocene times the main relief zones (white cross) were located above the Dora-Maira massif and in Provence. These two main relief zones feed the Barrême basin while the Tertiary Piedmont basin and Torino Hills are only fed by the internal Alpine relief with blueschist clasts. (b) During late Oligocene times the activation of the Briançonnais thrust front allows the formation of a westward relief zone feeding the Barrême basin while in the internal Alpine zone the persistence of high relief feeds the Tertiary Piedmont basin and the Torino Hills with eclogitic clasts.
6.d. Implications for the evolution of the internal western Alps
Looking for the potential source areas of the sand and pebbles of antigorite-bearing serpentinites in the Barrême basin molasse deposits, we took into account different constraints relative to the metamorphic grade, and structural and orogenic history of the internal zone, but also sedimentological constraints relative to the Barrême basin stratigraphy. The source we are looking for must present lithologies such as those presently exposed in the Piedmont zone. As a working hypothesis we consider a generally N–S-trending, elevated, subaerially exposed accretionary wedge (Platt, Reference Platt1986, Reference Platt1987, Reference Platt1993), which experienced syncollisional shortening and unroofing during early Oligocene times (Fig. 8a). Its shape probably corresponded to the present-day internal zone, but extended outwards to the presently exposed exotic flysch nappes (Fig. 1b, c). The pile of shallow-dipping nappes displays intense fold and thrust structures, due to early Oligocene shortening, especially westward-directed structures (Fig. 8). Nevertheless, the whole structure was inherited from the Palaeogene accretionary wedge and displays a general metamorphic gradient. The base of the nappe pile corresponds to Subbriançonnais and Briançonnais nappes with HP (lawsonite-bearing) greenschist metamorphism, being covered by the Piedmont nappes inherited from the Palaeogene accretionary wedge. In the frontal part, the LT-blueschist-facies unit cropped out, while HT-blueschists were already exhumed (Fig. 8) in the steep rear part of the internal wedge (Yamato et al. Reference Yamato, Agard, Burov, Le Pourhiet, Jolivet and Tiberi2007).
The location of eclogitic units within the early Oligocene structure deserves specific discussion. Geometrically, they possibly could have topped this wedge, cropping out along the highest parts of the corresponding mountain range, knowing that the antigorite-bearing clasts may as well originate from both blueschist- and eclogite-bearing nappes. Up to now no evidence for the erosion of eclogites was found in the Barrême basin. Presently, eclogites are only exposed along the Italian flank of the Alps forming the present Monviso massif (Lombardo et al. Reference Lombardo, Nervo, Compagnoni, Messiga, Kienast, Mével, Fiora, Piccardo and Lanza1978; Schwartz et al. Reference Schwartz, Lardeaux, Guillot and Tricart2000). Further to the north they crop out in the high Maurienne valley (Rolland et al. Reference Rolland, Lardeaux, Guillot and Nicollet2000) and in the Zermatt-Saas zone (Barnicoat, Reference Barnicoat1985). In the opposite direction, and also very far away, eclogites crop out in the Voltri group of Liguria, where their first detrital reworking in the Tertiary Piedmont basin started during early Oligocene times (Federico et al. Reference Federico, Capponi, Crispini and Scambelluri2004). Concerning the timing of exhumation along our transect, eclogite boulders do not appear in the Torino Hills conglomerates, at the internal foot of the mountain range, before early Miocene times (Polino, Ruffini & Ricci, Reference Polino, Ruffini and Ricci1991). Fission-track analyses in the Monviso massif confirm that these eclogites were exhumed later than Miocene times and were structurally located beneath the Schistes lustrés accretionary wedge (Schwartz et al. Reference Schwartz, Lardeaux, Tricart, Guillot and Labrin2007; Yamato et al. Reference Yamato, Agard, Burov, Le Pourhiet, Jolivet and Tiberi2007). By early Oligocene times the eclogitic unit forming the serpentinite subduction channel was not yet being exhumed along the studied transect (e.g. Guillot et al. Reference Guillot, Hattori, Agard, Schwartz, Vidal, Lallemand and Funiciello2009). The only exposed HP–LT metamorphic rocks were the blueschists coming from the Schistes lustrés accretionary subduction wedge. The variation of serpentine species from antigorite to lizardite between early Oligocene and late Oligocene/early Miocene times in the Barrême basin is interpreted as a shift of the palaeo-relief through time enhancing drainage system modifications. During early Oligocene times the prominent relief was probably located above the HT-blueschist zone (Fig. 8a) and progressively a second zone of relief developed westwards above the LT-blueschist zone (Fig. 8b). Nevertheless, in the lower Miocene Torino Hills conglomerates, the abundant arrival of decimetric to metric blocks of blueschists, with serpentinites representing more than 50% of the detrital sediments, and the first arrival of eclogites (2%) (Polino, Ruffini & Ricci, Reference Polino, Ruffini and Ricci1991) suggests that rapid surface uplift enhanced erosion and created high relief persisting in the vicinity of the present-day Torino Hills.
Thus we propose the following scenario to explain the mineralogical evolution recorded on both flanks of the young Alpine orogenic system during Oligocene and early Miocene times. The transition from subduction to collision by early Oligocene times induced the rapid and prominent aerial exposure of the eastern part of the HP–LT accretionary wedge (Fig. 8a) and predominantly fed the external molasse basins. From early Oligocene to early Miocene times, the relief shifted westwards due to the uplift of the frontal part of the collisional wedge (Briançonnais zone) in response to the activation of the Briançonnais thrust front (Fig. 8b). This renewed topography and dynamics enhanced erosion of the external part of the accretionary wedge (LT-blueschist). In the eastern part, the vertical indentation at the rear of the accretionary wedge related to the emplacement of the Ivrea body (Schmid & Kissling, Reference Schmid and Kissling2000) allowed the progressive exhumation of eclogites, and focused erosion fed the internal molasse basins in the Po plain (Torino Hills and Tertiary Piedmont basin in Fig. 1b). At that time (late Oligocene to early Miocene), the western Alps progressively acquired its doubly vergent structure (Fig. 8b).
7. Conclusions
Sands and conglomerates derived from blueschist-facies metamorphic rocks have some characteristic minerals such as glaucophane that are poorly resistant to transport or diagenesis in a sedimentary basin. Serpentinite, hitherto little studied, may be a more reliable marker of HP–LT source rocks in provenance studies. In order to quickly identify serpentine species in sedimentary rocks we propose to use Raman spectroscopy on single detrital clasts. We successfully applied this method to detrital serpentinite from the Barrême basin, in the western Alpine foreland. Mineralogical analysis by Raman spectroscopy shows that serpentinite pebbles from the lower Oligocene conglomerates in the Barrême basin consist entirely of antigorite, indicative of HT-blueschist-facies metamorphism.
During early Oligocene times, palaeocurrent indicators indicate a shift from a southern provenance (Provencal shelf and the Corso-Sardinian massif) to a northeastern provenance of detrital grains in the Barrême molasse basin. Thus the source of the lower Oligocene pebbles has to be searched for in the internal part of the Alps, SE of the Pelvoux massif. The arrival of such exotic clasts reveals the erosion of a suddenly raised mountain range in the internal Alpine zones along the same transect, the Piedmont cordillera, as a consequence of major shortening pulses. The coexistence of metamorphic pebbles (abundant antigorite clasts but also rare blueschist gabbros) with non-metamorphic radiolarite clasts constrains the lithologic composition and structure of the sediment source area. The presence of metamorphic ophiolitic material contradicts the classical view that the only source of exotic clasts in the lower Oligocene Barrême molasse basin is located in the Embrunais-Ubaye flysch nappes that glided in front of the Alpine wedge.
On the contrary we propose that the supply of exotic clasts to the Barrême basin is a direct consequence of the major syncollisional shortening experienced by the Palaeogene accretionary wedge in early Oligocene times, by 34–28 Ma, to the SE of the Pelvoux massif. The remnants of this refolded subduction wedge correspond to the present-day deeply eroded internal zone. The absence of eclogite clasts on both flanks of the early Oligocene cordillera means that eclogitic units were not exhumed during early Oligocene times along the Monviso transect, contrary to what is known along the more southerly Liguria transect.
The same mineralogical analysis by Raman spectroscopy was also conducted on serpentinite sand grains from the upper Oligocene/lower Miocene Grès verts of the Barrême basin. The analysed clasts consist of antigorite and dominant lizardite, to some part depending on grain size, indicative of a change through time in source rock lithology. This observation is compatible with a westward propagation of the cordillera with erosion of lesser metamorphic units during ongoing collision. Simultaneously, the increasing arrival of HP ophiolitic materials in the Po plain suggest the persistence of high relief in the internal zone, maintained by deep vertical indentation of the Ivrea body.
Acknowledgements
We gratefully acknowledge M. Malusà and G. Vignaroli for their careful suggestions and comments. This study was supported by the Agence Nationale de la Recherche project no. ANR-08-BLAN-0303-01 “Erosion and Relief Development in the Western Alps”.
