Tectono-sedimentary evolution of the Tertiary Piedmont Basin (NW Italy) within the Oligo–Miocene central Mediterranean geodynamics


Corresponding author: M. Maino, Dipartimento di Scienze della Terra e dell'Ambiente, Università degli Studi di Pavia, via Ferrata, 1 - 27100 Pavia, Italy. (matteo.maino@unipv.it)


[1] We analyze the tectono-sedimentary and thermochronometric constraints of the Tertiary Piedmont Basin (TPB) and its adjoining orogen, the Ligurian Alps, providing new insights on the basin evolution in response to a changing geodynamic setting. The geometry of the post-metamorphic faults of the Ligurian belt as well as the fault network that controlled the Oligo–Miocene TPB deposition has been characterized through a detailed structural analysis. Three main faulting stages have been distinguished and dated thanks to the relationships among faults and basin stratigraphy and thermochronometric data. The first stage (F1, Rupelian–Early Chattian) is related to the development of extensional NNW-directed faults, which controlled the exhumation of the orogen and the deposition of nearshore clastics. During the Late Chattian, the basin drowning is marked by mudstones and turbidites, which deposition was influenced by the second faulting stage (F2). This phase was mainly characterized by NE- to ENE-striking faults developed within a transtensional zone. Since the Miocene, the whole area was dominated by transpressive tectonics. The sedimentation was represented by a condensed succession followed by a very thick, turbiditic complex. At the regional scale, this succession of events reflects the major geodynamic reorganization in the central Mediterranean region during the Oligo–Miocene times, induced by the late-collisional processes of the Alps, by the eastward migration of the Apennines subduction and by the opening of extensional basins (i. e., the Liguro–Provençal Ocean).

1 Introduction

[2] Within a foreland system, basins may form ahead of the active pro- or retro-front of the fold-and-thrust belts [Allen and Allen, 1990; DeCelles and Giles, 1996]. These basins are closely controlled by the tectonic processes that result from the interplay of orogen uplift, flexurally-triggered basin subsidence, and eustacy [e. g., Catuneanu, 2004; Cederbom et al., 2004; Allen, 2008; Cloetingh and Negendank, 2009]. Understanding the evolution of these basins may be difficult, especially when an evolving geodynamic setting promotes multiple stages of deformation, severely modifying the primary structural features of the orogen-basin system. However, the integration of chronological constraints (both biostratigraphic and thermochronometric) with detailed stratigraphic and structural analyses helps to determine the timing and the rates of deformation stages describing the tectonic evolution of the orogen-basin system.

[3] This paper is concerned with the Tertiary Piedmont Basin (TPB; Figure 1), which is an enigmatic basin that evolved in response to a changing tectonic setting. According to the present literature [e. g., Gelati and Gnaccolini, 1998; Schmid and Kissling, 2000], it is generally regarded as a wedge-top-basin located on top of the junction between the two main orogenic systems of the central Mediterranean: the Western Alps, characterized by a main westward tectonic nappe-stacking [e. g., Butler et al., 1986; Platt et al., 1989a; Ford et al., 2006; Dumont et al., 2011; Kissling et al., 2012], and the east/northeastward verging Northern Apennines [e. g., Patacca et al., 1990; Jolivet and Faccenna, 2000]. The TPB sedimentation occurred in the Oligo–Miocene during three main tectonic episodes, the exhumation of the Ligurian sector of the Western Alps, the opening of the Liguro-Provençal basin, and the formation of the Apennines thrust belt.

Figure 1.

Simplified present-day tectonic map of the Central-Western Mediterranean region (from Faccenna et al. [2004] and references within), showing the location of the study area. LN: Ligurian knot; Pan. basin: Pannonian Basin; TPB: Tertiary Piedmont Basin,

[4] The complex evolution of the TPB is well documented, although conclusions on its evolution are often contradictory as its origin and development is explained in terms of either compressional or extensional tectonics [e. g., Lorenz, 1969; 1984; Hoogerduijn Strating et al., 1991, Laubscher et al., 1992; Di Giulio and Galbiati, 1995; Mutti et al., 1995; Gelati and Gnaccolini, 1998; Felletti, 2002; Carrapa et al., 2003a; Gelati and Gnaccolini, 2003; Carrapa and Garcia Castellanos, 2005; Maffione et al., 2008; Naylor and Sinclair, 2008; Vignaroli et al., 2008, 2009, 2010]. A huge number of multidisciplinary studies have been performed in the TPB during the last decades, including structural and sedimentologic analyses [e. g., Lorenz, 1969, 1984; Mutti et al., 1995; D'Atri et al., 1997, 2002; Gelati and Gnaccolini, 1998; 2003; Carrapa et al., 2003a; Vignaroli et al., 2008; Mosca et al., 2010], thermochronology [Barbieri et al., 2003; Carrapa et al., 2003b, 2004; Bertotti et al., 2006], paleomagnetism [Carrapa et al., 2003a; Maffione et al., 2008], and numerical and statistical modeling [Felletti, 2004a, 2004b; Carrapa and Garcia Castellanos, 2005; Bertotti and Mosca, 2008]. Despite this abundant information, chronological constraints of the tectonic stages experienced by the basin, in particular of the early phases, need to be strengthened. These uncertainties prevent an accurate determination of the mechanisms that drove to the formation of the basin and hamper the formulation of appropriate models describing the tectonic evolution of the region. This work focuses on the SW part of the TPB (Figure 2), along its boundary with the orogen, where the tectonic features associated with the early stages of the basin formation are still preserved. Extensive sedimentary and structural investigations have been combined with the structural analysis of the Ligurian orogen post-metamorphic fault-pattern and compared with neighboring areas data. Geometric relationships between sedimentary rocks and tectonic structures as well as thermochronologic data are used to constrain the timing of the deformation phases.

Figure 2.

Geological map of the Ligurian Alps and Tertiary Piedmont Basin with the distribution of the main late Alpine faults. Ba: Bagnasco basin; Gtz: Grognardo thrust zone; PF: Pietra di Finale; PFT: Penninic frontal thrust; Sa: Sassello basin; SL: Stura line; SF: Scrivia fault; SV: Sestri-Voltaggio fault; VVL: Villalvernia-Varzi line. The position of the analyzed structural stations is shown: the exact location of the stations is given in Table 2.

[5] We present here an improved basin analysis based on a precise reconstruction of the tectonic phases experienced by the TPB and the adjoining Ligurian Alps. The depositional and structural features reflect the deep processes controlling the linked evolution of the Alps and Apennines orogenic systems. The new data lead to a critical evaluation of the various geodynamic models adopted for this region.

2 Geological Background

2.1 Regional Framework

[6] The central/western Mediterranean area features a geodynamic complexity caused by the relative movements of three main plates (Africa, Adria, and Europe) and several interposed oceanic basins, which led to the formation of two arcuate orogenic systems (i. e., the Alps-Dinarides and the Apennines-Maghrebides belts; Figure 1).

[7] Since the Cretaceous, the Piedmont-Ligurian Ocean was consumed in a south-east to south directed subduction zone [e. g., Stampfli and Marchant, 1997]. After the Eocene–Oligocene collision, the indentation of the Adriatic and the European plates leads to the building of the Alpine chain in the central Mediterranean [e. g., Platt et al., 1989a; Ford et al., 2006]. Starting from the Oligocene, several extensional basins (i. e., the Algerian, Ligurian-Provençal Alboran, and Tyrrhenian basins; Figure 1) originated above a previously thickened continental crust [De Voogd et al., 1991; Jolivet and Faccenna, 2000; Roca, 2001]. These basins obliquely cross-cuts the contractional structures of the Betics, Iberian Chain, Pyrenees, and Alps, suggesting an independent origin with respect to these orogens [Carminati et al., 1998]. There is a general consensus to consider these basins as the results of back-arc extension related to the W-dipping Apennines-Maghrebide subduction [e. g., Boccaletti et al., 1980; Castellarin, 2001]. Nevertheless, the origin of this subduction is a hardly-debated matter, as it implies the underthrusting of a composite crust (the Adria and Africa continents and the interposed Neotethys Ocean, actually represented by the Ionian basin; Figure 1) below the European plate. It was interpreted either as a subduction with a single permanent polarity since Late Cretaceous time [e. g., Jolivet et al., 1998; Faccenna et al., 2004] or as the result of an Eocene flip of an earlier Cretaceous SE-directed Alpine subduction [e. g., Elter and Pertusati, 1973; Michard et al., 2002; Handy et al., 2010]. In any case, since the Oligocene onward, the retreat of the slab resulted in a progressive trench rotation from a NE-SW direction (i. e., parallel to the European passive margin) to an N-S one. This geodynamic process was contemporaneous with the eastward Apennines nappe stacking and migration [Castellarin, 2001]. During the Early–Middle Miocene, the progressive eastward propagation of the Apennines arc, associated with the oceanic spreading of the Liguro-Provençal and Algerian basins, led to the anticlockwise drifting of the Corsica-Sardinia block [e. g., Alvarez et al., 1974; Rosenbaum et al., 2002]. From the Late Miocene onward, extension shifted eastward to the Tyrrhenian Sea (Figure 1), leading to the present-day configuration of the Apennines chain [Malinverno and Ryan, 1986; Ciarcia et al., 2012].

[8] The Alps-Apennines evolution is characterized by the onset of several foreland basins along both the pro- and the retro-side of the orogenic belts [e. g., Catanzariti et al., 1996; Sinclair, 1997; Schlunegger, 1999; Elter et al., 1999; Di Giulio et al., 2001; Cibin et al., 2003; Ford and Lickorish, 2004; Naylor and Sinclair, 2008]. The Tertiary Piedmont Basin is located above the connection of the two orogenic systems (the “Ligurian knot” of Laubscher et al. [1992]; Figure 1), thus recording the interplay between the Alps-Apennines belts. The Upper Eocene–Miocene sedimentary succession of the TPB is mainly formed by clastic deposits resting on the Pre-Cenozoic substratum of the Ligurian Alps in the west and the Ligurian units of the Northern Apennines to the east (Figure 2). The TPB is divided in four main paleogeographic domains: from west to east, the Monregalese High, the Langhe basin, the Alto Monferrato High and the Borbera-Curone basin (Figure 3). The paper focuses on the Langhe basin, which represents the south-western sector of the TPB.

Figure 3.

(a) Study area (square) within the Langhe Basin. (b) General overview of stratigraphy of the Oligo–Miocene basin fill. A, B1–B6, and C1–C3 indicate the Oligo–Miocene depositional sequences defined by Gelati and Gnaccolini [1998]. Abbreviations: Cassinasco Fm. (csi), Cortemilia Fm. (com), siliceous lithozone (sl), Noceto turbidite systems (nts), Castelnuovo Bric la Croce turbidite system (bcts), Cengio and Retano turbidite system (cts), Millesimo body (ml), muddy framework (mud), Molare Fm (mor), pre-Cenozoic bedrock (s).

2.2 Structure of the TPB-Ligurian Alps System

[9] The Ligurian Alps (Figure 2) are constituted by imbricated tectonic units belonging to different palaeogeographic domains [Vanossi et al., 1986; Seno et al., 2005a, 2005b]: the Briançonnais and Prepiedmont (belonging to the European continental margin) and the Piedmont-Ligurian. The Briançonnais and Prepiedmont units show a stratigraphic sequence composed of a Paleozoic basement [Maino et al., 2012b] comparable with other basement rocks belonging to the “souther Variscan realm” [Rossi et al., 2009; Casini et al., 2012], a Permian volcano-sedimentary succession [Dallagiovanna et al., 2009] and a Meso–Cenozoic cover with huge sedimentary gaps [Decarlis and Lualdi, 2008, 2009, 2011]. The Piedmont-Ligurian domain includes metaophiolite and metasedimentary units, which are widely exposed in the eastern part of the Ligurian belt (i. e., the Voltri massif [Capponi and Crispini, 2002]; Figure 2). A detached part of the oceanic cover, the turbiditic successions of the Helminthoid Flysch, has been shifted far from its original basement, and now rests at the south-westernmost sector of the chain [Di Giulio and Galbiati, 1991; Ford et al., 1999].

[10] To the east, the sub-vertical Sestri-Voltaggio Fault separated the high pressure (HP) metamorphic units of the Voltri Massif from three blueschist to pumpellyite-actinolite facies tectonic units (i. e., the Sestri-Voltaggio Zone, Figure 2), which are in turn tectonically overlaid by the low- to non-metamorphic sedimentary successions of the Apennines chain [Cortesogno and Haccard, 1984]. These are here represented by the Albian flysch sequences and the Late Cretaceous–Eocene marls of the “Ligurian units”.

[11] During the Alpine orogenic events, the rocks of the Ligurian Alps experienced a polyphase deformational evolution linked to subduction-collisional events [Vanossi et al., 1986; Seno et al., 2005a]. The main deformational phases and the associated metamorphic conditions are summarized in Table 1.

Table 1. Deformation Phases in the Continental Briançonnais Domain (Brianç) and in the Oceanic Piedmont-Ligurian Units (Voltri) of the Ligurian Alpsa
Deformation phaseAreaFabricMetamorphismAgeReferences
  1. aThe late-orogenic deformational phases (D4–D5) are coeval with the deposition of the TPB sedimentary succession.
Pre-D1VoltriIsoclinal similar folds (unconstrained vergence)eclogite-blueschist-faciesEarly/Middle EoceneCapponi and Crispini [2002]; Federico et al., 2005; [2007]
D1 (D1/D2 in Capponi and Crispini [2002])VoltriTight, sub-isoclinal folds; nappe stacking (unconstrained vergence)Na-amphibole greenschist/greenschist-facies s.s.Middle/Late EoceneCapponi and Crispini [2002]; Federico et al. [2005]; Vignaroli et al. [2005]
D1BriançSW-verging sub-isoclinal folds and nappe stackingUp to blueschist faciesBartonian–PriabonianVanossi et al. [1986]; Seno et al. [2005a]; Bonini et al. [2010]
D2 (EPF in Capponi and Crispini [2002])VoltriExtensional W-verging shear bandsgreenschist-faciesPriabonian–Early OligoceneHoogerduijn Strating [1994]; Capponi and Crispini [2002]; Vignaroli et al. [2009, 2010]
D2 (RSZ in Capponi and Crispini [2002])VoltriSW verging reverse shear zonesgreenschist-facies to low greenschist-faciesPriabonian–Early OligoceneCapponi and Crispini [2002]
D2BriançN/NE-verging back-folds and back-thrustsUp to greenschist-faciesPriabonian–Early OligoceneVanossi et al. [1986]; Seno et al. [2005a]; Bonini et al. [2010]
D3VoltriChevron folds and semi-brittle extensional W-verging shearingLow greenschist-faciesEarly OligoceneCapponi and Crispini [2002]; Vignaroli et al. [2009]
D3BriançS-verging chevron or kink foldsNon-metamorphicEarly OligoceneSeno et al. [2005a]; Bonini et al. [2010]
D4VoltriNormal and strike-slip faultsNon-metamorphicEarly OligoceneVignaroli et al. [2009]
D5 (D4 in Federico et al. [2009])VoltriOpen folds, thrust and strike-slip faultsNon-metamorphicEarly MioceneCapponi and Crispini [2002]; Crispini et al. [2009]; Federico et al. [2009]
D4/D5 (D4 in Bonini et al. [2010])BriançOpen folds, normal and strike-slip faultsNon-metamorphicOligocene–MioceneVanossi et al. [1994]; Seno et al. [2005a]; Bonini et al. [2010]

[12] In response to the Late Cretaceous–Eocene subduction of the Piedmont-Ligurian Ocean under the Adria plate, the rocks of the Piedmont-Ligurian and Briançonnais domains experienced a highly variable metamorphic re-crystallization (from greenschists to eclogite facies), depending on the position within either the subduction channel or the orogenic wedge [e. g., Ernst, 1981; Messiga and Scambelluri, 1991; Goffé et al., 2004]. During the Middle/Late Eocene, these units were displaced and thrusted toward the foreland (SW; D1 phase) forming a nappe pile [Seno et al., 2005a], constituted by, from bottom to top, the Briançonnais, the Prepiedmont, and the Piedmont-Ligurian units, all stacked onto the Dauphinois domain [Dallagiovanna et al., 1997; Seno et al., 2005b; Bonini et al., 2010]. The following (Late Eocene–Early Oligocene) retrogressive metamorphism recorded by the Piedmont-Ligurian and Briançonnais units was achieved during the last ductile deformational phases (D2–D3; Table 1) leading the exhumation of the deep-seated rocks presently cropping out in the eastern part of the belt, i. e., the Voltri Massif. The structural interpretations of these phases are debated, as they have been related to polyphase contractional regime [e. g., Vanossi et al., 1986; Capponi and Crispini, 2002; Seno et al., 2005a] or extensional tectonics [Hoogerduijn Strating, 1994; Vignaroli et al., 2009].

[13] During the Oligocene, the unmetamorphosed Helminthoid Flysch (Figure 2) was emplaced by a gravitational gliding onto the Briançonnais/Dauphinois boundary [Kerckhove, 1969, Ford et al., 1999; Corsini et al., 2004; Seno et al., 2005a, 2005b). The final emplacement of the Helminthoid Flysch followed a NW- to W-ward Late Cretaceous–Eocene subaqueous translation of about 200 km due to its involvement in an intra-oceanic accretionary prism [Di Giulio and Galbiati, 1991; Seno et al., 2005a]. The Oligocene gravity-driven detachment of the Helminthoid Flysch resulted in passive folding, with folds truncated at the base, due to shear-flow perturbations over basal irregularities [Merle and Brun, 1984; Merizzi and Seno, 1991].

[14] Since the Oligocene, in the metamorphic basement, deformation in frictional conditions brought to the development of the D4–D5 systems of cataclastic-gouge fault zones [Table 1; Vanossi et al., 1994; Crispini et al., 2009; Federico et al., 2009; Vignaroli et al., 2009; Bonini et al., 2010]. The stress/strain regime and the timing of these phases are poorly resolved as they resulted from the overlap of several tectonic events, including the Early–Middle Miocene oceanic spreading of the Liguro-Provençal basin, the coeval 50° of vertical rotation [Maffione et al., 2008] of the Ligurian Alps-TPB system [Vanossi et al., 1994], and the drifting of the Corsica-Sardinia block [Gattacceca et al., 2007].

[15] The late-orogenic evolution (D4–D5) of the Ligurian Alps is coeval with the deposition of the TPB sedimentary succession. The TPB is characterized by the lack of major tectonic structures, whereas small scale structural features, such as folds and faults ranging from hundreds of meters to centimeters are common. Several authors documented synsedimentary folds and thrust zones in the Oligocene–Serravallian sediments [Perotti, 1985; Fossati et al., 1988; Gelati and Gnaccolini, 1998; Bernini and Zecca, 1990; D'Atri et al., 1997, 2002; Marroni et al., 2002; Carrapa et al., 2003a; Piana et al., 2006], and decametric to centimetric-scale normal and strike-slip fault systems have been detected on the whole of the basin succession [e. g., Mutti et al., 1995; Gelati and Gnaccolini, 1998; Felletti, 2002; Carrapa et al., 2003a; Vignaroli et al., 2009].

2.3 Stratigraphy of the Langhe Basin

[16] The Langhe Basin overlies a series of Alpine nappes that were already emplaced in the Late Eocene. The basin extends over an area of about 1800 km2 and hosts an Oligo–Miocene succession more than 4000 m thick [Gelati and Gnaccolini, 1998, 2003; Forcella et al., 1999]. The formation of the basin in the Early Oligocene is recorded by alluvial fan and fan-delta deposits, followed by transgressive shallow-marine sediments (Molare Formation, depositional sequence A; Figure 3b, [Gelati and Gnaccolini, 2003]). During this time, some topographical “highs” were directly covered by small coral reefs [Franceschetti, 1967; Gnaccolini, 1978; Fravega et al., 1987].

[17] Drowning of the basin in the Late Oligocene delineated a complex shape of the south-western margin and originated an array of confined depocenters at different elevations with respect to the basin floor. The drowning is marked by the sudden change in the deposition from the shallow marine sediments of the Molare Fm. to the hemipelagic mudstones of the “Muddy Framework” [Gelati and Gnaccolini, 2003], with interbedded lenticular bodies of sandstone and conglomerate (Figure 3b). On the whole, these sediments form the Rocchetta-Monesiglio Formation (Late Rupelian–Burdigalian) and are organized into six depositional sequences (B1–B6; Figure 3b, [Gelati and Gnaccolini, 2003]). The lenticular sandstone bodies belong to different turbidite-systems that fill local depocenters close to the western and eastern margins (in present-day coordinates) of the Langhe basin.

[18] In the Late Burdigalian–Earliest Langhian, the basin acquired a more homogeneous, SE–NW elongated shape (Sequence C1, Cortemilia Formation; [Gelati and Gnaccolini, 2003]). The Middle Miocene of the Langhe basin is represented by five depositional sequences (C2–C6; Figure 3b). They are mostly formed by high-density flow sandstones of the Cassinasco Formation and, toward the top, by hemypelagic pelites of the Murazzano Formation.

2.4 Thermochronometric Data

[19] In the last 10 years, many thermochronometric data both from the metamorphic rocks of the Ligurian Alps and from clasts in the TPB sediments have been published [Figure 4a; Barbieri et al., 2003; Carrapa et al., 2003b; Carrapa et al., 2004; Federico et al., 2005; Bertotti et al., 2006; Vignaroli et al., 2010; Maino et al., 2012a].

Figure 4.

(a) Tectonic map of the Ligurian Alps and Tertiary Piedmont Basin with a compilation of published thermochronometric data. (b) Published 40Ar/39Ar, zircon/apatite fission tracks and (U-Th)/He ages plotted against distance—d—from the boundary between the Ligurian Alps and the TPB; ZHe (green; [Maino, 2012a]), AFT (light blue ages from the Molare Formation samples; dark blue from the basement [Barbieri et al., 2003]), ZFT (red; [Vance, 1999]) AHe (light gray from the TPB succession, dark gray from the basement; [Bertotti et al., 2006]) and 40Ar/39Ar detrital ages from the Molare Fm. (light gray square; [Barbieri et al., 2003; Carrapa et al., 2004]). Note that the ZHe, ZFT, AFT, and the detrital 40Ar/39Ar ages from samples collected close to the stratigraphic boundary between TPB and metamorphic basement are identical (34–26 Ma) with the Early Oligocene biostratigraphic age of the Molare Fm. [Gelati and Gnaccolini, 1998; Seno et al., 2010]. ZHe ages young away from the basin boundary (to the south).

[20] White mica 40Ar/39Ar indicates that the eclogite-blueschist-facies (HP) metamorphism occurred between 49 and 40 Ma in the Voltri Massif [Federico et al., 2005, 2007]. Phengite 40Ar/39Ar (32.9 ± 0.8 Ma [Federico et al., 2005]) and U-Pb sensitive high-resolution ion microprobe (SHRIMP) dating on a zircon rim (33.8 ± 0.8 Ma [Vignaroli et al., 2010]) dated the greenschist-facies metamorphic stage within the Voltri metasediments. Considering the biostratigraphic age (Early Rupelian, ~33.9–30 Ma; [Gelati and Gnaccolini, 1998]) of the TPB sediments overlying the sampling sites, fast (3–5 mm/yr) exhumation rates have been proposed during the Oligocene time [Federico et al., 2005]. Remarkably higher exhumation rates (in the order of several cm/yr) have been proposed by Rubatto and Scambelluri [2003] based on young radiometric ages for the baric peak of HP metamorphism (33.6 ± 1 Ma from U/Pb dating on baddeleyite from eclogites).

[21] The data from the metamorphic basement correspond with the 40Ar/39Ar ages (~32–50 Ma) from detrital micas collected in the basal deposits of the TPB, which derived from the erosion of the metamorphic units of the Ligurian Alps [Barbieri et al., 2003, Carrapa et al., 2004]. The presence of a ~34–32 Ma 40Ar/39Ar ages from the ophiolitic pebbles of the Rupelian (~33.9–28.4 Ma) Molare Formation (stratigraphically overlying the Voltri Massif) supports the concept of a fast exhuming basement. 40Ar/39Ar ages of sediments from the present rivers draining the Ligurian and Western Alps also indicate that the metamorphic rocks were exhumed in rapid pulses prior to ca. 38 Ma and that relatively slow and continuous erosion occurred thereafter [Carrapa et al., 2003b].

[22] Zircon fission track (ZFT) [Vance, 1999] and zircon (U-Th)/He ages (ZHe) [Maino et al., 2012a] documented the time-temperature history of the Briançonnais and Prepiedmont domains of the Ligurian Alps. ZFT data from these domains cluster around 31 Ma, but some grains from low metamorphic units are much older (~60–254 Ma, Figure 4a; [Vance, 1999]). The old ages clearly indicate that the ZFT system was not reset everywhere during the Alpine phases; hence, not all the units of the Ligurian Alps experienced, during Alpine metamorphism, temperatures hotter than ~240 ± 25°C (approximate temperature of closure for the ZFT system is derived from natural samples [e. g., Brandon et al., 1998; Bernet, 2009]).

[23] The ZHe thermochronometer records the cooling of rocks on a range of temperatures that, depending on cooling rates and crystal size, varies from 210°C to 140°C [e. g., Reiners, 2005]. ZHe ranged from 32.2 ± 2.3 to 25.4 ± 1.6 Ma (Figure 4a). The ZHe data from the present-day exposed stratigraphic boundary with the TPB (figure 4) indicate that the metamorphic rocks were at ~200°C between 32.2 ± 2.3 and 29.1 ± 1.1 Ma. The ZFT and ZHe ages are within the error of the biostratigraphic age (Early Rupelian, ~33.9–28.4 Ma) of the Molare Formation. These data imply a Late Rupelian average cooling rate of more than 100°C/Myr and an apparent (ultra-) fast exhumation (6.8–1.3 mm/yr) through the shallow crust [Maino et al., 2012a].

[24] Apatite fission track (AFT) analyses from two samples of Briançonnais basement and one of the overlying Molare Fm. (Figure 4a) indicate a rapid cooling between 120° and 60°C at ~26 Ma, which is an age slightly younger than the depositional age of sediments [Barbieri et al., 2003]. These data seem to suggest a resetting of the AFT system by post depositional thermal overprinting as a result of burial following subsidence in the TPB [see Bertotti et al., 2006]. Otherwise, other thermal indicators, such as vitrinite reflectance and thermal alteration index on palynomorphs, indicate paleo-temperatures lower than 100°C for the Molare sediments discounting the possibility of a completely reset of the AFT system [Barbieri et al., 2003]. Moreover, in the Voltri massif, a limited number of AFT data yielded dispersed ages comprised between 2.5 ± 0.6 and 23.9 ± 4.9 Ma [Vignaroli et al., 2010]. On the whole, the AFT data set is insufficient to derive a valid interpretation of the Miocene thermal history of the TPB. They may suggest an inhomogeneous and partial resetting of the AFT system by post-depositional thermal overprinting as a result of Oligocene–Late Miocene burial due to TPB subsidence and final Pliocene–Quaternary uplift.

[25] Bertotti et al. [2006] carried out apatite (U-Th)/He (AHe) analyses both in the basement and through all the TPB succession. Most of the data indicate a general cooling of the basin succession between 70° and 40° C at ~14–11 Ma, but some samples from the stratigraphic boundary between the basement and the Molare Fm. yielded contrasting results (~23 and ~14 Ma, respectively) despite their proximity (Figure 4a). The AHe data cannot be explained unequivocally: only the post-14 Ma inversion from subsidence to basin exhumation can be reliably interpreted.

3 Methods

3.1 Stratigraphic Analysis

[26] In the frame of the National Project for survey, the 1:50000 Italian Geological Map, a preliminary geological survey of the study area at the scale of 1:10000 [Seno et al., 2010] provided new data on the Tertiary turbidite systems organization (Figures 5 and 6). Their stratigraphic framework has been refinished on the base of km-scale correlations of their component bed-sets and on the contour map of their pinch-outs [Bersezio et al., 2005, 2009; Felletti and Bersezio, 2010a, 2010b]. In this way, the basin geometry was defined and the depocentral areas were distinguished from marginal settings. Several stratigraphic sections were physically correlated [Bersezio et al., 2005, 2009; Felletti and Bersezio, 2010a, 2010b] and measured bed-by-bed.

Figure 5.

Geological-structural sketch of the mapped area at the orogen-basin boundary, with the distribution of the late Alpine faults. RV: Roccavignale. Field data are collected in the framework of the CARG project (Cairo Montenotte sheet of the geological map of Italy at scale 1.50000; [Seno et al., 2010]). The complete geological map (with more details) is consultable online at http://www.isprambiente.gov.it/Media/carg/liguria.html.

Figure 6.

Geological-structural map of the Mioglia area (see Figure 2) with the distribution of the late Alpine faults (modified from [Bernini and Zecca, 1990]). Legend as in Figure 5.

[27] Micropaleontological samples have been collected within key horizons traced through the basin. The calibration of the bio-magnetostratigraphic events is according to Gradstein et al. [2004].

3.2 Structural Analysis

[28] Until now, a detailed structural analyses of the TPB or the Ligurian Alps has only been performed at a local scale [e. g., Bernini and Zecca, 1990; Carrapa et al., 2003a; Piana et al., 2006] and no attempt has been made to quantify and classify the brittle fault network characterizing the whole orogen-basin system. We have performed a field-based detailed structural analysis of the brittle structures from the southern part of the Langhe basin and the western sector of the Ligurian Alps (Figures 2, 5, and 6). These data, integrated with the sedimentological investigation carried out at the present-day exposed erosional orogen-basin boundary, provide constraints on the overall evolution of the orogen-basin system. Our approach is based on the systematic measurement of fault populations, including fault plane orientation, striae and slickenside orientation, and shear sense, classified on the basis of the reliability criterion of the shear sense. Structural analyses in the basement were performed in 17 well-exposed sites (structural stations, hereafter SS; Table 2).

Table 2. Location of the 17 Structural Stations in the Ligurian Alps
N° of Structural StationLithologyLongitudeLatitude

[29] The structural data have been computed in order to obtain paleostrain analysis [Malusà et al., 2009]. Details on the method applied on both the field-based analysis and the data computing are provided in sections A1 and A2, respectively.

3.2.1 Dating the Brittle Structures

[30] The chronology of the faulting phases is primarily derived from the TPB area, where the activity of the fault-families is deduced from the relationships between deformation and stratigraphy. Dating brittle faulting in the metamorphic rocks is more difficult but can be extrapolated by integrating the rock-fault type analysis with the thermochronometric data [Malusà et al., 2009], exploiting the fact that the thermal structure of the crust exerts a primary control on the fault-rocks types which are produced during deformation [e. g., Sibson, 1977; Scholz, 2002]. The change from plastic to cataclastic deformation mechanisms is temperature dependent; in silicate rocks, it is referred to a large temperature range between ~450°C and ~120°C, depending on strain, mineralogy, structural inheritance, and fluid flows [e.g., Kohlstedt et al., 1995; Scholz, 2002]. The very complex tectonic processes, the lithological variations, and the extremely various strain rates active during the multiphase orogenic evolution of the Alps do not allow to extrapolate from a single fault the age for the occurrence of brittle (elastico-frictional; [Sibson, 1977]) deformation in all the belt. Thermochronometric dating on multi-locations, however, provides the timing when the study region cooled through a given temperature. In particular, the zircon fission track and zircon (U-Th)/He techniques with their closure temperature (Tc) of ~240 ± 25°C [Brandon et al., 1998; Bernet, 2009] and 140–210°C [Reiners, 2005] have the potential to date the post-metamorphic tectonic deformations characterized by purely cataclastic faulting. Therefore, the ZFT and the ZHe ages from the metamorphic units of the Ligurian Alps, coupled with 40Ar/39Ar, AHe, and AFT data, have been used to constrain the activity of the faulting phases in the basement.

4 Sedimentology and Stratigraphy

[31] Geometry, stratigraphic relationships, facies distribution, and palaeocurrent directions indicate that the sandstone bodies of the Molare Formation (depositional sequence A) and of the Rocchetta-Monesiglio Formation (depositional sequences B1–B6) were deposited in a basin controlled by very active tectonics [Gelati and Gnaccolini, 2003]. This tectonic activity is documented by the regional stratigraphic architecture that record migrations of the source areas and depocenters with time, and by the complex lateral and vertical distribution of terrigenous depositional systems mainly controlled by structurally-induced submarine topography. These aspects are described in the following paragraphs.

4.1 Rupelian (~34–28 Ma)

[32] At this time, the basement was being eroded and part of it was covered by clastic sediments of the Molare Formation (Sequence A).

[33] The unconformable and transgressive Molare Fm. consists of flood-dominated alluvial fan and fan-delta systems mainly composed of conglomerate and pebbly-sandstone facies. Conglomerate clasts range from centimeter to meter in size, and their lithology reflects the composition of the neighboring-outcropping pre-Cenozoic substratum (Figure 7a). Sandstone is mainly composed of abundant quartz, quartz-mica bearing lithic fragments, and carbonate grains. QFL + CE (Q = quartz, F = feldspars, L + CE = total fine-grained lithic fragments plus carbonate rock fragments) analysis [Gelati and Gnaccolini, 2003] on sandstone sampled near Millesimo revealed a high content on lithic fragments (L ≈ 70–90%) mainly derived from quartzite, migmatite, orthogneiss, micaschist, phyllades, and mafics. Carbonate fragments are also very frequent (30–40% on total QFL + CE). Thus again, the very heterogeneous pre-Cenozoic substratum is itself the main source for the Molare Fm. Most of these sediments were deposited through a debris flow mechanism; they show substantial lateral variation in both thickness and facies, reflecting a deposition on a surface of pronounced relief. The flood-dominated alluvial fan and fan-delta deposits are gradationally replaced by highly burrowed shallow marine fossiliferous sandstones and pebbly sandstones. Because of the intense faulting activity during sedimentation, thickness, composition, and paleocurrent directions are laterally highly variable (Figure 3b). The Molare Formation has been generally ascribed to the Lower Oligocene (Rupelian, e. g., Gelati and Gnaccolini [1998]). Nevertheless, in the study area, the passage with the overlying Rocchetta-Monesiglio Formation is diachronous: the pelites overlying the Molare sandstones contains planktonic foraminifera belonging to the Paragloborotalia opima opima biozone (probably to the upper part: IFP21b) toward the east and to the Globigerina ciperoensis (IFP22) in the western sector [Seno et al., 2010].

Figure 7.

(a) Conglomerate clasts of the Millesimo body (Early Chattian) ranging from centimeter to meter in size; their lithology reflects the composition of the neighboring-outcropping pre-Cenozoic substratum: orthogneiss (Og), migmatite (M), carbonate (Ca), serpentinite (Se), quartzite (Qz). (b–d) On-lap terminations (red arrows) of Cengio (cg) and Castelnuovo-Bric la Croce (cb) turbidites (ascribed to the Middle–Late Chattian) onto lateral and frontal confining slopes (he: hemipelagic marlstone); these turbidite systems were developed during stages of relative quiescence of intrabasinal tectonics and were accommodated into a progressively westward-widening depositional area under the confining effect of marginal slopes.

4.2 Early Chattian (~28–26 Ma)

[34] At the beginning of the Late Oligocene, the emplacement of widespread purely marine sediments testifies a generalized deepening of the basin, associated with a change in the provenance area and in the denudation rates. In the lower part of the Rocchetta-Monesiglio Formation (depositional sequence B1), the submarine topography is controlled by NW-oriented, high-angle, basement-involving normal faults and flexures. The sedimentation was strongly confined in small, isolated fault-bounded basins, which were infilled by relatively small and lenticular submarine fan systems. The compositional analysis on sandstone of the B1 sequence sedimentary bodies shows different results for the northeastern sector of the basin and the southwestern one [Gelati and Gnaccolini, 2003]. While toward the north, a sudden enrichment in lithic fragments and especially in serpentinite clasts has been recorded (QFL + CE: 40% < L < 60%; serpentinite + metabasites fragments between 27% and 50% over the total); in the southern sector, lithics are less abundant (17% < L < 35%) and mafic clasts are rarer (serpentinite clasts between 2% and 10%). These compositional analyses accounted for different source areas in the northern sector of the basin (the N–NW Asti-Cuneo area) and in the southern one (the S–SW Brianconnais basement and cover plus the Molare Fm.).

[35] In the southeastern area, the Millesimo body (ascribed to the Early Chattian [Gelati and Gnaccolini, 1998; Seno et al., 2010]) is a typical example of this phase of deposition (Figures 5 and 8a–8c). It is mainly composed of channel-fill conglomeratic sandstones and conglomerates that are supposed to be the infill of a submarine depression surrounded by topographic highs with steep, fault-controlled slopes. The fault-bounded depositional setting, the textural coarseness, and facies association suggest that the clasts of the Millesimo body are reworked material from the contemporaneous alluvial fan and fan delta systems, which were probably developing on adjacent structural highs. These alluvial systems were fed by very young, high-gradient streams, which where draining small, high-relief and structurally fragmented areas. As a result, these alluvial and nearshore systems were small in size and mainly composed of very coarse-grained clastics.

Figure 8.

(a) Decametric channelized unit of Millesimo body (ml; Early Chattian ~28–26 Ma—Sequence B1, Rocchetta Monesiglio Fm); (b) conglomerate and sandstone facies filling a submarine valley (Millesimo body) that is cut into upper bathial mudstones (he: hemipelagic marlstone) deposited on the southern TPB slope. (c) Metric dislocations in the Millesimo unit by the F1 phase faults (Early Chattian ~28–26 Ma). (d) Synsedimentary pelitic slump-sheets along the gentle ramps generated by the open folding (Aquitanian–Early Burdigalian, ~23–18 Ma). The contractional structures detected in the study area suggest a general E–W main shortening, (e) ENE–WSW to NE–SW synsedimentary normal faulting (F2 phase) associated to the top of the sequence B4 (Middle–Late Chattian). The F2 faults systematically dislocate the previous F1 structures. These faults are sealed by the Aquitanian siliceous lithozone (sl).

[36] Because of the rapid deepening of the Molare-Rocchetta basin, the transgression toward the southwest did not develop a shelf area. Large volumes of the Molare alluvial fan and fan-delta coarse clastics underwent failures on the steep slopes bounding the highs and were resedimented in the adjacent submarine depression through gravity flow. These relatively small-volume flows were mostly loaded with coarse clastics. Initially, they start as debris flows, evolving downcurrent to highly concentrated turbidity currents. These flows deposited thick-bedded conglomeratic sandstones and conglomerates that display extensive channeling with the absence of lobe deposits. Slump facies are locally present.

4.3 Middle–Late Chattian (~26–23 Ma)

[37] The Middle Chattian was characterized by a fundamental change in the tectono-sedimentary features. The general depositional environment might have been similar to the delta-fed submarine ramp and lobe system or to the deep-water slope-type fan delta. The structural control on the submarine topography of the basin became subtler and can be detected only through detailed stratigraphical and sedimentological analysis. The depositional environment was characterized by a markedly flatter geometry than for the older units (e. g., Millesimo body), and volumes of sediments are considerably larger. The Retano, Cengio, Castelnuovo-Bric la Croce, and Noceto turbidites (ascribed to the Middle–Late Chattian [Gelati et al., 2010; Seno et al., 2010]; depositional sequences B2–B5, Figures 3b and 5; [Gelati and Gnaccolini, 1998]) developed during stages of relative quiescence of intrabasinal tectonics. The systems were fed by the same source area and were accommodated into a progressively westward-widening depositional area, under the confining effect of marginal slopes (Figure7b–7d). These fault-controlled slopes are mainly NE–SW to ENE–WSW oriented (Figure 8e).

[38] The sandstone composition still markedly differ from the north and the south sector during the deposition of the B2 and B3 sequences, being enriched in mafic clast toward the north and in extrabasinal carbonate clasts toward the south (initially with a low percentage of quartz + mica fragments). Lithic fragments (quartz + mica) occurrence rise up during the B4 sequence deposition in the south western part of the basin; this trend gradually expanded toward the north during the successive depositional sequences (B5 and B6) with a progressive decrease in serpentinite clasts. This accounts for a progressive homogenization of sediments during the later stages of deposition and for an enhanced influence of recycling processes [Gelati and Gnaccolini, 2003].

4.4 Aquitanian–Early Burdigalian (~23–18 Ma)

[39] This time interval records the development of gentle folds that create structural high and relatively low areas. These folds (with axes ~N–S directed) affected hemipelagic mudstones and sandstone bodies belonging to the upper part of the Rocchetta-Monesiglio Formation (depositional sequence B6 [Gelati and Gnaccolini, 1998]). In the Ceva area (Figure 5), a wide anticline of pelites is bounded at its top by an extensive slump scar, overlaid by siliceous sediments. Pelitic slump-sheets, derived at least in part from submarine denudation of the sediment originally present on top of the anticline, now flank the structure (Figure 8d). In this time interval, the sedimentation is represented by siliceous sediments, a condensed succession typically composed of well-lithified mudstone and crypto- and micro-crystalline quartz-rich siltstone, which rhythmically alternate with mm- to cm-thick fine to very fine sandstones, with local faint ripple laminations (depositional sequence B6). This condensed succession shows sudden thickness variations; it is 10–50 m thick on average (up to 150 m thick) and invariably separate pre-Miocene turbiditic sandstone bodies from overlying turbidite systems, mostly deposited during Burdigalian [Gelati et al., 2010; Seno et al., 2010]). The lower boundary of the siliceous zone is almost transitional. The upper boundary, which is marked by abrupt facies changes, displays clear unconformable contacts that probably correspond to the Burdigalian unconformity [Mutti et al., 1995].

4.5 Late Burdigalian–Present Day (~18–0 Ma)

[40] The post-Early Burdigalian sediments (depositional sequence B6-upper lithozone and C1–C6 [Gelati and Gnaccolini, 1998; Gelati et al., 2010]) of the Langhe basin crop out mainly north to the study area, and they were not analyzed in detail (Figures 2 and 5). The sedimentation was achieved in a highly subsiding asymmetric trough (E–W oriented) that was infilled with a very thick basin-wide succession of westerly-derived turbiditic sandstones and pelites (Cortemilia and Cassinasco formations, depositional sequences C1–C6; Figure 3b). In the western sector of the Langhe basin, the Late Burdigalian-Serravallian succession is characterized by a sand-rich wedge with a thickness in excess of 2000 m that markedly tapers toward SE. The analysis on sandstones belonging to C1–C6 sequences [Gelati and Gnaccolini, 2003] shows similar composition to those of the sedimentary bodies of the B6 sequence, being enriched in mica + quartz lithics, extrabasinal carbonate grains, and progressively depleted in serpentinite clasts. This accounts for a progressive shift of the source area toward the SW Brianconnais domain.

5 Post-Metamorphic Fault Network

5.1 TPB Fault Network

[41] In the TPB, km-length faults are rarely exposed; the most important exposed structure is the Villalvernia-Varzi line (Figures 2 and 3a), which during the Oligocene–Early Miocene acted as a sinistral strike-slip fault dividing the Borbera-Curone basin from the Apennines [e. g., Di Giulio and Galbiati, 1995; Felletti, 2002]. Minor faults are abundant and show two prominent directions: NNW–SSE and ENE–WSW (Figure 9a). These last are the most abundant within all the stratigraphic succession (Figure 9a) but, considering only the sequence A–B1, the NNW-striking faults are predominant (Figure 9b). Each mapped fault is characterized by a centimeters- to meters-wide cataclastic zone (Figure 10a), mostly constituted by fault breccias, whereas gouges are predominant in the fine-grained deposits such as marls or pelites. Faults are frequently associated with synsedimentary flexural folds, with axes parallel to the direction of the faults. Locally, this folding develops an inter-stratum extensional shear zone characterized by genetically related small-scale structures.

Figure 9.

Stereonets of the faults mapped in the Tertiary Piedmont Basin and Ligurian Alps. F1 and F2 indicate two distinct faulting events. F1 mapped within: (a) the whole of the TPB sedimentary succession considered in this work; (b) only in the depositional sequences A–B1. The faults trend preferentially NNW and ENE. In the sequence A–B1, the NNW-striking faults are predominant. (c, d) Direction of the normal F2 faults, mapped within all the basin, divided on the basis of their strike-slip component. The right-lateral F2 faults trend preferentially NE, while the left-lateral F2 faults are mainly ENE-oriented. (e) all faults mapped in the metamorphic basement; (f) only the basement faults where the kinematics has been constrained with a good degree of confidence. (g, h) Direction of the normal F2 faults from the basement divided on the basis of their strike-slip component.

Figure 10.

(a) Exposure of fault breccia composed of cm-thick angular fragments between the conglomerates of the Molare Formation and orthogneisses. (b) Decametric fault damage zone composed of breccias and pervasively fractured quartzites. The sketch shows the cross-cutting relationships between the older NNW-striking F1 faults and the younger NE to NNE F2 faults. (c) NNW fault system (F1): conjugate normal faults in dolomites. (d) Calcite slickenfibers on a fault facet showing (e) left-lateral/normal movement. (f, g) En echelon veins system associated with conjugate F1 faults cutting the S1 + S2 schistosity related to the ductile phases (D1 and D2). (h) Tectonic breccia in orthogneiss. (i, l) Thin sections from outcrop showed in Figure 10h, viewed in transmitted light. Cataclastic process controlled the grain sized reduction of quartz (Q) and feldspars (F) with different percentage of cement precipitation (quartz). (m) Exposure of the tectonic breccia with cm-thick angular fragments within orthogneisses.

[42] Geological survey of the area evidenced that the NNW-striking faults developed only in the Late Rupelian–Early Chattian depositional sequences A and B1 and are sealed by the Late Chattian deposits (Figures 5 and 6). On the contrary, the NE- or ENE-striking faults are present within all the Oligocene sedimentary succession (depositional sequence A–B5) and are sealed by the Aquitanian (B6) deposits (Figures 5, 6, and 8e). These field relationships demonstrate that the NNW-striking faults are older than the ENE-striking ones, indicating the existence of two chronologically distinct faulting events (F1 and F2).

[43] The NNW-striking faults (F1) bound horsts, grabens, and tilted blocks (Figure 8c). Bed displacement, drag folds, conjugate fault sets, and striations indicate that 92% of the NNW-striking faults are normal (84% with less than 30° between fault dip and lineation plunge), with oblique motion (56% left-lateral and 44% right-lateral; pitch comprised between 85° and 55°; Figure 11a). These faults show throws ranging from 0.5 to 40 m. Eastern to the study area, N–S to NW–SE striking normal faults affecting the Molare Fm. (Sequence A) have been also recognized along the boundary with the Voltri Massif [Vignaroli et al., 2009]

Figure 11.

Equal-area projections (lower hemisphere) showing fault-slip data for (a) the faults mapped in the TPB (Figures 5 and 6), grouped on the basis of the depositional sequence that they cut and (b) from the structural stations (SS) located in the metamorphic basement (see Figure 2). The fault planes are represented as great circles, the arrows indicate the direction of the relative movement of the hanging wall of the faults. Fi indicates reverse faults or thrusts. N indicates the number of measurements. In Figure11a, data from the Mioglia area (Figure 6) include those from Bernini and Zecca [1990]. Projection of D5 fold axes shows data from both the TPB (white squares) and the metamorphic basement (black squares). In Figure11b, the structural stations where two families of faults have been distinguished on the basis of cross-cutting relationships are divided in two diagrams (A and B). Only the structural stations, where the fault measurements have been carried out with a high quality of confidence in the inferred sense of movement, are reported. White squares indicate poles to the extensional veins.

[44] In the Roccavignale area (Figure 5), few high-angle NNW-striking faults, related to a conjugate system, show reverse activity (Fi; Figure 11a). Moving toward these faults, the bedding of the deposits increases the dip from 10–20° to 40–50°. Within the gage layers, typical fabric associations are represented by conjugate Riedel (R, R', P). These planes are pervasive and compose of duplex arrays with inverse shear sense. Overall, the geometry of the shear bands suggests that these faults developed in a transpression context, which involved oblique dextral simple shear and contraction components. However, the apparent bed displacement associated with these faults displays a normal sense of movement. This misfit suggests that these contraction structures were born as normal faults than experienced tectonic inversion. The geological survey of the area shows that also the reverse faults are truncated by ENE-striking faults (Figure 5).

[45] Similar structures are well exposed further east, in the Mioglia area (Figure 6). Here, the Rupelian–Early Chattian sediments of the Molare Formation and the lower part of the Rocchetta-Monesiglio Formation (depositional sequence A–B1) are involved into NNW-trending folds (e. g,. the Mioglia flexure) and thrusts [Bernini and Zecca, 1990]. The orientation of the horses of recognized duplex structures indicates a main NNE–SSW shortening with a rather important dextral component (Figure 11a). We interpret the righ-lateral-reverse faults as reworked inherited normal faults related to the F1 tectonic phase, as suggested by Mutti et al. [1995]. Also, these structures are dislocated by ENE-striking normal faults (Figure 6) involving the complete Oligocene succession (depositional sequence A–B5).

[46] The NE- to ENE-striking faults are present in the whole of the studied sedimentary succession, with the exception of the Aquitanian deposits (deposition sequence B6; Figures 8e). They systematically truncate the NNW-striking faults and display throws varying from centimeters to tens of meters. Accommodation folds are frequent, suggesting synsedimentary faulting. The ENE-striking faults show a prevailing extensional kinematics (85% with pitch >45°) associated with an important strike-slip component (15% of faults have pitch between 45° and 20°; 48% between 60° and 45°; 18% between 60 and 70°) (Figure 11a). On the basis of their transcurrent component, these faults may be divided in two sets: predominant (~70%) ENE–WSW left-lateral faults and subordinate (~30%) NE–SW right-lateral faults (Figures 9c and 9d). In the study area, the two groups of faults coexist, and the field relationships suggest that they were contemporaneous (Figures 5 and 6).

[47] The predominant extensional character of both the NW–SE (F1) and ENE–WSW (F2) faults is supported by the strain analysis: the compression (P) axes are nearly vertical in the whole of the study area, and the tension (T) axes are sub-horizontal or gently dipping (Figures 12a). The T axes of the F1 faults show a well-defined ENE–WSW incremental extension. The distribution of the T axes of the F2 faults points to a main NNW–SSE extension.

Figure 12.

Average kinematic solutions of the fault-slip data (Figure 11) from the (a) TPB and (b) the Ligurian Alps. P and T axes (light and dark gray circles, respectively) and global incremental strain axes (big, medium, and little white squares indicate the incremental shortening, intermediate and extension axes, respectively). Best fit domains according to the right dihedral analysis are indicated with light gray (shortening) and dark gray (extension).

[48] The strain analysis suggests that the NNW-trending contractional structures in the Roccavignale and Mioglia areas can be explained with the occurrence of a NNE–SSW shortening active before the F2 phase (Figure 12a).

5.2 Post-Metamorphic Fault Network of the Ligurian Alps

[49] Major brittle structures are mainly located on the western and eastern boundary of the Ligurian orogen, whereas they are rare in the central part of the chain; the most important of these structures are the Stura shear zone to the west and the Sestri-Voltaggio Fault to the east (Figure 2). The WNW-striking Stura shear zone comprises several related faults, as the Stura and Preit lines, and has been acting as a sinistral strike-slip fault, accumulating about 40–50 km of inferred displacement, since the Oligocene [Ricou, 1981; Giglia et al., 1996]. To the east, the Sestri-Voltaggio Fault represents a km-scale N–S discontinuity that played different roles in different stages of the tectonic evolution of the area [Crispini and Capponi, 2001; Vignaroli et al., 2009]. Generated as a mylonitic shear zone during the metamorphic deformational phases (D1–D3, Table 1), it records an Oligo–Miocene brittle dextral strike-slip kinematics [Capponi et al., 2009; Federico et al., 2009].

[50] At the outcrop scale, different assemblages of fault rocks with heterogeneous distribution and well-defined geometric relationships were recognized in the study area: (i) mylonitic rocks, (ii) cataclasites, and (iii) fault breccia and gouges. The deformational styles of the fault rocks can be correlated with the occurrence of the Alpine tectonic phases in different crustal conditions.

[51] Anastomosing mylonites are the oldest fault rocks; they are associated with deformations occurred under viscous regime [Schmid and Handy, 1991] associated with the development of blueschist- or greenschist-facies metamorphic recrystallization in the Briançonnais units (D1 and D2 phases, Table 1).

[52] Cohesive, fine-grained cataclasites are characterized by the development of quartz-chlorite-epidote-phyllosilicate slickensides in silicate rocks related to low greenschist facies metamorphic recrystallization (D3 phase). These latter fault-rocks are important indicators of the transition from plastic flow to cataclastic faulting [Rutter, 1986; Schmid and Handy, 1991]. Array of planar and sigmoidal tension gashes filled with quartz are often found close to these fault zone. Cohesive cataclasites preserved in outcrop are quite rare, as they are often overprinted by later structures.

[53] Cataclastic breccia (mainly incohesive) and gouges are found in multiple anastomosing layers of normal faults (D4 and D5 deformation phases; Table 1); they are characterized by the absence of metamorphic recrystallization and are generally associated with closely-spaced joint systems, especially within gneiss, quartzites, and dolomites (Figures 10b and 10h). Arrays of calcite- or quartz-filled veins are commonly found in limestones and schists (Figures 10f and 10g). Fault zone ranges from millimeters to tens of meters (Figures 10b, 10h and 10m). Faults rocks vary within the different lithology. Most of the fault rocks show > or >> 30% of large clasts (>2 mm) derived from grain size reduction with variable percentage of cement precipitation (Figures 10i and 10l). Following Woodcock and Mort [2008], they can be classified as chaotic to crackle breccias. Fault gouges, composed of incohesive fine-grained material, are common.

[54] The basement rock exposures often display multidirectional sets of conjugate, mainly normal, faults (Figures 10b, 10c and 13). These are steep in gneiss, quartzites, dolomites, or marbles, whereas they tend to listric geometry in calcschists and quartz-schists. The dominant extensional character of these faults is testified by displaced beds and micro-structures, such as striae, slickensides, and en échelon tension gashes (planar or sigmoidal). Less frequent strike-slip faults (mainly sinistral) are also present (Figures 10d and 10e). The basement faults show two main directions, similar to those found in the TPB: NNW- and NE- to E-striking (Figures 9e–9h). The trend of all the more than 1500 mapped faults (Figure 9e) is consistent with the strike direction of the kinematically constrained faults (n = 326, Figure 9f), which are hereafter used for the computation.

Figure 13.

(a) Exposure of quartzites showing a NNW-striking set of normal faults (F1) dissected by an ENE-striking normal/left lateral fault (F2). (b–d) Aerial photographs of well-exposed Marguareis (Figure 13b), Mongioe (Figure 13c), and Brignola (Figure 13d) areas (Briançonnais domain), with the main fault underlined with solid (F1) or dotted (F2) black lines.

[55] The cross-cutting relationships (Figure 13) recognized in most sites both at the outcrop and map-scale (SS 1, 4, 5, 6, 7, 10, 11, 12, 14, 15, 17) identify two fault-families, which are associated with chronologically different faulting episodes, although in a minor number of structural stations (SS 2, 3, 8, 9, 13, 16) the relative chronology is difficult to solve. The uncertainty is probably due to the heterogeneity of poly-deformed rocks as well as to the existence of preexisting discontinuities. However, the fault-distribution and the cross-cutting relationships at the map-scale (Figures 5 and 13) confirm an older age for the NNW–SSE oriented faults.

[56] The first and older family group (F1) is composed for the 84% of NNW-striking normal faults (88% with pitch >60°; Figure 12b). The second group (F2) is constituted by NE to ENE-oriented mainly normal faults (17% with pitch <45°; 38% with pitch comprised between 45° and 60°; 30% between 60° and 70°; 15% >70°). The 55% of these F2 faults have pitch <60° indicating a strong strike-slip component; ~63% of these faults are ENE left-lateral and ~37% are NE–SW right-lateral faults (Figures 9g, 9h, and 12b).

[57] About 10% of the analyzed faults are steeply dipping, mainly N-striking faults, and show reverse activity (Fi). These are, however, excluded from the analyses because their temporal relationship with the other faults could not be determined.

[58] The P and T axes distribution of the populations F1 and F2 shows a prevailing extension (Figure 12b). On the whole, and according to the fault analysis performed in the basin, the strain relationships associated with the two faulting phases in each structural station where two families have been distinguished display a main E–W extension for the first faulting phase, changing into NW–SE direction during the F2 phase.

5.3 Folding

[59] The entire Ligurian Alps-TPB system is affected by up to 10 km long wavelength parallel open folds (D5 structures; [Seno et al., 2005a, 2005b]). Parallel open folds with sub-vertical axial planes and ~N–S directed sub-horizontal axes also affect the Aquitanian–Early Burdigalian deposits of the TPB (Figure 11a). D5 folds generated asymmetric dome and basin interference patterns in the Ligurian basement [Bonini et al., 2010].

6 Discussion

6.1 Mechanisms of Basin Formation and Evolution

[60] Because of the overprinting and reactivation of extensional and compressive structures, the stress/strain regime is poorly resolved, and the age of the deformations in the TPB is poorly constrained. Extension has been related to a first period of subsidence caused by the Oligocene opening of the Liguro-Provençal basin [Mutti et al., 1995; Gelati and Gnaccolini, 1998]. The same authors suggest a Late Oligocene–Early Miocene inversion from an extensional to a compressional stress field, related to the Corsica-Sardinia drifting and to the thrust-activity of the Southern Alps.

[61] Other works [e. g., Carrapa et al., 2003a; Mosca et al., 2010] suggest that the TPB evolution was dominated by compression or transpression from the Oligocene until post-Pliocene time, with extension playing a minor role.

[62] The role played by the extensional regime in the evolution of the TPB is, however, controversial. On the basis of Anisotropy of magnetic susceptibility (AMS), Maffione et al. [2008] propose that N–S synsedimentary extension controlled the formation of the TPB, which acted as a basin passively carried on top of displacing nappes. The compressive structures found within the basin are here interpreted as gravitational slumps or post-Tortonian in age. However, other AMS analyses have been interpreted as suggesting a NE–SW to NW–SE shortening, which was active since the Oligocene [Carrapa et al. 2003a].

[63] Our approach integrates structural and stratigraphic analysis on the TPB-Ligurian Alps boundary and suggests that during the Oligo–Miocene the orogen-basin system evolved through three distinct tectonic phases.

6.1.1 Rupelian–Early Chattian (~34–26 Ma)

[64] Thermochronometric data indicate that, during the Rupelian–Early Chattian times, the metamorphic rocks of the Ligurian Alps were rapidly exhumed and exposed [Barbieri et al., 2003; Maino et al., 2012a]. At this time, the exhumation of the Briançonnais-Prepiedmont domain was controlled by a NNW-striking fault-system, related to the first brittle deformation phase (F1), which took place during deposition of the depositional sequences A–B1 (Figures 14a and 14b). This is the first brittle deformation phase (D4), and it overprints the older structures related to the metamorphic recrystallization (D1–D3 phases). Strain analyses of the F1 structures are compatible with a main ENE–WSW stretching (Figures 12). Similar E–W trending maximum extension direction has been derived from the coeval brittle structures recognized in the Voltri massif [Vignaroli et al., 2009].

Figure 14.

Simplified palaeogeographic models and basin evolution for the study area. Mi: Millesimo; Ce: Ceva. (a) Rupelian (~34–28 Ma—Sequence A): deposition of the Molare Formation. Depositional environments: flood-dominated alluvial-fan and fan-delta systems (1) and coastal plain (2). Activation of NNW-directed extensional faults (tectonic phase F1). (b) Early Chattian (~28–26 Ma—Sequence B1): deposition of channel-fill conglomeratic sandstones and conglomerates of the Millesimo body (Rocchetta-Monesiglio Fm.). Depositional environments: uplifted chain (3), alluvial plain (4), flood-dominated alluvial-fan and fan-delta systems (5), remobilized shelf (6), and deep-water channel-fill deposit (7). NNW extensional faulting still active but locally associated to reverse faulting (Fi). (c) Late Chattian (~26–23 Ma—Sequence B2–B5): deposition of the Retano, Cengio, Castelnuovo-Bric la Croce and Noceto turbidite systems (Rocchetta Monesiglio Fm.), and of the Cima della Costa Unit. Depositional environments: delta-fed submarine ramp system (8) and turbiditic lobes (9). Activation of NE to ENE-directed extensional faults (F2). (d) Aquitanian–Early Burdigalian (~23–18 Ma—Sequence B6): Condensed sedimentation of siliceous marls alternating with mm- to cm-thick fine to very fine sandstones (Rocchetta Monesiglio Fm.). Depositional environments: siliceous low-rate sedimentation area (10) and shelf with higher-rate deposition (11). Formation of long wavelength open folds. (e) Burdigalian (~20–16—Sequence B6): deposition of the upper turbiditic lobe of the B6 depositional sequence, which represents the transition to the more conspicuous turbiditic sequence C.

[65] During the Rupelian, the deposition of the Molare Fm. coarse-grained clastics marks the beginning of the marine transgression that migrated from NE to SW. The subsidence-driven transgression between the alluvial and the near-shore deposits of the Molare Formation and the overlying Early Chattian mudstone of the Rocchetta Monesiglio Fm. marks a general deepening of the TPB (Figures 14a and 14b). This transgression is younging toward south and west, from IFP 20 biostratigraphic zone (~33–31 Ma) in the north-eastern TPB, IFP 21b (~28–27 Ma) in the study area, to IFP 22 (~26–24 Ma) zone in the westernmost part of the basin [Gelati and Gnaccolini, 1998; Seno et al., 2010]. Such diachronous transgression is interpreted as the result of a Late Rupelian–Early Chattian westward migrating subsidence. Subsidence analyses from different localities of the TPB indicate that the vertical movements began in the Early Oligocene and continued throughout the Miocene [Carrapa et al., 2003b]. The subsidence curves indicate the first (Late Rupelian–Early Chattian) moderate pulse <1 km that prevalently acted in the southernmost part of the basin (Ceva-Cairo Montenotte; Figure 5). This early phase of subsidence is closely associated with the development of the extensional F1 structures in the basin. The F1 faults are characterized by modest length and throw ranging from centimeters to several tens of meters. The overall extension accommodated by these faults is of the order of several hundred meters, in agreement with the moderate subsidence rate calculated for this period [Carrapa et al., 2003a]. The orientation and structural characteristics of the F1 faults, therefore, support the interpretation that subsidence in the TPB during the Rupelian–Early Chattian was driven by extension [Mutti et al., 1995].

[66] Locally, some NNW-striking reverse faults, thrusts, and NE-verging folds (Fi; Figures 11a) have been detected in the lower stratigraphic succession (depositional sequence A–B1). The Fi faults have been interpreted as originated from the (Late Chattian?) inversion of early (Rupelian) extensional structures as already proposed by [Mutti et al., 1995]

6.1.2 Late Chattian (~26–23 Ma)

[67] In the Late Chattian, the deposition of the B2–B5 sequence marks the formation of the Langhe basin [Gelati and Gnaccolini, 1998]. The depocenters of the turbidite systems progressively shifted westward. Volume and facies suggest that these systems were fed by clastics source areas (with similar composition) which were considerably richer in sand than those which had supplied the coarse-grained clastics of the Millesimo body lower in the succession (Figure 14c). Such conditions were probably reached in periods during which subsidence was less pronounced.

[68] The turbidite systems were affected by ENE–WSW to NE–SW synsedimentary faulting (F2). The F2 faults systematically dislocate the previous F1 and Fi structures (Figure 14c). A related NW–SE extension is suggested not only by the strain analysis (Figure 12) but also by the progressive westward migration of the depocenters.

[69] In map view (Figures 13b–13d), the F2 faults show orientations comparable with the model of shear fracture orientation in non-coaxial deformation: specifically, these structures comprises purely extensional faults (mainly NE-oriented), synthetic Riedel fractures (R) with a pronounced ENE–WSW orientation (Figures 9d and 9h), and conjugate antithetical NNE-oriented Riedel fractures (R') (Figures 9c and 9g). This geometry is compatible with a regional sinistral strike-slip shear zone.

[70] The Fi faults recognized in the Sequence A–B1 are coherent with this tectonic scenario. They may be interpreted as localized restraining areas within the regional Late Chattian left-lateral transtensive zone. This consideration is supported by the deep basement/basin geometry illustrated by seismic lines [Mosca et al., 2010]. Indeed, in the Cuneo area, localized transpressive structures, involving both the basement and the Oligocene strata of the TPB, were detected: strike-slip faults with flower geometry influenced the deposition of the Oligocene sub-basins and are sealed by Miocene sediments. Strike-slip tectonics is furthermore testified in the Voltri massif, where the Sestri-Voltaggio fault system shows dextral kinematics since the Chattian [Capponi et al., 2009; Federico et al., 2009].

6.1.3 Aquitanian–Serravallian (~23–12 Ma)

[71] The active extensional and transtensional faulting documented in the Oligocene deposits dies away at the beginning of the Early Miocene. During the Aquitanian–Early Burdigalian, open folds developed both in the basin and in the metamorphic basement (D5 phase; Figure 11a). The siliceous lithozone (sequence B6—lower lithozone), characterized by a reduced sedimentation rate, reached its maximum thickness into the gentle depression of the synclines (Figure 14d). These contractional structures suggest a general NE–SW main shortening, which is also testified by the Aquitanian E–NE verging thrust of metamorphic basement onto Oligocene deposits, cropping out in the Alto Monferrato area (Grognardo thrust zone, Figure 2; [Piana et al., 2006]). Transpressional tectonics is furthermore documented in the Voltri massif, particularly for the Sestri-Voltaggio fault system [Capponi and Crispini, 2002, Capponi et al., 2009; Federico et al., 2009], and in the northern sector of the TPB [Mosca et al., 2010].

[72] The successive Late Burdigalian sedimentation of the upper lithozone of the B6 sequence was achieved in a highly subsiding trough SE–NW oriented, corresponding to the asymmetric troughs derived from the Aquitanian–Early Burdigalian folding (Figure 14e).

[73] The post-Early Burdigalian succession (depositional sequences C1–C6) is characterized by the presence of E–W to NE–SW-oriented extensional faults [Carrapa et al., 2003a; Gelati et al., 2010]. These faults are widespread but are systematically associated with small displacements. Although extensional tectonics is everywhere diffused, synsedimentary NW-trending contractional structures developed within Langhian-Serravallian deposits. The best exposed examples are the Ciglie anticline, unconformably overlapped by Mid-Langhian sediments, and the Bossola Pass structures characterized by reverse faults and folds involving Middle Miocene formations [Carrapa et al., 2003a]. On the whole, the extensional and contractional structures indicate ~NW–SE extension coeval with NE–SW shortening.

[74] During the Late Burdigalian–Serravallian, the subsidence accelerated over the entire TPB (vertical movements >3 km; [Carrapa et al., 2003a]). The calculated subsidence rates indicate a short period (17.5–15.5 Ma) of main subsidence (with rates ≥1 mm/yr) followed by a considerable slowdown of the rates (≤0,5 mm/yr). The low displacement associated with the detected normal faults and the absence of major structures clearly indicates that extension alone cannot justify the strong Late Burdigalian–Serravallian subsidence. In absence of important stretching, the Middle Miocene high values of subsidence may be explained by flexural response to thrust loading [Carrapa et al., 2003a; Carrapa and Garcia Castellanos, 2005]. Subsidence of the basin was mainly achieved under an overall transpressional stress regime associated with the convergence of the multi-vergent thrust-systems Apennines and Southern Alps converging under the TPB [Mosca et al., 2010].

[75] From the Late Miocene, the TPB succession experienced NW–SE directed compression, producing synsedimentary folding structures [Carrapa et al., 2003a]. Plio–Pleistocene uplift is responsible for the present day TPB morphology and elevation [e. g., Lorenz, 1984], characterized by gentle hills up to 800 m. During this stage, the inversion from subsidence to exhumation occurred, as testified by themochronometric data. Indeed, despite some inconsistency, apatite (U-Th)/He (AHe) analyses of Bertotti et al. [2006] from the Molare and Rocchetta-Monesiglio formations record cooling between 70° and 40° C at 11–14 Ma, which may be interpreted as the northward migrating basin exhumation through the time [Bertotti and Mosca, 2008]. These data are in agreement with the sedimentary record of the Langhian–Messinian succession that defines an overall regression [Gelati and Gnaccolini, 2003].

6.2 The Ligurian “Knot” in the Frame of Alpine-Apennines Tectonics

[76] Given its peculiar position at the Alps-Apennines junction, the TPB basin provides unique constraints for understanding the Late Neogene geodynamic setting of the central Mediterranean region.

6.2.1 Pre-Oligocene

[77] The Neogene evolution of the TPB basin is influenced by the pre-Oligocene tectonic scenario resulted from the early Alpine orogenic phases. During the Paleocene–Eocene, a wide range of units from both the Piedmont-Ligurian oceanic domain and the European continental margin (e. g., Briançonnais) were gradually involved in the subduction channel between Adria and Europe plates [e. g., Schmid et al., 1996; Stampfli and Marchant, 1997; Ford et al., 2006; Dumont et al., 2011]. Geochronological and metamorphic data from Western and Central Alps indicate that high-pressure metamorphism propagated from the internal (SE) orogenic zones (the oceanic domain) toward the outer (NW) ones (the European basement) until ca. 35 Ma [Schmid et al., 1996; Rosenbaum and Lister, 2005; Berger and Bousquet, 2008; Bousquet et al., 2008 and references within]. Also the rocks of the Ligurian Alps record the forelandward shift of the subduction zone during progressive accretion of the overriding plate: the Voltri oceanic units experienced eclogite-bluschist facies metamorphism between 49 and 40 Ma [Federico et al., 2005]. As indicated by the Lutetian age (~48–40 Ma) of the syn-orogenic sediments [Vanossi et al., 1986; Cabella et al., 1991; Dallagiovanna, 1995], the blueschists Ligurian Briançonnais units were involved in the subduction later than the oceanic rocks.

[78] The onset of metamorphism corresponds to the NW-ward obduction of part of the oceanic accretionary wedge (the Helminthoid Flysch, [Merle and Brun, 1984; Merizzi and Seno, 1991]) and with the deposition of syn-orogenic sediments into the European foreland basin related to the flexural loading by the Adriatic wedge [Ford et al., 2006]. These deposits record a diachronous marine transgression that migrated toward W–NW during the Middle to Late Eocene [Ford et al., 1999]. The early Alpine deformation is mainly northwestward directed [Malavieille et al., 1984; Platt et al., 1989b; Dumont et al., 2011] and has been rearranged by rotation during the Neogene bending of the arc [Collombet et al., 2002; Maffione et al., 2008]. Also in the Ligurian Alps, a NW to W-directed tectonic transport can be assumed, if the early Alpine transport directions (~SW-directed in present-day coordinates [Vanossi et al., 1986]) are restored to their pre-Late Oligocene position [Maffione et al., 2008]. Consistently, all the data indicate a ~NW-ward propagating setting before the early Oligocene and are coherent with the Eocene SE–NW directed relative motion of Adria and Europe reconstructed by palaeomagnetic and structural investigations [Figure 15a; Schmid and Kissling, 2000; Handy et al., 2010]. The western lateral termination of the Adria plate can be fixed in correspondence of the Ligurian Alps as the Eocene foreland flexural basin did not extend westward of the SE France [Ford et al., 2006; Dumont et al., 2011].

Figure 15.

Geodynamic reconstruction of the central-western Mediterranean region during the (a) Late Eocene, (b) Late Rupelian, (c) Late Chattian, and (d) Langhian times. Modified from Jolivet and Faccenna [2000] and Handy et al. [2010]. Adria and Africa motion paths are from Handy et al. [2010] and references within. LP: Liguro-Provençal basin; A: Algerian basin.

[79] Southern of the Ligurian area, the NW dipping Apenninic subduction zone extended for more than 1500 km until southern Iberia [Jolivet and Faccenna, 2000; Faccenna et al., 2004]. It consumed the Neotethyan Ocean eastward the Corsica, Sardinia, and Calabria blocks. Following the reconstruction of Lacombe and Jolivet [2005], Vignaroli et al. [2008], and Dumont et al. [2011], a major N–S sinistral transform boundary connecting the E-dipping Alpine and the NW-dipping Apennines subductions can be speculated east of Liguria and Corsica (Figure 15a).

6.2.2 Rupelian–Early Chattian (~34–26 Ma)

[80] At around 35 Ma, the motion of the Adria plate changed from NW-ward to WNW-ward (Figure 15b; [Schmid and Kissling, 2000; Handy et al., 2010]). The Western Alps become a zone of frontal collision between the Ivrea body (as frontal portion of Adria) and the European margin [Ford et al., 2006]. Since the Oligocene, the ~NNE–SSW trending curved shape of the Western Alps was forming [Collombet et al., 2002; Dumont et al., 2011]: it shows a general westward polarity related to the continuous SE–NW continent–continent collision between Adria and Europe, associated with indentation of the Ivrea body [Schmid and Kissling, 2000], whose southern termination was in correspondence of the Ligurian Alps [Dumont et al., 2011]. Consistently, kinematic data from Western Alps (Figure 16a) indicate a general NE–SW to NW–SE compression (moving from south to north) in the external units (Dauphionis; [Platt et al., 1989a; Dumont et al., 2011]). In the internal zones, the fossil accretionary wedge experienced a change from Late Eocene–Early Rupelian E–W ductile-semibrittle extensional shearing [Schwartz et al., 2009] to Late Rupelian–Early Miocene NW–SE extension coexisting with sub-horizontal NE–SW shortening and important right-lateral displacement along the major faults [Sue and Tricart, 2003; Tricart et al., 2004; Champagnac et al., 2006; Malusà et al., 2009; Perrone et al., 2010].

Figure 16.

Schematic tectonic sketches showing the Oligo–Miocene geodynamic evolution of the Western Alps and the Liguro-Provençal basin. (a) Late Rupelian: the Ligurian Alps-TPB and the Sardinia-Corsica block are restored to their Oligocene position prior the 50° of Miocene rotation [Gattacceca et al., 2007; Maffione et al., 2008]. The main direction of extension in the Ligurian Alps is orogen-parallel (NW–SE) and kinematically compatible with the coeval extension generated by the opening of the Liguro-Provençal basin. At this time, the Ligurian Alps record a fast exhumation, probably induced by tectonic denudation of the metamorphic chain via shallower detachment of the Helminthoid Flysch (HF). The Sestri-Voltaggio fault zone (SV; [Vignaroli et al., 2009]) separates the ophiolitic domain of the Ligurian Alps from the Appenines. (b) Late Chattian: the Ligurian Alps-TPB system was involved in a large scale left-lateral shear zone, accommodating the different motion of the west-directed Alps and north-east verging Apennines. In the box, an idealized distribution of tectonic elements associated with a wrench/strike-slip fault system is shown as comparison with the structures detected in the study area. The Sestri-Voltaggio and Villalvernia-Varzi lines acted as strike-slip faults [Di Giulio and Galbiati, 1995; Felletti, 2002; Capponi et al., 2009] (c) Early Miocene: the compressive front of the Western Alps is propagating toward the external domains while the inner zone of the chain was dominated by orogen parallel extension. The Corsica-Sardinia block accomplished their counterclockwise rotation of 50° [Gattacceca et al., 2007] inducing a comparable rotation in the Ligurian Alps-TPB system, which was affected by transpression. During this phase, the TPB acted as a strongly subsiding piggyback basin above rotating thrust sheets. Kinematic data of the Western Alps are from Perrone et al. [2010] and references within. HF: Helminthoid Flysch. Sa: Southern Alps. SL: Stura line; SA: Southern Alps; SV: Sestri-Voltaggio fault; TPB: Tertiary Piedmont Basin. VVL: Villalvernia-Varzi line. See text for a complete discussion of the evolutionary scenario.

[81] The Early Oligocene also corresponds to the initiation of the NE–SW trending Liguro-Provençal rifting [De Voogd et al., 1991; Séranne, 1999; Roca, 2001; Rollet et al., 2002]. This is considered as a back-arc basin generated from the southeastward roll-back of the Apennines subduction [e. g., Réhault et al., 1984; Jolivet and Faccenna, 2000]. The related extension developed along the ~NE trending axis in the western Mediterranean region (S France and E Spain). Extensional faults dissected pre-Alpine basements (e. g., in the western Corsica, Provence) and areas previously affected by compressive deformation (e. g., Pyrenees, the Iberian chain, Figure 15b). This extensional deformation caused crustal thinning and was subsequently associated with Early Miocene emplacement of oceanic crust actually exposed in the central part of the basin [De Voogd et al., 1991].

[82] In the Liguria area, the Rupelian marks the beginning of the TPB deposition; the early sediments record the dismantling of the rapidly exhuming orogen associated with moderate rate of basin subsidence. The deposition occurred in the retro-foreland of the Ligurian Alps and was controlled by NNW-directed mainly extensional faults (F1) at high angle with the previous contractional structures of the metamorphic basement. The F1 faults represent the first brittle deformation phase (D4) developed in the metamorphic basement: the general pattern of mesoscale strain axes indicates a regional E–W to NE–SW extension (considering present-day coordinates). Paleomagnetic data reveal ~50° of counterclockwise rotation for the TPB and the underlying Ligurian basement in the Aquitanian–Serravallian times [Maffione et al., 2008]. If basin and basement are restored to their Early Oligocene position, the inferred kinematics from the F1 faults is in agreement with the regional strain of the coeval Liguro-Provençal basin (Figure 16a).

[83] These extensional faults show continuity with the structures associated with the Liguro-Provençal rifting in the SE France (Provence). The close temporal, spatial, and geometrical relationship between the onset of mainly NE–SW trending extensional structures and the TPB deposition suggests that the regional extension in the orogen-basin system was induced by the rifting dynamics as already envisaged by Vanossi et al. [1994], Mutti et al. [1995], and Gelati and Gnaccolini [1998]. The occurrence of the strongly thickened Alpine crust (i. e., the Ligurian Alps) blocked northward by the ongoing indentation of the Ivrea body, probably represented a physical barrier for the propagation of the rift, which no longer developed northern to the Liguria. Therefore, following the model of rifting evolution of Lavier and Manatschal [2006], the Ligurian Alps represented a stretching area ahead of the Liguro-Provençal rifting south to the tip of the Ivrea body (Figure 16a).

[84] The presence of an extensional domain in the Voltri Massif has been recently suggested by Vignaroli et al. [2008, 2009, 2010]. In these works, a Late Eocene–Oligocene syn-greenschist unroofing of the HP oceanic units via extension has been proposed. Extensional tectonics operated by the reactivation of previous compressional structures, with a progression from plastic to frictional deformation mechanisms. The Voltri Massif undergone extensional denudation through shear localization along major plastic to brittle shear zones separating the Voltri Massif from the Apennines in the east and from the Alpine Briançonnais domain in the west.

[85] This model accounts for the extensional character of the ductile-brittle and brittle structure identified both in the Voltri Massif and Briançonnais units (Table 1). It is alternative to the classical interpretation of the exhumation of HP oceanic rocks caused by subduction channel processes associated with polyphase contractional regime [e. g., Federico et al., 2007]. Moreover, the emphasis of the role of extension during the evolution of the Voltri Massif [Hoogerduijn Strating, 1994; Vignaroli et al., 2008, 2010] can explain the fast Eocene–earliest Oligocene exhumation (3–5 mm/yr; Federico et al. [2005]) recorded by the ophiolitic rocks from greenschist facies stage condition to the surface and supported by the thermochronometric data from the TPB [Barbieri et al., 2003; Carrapa et al., 2004].

[86] Low-temperature thermochronometric data indicate that also the continental units record fast (1.3–6.8 mm/yr), tectonically controlled, Oligocene exhumation through the uppermost (4–5 km) crust [Maino et al. [2012a]. These data support the concept that, at least in the shallowest crust, not only the oceanic rocks were affected by rapid exhumation. The very high rates recorded by the Briançonnais rocks suggest that the fast exhumation was not limited to the Voltri Massif. This implies that, in the uppermost 4–5 km, the entire Ligurian orogen experienced a relatively homogenous denudation without significant difference between oceanic and continental units.

[87] This scenario can be explained by tectonic denudation via extensional detachment: Maino et al. [2012a] suggests a westward (in Oligocene coordinates) sliding of the Helminthoid Flysch (Figure 16a). During this translation, about 4–5 km-thick rock mass was progressively removed from the metamorphic chain, resulting in the high cooling/exhumation rates calculated from the thermochronometric data. The shallow extensional detachment plane (dipping <30°) is represented by a plastic shear zone (Lambeaux des charriage, [Lanteaume, 1968]), constituted by a clay-rich tectonosedimentary mélange, now resting between the Dauphinois/Briançonnais units and the Helminthoid Flysch [Lanteaume et al., 1990; Ford et al., 1999].

[88] Summarizing, the whole of data are coherent with the presence in Liguria of an Early Oligocene extensional domain that promoted a two-stage exhumation process: a syn-greenschist exhumation of HP rocks associated with the coupling of oceanic (Voltri Massif) and continental rocks [Vignaroli et al., 2008, 2009, 2010] and a subsequent tectonic denudation of the entire metamorphic basement through shallow detachment [Maino et al., 2012a]. The extensional regime caused crustal thinning and generated the space for the TPB sedimentation. This tectonic scenario is in agreement with the proposed position of the Ligurian Alps (Figures 15b and 16a) ahead of the Liguro-Provençal rifting and connecting the opposite propagation of the Apennines and Alpine arcs [Vignaroli et al., 2008].

6.2.3 Late Chattian (~26–23 Ma)

[89] The Late Chattian coincides with the switch from extensional to transtensional regime in the Ligurian Alps-TPB, marked by the development of the F2 fault system.

[90] At this time, the compressive thrust front of the Apennines were migrating outward while the deep indentation of the Ivrea body under the Western Alps represents a stop for the N Apennines propagation (Figures 15c). The eastward migration of the fold-and-thrust belt of the Apennines is testified by the coeval migration of the related foredeeps [Ricci Lucchi, 1986]. This translation resulted in a curved belt reflecting the convexity of the subduction trench. This shape is due to the presence of easily subductable oceanic Ionian crust in the central part of the Apennines-Maghrebide subduction system (Figure 15c), with respect to the continental subduction of Adriatic crust below the Europe in the north [Jolivet and Faccenna, 2000]. The Northern Apennines were therefore characterized by relatively slow north-eastward translation, hampered by the complex interaction between Adria and Europe. In this context, the junction between the two belts, i. e., the Ligurian Alps, suffered large-scale sinistral transtension accommodating the different motions of the Alps and Apennines. The F2 fault-network shows orientations comparable to a widely accepted model of left-lateral shear zone (Figure 16b). This geometry, therefore, indicate that, since the Late Chattian, the TPB-Ligurian Alps system acted as a pluri-kilometric sinistral strike-slip shear zone. In this context, the Sestri-Voltaggio line is interpreted as a dextral Riedel fracture as proposed by Capponi et al. [2009].

[91] The Late Chattian transtensional phase in the Ligurian Alps-TPB system represents the transition from Early Oligocene rifting-related extension to the Early Miocene rotation-related transpression. It marks the period when the Apennines start to “pull” the southern Western Alps northeastward, producing a large shear zone (i. e., the Ligurian Alps-TPB) between the two chains.

6.2.4 Aquitanian–Serravallian (~23–12 Ma)

[92] During the Early–Middle Miocene, the Ligurian Alps-TPB system was dominated by ~NE–SW shortening and coeval ~NW–SE extension. This transpressive phase (D5) is coeval with a period of major assessment of the central Mediterranean area (Figure 15d): in the Liguro-Provençal back-arc basin, the thinning of the continental lithosphere resulted in the formation of new oceanic crust between ca. 20 and 15 Ma [Séranne, 1999; De Voogd et al., 1991]. Synchronously with the oceanization of the basin, the Corsica-Sardinia block broke away from the European margin experiencing an anticlockwise rotation of ~60° [Gattacceca et al., 2007]. This rotation was driven by the eastward retreat of the Adriatic-Ionian composite slab [Jolivet and Faccenna, 2000; Castellarin, 2001; Faccenna et al., 2004]. The Ionian oceanic subducting crust increased the curvature of the central part of the subduction trench. This deep structure resulted in a highly curvilinear fold-and-thrust belt characterized by mainly E-vergent structures. Differently, the presence of subducting continental Adriatic crust in the northernmost part of the subduction and the proximity of the retro-front of the Alps impeded a free propagation of the northern Apennines (Figures 15d and 16c). Here, the north-verging thrust-fronts (Monferrato arc) progressively approached the south-verging structures of the Southern Alps beneath the TPB [Mosca et al., 2010]. In this period, the TPB experienced an acceleration of the subsidence rates [Carrapa et al., 2003a], likely related to a flexural response to thrust loading [Carrapa and Garcia Castellanos, 2005]. The increasing load caused by the N-migrating Apennines thrusts approaching the Southern Alps front provided flexural tilting accommodating the high value of subsidence recorded in the TPB during the Middle Miocene. Therefore, the most plausible reason for the flexural subsidence in the TPB is the downward pull of the Apennines slab.

[93] The indentation of the Adriatic plate against the Europe [Schmid and Kissling, 2000] preserves an overall convergent framework in the Western Alps still during the Early Miocene. In the internal zones, transpressional tectonics was accommodated by ca. orogen-perpendicular compression and orogen-parallel extension (~N–S) associated with dextral strike-slip movements [e. g., Sue and Tricart, 2003; Tricart et al., 2004; Champagnac et al., 2006; Baietto et al., 2009; Perrone et al., 2010; Sanchez et al., 2011]; while in the foreland, outward-verging thrust sequences developed [e. g., Ford et al., 2006] (Figure 16c).

[94] In this geodynamic framework, the Ligurian Alps-TPB system was forced to rotate of ~50° [Maffione et al., 2008] and migrate east/northeastward. This rotation was possible because the Ligurian segment of the Western Alps was not hampered by the presence of the Ivrea body as in the Western Alps [Schmid and Kissling, 2000; Dumont et al., 2011]. The Ligurian Alps acted as a free junction between the Western Alps and the Northern Apennines (Figures 15d and 16c). The shortening directions derived from the structural analyses are consistent with the orientation of a regional left-lateral shear zone, producing t he oroclinal bending of the Ligurian orogen (Figure 16c).

[95] Finally, from the Late Miocene, the mainly NW–SE directed compression recorded in the TPB resulted in the final Plio–Pleistocene exhumation of the basin. At this time, the regional tectonic evolution is associated with the N- and NE-ward translation of the Ligurian units of the Apennines and the TPB successions onto the Insubric foreland, along the Padan thrust front [e. g., Pieri and Groppi, 1981].

7 Conclusion

[96] Our research addresses field-based structural and stratigraphic investigations to the Tertiary Piedmont Basin (TPB) and the Ligurian Alps, which represent the junction between the two major orogenic systems of the Europe, the Alps-Dinarides and the Apennines-Maghrebides belts. The data here presented allow proposing a complete evolutionary scenario for the Oligo–Miocene history of the TPB: born as a retro-foreland basin of the Ligurian Alps, it records Late Rupelian–Early Chattian extension which also controlled the coeval tectonic exhumation and denudation of the metamorphic basement. This extensional regime was generated ahead the apical closure of the Liguro-Provençal rifting, and it is associated with back-arc crustal thinning caused by the roll-back of the Apennines subduction.

[97] During the Late Chattian, the TPB evolved as a foreland basin located above a transtensional zone connecting the opposite movements of the Alpine and Apennines arcs. Since the Early Miocene, the TPB acted as a strongly subsiding piggy-back basin above rotating thrust sheets associated with the regional rotation caused by the oceanic spreading of the Liguro-Provençal basin, the Corsica-Sardinia drifting, and the eastward retreat of the Apenninic slab.

[98] On the whole, during the Oligocene–Miocene, the TPB-Ligurian Alps system represented the pivot around which the Western Alps, the Liguro-Provençal basin, the Corsica-Sardinia block, and the Apennines moved controlling the evolution of the central Mediterranean area.

Appendix A: Methodological Approach

A1. Structural Analysis

[99] Meso-structural analysis and mapping of brittle structures have been carried out on both the metamorphic chain of the Ligurian Alps and the sedimentary succession of the TPB (Figures 2 and 5).

[100] Structural analyses in the basement were performed in outcrops with homogeneous lithology, tectono-metamorphic evolution, fault morphology, rheology, thickness of fault-rocks, and kinematic indicators. More than 1500 faults have been cataloged in more than 60 sites; here we only report those measurements for which the sense of movement has been indubitably identified; for the Briançonnais and Prepiedmont units, we report 326 fault/striation pairs from 17 well-exposed sites (structural stations, hereafter SS; Table 2). Each site corresponds to between 12 and 33 measurements on a single outcrop or on a section no longer than 60 m. Field observations have been integrated with the orientation data, morphology, dimension, and spacing of faults. In this work, we used the fault rocks classification of Woodcock and Mort [2008], which assumes the coarse grain-size as the most relevant criterion to distinguish the fault breccias. Families of faults have been distinguished on the basis of genetically related small-scale structures, superimposition of movements, and by the cross-cutting relationships. Most of our measurement sites are along well exposed natural outcrops at high altitude (>2000 m) or within open quarries, where faults are completely exposed; at low elevations where the vegetation masks the rocks, faults are difficult to trace and sense of movement difficult to be determined. Field data have been compared with observation on aerial photographs of well outcropping basement rocks (Figures 11b–11d).

[101] Within the basin, stratigraphic markers allow faults and their relative movement to be more easily detected. Here, we have recorded 196 faults and their unequivocal sense of movement. The cross-cutting relationships are often difficult to investigate, and the fault hierarchy has been deduced by the occurrence/absence of families of faults distinguished by their prevailing direction.

A2. Strain Analysis

[102] Our field work is based on the systematic measurement of fault populations, including fault plane orientation, striae and slickenside orientation, and shear sense, classified on the basis of the reliability criterion of the shear sense (certain, probable, uncertain). We decided not to resolve the paleostress field because the determination of the directions of principal stress axes from the field measurement of striated faults (stress inversion method) can be considered reliable only in rocks that experienced deformation under a stress field that had to be homogeneous both spatially and temporally [e. g., Angelier, 1990; Twiss and Unruh, 1998]. These assumptions are not valid for many of the rocks exposed in the study area and, in particular, for the basement rocks, affected by multiple deformations or reactivations of preexisting discontinuities [Harris and Cobbold 1984; Flodin and Aydin 2004]. For these reasons, we performed paleostrain analysis, which represents a qualitative determination of the incremental strain ellipsoid for each station [e. g., Malusà et al., 2009]. The strain analysis has been performed through the graphical construction of the principal incremental shortening (P) and extension (T) axes for each population of faults. The P and T axes concentration provides an approximate orientation of the strain axes [Marrett and Allmendinger, 1990; Twiss and Unruh, 1998]. If the P-T results indicate a single maximum concentration which can be related to a single tectonic event, we compared the P-T analyses with the field of incremental shortening and extension deriving from the right dihedra method of Angelier and Mecheler [1977]. All the strain analyses have been computed using T-TECTO 3.0 computer program (Žalohar and Vrabec [2008], available at http://www2.arnes.si/~jzaloh/t-tecto_homepage.htm).


[103] We would like to thank C. Persano, G. Dallagiovanna, G. Ghibaudo, and A. Di Giulio for an early reading of the manuscript and useful suggestions. The editors O. Oncken and C. Faccenna and the reviewers H. J. Gawlick, F. Massari, and G. Vignaroli are gratefully acknowledged for substantial revisions and constructive suggestions that greatly improved the paper. The research was supported by PRIN grants.