tectono-sedimentary evolution of the tertiary piedmont basin (nw italy) within the oligo-miocene...

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

Matteo Maino,1 Alessandro Decarlis,2 Fabrizio Felletti,3 and Silvio Seno1

Received 12 October 2012; revised 18 April 2013; accepted 23 April 2013.

[1] We analyze the tectono-sedimentary and thermochronometric constraints of the TertiaryPiedmont Basin (TPB) and its adjoining orogen, the Ligurian Alps, providing new insights onthe basin evolution in response to a changing geodynamic setting. The geometry of thepost-metamorphic faults of the Ligurian belt as well as the fault network that controlled theOligo–Miocene TPB deposition has been characterized through a detailed structural analysis.Three main faulting stages have been distinguished and dated thanks to the relationshipsamong faults and basin stratigraphy and thermochronometric data. The first stage(F1, Rupelian–Early Chattian) is related to the development of extensional NNW-directedfaults, 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, whichdeposition was influenced by the second faulting stage (F2). This phase was mainlycharacterized by NE- to ENE-striking faults developed within a transtensional zone. Since theMiocene, the whole area was dominated by transpressive tectonics. The sedimentation wasrepresented by a condensed succession followed by a very thick, turbiditic complex. At theregional scale, this succession of events reflects the major geodynamic reorganization in thecentral Mediterranean region during the Oligo–Miocene times, induced by the late-collisionalprocesses of the Alps, by the eastward migration of the Apennines subduction and by theopening of extensional basins (i. e., the Liguro–Provençal Ocean).

Citation: Maino,M., A. Decarlis, F. Felletti, and S. Seno (2013), Tectono-sedimentary evolution of the Tertiary PiedmontBasin (NW Italy) within the Oligo–Miocene central Mediterranean geodynamics, Tectonics, 32, doi:10.1002/tect.20047.

1. Introduction

[2] Within a foreland system, basins may form ahead of theactive pro- or retro-front of the fold-and-thrust belts [Allenand Allen, 1990; DeCelles and Giles, 1996]. These basinsare closely controlled by the tectonic processes that resultfrom the interplay of orogen uplift, flexurally-triggered basinsubsidence, and eustacy [e. g., Catuneanu, 2004; Cederbomet al., 2004; Allen, 2008; Cloetingh and Negendank, 2009].Understanding the evolution of these basins may be difficult,especially when an evolving geodynamic setting promotesmultiple stages of deformation, severely modifying theprimary structural features of the orogen-basin system.However, the integration of chronological constraints (bothbiostratigraphic and thermochronometric) with detailed

stratigraphic and structural analyses helps to determine thetiming and the rates of deformation stages describing thetectonic evolution of the orogen-basin system.[3] This paper is concerned with the Tertiary Piedmont

Basin (TPB; Figure 1), which is an enigmatic basin thatevolved in response to a changing tectonic setting. Accordingto the present literature [e. g., Gelati and Gnaccolini, 1998;Schmid and Kissling, 2000], it is generally regarded as awedge-top-basin located on top of the junction between thetwo main orogenic systems of the central Mediterranean: theWestern Alps, characterized by a main westward tectonicnappe-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 dur-ing three main tectonic episodes, the exhumation of theLigurian sector of the Western Alps, the opening of theLiguro-Provençal basin, and the formation of the Apenninesthrust belt.[4] The complex evolution of the TPB is well documented,

although conclusions on its evolution are often contradictoryas its origin and development is explained in terms of eithercompressional 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;

1Dipartimento di Scienze della Terra e dell’Ambiente, Università degliStudi di Pavia, Pavia, Italy.

2IPGS-EOST, Université de Strasbourg, CNRS UMR 7516, Strasbourg,France.

3Dipartimento di Scienze della Terra, Università degli Studi di Milano,Milano, Italy.

Corresponding author: M. Maino, Dipartimento di Scienze della Terra edell’Ambiente, Università degli Studi di Pavia, via Ferrata, 1 - 27100 Pavia,Italy. ([email protected])

©2013. American Geophysical Union. All Rights Reserved.0278-7407/13/10.1002/tect.20047

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TECTONICS, VOL. 32, 1–27, doi:10.1002/tect.20047, 2013

Gelati and Gnaccolini, 1998; Felletti, 2002; Carrapa et al.,2003a; Gelati and Gnaccolini, 2003; Carrapa and GarciaCastellanos, 2005; Maffione et al., 2008; Naylor andSinclair, 2008; Vignaroli et al., 2008, 2009, 2010]. A hugenumber of multidisciplinary studies have been performed inthe TPB during the last decades, including structural and sed-imentologic 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; Moscaet al., 2010], thermochronology [Barbieri et al., 2003;Carrapa et al., 2003b, 2004; Bertotti et al., 2006], paleomag-netism [Carrapa et al., 2003a;Maffione et al., 2008], and nu-merical and statistical modeling [Felletti, 2004a, 2004b;Carrapa and Garcia Castellanos, 2005; Bertotti andMosca, 2008]. Despite this abundant information, chronolog-ical constraints of the tectonic stages experienced by thebasin, in particular of the early phases, need to be strength-ened. These uncertainties prevent an accurate determinationof the mechanisms that drove to the formation of the basinand hamper the formulation of appropriate models describingthe tectonic evolution of the region. This work focuseson the SW part of the TPB (Figure 2), along its boundarywith the orogen, where the tectonic features associated withthe early stages of the basin formation are still preserved.Extensive sedimentary and structural investigations havebeen combined with the structural analysis of the Ligurianorogen post-metamorphic fault-pattern and compared withneighboring areas data. Geometric relationships betweensedimentary rocks and tectonic structures as well asthermochronologic data are used to constrain the timing ofthe deformation phases.[5] We present here an improved basin analysis based on a

precise reconstruction of the tectonic phases experienced bythe TPB and the adjoining Ligurian Alps. The depositionaland structural features reflect the deep processes controllingthe linked evolution of the Alps and Apennines orogenic

systems. The new data lead to a critical evaluation of thevarious geodynamic models adopted for this region.

2. Geological Background

2.1. Regional Framework

[6] The central/western Mediterranean area features ageodynamic complexity caused by the relative movementsof three main plates (Africa, Adria, and Europe) and severalinterposed oceanic basins, which led to the formation oftwo arcuate orogenic systems (i. e., the Alps-Dinarides andthe Apennines-Maghrebides belts; Figure 1).[7] Since the Cretaceous, the Piedmont-Ligurian Ocean

was consumed in a south-east to south directed subductionzone [e. g., Stampfli and Marchant, 1997]. After theEocene–Oligocene collision, the indentation of the Adriaticand the European plates leads to the building of the Alpinechain in the central Mediterranean [e. g., Platt et al., 1989a;Ford et al., 2006]. Starting from the Oligocene, severalextensional basins (i. e., the Algerian, Ligurian-ProvençalAlboran, and Tyrrhenian basins; Figure 1) originated abovea previously thickened continental crust [De Voogd et al.,1991; Jolivet and Faccenna, 2000; Roca, 2001]. Thesebasins obliquely cross-cuts the contractional structures ofthe Betics, Iberian Chain, Pyrenees, and Alps, suggestingan independent origin with respect to these orogens[Carminati et al., 1998]. There is a general consensus to con-sider these basins as the results of back-arc extension relatedto 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, asit implies the underthrusting of a composite crust (the Adriaand Africa continents and the interposed Neotethys Ocean,actually represented by the Ionian basin; Figure 1) belowthe European plate. It was interpreted either as a subductionwith a single permanent polarity since Late Cretaceous time

Figure 1. Simplified present-day tectonic map of the Central-Western Mediterranean region (fromFaccenna et al. [2004] and references within), showing the location of the study area. LN: Ligurian knot;Pan. basin: Pannonian Basin; TPB: Tertiary Piedmont Basin,

MAINO ET AL.: EVOLUTION OF THE TERTIARY PIEDMONT BASIN

2

[e. g., Jolivet et al., 1998; Faccenna et al., 2004] or as theresult of an Eocene flip of an earlier Cretaceous SE-directedAlpine subduction [e. g., Elter and Pertusati, 1973;Michard et al., 2002; Handy et al., 2010]. In any case, sincethe Oligocene onward, the retreat of the slab resulted in aprogressive 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 theeastward Apennines nappe stacking and migration[Castellarin, 2001]. During the Early–Middle Miocene, theprogressive eastward propagation of the Apennines arc,associated with the oceanic spreading of the Liguro-Provençal and Algerian basins, led to the anticlockwisedrifting of the Corsica-Sardinia block [e. g., Alvarez et al.,1974; Rosenbaum et al., 2002]. From the Late Mioceneonward, extension shifted eastward to the Tyrrhenian Sea(Figure 1), leading to the present-day configuration of

the Apennines chain [Malinverno and Ryan, 1986; Ciarciaet al., 2012].[8] The Alps-Apennines evolution is characterized by the

onset of several foreland basins along both the pro- and theretro-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 andLickorish, 2004; Naylor and Sinclair, 2008]. The TertiaryPiedmont Basin is located above the connection of the twoorogenic systems (the “Ligurian knot” of Laubscher et al.[1992]; Figure 1), thus recording the interplay between theAlps-Apennines belts. The Upper Eocene–Miocene sedi-mentary succession of the TPB is mainly formed by clasticdeposits resting on the Pre-Cenozoic substratum of theLigurian Alps in the west and the Ligurian units of theNorthern Apennines to the east (Figure 2). The TPB isdivided in four main paleogeographic domains: from west

Figure 2. Geological map of the Ligurian Alps and Tertiary Piedmont Basin with the distribution of themain 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 locationof the stations is given in Table 2.


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to east, the Monregalese High, the Langhe basin, the AltoMonferrato High and the Borbera-Curone basin (Figure 3).The paper focuses on the Langhe basin, which representsthe south-western sector of the TPB.

2.2. Structure of the TPB-Ligurian Alps System

[9] The Ligurian Alps (Figure 2) are constituted by imbri-cated tectonic units belonging to different palaeogeographicdomains [Vanossi et al., 1986; Seno et al., 2005a, 2005b]:the Briançonnais and Prepiedmont (belonging to theEuropean continental margin) and the Piedmont-Ligurian.The Briançonnais and Prepiedmont units show a strati-graphic sequence composed of a Paleozoic basement[Maino et al., 2012b] comparable with other basement rocksbelonging to the “souther Variscan realm” [Rossi et al., 2009;Casini et al., 2012], a Permian volcano-sedimentary succes-sion [Dallagiovanna et al., 2009] and a Meso–Cenozoiccover with huge sedimentary gaps [Decarlis and Lualdi,2008, 2009, 2011]. The Piedmont-Ligurian domain includesmetaophiolite and metasedimentary units, which are widelyexposed in the eastern part of the Ligurian belt (i. e., theVoltri massif [Capponi and Crispini, 2002]; Figure 2). Adetached part of the oceanic cover, the turbiditic successionsof the Helminthoid Flysch, has been shifted far from its orig-inal basement, and now rests at the south-westernmost sectorof the chain [Di Giulio and Galbiati, 1991; Ford et al., 1999].[10] To the east, the sub-vertical Sestri-Voltaggio Fault sep-

arated the high pressure (HP) metamorphic units of the VoltriMassif from three blueschist to pumpellyite-actinolite faciestectonic 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 representedby 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 evolu-tion linked to subduction-collisional events [Vanossi et al.,1986; Seno et al., 2005a]. The main deformational phasesand the associated metamorphic conditions are summarizedin Table 1.[12] In response to the Late Cretaceous–Eocene subduction

of the Piedmont-Ligurian Ocean under the Adria plate, therocks of the Piedmont-Ligurian and Briançonnais domains ex-perienced a highly variable metamorphic re-crystallization(from greenschists to eclogite facies), depending on the posi-tion within either the subduction channel or the orogenicwedge [e. g., Ernst, 1981; Messiga and Scambelluri, 1991;Goffé et al., 2004]. During theMiddle/Late Eocene, these unitswere displaced and thrusted toward the foreland (SW; D1phase) forming a nappe pile [Seno et al., 2005a], constitutedby, from bottom to top, the Briançonnais, the Prepiedmont,and the Piedmont-Ligurian units, all stacked onto theDauphinois domain [Dallagiovanna et al., 1997; Seno et al.,2005b; Bonini et al., 2010]. The following (Late Eocene–Early Oligocene) retrogressive metamorphism recorded bythe Piedmont-Ligurian and Briançonnais units was achievedduring the last ductile deformational phases (D2–D3;Table 1) leading the exhumation of the deep-seated rocks pres-ently cropping out in the eastern part of the belt, i. e., the

Figure 3. (a) Study area (square) within the Langhe Basin. (b) General overview of stratigraphy of theOligo–Miocene basin fill. A, B1–B6, and C1–C3 indicate the Oligo–Miocene depositional sequencesdefined 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).


4

Voltri Massif. The structural interpretations of these phasesare debated, as they have been related to polyphase contrac-tional regime [e. g., Vanossi et al., 1986; Capponi andCrispini, 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 gravita-tional 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 theHelminthoid Flysch followed a NW- to W-ward LateCretaceous–Eocene subaqueous translation of about 200 kmdue to its involvement in an intra-oceanic accretionary prism[Di Giulio and Galbiati, 1991; Seno et al., 2005a]. TheOligocene gravity-driven detachment of the HelminthoidFlysch resulted in passive folding, with folds truncated at thebase, 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 develop-ment of the D4–D5 systems of cataclastic-gouge fault zones[Table 1; Vanossi et al., 1994; Crispini et al., 2009; Federicoet al., 2009; Vignaroli et al., 2009; Bonini et al., 2010]. Thestress/strain regime and the timing of these phases are poorlyresolved as they resulted from the overlap of several tectonicevents, including the Early–Middle Miocene oceanic spread-ing of the Liguro-Provençal basin, the coeval 50� of verticalrotation [Maffione et al., 2008] of the Ligurian Alps-TPBsystem [Vanossi et al., 1994], and the drifting of theCorsica-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 sedimentarysuccession. The TPB is characterized by the lack of majortectonic structures, whereas small scale structural features,such as folds and faults ranging from hundreds of meters tocentimeters are common. Several authors documentedsynsedimentary 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; Carrapaet al., 2003a; Piana et al., 2006], and decametric tocentimetric-scale normal and strike-slip fault systems have beendetected on the whole of the basin succession [e. g.,Mutti et al.,1995; Gelati and Gnaccolini, 1998; Felletti, 2002; Carrapaet al., 2003a; Vignaroli et al., 2009].

2.3. Stratigraphy of the Langhe Basin

[16] The Langhe Basin overlies a series of Alpine nappesthat were already emplaced in the Late Eocene. The basinextends over an area of about 1800 km2 and hosts anOligo–Miocene succession more than 4000m thick [Gelatiand Gnaccolini, 1998, 2003; Forcella et al., 1999]. Theformation of the basin in the Early Oligocene is recordedby alluvial fan and fan-delta deposits, followed by transgres-sive shallow-marine sediments (Molare Formation, deposi-tional sequence A; Figure 3b, [Gelati and Gnaccolini,2003]). During this time, some topographical “highs” weredirectly covered by small coral reefs [Franceschetti, 1967;Gnaccolini, 1978; Fravega et al., 1987].[17] Drowning of the basin in the Late Oligocene delin-

eated a complex shape of the south-western margin andoriginated an array of confined depocenters at differentT

able1.

DeformationPhasesin

theContin

entalB

riançonnaisDom

ain(Brianç)

andin

theOceanicPiedm

ont-LigurianUnits(V

oltri)of

theLigurianAlpsa

Deformationphase

Area

Fabric

Metam

orphism

Age

References

Pre-D

1Voltri

Isoclin

alsimilarfolds(unconstrained

vergence)

eclogite-blueschist-facies

Early/M

iddleEocene

Capponi

andCrispini[2002];F

edericoetal.,2005;[2007]

D1(D

1/D2in

Capponi

andCrispini[2002])

Voltri

Tight,sub-isoclinalfolds;nappestacking

(unconstrained

vergence)

Na-am

phibolegreenschist/

greenschist-facies

s.s.

Middle/LateEocene

Capponi

andCrispini[2002];Federicoetal.[2005];

Vignarolietal.[2005]

D1

Brianç

SW-verging

sub-isoclin

alfoldsandnappestacking

Upto

blueschistfacies

Bartonian–

Priabonian

Vanossietal.[1986];Seno

etal.[2005a];

Boninietal.[2010]

D2(EPFin

Capponi

andCrispini[2002])

Voltri

ExtensionalW-verging

shearbands

greenschist-facies

Priabonian–Early

Olig

ocene

HoogerduijnStratin

g[1994];C

apponiandCrispini[2002];Vignaroli

etal.[2009,2

010]

D2(RSZin

Capponi

andCrispini[2002])

Voltri

SW

vergingreverseshearzones

greenschist-facies

tolow

greenschist-facies

Priabonian–Early

Olig

ocene

Capponi

andCrispini[2002]

D2

Brianç

N/NE-verging

back-folds

andback-thrusts

Upto

greenschist-facies

Priabonian–Early

Olig

ocene


etal.[2005a];B

oninieta

l.[2010]

D3

Voltri

Chevron

foldsandsemi-brittleextensionalW

-vergingshearing

Low

greenschist-facies

Early

Olig

ocene

Capponi

andCrispini[2002];Vignarolietal.[2009]

D3

Brianç

S-verging

chevronor

kink

folds

Non-m

etam

orphic

Early

Olig

ocene

Seno

etal.[2005a];B

oninietal.[2010]

D4

Voltri

Normalandstrike-slip

faults

Non-m

etam

orphic

Early

Olig

ocene

Vignarolietal.[2009]

D5(D

4in

Federico

etal.[2009])

Voltri

Openfolds,thrustandstrike-slip

faults

Non-m

etam

orphic

Early

Miocene

Capponi

andCrispini[2002];C

rispinieta

l.[2009];

Federicoetal.[2009]

D4/D5(D

4in

Bonini

etal.[2010])

Brianç

Openfolds,norm

alandstrike-slip

faults

Non-m

etam

orphic

Olig

ocene–Miocene


etal.[2005a];B

oninieta

l.[2010]

a The

late-orogenicdeform

ationalp

hases(D

4–D5)

arecoevalwith

thedepositio

nof

theTPBsedimentary

succession.


5

elevations with respect to the basin floor. The drowning ismarked by the sudden change in the deposition from theshallowmarine sediments of theMolare Fm. to the hemipelagicmudstones of the “Muddy Framework” [Gelati and Gnaccolini,2003], with interbedded lenticular bodies of sandstone and con-glomerate (Figure 3b). On the whole, these sediments form theRocchetta-Monesiglio Formation (Late Rupelian–Burdigalian)and are organized into six depositional sequences (B1–B6;Figure 3b, [Gelati and Gnaccolini, 2003]). The lenticularsandstone bodies belong to different turbidite-systems that fill

local depocenters close to the western and eastern margins (inpresent-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 repre-sented by five depositional sequences (C2–C6; Figure 3b).They are mostly formed by high-density flow sandstones ofthe Cassinasco Formation and, toward the top, by hemypelagicpelites of the Murazzano Formation.

Figure 4. (a) Tectonic map of the Ligurian Alps and Tertiary Piedmont Basin with a compilation ofpublished thermochronometric data. (b) Published 40Ar/39Ar, zircon/apatite fission tracks and (U-Th)/Heages 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 fromthe 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 detrital40Ar/39Ar ages from samples collected close to the stratigraphic boundary between TPB and meta-morphic basement are identical (34–26Ma) with the Early Oligocene biostratigraphic age of theMolare Fm. [Gelati and Gnaccolini, 1998; Seno et al., 2010]. ZHe ages young away from the basinboundary (to the south).


6

2.4. Thermochronometric Data

[19] In the last 10 years, many thermochronometric databoth from the metamorphic rocks of the Ligurian Alps andfrom 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].[20] White mica 40Ar/39Ar indicates that the eclogite-

blueschist-facies (HP) metamorphism occurred between 49and 40Ma in the Voltri Massif [Federico et al., 2005, 2007].Phengite 40Ar/39Ar (32.9� 0.8Ma [Federico et al., 2005])and U-Pb sensitive high-resolution ion microprobe (SHRIMP)dating on a zircon rim (33.8� 0.8Ma [Vignaroli et al., 2010])dated the greenschist-facies metamorphic stage within theVoltri metasediments. Considering the biostratigraphic age(Early Rupelian, ~33.9–30Ma; [Gelati and Gnaccolini,1998]) of the TPB sediments overlying the sampling sites, fast(3–5mm/yr) exhumation rates have been proposed during theOligocene time [Federico et al., 2005]. Remarkably higherexhumation rates (in the order of several cm/yr) have beenproposed by Rubatto and Scambelluri [2003] based on youngradiometric ages for the baric peak of HP metamorphism(33.6� 1Ma from U/Pb dating on baddeleyite from eclogites).[21] The data from the metamorphic basement correspond

with the 40Ar/39Ar ages (~32–50Ma) from detrital micas col-lected in the basal deposits of the TPB, which derived fromthe erosion of the metamorphic units of the Ligurian Alps[Barbieri et al., 2003, Carrapa et al., 2004]. The presenceof a ~34–32Ma 40Ar/39Ar ages from the ophiolitic pebblesof the Rupelian (~33.9–28.4Ma) Molare Formation(stratigraphically overlying the Voltri Massif) supports theconcept of a fast exhuming basement. 40Ar/39Ar ages ofsediments from the present rivers draining the Ligurian andWestern Alps also indicate that the metamorphic rocks wereexhumed in rapid pulses prior to ca. 38Ma and that relativelyslow and continuous erosion occurred thereafter [Carrapaet al., 2003b].

[22] Zircon fission track (ZFT) [Vance, 1999] and zircon(U-Th)/He ages (ZHe) [Maino et al., 2012a] documentedthe time-temperature history of the Briançonnais andPrepiedmont domains of the Ligurian Alps. ZFT data fromthese domains cluster around 31Ma, but some grains fromlow metamorphic units are much older (~60–254Ma,Figure 4a; [Vance, 1999]). The old ages clearly indicate thatthe ZFT system was not reset everywhere during the Alpinephases; hence, not all the units of the Ligurian Alps experi-enced, during Alpine metamorphism, temperatures hotter than~240� 25�C (approximate temperature of closure for the ZFTsystem 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 oncooling rates and crystal size, varies from 210�C to 140�C[e. g., Reiners, 2005]. ZHe ranged from 32.2� 2.3 to25.4� 1.6Ma (Figure 4a). The ZHe data from the present-dayexposed stratigraphic boundary with the TPB (figure 4)indicate that the metamorphic rocks were at ~200�C between32.2� 2.3 and 29.1� 1.1Ma. The ZFT and ZHe ages arewithin the error of the biostratigraphic age (Early Rupelian,~33.9–28.4Ma) of the Molare Formation. These data imply aLate Rupelian average cooling rate of more than 100�C/Myrand an apparent (ultra-) fast exhumation (6.8–1.3mm/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 MolareFm. (Figure 4a) indicate a rapid cooling between 120� and60�C at ~26Ma, which is an age slightly younger than thedepositional age of sediments [Barbieri et al., 2003]. Thesedata seem to suggest a resetting of the AFT system bypost depositional thermal overprinting as a result of burialfollowing subsidence in the TPB [see Bertotti et al., 2006].Otherwise, other thermal indicators, such as vitrinite reflec-tance and thermal alteration index on palynomorphs, indicatepaleo-temperatures lower than 100�C for the Molaresediments discounting the possibility of a completely reset of

Figure 5. Geological-structural sketch of the mapped area at the orogen-basin boundary, with thedistribution of the late Alpine faults. RV: Roccavignale. Field data are collected in the framework of theCARG 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.


7

the AFT system [Barbieri et al., 2003].Moreover, in the Voltrimassif, a limited number of AFT data yielded dispersed agescomprised between 2.5� 0.6 and 23.9� 4.9Ma [Vignaroliet al., 2010]. On the whole, the AFT data set is insufficientto derive a valid interpretation of the Miocene thermal historyof the TPB. They may suggest an inhomogeneous and partialresetting of the AFT system by post-depositional thermaloverprinting as a result of Oligocene–Late Miocene burialdue 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 TPBsuccession. Most of the data indicate a general cooling of thebasin succession between 70� and 40� C at ~14–11Ma, butsome samples from the stratigraphic boundary between thebasement and the Molare Fm. yielded contrasting results (~23and ~14Ma, respectively) despite their proximity (Figure 4a).The AHe data cannot be explained unequivocally: only thepost-14Ma inversion from subsidence to basin exhumationcan be reliably interpreted.

3. Methods

3.1. Stratigraphic Analysis

[26] In the frame of the National Project for survey, the1:50000 Italian Geological Map, a preliminary geological sur-vey of the study area at the scale of 1:10000 [Seno et al., 2010]provided new data on the Tertiary turbidite systems organiza-tion (Figures 5 and 6). Their stratigraphic framework has beenrefinished on the base of km-scale correlations of their compo-nent 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 thedepocentral 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.[27] Micropaleontological samples have been collected

within key horizons traced through the basin. The calibrationof the bio-magnetostratigraphic events is according toGradstein et al. [2004].

3.2. Structural Analysis

[28] Until now, a detailed structural analyses of the TPBor the Ligurian Alps has only been performed at a local scale[e. g., Bernini and Zecca, 1990; Carrapa et al., 2003a; Pianaet al., 2006] and no attempt has been made to quantify andclassify the brittle fault network characterizing the wholeorogen-basin system. We have performed a field-baseddetailed structural analysis of the brittle structures from thesouthern part of the Langhe basin and the western sector ofthe Ligurian Alps (Figures 2, 5, and 6). These data, integratedwith the sedimentological investigation carried out at thepresent-day exposed erosional orogen-basin boundary,provide constraints on the overall evolution of the orogen-basin system. Our approach is based on the systematicmeasurement of fault populations, including fault planeorientation, striae and slickenside orientation, and shearsense, classified on the basis of the reliability criterion ofthe shear sense. Structural analyses in the basement wereperformed in 17 well-exposed sites (structural stations, here-after SS; Table 2).[29] The structural data have been computed in order to

obtain paleostrain analysis [Malusà et al., 2009]. Detailson the method applied on both the field-based analysisand the data computing are provided in sections A1 andA2, 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 deforma-tion and stratigraphy. Dating brittle faulting in the metamorphicrocks is more difficult but can be extrapolated by integratingthe rock-fault type analysis with the thermochronometric data[Malusà et al., 2009], exploiting the fact that the thermal struc-ture of the crust exerts a primary control on the fault-rockstypes which are produced during deformation [e. g., Sibson,1977; Scholz, 2002]. The change from plastic to cataclasticdeformation mechanisms is temperature dependent; in silicaterocks, it is referred to a large temperature range between~450�C and ~120�C, depending on strain, mineralogy, struc-tural inheritance, and fluid flows [e.g., Kohlstedt et al., 1995;Scholz, 2002]. The very complex tectonic processes, thelithological variations, and the extremely various strain ratesactive during the multiphase orogenic evolution of the Alps

Figure 6. Geological-structural map of the Mioglia area (seeFigure 2) with the distribution of the late Alpine faults (modi-fied from [Bernini and Zecca, 1990]). Legend as in Figure 5.

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

N� of Structural Station Lithology Longitude Latitude

1 limestones 7�4005400 44�10032002 limestones/dolomites 7�4603200 44�09050003 rhyolites 7�4303100 44�12053004 limestones 7�4802500 44�06054005 quartzites 7�4505600 44�11036006 limestones/dolomites 7�4902200 44�13023007 limestones 7�5501300 44�11038008 orthogneisses/schists 7�5405400 44�15029009 rhyolites 7�5000500 44�100080010 quartzites 8�1005500 44�120110011 dolomites 8�0104900 44�170870012 orthogneisses 8�1002400 44�210310013 dolomites 8�0501700 44�150230014 dolomites 8�1503900 44�150420015 limestones 8�0401500 44�050480016 quartzites 8�0000400 44�100190017 limestones 8�2504100 44�1103200


8

do not allow to extrapolate from a single fault the age for theoccurrence of brittle (elastico-frictional; [Sibson, 1977]) defor-mation in all the belt. Thermochronometric dating on multi-locations, however, provides the timing when the study regioncooled through a given temperature. In particular, the zirconfission track and zircon (U-Th)/He techniques with their clo-sure temperature (Tc) of ~240� 25�C [Brandon et al., 1998;Bernet, 2009] and 140–210�C [Reiners, 2005] have the poten-tial to date the post-metamorphic tectonic deformations char-acterized by purely cataclastic faulting. Therefore, the ZFTand the ZHe ages from the metamorphic units of theLigurian Alps, coupled with 40Ar/39Ar, AHe, and AFT data,have been used to constrain the activity of the faulting phasesin the basement.

4. Sedimentology and Stratigraphy

[31] Geometry, stratigraphic relationships, facies distribu-tion, and palaeocurrent directions indicate that the sandstonebodies of the Molare Formation (depositional sequence A)and of the Rocchetta-Monesiglio Formation (depositionalsequences B1–B6) were deposited in a basin controlled byvery active tectonics [Gelati and Gnaccolini, 2003]. Thistectonic activity is documented by the regional stratigraphicarchitecture that record migrations of the source areas anddepocenters with time, and by the complex lateral and verti-cal distribution of terrigenous depositional systems mainlycontrolled by structurally-induced submarine topography.These aspects are described in the following paragraphs.

4.1. Rupelian (~34–28Ma)

[32] At this time, the basement was being eroded and partof it was covered by clastic sediments of the MolareFormation (Sequence A).[33] The unconformable and transgressive Molare Fm.

consists of flood-dominated alluvial fan and fan-deltasystems mainly composed of conglomerate and pebbly-sandstone facies. Conglomerate clasts range from centimeterto meter in size, and their lithology reflects the compositionof the neighboring-outcropping pre-Cenozoic substratum(Figure 7a). Sandstone is mainly composed of abundantquartz, quartz-mica bearing lithic fragments, and carbonategrains. QFL +CE (Q = quartz, F = feldspars, L + CE = totalfine-grained lithic fragments plus carbonate rock fragments)analysis [Gelati and Gnaccolini, 2003] on sandstone samplednear Millesimo revealed a high content on lithic fragments(L� 70–90%) mainly derived from quartzite, migmatite,orthogneiss, micaschist, phyllades, and mafics. Carbonate frag-ments are also very frequent (30–40% on total QFL+CE). Thusagain, the very heterogeneous pre-Cenozoic substratum is itselfthe main source for the Molare Fm. Most of these sedimentswere deposited through a debris flow mechanism; they showsubstantial lateral variation in both thickness and facies,reflecting a deposition on a surface of pronounced relief. Theflood-dominated alluvial fan and fan-delta deposits are grada-tionally replaced by highly burrowed shallow marine fossilif-erous sandstones and pebbly sandstones. Because of theintense faulting activity during sedimentation, thickness,composition, and paleocurrent directions are laterally highlyvariable (Figure 3b). The Molare Formation has been gener-ally ascribed to the Lower Oligocene (Rupelian, e. g., Gelatiand Gnaccolini [1998]). Nevertheless, in the study area, the

passage with the overlying Rocchetta-Monesiglio Formationis diachronous: the pelites overlying theMolare sandstones con-tains planktonic foraminifera belonging to theParagloborotaliaopima opima biozone (probably to the upper part: IFP21b)toward the east and to the Globigerina ciperoensis (IFP22) inthe western sector [Seno et al., 2010].

4.2. Early Chattian (~28–26Ma)

[34] At the beginning of the Late Oligocene, the emplace-ment of widespread purely marine sediments testifies a gen-eralized deepening of the basin, associated with a change inthe provenance area and in the denudation rates. In the lowerpart of the Rocchetta-Monesiglio Formation (depositionalsequence B1), the submarine topography is controlled byNW-oriented, high-angle, basement-involving normal faultsand flexures. The sedimentation was strongly confined insmall, isolated fault-bounded basins, which were infilled byrelatively small and lenticular submarine fan systems. Thecompositional analysis on sandstone of the B1 sequence sed-imentary bodies shows different results for the northeasternsector of the basin and the southwestern one [Gelati andGnaccolini, 2003]. While toward the north, a sudden enrich-ment in lithic fragments and especially in serpentinite clastshas been recorded (QFL+CE: 40%<L< 60%; serpentinite +metabasites fragments between 27% and 50% over thetotal); in the southern sector, lithics are less abundant(17%<L< 35%) and mafic clasts are rarer (serpentiniteclasts between 2% and 10%). These compositional analysesaccounted for different source areas in the northern sectorof the basin (the N–NW Asti-Cuneo area) and in the south-ern one (the S–SW Brianconnais basement and cover plusthe Molare Fm.).[35] In the southeastern area, the Millesimo body (ascribed

to the Early Chattian [Gelati and Gnaccolini, 1998; Senoet al., 2010]) is a typical example of this phase of deposition(Figures 5 and 8a–8c). It is mainly composed of channel-fillconglomeratic sandstones and conglomerates that are sup-posed to be the infill of a submarine depression surroundedby topographic highs with steep, fault-controlled slopes.The fault-bounded depositional setting, the textural coarse-ness, and facies association suggest that the clasts of theMillesimo body are reworked material from the contempora-neous alluvial fan and fan delta systems, which were proba-bly developing on adjacent structural highs. These alluvialsystems were fed by very young, high-gradient streams,which where draining small, high-relief and structurallyfragmented areas. As a result, these alluvial and nearshoresystems were small in size and mainly composed of verycoarse-grained clastics.[36] Because of the rapid deepening of the Molare-Rocchetta

basin, the transgression toward the southwest did notdevelop a shelf area. Large volumes of the Molare alluvialfan and fan-delta coarse clastics underwent failures on thesteep slopes bounding the highs and were resedimented inthe adjacent submarine depression through gravity flow.These relatively small-volume flows were mostly loadedwith coarse clastics. Initially, they start as debris flows,evolving downcurrent to highly concentrated turbiditycurrents. These flows deposited thick-bedded conglomer-atic sandstones and conglomerates that display extensivechanneling with the absence of lobe deposits. Slump faciesare locally present.


9

4.3. Middle–Late Chattian (~26–23Ma)

[37] The Middle Chattian was characterized by a funda-mental change in the tectono-sedimentary features. Thegeneral depositional environment might have been similarto the delta-fed submarine ramp and lobe system or to thedeep-water slope-type fan delta. The structural control onthe submarine topography of the basin became subtler andcan be detected only through detailed stratigraphical andsedimentological analysis. The depositional environmentwas characterized by a markedly flatter geometry than forthe older units (e. g., Millesimo body), and volumes ofsediments are considerably larger. The Retano, Cengio,Castelnuovo-Bric la Croce, and Noceto turbidites (ascribedto 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 ofrelative quiescence of intrabasinal tectonics. The systems were

fed by the same source area and were accommodated into aprogressively westward-widening depositional area, underthe confining effect of marginal slopes (Figure7b–7d). Thesefault-controlled slopes are mainly NE–SW to ENE–WSWoriented (Figure 8e).[38] The sandstone composition still markedly differ from

the north and the south sector during the deposition of theB2 and B3 sequences, being enriched in mafic clast towardthe north and in extrabasinal carbonate clasts toward thesouth (initially with a low percentage of quartz +mica frag-ments). Lithic fragments (quartz +mica) occurrence rise upduring the B4 sequence deposition in the south western partof the basin; this trend gradually expanded toward the northduring the successive depositional sequences (B5 and B6)with a progressive decrease in serpentinite clasts. Thisaccounts for a progressive homogenization of sedimentsduring the later stages of deposition and for an enhanced influ-ence of recycling processes [Gelati and Gnaccolini, 2003].

Figure 7. (a) Conglomerate clasts of theMillesimo body (Early Chattian) ranging from centimeter to meter insize; 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–LateChattian) onto lateral and frontal confining slopes (he: hemipelagic marlstone); these turbidite systems weredeveloped during stages of relative quiescence of intrabasinal tectonics and were accommodated into a progres-sively westward-widening depositional area under the confining effect of marginal slopes.


10

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

[39] This time interval records the development of gentlefolds that create structural high and relatively low areas.These folds (with axes ~N–S directed) affected hemipelagicmudstones and sandstone bodies belonging to the upperpart of the Rocchetta-Monesiglio Formation (depositionalsequence B6 [Gelati and Gnaccolini, 1998]). In the Ceva area(Figure 5), a wide anticline of pelites is bounded at its top byan extensive slump scar, overlaid by siliceous sediments.Pelitic slump-sheets, derived at least in part from submarinedenudation of the sediment originally present on top of theanticline, now flank the structure (Figure 8d). In this time inter-val, the sedimentation is represented by siliceous sediments, acondensed succession typically composed of well-lithifiedmudstone and crypto- and micro-crystalline quartz-rich silt-stone, which rhythmically alternate with mm- to cm-thick fine

to very fine sandstones, with local faint ripple laminations(depositional sequence B6). This condensed succession showssudden thickness variations; it is 10–50m thick on average(up to 150m thick) and invariably separate pre-Mioceneturbiditic sandstone bodies from overlying turbidite systems,mostly deposited during Burdigalian [Gelati et al., 2010;Seno et al., 2010]). The lower boundary of the siliceous zoneis almost transitional. The upper boundary, which is markedby abrupt facies changes, displays clear unconformable con-tacts that probably correspond to the Burdigalian unconfor-mity [Mutti et al., 1995].

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

[40] The post-Early Burdigalian sediments (depositionalsequence B6-upper lithozone and C1–C6 [Gelati andGnaccolini, 1998; Gelati et al., 2010]) of the Langhe basin

Figure 8. (a) Decametric channelized unit of Millesimo body (ml; Early Chattian ~28–26Ma—SequenceB1, 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 thesouthern TPB slope. (c) Metric dislocations in the Millesimo unit by the F1 phase faults (Early Chattian~28–26Ma). (d) Synsedimentary pelitic slump-sheets along the gentle ramps generated by the open folding(Aquitanian–Early Burdigalian, ~23–18Ma). The contractional structures detected in the study areasuggest 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 systematicallydislocate the previous F1 structures. These faults are sealed by the Aquitanian siliceous lithozone (sl).


11

crop out mainly north to the study area, and they were notanalyzed in detail (Figures 2 and 5). The sedimentation wasachieved in a highly subsiding asymmetric trough (E–Woriented) that was infilled with a very thick basin-wide succes-sion of westerly-derived turbiditic sandstones and pelites(Cortemilia and Cassinasco formations, depositional sequencesC1–C6; Figure 3b). In the western sector of the Langhe basin,the Late Burdigalian-Serravallian succession is characterizedby a sand-rich wedge with a thickness in excess of 2000mthat markedly tapers toward SE. The analysis on sandstonesbelonging to C1–C6 sequences [Gelati and Gnaccolini, 2003]shows similar composition to those of the sedimentary bodiesof the B6 sequence, being enriched in mica +quartz lithics,extrabasinal carbonate grains, and progressively depleted inserpentinite clasts. This accounts for a progressive shift of thesource 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; themost important exposed structure is the Villalvernia-Varziline (Figures 2 and 3a), which during the Oligocene–EarlyMiocene acted as a sinistral strike-slip fault dividing theBorbera-Curone basin from the Apennines [e. g., Di Giulioand Galbiati, 1995; Felletti, 2002]. Minor faults are abun-dant and show two prominent directions: NNW–SSE andENE–WSW (Figure 9a). These last are the most abundantwithin all the stratigraphic succession (Figure 9a) but, con-sidering only the sequence A–B1, the NNW-striking faultsare predominant (Figure 9b). Each mapped fault is character-ized by a centimeters- to meters-wide cataclastic zone(Figure 10a), mostly constituted by fault breccias, whereasgouges are predominant in the fine-grained deposits such

as marls or pelites. Faults are frequently associated withsynsedimentary flexural folds, with axes parallel to the direc-tion of the faults. Locally, this folding develops an inter-stratum extensional shear zone characterized by geneticallyrelated small-scale structures.[42] Geological survey of the area evidenced that the NNW-

striking faults developed only in the Late Rupelian–EarlyChattian depositional sequences A and B1 and are sealed bythe Late Chattian deposits (Figures 5 and 6). On the contrary,the NE- or ENE-striking faults are present within all theOligocene sedimentary succession (depositional sequenceA–B5) and are sealed by the Aquitanian (B6) deposits(Figures 5, 6, and 8e). These field relationships demonstratethat the NNW-striking faults are older than the ENE-strikingones, indicating the existence of two chronologically distinctfaulting 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% ofthe NNW-striking faults are normal (84% with less than30� between fault dip and lineation plunge), with obliquemotion (56% left-lateral and 44% right-lateral; pitch com-prised between 85� and 55�; Figure 11a). These faults showthrows ranging from 0.5 to 40m. Eastern to the study area,N–S to NW–SE striking normal faults affecting the MolareFm. (Sequence A) have been also recognized along the bound-ary with the Voltri Massif [Vignaroli et al., 2009][44] In the Roccavignale area (Figure 5), few high-angle

NNW-striking faults, related to a conjugate system, showreverse activity (Fi; Figure 11a). Moving toward these faults,the bedding of the deposits increases the dip from 10–20� to40–50�. Within the gage layers, typical fabric associationsare represented by conjugate Riedel (R, R’, P). These planesare pervasive and compose of duplex arrays with inverse

Figure 9. Stereonets of the faults mapped in the Tertiary Piedmont Basin and Ligurian Alps. F1 and F2indicate two distinct faulting events. F1 mapped within: (a) the whole of the TPB sedimentary successionconsidered in this work; (b) only in the depositional sequences A–B1. The faults trend preferentially NNWand ENE. In the sequence A–B1, the NNW-striking faults are predominant. (c, d) Direction of the normalF2 faults, mapped within all the basin, divided on the basis of their strike-slip component. The right-lateralF2 faults trend preferentially NE, while the left-lateral F2 faults are mainly ENE-oriented. (e) all faultsmapped in the metamorphic basement; (f) only the basement faults where the kinematics has beenconstrained with a good degree of confidence. (g, h) Direction of the normal F2 faults from the basementdivided on the basis of their strike-slip component.


12

shear sense. Overall, the geometry of the shear bands sug-gests that these faults developed in a transpression context,which involved oblique dextral simple shear and contractioncomponents. However, the apparent bed displacement asso-ciated with these faults displays a normal sense of movement.This misfit suggests that these contraction structures wereborn as normal faults than experienced tectonic inversion.The geological survey of the area shows that also the reversefaults 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 Chattiansediments of the Molare Formation and the lower part ofthe Rocchetta-Monesiglio Formation (depositional sequenceA–B1) are involved into NNW-trending folds (e. g,. theMioglia flexure) and thrusts [Bernini and Zecca, 1990]. Theorientation of the horses of recognized duplex structures indi-cates a main NNE–SSW shortening with a rather important

dextral component (Figure 11a). We interpret the righ-lateral-reverse faults as reworked inherited normal faultsrelated to the F1 tectonic phase, as suggested by Mutti et al.[1995]. Also, these structures are dislocated by ENE-strikingnormal faults (Figure 6) involving the complete Oligocenesuccession (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 theAquitanian deposits (deposition sequence B6; Figures 8e).They systematically truncate the NNW-striking faults anddisplay throws varying from centimeters to tens of meters.Accommodation folds are frequent, suggesting synsedimentaryfaulting. The ENE-striking faults show a prevailing extensionalkinematics (85% with pitch >45�) associated with an impor-tant strike-slip component (15% of faults have pitch between45� and 20�; 48% between 60� and 45�; 18% between 60and 70�) (Figure 11a). On the basis of their transcurrent

Figure 10. (a) Exposure of fault breccia composed of cm-thick angular fragments between the conglom-erates of the Molare Formation and orthogneisses. (b) Decametric fault damage zone composed of brecciasand pervasively fractured quartzites. The sketch shows the cross-cutting relationships between the olderNNW-striking F1 faults and the younger NE to NNE F2 faults. (c) NNW fault system (F1): conjugate nor-mal 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 tothe ductile phases (D1 and D2). (h) Tectonic breccia in orthogneiss. (i, l) Thin sections from outcropshowed in Figure 10h, viewed in transmitted light. Cataclastic process controlled the grain sized reductionof quartz (Q) and feldspars (F) with different percentage of cement precipitation (quartz). (m) Exposure ofthe tectonic breccia with cm-thick angular fragments within orthogneisses.


13

Figure 11. Equal-area projections (lower hemisphere) showing fault-slip data for (a) the faults mapped in theTPB (Figures 5 and 6), grouped on the basis of the depositional sequence that they cut and (b) from the structuralstations (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 reversefaults 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 (whitesquares) and the metamorphic basement (black squares). In Figure11b, the structural stations where two familiesof faults have been distinguished on the basis of cross-cutting relationships are divided in two diagrams (A andB).Only the structural stations, where the fault measurements have been carried out with a high quality of confidencein the inferred sense of movement, are reported. White squares indicate poles to the extensional veins.


14

component, these faults may be divided in two sets: predomi-nant (~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 thefield 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 bythe strain analysis: the compression (P) axes are nearly verti-cal in the whole of the study area, and the tension (T) axes aresub-horizontal or gently dipping (Figures 12a). The T axes ofthe F1 faults show a well-defined ENE–WSW incrementalextension. The distribution of the T axes of the F2 faultspoints to a main NNW–SSE 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–SSWshortening active before the F2 phase (Figure 12a).

5.2. Post-Metamorphic Fault Network of theLigurian Alps

[49] Major brittle structures are mainly located on the west-ern and eastern boundary of the Ligurian orogen, whereasthey are rare in the central part of the chain; the most impor-tant of these structures are the Stura shear zone to the westand the Sestri-Voltaggio Fault to the east (Figure 2). TheWNW-striking Stura shear zone comprises several relatedfaults, as the Stura and Preit lines, and has been acting as asinistral strike-slip fault, accumulating about 40–50 km ofinferred displacement, since the Oligocene [Ricou, 1981;Giglia et al., 1996]. To the east, the Sestri-Voltaggio Faultrepresents a km-scale N–S discontinuity that played different

Figure 12. Average kinematic solutions of the fault-slip data (Figure 11) from the (a) TPB and (b) theLigurian 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 (short-ening) and dark gray (extension).


15

roles in different stages of the tectonic evolution of thearea [Crispini and Capponi, 2001; Vignaroli et al., 2009].Generated as a mylonitic shear zone during the metamor-phic deformational phases (D1–D3, Table 1), it records anOligo–Miocene brittle dextral strike-slip kinematics [Capponiet al., 2009; Federico et al., 2009].[50] At the outcrop scale, different assemblages of fault rocks

with heterogeneous distribution and well-defined geometricrelationships were recognized in the study area: (i) myloniticrocks, (ii) cataclasites, and (iii) fault breccia and gouges. Thedeformational styles of the fault rocks can be correlatedwith the occurrence of the Alpine tectonic phases in differentcrustal conditions.[51] Anastomosing mylonites are the oldest fault rocks; they

are associatedwith deformations occurred under viscous regime[Schmid and Handy, 1991] associated with the development ofblueschist- or greenschist-facies metamorphic recrystallizationin the Briançonnais units (D1 and D2 phases, Table 1).[52] Cohesive, fine-grained cataclasites are characterized by

the development of quartz-chlorite-epidote-phyllosilicate slick-ensides in silicate rocks related to low greenschist facies meta-morphic recrystallization (D3 phase). These latter fault-rocksare important indicators of the transition from plastic flow tocataclastic faulting [Rutter, 1986; Schmid and Handy, 1991].Array of planar and sigmoidal tension gashes filled with quartzare often found close to these fault zone. Cohesive cataclasitespreserved in outcrop are quite rare, as they are often overprintedby later structures.[53] Cataclastic breccia (mainly incohesive) and gouges are

found in multiple anastomosing layers of normal faults (D4and D5 deformation phases; Table 1); they are characterized

by the absence of metamorphic recrystallization and are gener-ally associated with closely-spaced joint systems, especiallywithin gneiss, quartzites, and dolomites (Figures 10b and10h). Arrays of calcite- or quartz-filled veins are commonlyfound in limestones and schists (Figures 10f and 10g). Faultzone 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(>2mm) derived from grain size reduction with variablepercentage of cement precipitation (Figures 10i and 10l).Following Woodcock and Mort [2008], they can be classifiedas chaotic to crackle breccias. Fault gouges, composedof 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 incalcschists and quartz-schists. The dominant extensionalcharacter of these faults is testified by displaced beds andmicro-structures, such as striae, slickensides, and en échelontension gashes (planar or sigmoidal). Less frequent strike-slipfaults (mainly sinistral) are also present (Figures 10d and 10e).The basement faults show two main directions, similar to thosefound 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 kinemat-ically constrained faults (n= 326, Figure 9f), which are hereafterused for the computation.[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

Figure 13. (a) Exposure of quartzites showing a NNW-striking set of normal faults (F1) dissected by anENE-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 mainfault underlined with solid (F1) or dotted (F2) black lines.


16

are associated with chronologically different faulting epi-sodes, although in a minor number of structural stations(SS 2, 3, 8, 9, 13, 16) the relative chronology is difficult tosolve. The uncertainty is probably due to the heterogeneity ofpoly-deformed rocks as well as to the existence of preexistingdiscontinuities. 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 byNE 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 theseF2 faults have pitch<60� indicating a strong strike-slip com-ponent; ~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 becausetheir temporal relationship with the other faults could notbe determined.[58] The P and T axes distribution of the populations F1

and F2 shows a prevailing extension (Figure 12b). On thewhole, and according to the fault analysis performed in thebasin, the strain relationships associated with the two faultingphases in each structural station where two families havebeen distinguished display a main E–Wextension for the firstfaulting phase, changing into NW–SE direction during theF2 phase.

5.3. Folding

[59] The entire Ligurian Alps-TPB system is affected byup to 10 km long wavelength parallel open folds (D5 struc-tures; [Seno et al., 2005a, 2005b]). Parallel open folds withsub-vertical axial planes and ~N–S directed sub-horizontalaxes also affect the Aquitanian–Early Burdigalian depositsof the TPB (Figure 11a). D5 folds generated asymmetricdome and basin interference patterns in the Ligurian base-ment [Bonini et al., 2010].

6. Discussion

6.1. Mechanisms of Basin Formation and Evolution

[60] Because of the overprinting and reactivation of exten-sional and compressive structures, the stress/strain regime ispoorly resolved, and the age of the deformations in the TPBis poorly constrained. Extension has been related to a firstperiod of subsidence caused by the Oligocene opening ofthe Liguro-Provençal basin [Mutti et al., 1995; Gelati andGnaccolini, 1998]. The same authors suggest a LateOligocene–Early Miocene inversion from an extensional toa compressional stress field, related to the Corsica-Sardiniadrifting 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 dominatedby compression or transpression from the Oligocene untilpost-Pliocene time, with extension playing a minor role.[62] The role played by the extensional regime in the evo-

lution of the TPB is, however, controversial. On the basisof Anisotropy of magnetic susceptibility (AMS), Maffioneet al. [2008] propose that N–S synsedimentary extension

controlled the formation of the TPB, which acted as a basinpassively carried on top of displacing nappes. The compres-sive structures found within the basin are here interpretedas gravitational slumps or post-Tortonian in age. However,other AMS analyses have been interpreted as suggesting aNE–SW to NW–SE shortening, which was active since theOligocene [Carrapa et al. 2003a].[63] Our approach integrates structural and stratigraphic

analysis on the TPB-Ligurian Alps boundary and suggeststhat during the Oligo–Miocene the orogen-basin systemevolved through three distinct tectonic phases.6.1.1. Rupelian–Early Chattian (~34–26Ma)[64] Thermochronometric data indicate that, during the

Rupelian–Early Chattian times, the metamorphic rocks ofthe Ligurian Alps were rapidly exhumed and exposed[Barbieri et al., 2003; Maino et al., 2012a]. At this time,the exhumation of the Briançonnais-Prepiedmont domainwas controlled by a NNW-striking fault-system, related tothe first brittle deformation phase (F1), which took placeduring deposition of the depositional sequences A–B1(Figures 14a and 14b). This is the first brittle deformationphase (D4), and it overprints the older structures related tothe metamorphic recrystallization (D1–D3 phases). Strainanalyses of the F1 structures are compatible with a mainENE–WSW stretching (Figures 12). Similar E–W trendingmaximum extension direction has been derived from thecoeval brittle structures recognized in the Voltri massif[Vignaroli et al., 2009].[65] During the Rupelian, the deposition of the Molare Fm.

coarse-grained clastics marks the beginning of the marinetransgression that migrated from NE to SW. The subsidence-driven transgression between the alluvial and the near-shoredeposits of the Molare Formation and the overlying EarlyChattian mudstone of the Rocchetta Monesiglio Fm. marks ageneral deepening of the TPB (Figures 14a and 14b). Thistransgression is younging toward south and west, from IFP20 biostratigraphic zone (~33–31Ma) in the north-easternTPB, IFP 21b (~28–27Ma) in the study area, to IFP 22(~26–24Ma) zone in the westernmost part of the basin[Gelati and Gnaccolini, 1998; Seno et al., 2010]. Suchdiachronous transgression is interpreted as the result of aLate Rupelian–Early Chattian westwardmigrating subsidence.Subsidence analyses from different localities of the TPB indi-cate that the vertical movements began in the Early Oligoceneand continued throughout the Miocene [Carrapa et al.,2003b]. The subsidence curves indicate the first (LateRupelian–Early Chattian) moderate pulse <1 km that preva-lently acted in the southernmost part of the basin (Ceva-Cairo Montenotte; Figure 5). This early phase of subsidenceis closely associated with the development of the extensionalF1 structures in the basin. The F1 faults are characterized bymodest length and throw ranging from centimeters to severaltens of meters. The overall extension accommodated by thesefaults is of the order of several hundred meters, in agreementwith the moderate subsidence rate calculated for this period[Carrapa et al., 2003a]. The orientation and structural charac-teristics of the F1 faults, therefore, support the interpretationthat subsidence in the TPB during the Rupelian–EarlyChattian 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 detectedin the lower stratigraphic succession (depositional sequence


17

Figure 14. Simplified palaeogeographic models and basin evolution for the study area. Mi: Millesimo;Ce: Ceva. (a) Rupelian (~34–28Ma—Sequence A): deposition of the Molare Formation. Depositionalenvironments: flood-dominated alluvial-fan and fan-delta systems (1) and coastal plain (2). Activation ofNNW-directed extensional faults (tectonic phase F1). (b) Early Chattian (~28–26Ma—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-dominatedalluvial-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(~26–23Ma—Sequence B2–B5): deposition of the Retano, Cengio, Castelnuovo-Bric la Croce andNoceto turbidite systems (Rocchetta Monesiglio Fm.), and of the Cima della Costa Unit. Depositionalenvironments: delta-fed submarine ramp system (8) and turbiditic lobes (9). Activation of NE to ENE-directed extensional faults (F2). (d) Aquitanian–Early Burdigalian (~23–18Ma—Sequence B6):Condensed sedimentation of siliceous marls alternating with mm- to cm-thick fine to very fine sand-stones (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, whichrepresents the transition to the more conspicuous turbiditic sequence C.


18

A–B1). The Fi faults have been interpreted as originated fromthe (Late Chattian?) inversion of early (Rupelian) extensionalstructures as already proposed by [Mutti et al., 1995]6.1.2. Late Chattian (~26–23Ma)[67] In the Late Chattian, the deposition of the B2–B5

sequence marks the formation of the Langhe basin [Gelatiand Gnaccolini, 1998]. The depocenters of the turbiditesystems progressively shifted westward. Volume and faciessuggest that these systems were fed by clastics source areas(with similar composition) which were considerably richerin sand than those which had supplied the coarse-grainedclastics of the Millesimo body lower in the succession(Figure 14c). Such conditions were probably reached inperiods during which subsidence was less pronounced.[68] The turbidite systems were affected by ENE–WSW

to NE–SW synsedimentary faulting (F2). The F2 faults sys-tematically dislocate the previous F1 and Fi structures(Figure 14c). A related NW–SE extension is suggestednot only by the strain analysis (Figure 12) but also by theprogressive westward migration of the depocenters.[69] In map view (Figures 13b–13d), the F2 faults show

orientations comparable with the model of shear fractureorientation in non-coaxial deformation: specifically, thesestructures comprises purely extensional faults (mainly NE-oriented), synthetic Riedel fractures (R) with a pronouncedENE–WSW orientation (Figures 9d and 9h), and conjugateantithetical NNE-oriented Riedel fractures (R’) (Figures 9cand 9g). This geometry is compatible with a regional sinistralstrike-slip shear zone.[70] The Fi faults recognized in the Sequence A–B1 are

coherent with this tectonic scenario. They may be interpretedas localized restraining areas within the regional LateChattian left-lateral transtensive zone. This consideration issupported by the deep basement/basin geometry illustratedby seismic lines [Mosca et al., 2010]. Indeed, in the Cuneoarea, localized transpressive structures, involving both thebasement and the Oligocene strata of the TPB, were detected:strike-slip faults with flower geometry influenced the deposi-tion of the Oligocene sub-basins and are sealed by Miocenesediments. Strike-slip tectonics is furthermore testified inthe Voltri massif, where the Sestri-Voltaggio fault systemshows dextral kinematics since the Chattian [Capponiet al., 2009; Federico et al., 2009].6.1.3. Aquitanian–Serravallian (~23–12Ma)[71] The active extensional and transtensional faulting

documented in the Oligocene deposits dies away at thebeginning of the Early Miocene. During the Aquitanian–EarlyBurdigalian, open folds developed both in the basin andin the metamorphic basement (D5 phase; Figure 11a). Thesiliceous lithozone (sequence B6—lower lithozone), char-acterized by a reduced sedimentation rate, reached itsmaximum thickness into the gentle depression of the syn-clines (Figure 14d). These contractional structures suggesta general NE–SW main shortening, which is also testifiedby the Aquitanian E–NE verging thrust of metamorphicbasement onto Oligocene deposits, cropping out in theAlto Monferrato area (Grognardo thrust zone, Figure 2;[Piana et al., 2006]). Transpressional tectonics is further-more documented in the Voltri massif, particularly forthe Sestri-Voltaggio fault system [Capponi and Crispini,2002, Capponi et al., 2009; Federico et al., 2009], andin the northern sector of the TPB [Mosca et al., 2010].

[72] The successive Late Burdigalian sedimentation of theupper lithozone of the B6 sequence was achieved in a highlysubsiding trough SE–NW oriented, corresponding to theasymmetric troughs derived from the Aquitanian–EarlyBurdigalian folding (Figure 14e).[73] The post-Early Burdigalian succession (depositional

sequences C1–C6) is characterized by the presence of E–Wto NE–SW-oriented extensional faults [Carrapa et al.,2003a; Gelati et al., 2010]. These faults are widespread butare systematically associated with small displacements.Although extensional tectonics is everywhere diffused,synsedimentary NW-trending contractional structures devel-oped within Langhian-Serravallian deposits. The bestexposed examples are the Ciglie anticline, unconformablyoverlapped by Mid-Langhian sediments, and the BossolaPass structures characterized by reverse faults and foldsinvolving Middle Miocene formations [Carrapa et al.,2003a]. On the whole, the extensional and contractionalstructures indicate ~NW–SE extension coeval with NE–SW shortening.[74] During the Late Burdigalian–Serravallian, the subsi-

dence accelerated over the entire TPB (vertical movements>3 km; [Carrapa et al., 2003a]). The calculated subsidencerates indicate a short period (17.5–15.5Ma) of main subsi-dence (with rates ≥1mm/yr) followed by a considerableslowdown of the rates (≤0,5mm/yr). The low displacementassociated with the detected normal faults and the absenceof major structures clearly indicates that extension alonecannot justify the strong Late Burdigalian–Serravalliansubsidence. In absence of important stretching, the MiddleMiocene high values of subsidence may be explained by flex-ural response to thrust loading [Carrapa et al., 2003a;Carrapa and Garcia Castellanos, 2005]. Subsidence of thebasin was mainly achieved under an overall transpressionalstress regime associated with the convergence of the multi-vergent thrust-systems Apennines and Southern Alps con-verging under the TPB [Mosca et al., 2010].[75] From the LateMiocene, the TPB succession experienced

NW–SE directed compression, producing synsedimentaryfolding structures [Carrapa et al., 2003a]. Plio–Pleistoceneuplift is responsible for the present day TPB morphologyand elevation [e. g., Lorenz, 1984], characterized by gen-tle hills up to 800m. During this stage, the inversionfrom subsidence to exhumation occurred, as testified bythemochronometric data. Indeed, despite some inconsis-tency, apatite (U-Th)/He (AHe) analyses of Bertotti et al.[2006] from the Molare and Rocchetta-Monesiglio formationsrecord cooling between 70� and 40� C at 11–14Ma, whichmay be interpreted as the northward migrating basin exhuma-tion through the time [Bertotti and Mosca, 2008]. Thesedata are in agreement with the sedimentary record of theLanghian–Messinian succession that defines an overall regres-sion [Gelati and Gnaccolini, 2003].

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

[76] Given its peculiar position at the Alps-Apenninesjunction, the TPB basin provides unique constraints forunderstanding the Late Neogene geodynamic setting of thecentral Mediterranean region.


19

6.2.1. Pre-Oligocene[77] The Neogene evolution of the TPB basin is influenced

by the pre-Oligocene tectonic scenario resulted from the earlyAlpine orogenic phases. During the Paleocene–Eocene, a widerange of units from both the Piedmont-Ligurian oceanic domainand the European continental margin (e. g., Briançonnais)were gradually involved in the subduction channel betweenAdria and Europe plates [e. g., Schmid et al., 1996; Stampfliand Marchant, 1997; Ford et al., 2006; Dumont et al., 2011].Geochronological and metamorphic data from Western andCentral Alps indicate that high-pressure metamorphism propa-gated from the internal (SE) orogenic zones (the oceanicdomain) toward the outer (NW) ones (the European basement)until ca. 35Ma [Schmid et al., 1996; Rosenbaum and Lister,2005; Berger and Bousquet, 2008; Bousquet et al., 2008 andreferences within]. Also the rocks of the Ligurian Alps recordthe forelandward shift of the subduction zone during progres-sive accretion of the overriding plate: the Voltri oceanic unitsexperienced eclogite-bluschist facies metamorphism between49 and 40Ma [Federico et al., 2005]. As indicated by theLutetian age (~48–40Ma) of the syn-orogenic sediments[Vanossi et al., 1986; Cabella et al., 1991; Dallagiovanna,1995], the blueschists Ligurian Briançonnais units wereinvolved 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; Merizziand Seno, 1991]) and with the deposition of syn-orogenicsediments into the European foreland basin related to theflexural loading by the Adriatic wedge [Ford et al., 2006].These deposits record a diachronous marine transgressionthat migrated toward W–NW during the Middle to LateEocene [Ford et al., 1999]. The early Alpine deformation ismainly northwestward directed [Malavieille et al., 1984;Platt et al., 1989b; Dumont et al., 2011] and has beenrearranged by rotation during the Neogene bending of thearc [Collombet et al., 2002; Maffione et al., 2008]. Also inthe Ligurian Alps, a NW to W-directed tectonic transportcan 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 Oligoceneand are coherent with the Eocene SE–NW directed relative mo-tion of Adria and Europe reconstructed by palaeomagnetic andstructural investigations [Figure 15a; Schmid and Kissling,2000; Handy et al., 2010]. The western lateral termination ofthe Adria plate can be fixed in correspondence of the Ligurian

Figure 15. Geodynamic reconstruction of the central-western Mediterranean region during the (a) LateEocene, (b) Late Rupelian, (c) Late Chattian, and (d) Langhian times. Modified from Jolivet andFaccenna [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.


20

Alps as the Eocene foreland flexural basin did not extend west-ward of the SE France [Ford et al., 2006;Dumont et al., 2011].[79] Southern of the Ligurian area, the NW dipping

Apenninic subduction zone extended for more than 1500 kmuntil southern Iberia [Jolivet and Faccenna, 2000; Faccennaet al., 2004]. It consumed the Neotethyan Ocean eastwardthe Corsica, Sardinia, and Calabria blocks. Following thereconstruction of Lacombe and Jolivet [2005], Vignaroliet al. [2008], and Dumont et al. [2011], a major N–S sinistraltransform boundary connecting the E-dipping Alpine and theNW-dipping Apennines subductions can be speculated eastof Liguria and Corsica (Figure 15a).6.2.2. Rupelian–Early Chattian (~34–26Ma)[80] At around 35Ma, the motion of the Adria plate

changed from NW-ward to WNW-ward (Figure 15b;[Schmid and Kissling, 2000; Handy et al., 2010]). TheWestern Alps become a zone of frontal collision betweenthe Ivrea body (as frontal portion of Adria) and the Europeanmargin [Ford et al., 2006]. Since the Oligocene, the

~NNE–SSW trending curved shape of the Western Alps wasforming [Collombet et al., 2002;Dumont et al., 2011]: it showsa general westward polarity related to the continuous SE–NWcontinent–continent collision between Adria and Europe,associated with indentation of the Ivrea body [Schmid andKissling, 2000], whose southern termination was in corres-pondence of the Ligurian Alps [Dumont et al., 2011].Consistently, kinematic data from Western Alps(Figure 16a) indicate a general NE–SW to NW–SE com-pression (moving from south to north) in the external units(Dauphionis; [Platt et al., 1989a; Dumont et al., 2011]). Inthe internal zones, the fossil accretionary wedge experi-enced a change from Late Eocene–Early Rupelian E–Wductile-semibrittle extensional shearing [Schwartz et al.,2009] to Late Rupelian–Early Miocene NW–SE extensioncoexisting with sub-horizontal NE–SW shortening and im-portant right-lateral displacement along the major faults[Sue and Tricart, 2003; Tricart et al., 2004; Champagnacet al., 2006; Malusà et al., 2009; Perrone et al., 2010].

Figure 16. Schematic tectonic sketches showing the Oligo–Miocene geodynamic evolution of theWestern Alps and the Liguro-Provençal basin. (a) Late Rupelian: the Ligurian Alps-TPB and theSardinia-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 isorogen-parallel (NW–SE) and kinematically compatible with the coeval extension generated by the open-ing of the Liguro-Provençal basin. At this time, the Ligurian Alps record a fast exhumation, probablyinduced by tectonic denudation of the metamorphic chain via shallower detachment of the HelminthoidFlysch (HF). The Sestri-Voltaggio fault zone (SV; [Vignaroli et al., 2009]) separates the ophiolitic domainof the Ligurian Alps from the Appenines. (b) Late Chattian: the Ligurian Alps-TPB system was involved ina large scale left-lateral shear zone, accommodating the different motion of the west-directed Alps andnorth-east verging Apennines. In the box, an idealized distribution of tectonic elements associated with awrench/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 Alpsis propagating toward the external domains while the inner zone of the chain was dominated by orogenparallel 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 wasaffected by transpression. During this phase, the TPB acted as a strongly subsiding piggyback basin aboverotating thrust sheets. Kinematic data of the Western Alps are from Perrone et al. [2010] and referenceswithin. 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 completediscussion of the evolutionary scenario.


21

[81] The Early Oligocene also corresponds to the initiationof the NE–SW trending Liguro-Provençal rifting [De Voogdet al., 1991; Séranne, 1999; Roca, 2001; Rollet et al., 2002].This is considered as a back-arc basin generated fromthe 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 axisin the western Mediterranean region (S France and E Spain).Extensional faults dissected pre-Alpine basements (e. g., inthe western Corsica, Provence) and areas previously affectedby compressive deformation (e. g., Pyrenees, the Iberianchain, Figure 15b). This extensional deformation causedcrustal thinning and was subsequently associated with EarlyMiocene emplacement of oceanic crust actually exposed inthe 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 dis-mantling of the rapidly exhuming orogen associated withmoderate rate of basin subsidence. The deposition occurredin the retro-foreland of the Ligurian Alps and was controlledby NNW-directed mainly extensional faults (F1) at highangle with the previous contractional structures of the meta-morphic basement. The F1 faults represent the first brittledeformation phase (D4) developed in the metamorphic base-ment: the general pattern of mesoscale strain axes indicates aregional E–W to NE–SW extension (considering present-daycoordinates). Paleomagnetic data reveal ~50� of counter-clockwise rotation for the TPB and the underlying Ligurianbasement in the Aquitanian–Serravallian times [Maffioneet al., 2008]. If basin and basement are restored to theirEarly Oligocene position, the inferred kinematics from theF1 faults is in agreement with the regional strain of the coevalLiguro-Provençal basin (Figure 16a).[83] These extensional faults show continuity with the struc-

tures associated with the Liguro-Provençal rifting in the SEFrance (Provence). The close temporal, spatial, and geometri-cal relationship between the onset of mainly NE–SW trendingextensional structures and the TPB deposition suggests thatthe regional extension in the orogen-basin systemwas inducedby 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., theLigurian Alps) blocked northward by the ongoing indentationof the Ivrea body, probably represented a physical barrierfor the propagation of the rift, which no longer developednorthern to the Liguria. Therefore, following the modelof rifting evolution of Lavier and Manatschal [2006],the Ligurian Alps represented a stretching area ahead ofthe Liguro-Provençal rifting south to the tip of the Ivreabody (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 oceanicunits via extension has been proposed. Extensional tectonicsoperated by the reactivation of previous compressional struc-tures, with a progression from plastic to frictional deforma-tion mechanisms. The Voltri Massif undergone extensionaldenudation through shear localization along major plasticto brittle shear zones separating the Voltri Massif from theApennines in the east and from the Alpine Briançonnaisdomain in the west.

[85] This model accounts for the extensional character ofthe ductile-brittle and brittle structure identified both in theVoltri Massif and Briançonnais units (Table 1). It is alterna-tive to the classical interpretation of the exhumation of HPoceanic rocks caused by subduction channel processes asso-ciated with polyphase contractional regime [e. g., Federicoet al., 2007]. Moreover, the emphasis of the role of extensionduring the evolution of the Voltri Massif [HoogerduijnStrating, 1994; Vignaroli et al., 2008, 2010] can explainthe fast Eocene–earliest Oligocene exhumation (3–5mm/yr;Federico et al. [2005]) recorded by the ophiolitic rocks fromgreenschist facies stage condition to the surface andsupported 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 (1.3–6.8mm/yr),tectonically controlled, Oligocene exhumation through theuppermost (4–5 km) crust [Maino et al. [2012a]. These datasupport the concept that, at least in the shallowest crust, notonly the oceanic rocks were affected by rapid exhumation.The very high rates recorded by the Briançonnais rocks sug-gest that the fast exhumation was not limited to the VoltriMassif. This implies that, in the uppermost 4–5 km, the entireLigurian orogen experienced a relatively homogenous denu-dation without significant difference between oceanic andcontinental units.[87] This scenario can be explained by tectonic denudation

via extensional detachment: Maino et al. [2012a] suggestsa westward (in Oligocene coordinates) sliding of theHelminthoid Flysch (Figure 16a). During this translation,about 4–5 km-thick rock mass was progressively removedfrom the metamorphic chain, resulting in the high cooling/exhumation rates calculated from the thermochronometricdata. The shallow extensional detachment plane (dipping<30�) is represented by a plastic shear zone (Lambeaux descharriage, [Lanteaume, 1968]), constituted by a clay-richtectonosedimentary mélange, now resting between theDauphinois/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 extensionaldomain that promoted a two-stage exhumation process: asyn-greenschist exhumation of HP rocks associated withthe coupling of oceanic (Voltri Massif) and continental rocks[Vignaroli et al., 2008, 2009, 2010] and a subsequenttectonic denudation of the entire metamorphic basementthrough shallow detachment [Maino et al., 2012a]. Theextensional regime caused crustal thinning and generatedthe space for the TPB sedimentation. This tectonic scenariois in agreement with the proposed position of the LigurianAlps (Figures 15b and 16a) ahead of the Liguro-Provençalrifting and connecting the opposite propagation of theApennines and Alpine arcs [Vignaroli et al., 2008].6.2.3. Late Chattian (~26–23Ma)[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 indenta-tion of the Ivrea body under the Western Alps representsa stop for theNApennines propagation (Figures 15c). The east-ward migration of the fold-and-thrust belt of the Apennines


22

is testified by the coeval migration of the related foredeeps[Ricci Lucchi, 1986]. This translation resulted in a curved beltreflecting the convexity of the subduction trench. This shapeis due to the presence of easily subductable oceanic Ionian crustin the central part of the Apennines-Maghrebide subductionsystem (Figure 15c), with respect to the continental subductionof Adriatic crust below the Europe in the north [Jolivet andFaccenna, 2000]. The Northern Apennines were thereforecharacterized by relatively slow north-eastward translation,hampered by the complex interaction between Adria andEurope. In this context, the junction between the twobelts, i. e., the Ligurian Alps, suffered large-scale sinistraltranstension accommodating the different motions of theAlps and Apennines. The F2 fault-network shows orientationscomparable to a widely accepted model of left-lateral shearzone (Figure 16b). This geometry, therefore, indicate that,since the Late Chattian, the TPB-Ligurian Alps system actedas a pluri-kilometric sinistral strike-slip shear zone. In this con-text, the Sestri-Voltaggio line is interpreted as a dextral Riedelfracture as proposed by Capponi et al. [2009].[91] The Late Chattian transtensional phase in the Ligurian

Alps-TPB system represents the transition from EarlyOligocene rifting-related extension to the Early Miocenerotation-related transpression. It marks the period when theApennines start to “pull” the southern Western Alps north-eastward, producing a large shear zone (i. e., the LigurianAlps-TPB) between the two chains.6.2.4. Aquitanian–Serravallian (~23–12Ma)[92] During the Early–Middle Miocene, the Ligurian Alps-

TPB system was dominated by ~NE–SW shortening andcoeval ~NW–SE extension. This transpressive phase (D5)is coeval with a period of major assessment of the centralMediterranean area (Figure 15d): in the Liguro-Provençalback-arc basin, the thinning of the continental lithosphereresulted in the formation of new oceanic crust betweenca. 20 and 15Ma [Séranne, 1999; De Voogd et al., 1991].Synchronously with the oceanization of the basin, theCorsica-Sardinia block broke away from the Europeanmargin experiencing an anticlockwise rotation of ~60�[Gattacceca et al., 2007]. This rotation was driven by theeastward retreat of the Adriatic-Ionian composite slab[Jolivet and Faccenna, 2000; Castellarin, 2001; Faccennaet al., 2004]. The Ionian oceanic subducting crust increasedthe 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 Adriaticcrust in the northernmost part of the subduction and theproximity of the retro-front of the Alps impeded a free prop-agation of the northern Apennines (Figures 15d and 16c).Here, the north-verging thrust-fronts (Monferrato arc) pro-gressively approached the south-verging structures of theSouthern Alps beneath the TPB [Mosca et al., 2010]. In thisperiod, the TPB experienced an acceleration of the subsi-dence rates [Carrapa et al., 2003a], likely related to a flexuralresponse to thrust loading [Carrapa and Garcia Castellanos,2005]. The increasing load caused by the N-migratingApennines thrusts approaching the Southern Alps front pro-vided flexural tilting accommodating the high value of subsi-dence recorded in the TPB during the Middle Miocene.Therefore, the most plausible reason for the flexural subsi-dence in the TPB is the downward pull of the Apennines slab.

[93] The indentation of the Adriatic plate againstthe Europe [Schmid and Kissling, 2000] preserves anoverall convergent framework in the Western Alpsstill 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; Perroneet al., 2010; Sanchez et al., 2011]; while in the foreland,outward-verging thrust sequences developed [e. g., Fordet 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 pos-sible because the Ligurian segment of the Western Alps wasnot hampered by the presence of the Ivrea body as in theWestern Alps [Schmid and Kissling, 2000; Dumont et al.,2011]. The Ligurian Alps acted as a free junction betweenthe Western Alps and the Northern Apennines (Figures 15dand 16c). The shortening directions derived from the struc-tural analyses are consistent with the orientation of a regionalleft-lateral shear zone, producing t he oroclinal bending ofthe Ligurian orogen (Figure 16c).[95] Finally, from the Late Miocene, the mainly NW–SE

directed compression recorded in the TPB resulted in thefinal Plio–Pleistocene exhumation of the basin. At this time,the regional tectonic evolution is associated with the N- andNE-ward translation of the Ligurian units of the Apenninesand the TPB successions onto the Insubric foreland, alongthe Padan thrust front [e. g., Pieri and Groppi, 1981].

7. Conclusion

[96] Our research addresses field-based structural and strat-igraphic investigations to the Tertiary Piedmont Basin (TPB)and the Ligurian Alps, which represent the junction betweenthe two major orogenic systems of the Europe, the Alps-Dinarides and the Apennines-Maghrebides belts. The datahere presented allow proposing a complete evolutionaryscenario for the Oligo–Miocene history of the TPB: born asa retro-foreland basin of the Ligurian Alps, it records LateRupelian–Early Chattian extension which also controlledthe coeval tectonic exhumation and denudation of the meta-morphic basement. This extensional regime was generatedahead the apical closure of the Liguro-Provençal rifting,and it is associated with back-arc crustal thinning caused bythe roll-back of the Apennines subduction.[97] During the Late Chattian, the TPB evolved as a fore-

land basin located above a transtensional zone connectingthe opposite movements of the Alpine and Apennines arcs.Since the Early Miocene, the TPB acted as a strongly subsid-ing piggy-back basin above rotating thrust sheets associatedwith the regional rotation caused by the oceanic spreadingof 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 aroundwhich the Western Alps, the Liguro-Provençal basin, theCorsica-Sardinia block, and the Apennines moved control-ling the evolution of the central Mediterranean area.


23

Appendix A: Methodological Approach

A1. Structural Analysis

[99] Meso-structural analysis and mapping of brittle struc-tures have been carried out on both the metamorphic chain ofthe 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-metamorphicevolution, fault morphology, rheology, thickness of fault-rocks, and kinematic indicators. More than 1500 faults havebeen cataloged in more than 60 sites; here we only report thosemeasurements for which the sense of movement has beenindubitably identified; for the Briançonnais and Prepiedmontunits, 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 measurementson a single outcrop or on a section no longer than 60m.Field observations have been integrated with the orientationdata, morphology, dimension, and spacing of faults. In thiswork, we used the fault rocks classification of Woodcockand Mort [2008], which assumes the coarse grain-size as themost relevant criterion to distinguish the fault breccias.Families of faults have been distinguished on the basis ofgenetically related small-scale structures, superimposition ofmovements, and by the cross-cutting relationships. Mostof our measurement sites are along well exposed natural out-crops at high altitude (>2000m) or within open quarries,where faults are completely exposed; at low elevations wherethe vegetation masks the rocks, faults are difficult to trace andsense of movement difficult to be determined. Field data havebeen compared with observation on aerial photographs of welloutcropping 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 unequivocalsense of movement. The cross-cutting relationships are oftendifficult to investigate, and the fault hierarchy has beendeduced by the occurrence/absence of families of faultsdistinguished by their prevailing direction.

A2. Strain Analysis

[102] Our field work is based on the systematic measure-ment of fault populations, including fault plane orientation,striae and slickenside orientation, and shear sense, classifiedon the basis of the reliability criterion of the shear sense(certain, probable, uncertain). We decided not to resolve thepaleostress field because the determination of the directionsof principal stress axes from the field measurement of striatedfaults (stress inversion method) can be considered reliableonly in rocks that experienced deformation under a stressfield that had to be homogeneous both spatially and tempo-rally [e. g., Angelier, 1990; Twiss and Unruh, 1998]. Theseassumptions are not valid for many of the rocks exposed inthe study area and, in particular, for the basement rocks,affected bymultiple deformations or reactivations of preexistingdiscontinuities [Harris and Cobbold 1984; Flodin and Aydin2004]. For these reasons, we performed paleostrain analysis,which represents a qualitative determination of the incre-mental strain ellipsoid for each station [e. g., Malusà et al.,2009]. The strain analysis has been performed throughthe graphical construction of the principal incremental

shortening (P) and extension (T) axes for each populationof faults. The P and T axes concentration provides anapproximate orientation of the strain axes [Marrett andAllmendinger, 1990; Twiss and Unruh, 1998]. If the P-Tresults indicate a single maximum concentration which canbe related to a single tectonic event, we compared the P-Tanalyses with the field of incremental shortening and exten-sion deriving from the right dihedra method of Angelierand Mecheler [1977]. All the strain analyses have beencomputed using T-TECTO 3.0 computer program ( �Zaloharand Vrabec [2008], available at http://www2.arnes.si/~jzaloh/t-tecto_homepage.htm).

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

ReferencesAllen, P. A. (2008), From landscapes into geological history, Nature,451, 274–276.

Allen, P. A., and J. R. Allen (1990),Basin Analysis: Principles and Applications,Blackwell Sci., Oxford, U. K.

Alvarez, W., T. Cocozza, and F. C. Wezel, (1974), Fragmentation of Alpineorogenic belt by microplate dispersal, Nature, 248, 309–314.

Angelier, J. (1990), Inversion of field data in fault tectonics to obtain theregional stress – III. A new rapid direct inversion method by analyticalmeans, Geophys. J. Int., 103, 363–376.

Angelier, J., and P.Mecheler (1977), Sur unméthode graphique de recherche descontraintes principales égalment utilisable en tectonique et en séismologie: laméthode des dièdres droits, Bull. Soc. Geol. Fr., 19, 1309–1318.

Baietto, A., P. Perello, P. Cadoppi, G. Martinotti (2009), Alpine tectonicevolution and thermal water circulations of the Argentera Massif (South-Western Alps), Swiss J. Geosci., 102, 223–245.

Barbieri, C., B. Carrapa, A. Di Giulio, J. Wijbrans, G. Murrell (2003),Provenance of Oligocene synorogenic sediments of the Ligurian Alps(NW Italy): Inferences on belt age and cooling history, Int. J. Earth Sci.,92, 758–778.

Berger, J. P., and R. Bousquet (2008), Subduction-related metamorphismin the Alps: Review of isotopic ages based on petrology and theirgeodynamic consequences, Geol. Soc. Spec. Publ., 298, 117–144,doi:10.1144/SP298.7.

Bernet, M. (2009), A field-based estimate of the zircon fission-track closuretemperature, Chem. Geol., 259, 181–189.

Bernini, M., and M. Zecca (1990), Le deformazioni nelle Formazioni diMolare e di Rocchetta (Oligocene-Miocene inferiore) della regione diMioglia (SV) (Margine Sud del Bacino Terziario Piemontese), Atti Tic.Sc. Terra, 33, 1–10.

Bersezio, R., F. Felletti, and L. Micucci (2005), Statistical analysis of stratalpatterns and facies changes at the terminations of “turbiditic” sandstonebodies: The Oligocene Cengio Unit (Tertiary Piedmont Basin), GeoActa,4, 83–104.

Bersezio, R., F. Felletti, S. Riva, and L. Micucci (2009), Bed thickness andfacies trends of turbiditic sandstone bodies. Unravelling the effects ofbasin confinement, depositional processes and modes of sediment supply,in External Controls on Deep-Water Depositional Systems, vol. 92, editedby B. Kneller, O.J. Martinsen, and B. McCaffrey, pp. 303–324, Spec.Publ. SEPM Soc. Sediment. Geol., Tulsa, OK, USA.

Bertotti, G., and P. Mosca (2008), Late-orogenic vertical movements withinthe arc of the SW Alps and Ligurian Alps, Tectonophysics, 475, 117–127,doi:10.1016/j.tecto.2008.08.016.

Bertotti, G., P. Mosca, J. Juez-Larré, R. Polino, and T. Dunai (2006),Oligocene to present kilometres scale subsidence and exhumation of theLigurian Alps and the Tertiary Piedmont Basin (NW Italy) revealed byapatite (U–Th)/He thermochronology: Correlation with regional tectonics,Terra Nova, 18(1), 18–25.

Boccaletti, M., M. Coli, F. A. Decandia, E. Giannini, and A. Lazzarotto(1980), Evoluzione dell’Appennino Settentrionale secondo un nuovomodello strutturale, Mem. Soc. Geol. It., 25, 359–373.

Bonini, L., G. Dallagiovanna, and S. Seno (2010), The role of pre-existingfaults in the structural evolution of thrust systems: Insights from theLigurian Alps (Italy), Tectonophysics, 480 (1–4), 73–87.


24

Bousquet, R., R. Oberhänsli, B. Goffe, M.Wiederkehr, F. Koller, S. Schmid,R. Schuster, M. Engi, A. Berger, and G. Martinotti (2008), Metamorphismof metasediments at the scale of an orogen: A key to the tertiarygeodynamic evolution of the Alps, Geol. Soc. Spec. Publ., 298, 393–411,doi:10.1144/SP298.18.

Brandon, M. T., M. K. Roden-Tice, and J. I. Garver (1998), Late Cenozoicexhumation of the Cascadia accretionary wedge in the Olympic Mountains,northwest Washington State, Geol. Soc. Am. Bull., 110, 985–1009.

Butler, R. W. H., S. J. Matthews, and M. Parish (1986), The NW externalAlpine thrust belt and its implications for the geometry of the westernAlpine orogen, in Collision Tectonics, vol. 19, edited by M. P. Cowardand A. C. Ries, pp. 245–260, Geol. Soc. Spec. Publ., London.

Cabella, R., L. Cortesogno, G. Dallagiovanna, L. Gaggero, G. Lucchetti(1991), Metamorfismo a giadeite + quarzo in crosta continentale nelBrianzonese ligure, Atti Tic. Sc. Terra, 34, 43–54.

Capponi, G., and L. Crispini (2002), Structural and metamorphic signatureof Alpine tectonics in the Voltri Massif (Ligurian Alps, northwesternItaly), Ecl. Geol. Helv., 95, 31–42.

Capponi, G., L. Crispini, L. Federico, M. Piazza, and B. Fabbri (2009), LateAlpine tectonics in the Ligurian Alps: Constraints from the TertiaryPiedmont Basin conglomerates, Geol. J. 44, 211–244, doi:10.1002/gj.1140.

Carminati, E., M. J. R. Wortel, W. Spakman, and R. Sabadini (1998), Therole of slab detachment processes in the opening of the western-centralMediterranean basins: Some geological and geophysical evidence, EarthPlanet. Sc. Lett., 160(3–4), 651–665.

Carrapa, B., and D. Garcia Castellanos (2005), Western Alpine back-thrusting as subsidence mechanism in the Tertiary Piedmont Basin(NW Italy), Tectonophysics, 406, 197–212.

Carrapa, B., G. Bertotti, andW. Krijgsma (2003a), Subsidence, stress regimeand rotation(s) of a tectonically active sedimentary basin within theWestern Alps: The Tertiary Piedmont Basin (Alpine domain, NorthwestItaly), in Tracing Tectonic Deformation Using the Sedimentary Record,vol. 208, edited by T. McCann and A. Saintot, pp. 205–227, Geol.Soc. Spec. Publ., London.

Carrapa, B., J. Wijbrans, and G. Bertotti (2003b) Episodic exhumation in theWestern Alps, Geology, 31, 601–604.

Carrapa, B., A. Di Giulio, and J. Wijbrans (2004), The early stages of theAlpine collision: An image from the detrital thermochronology of UpperEocene–Lower Oligocene sediments in the Alps-Apennines knot area,Sediment. Geol., 171, 181–203.

Casini, L., S. Cuccurru, M. Maino, G. Oggiano, and M. Tiepolo (2012),Emplacement of the Arzachena Pluton (Corsica–Sardinia Batholith) andthe geodynamics of incoming Pangaea, Tectonophysics, 544–545, 31–49,doi:10.1016/j.tecto.2012.03.028.

Castellarin, A. (2001), Alps–Apennines and Po Plain-frontal Apennines rela-tions, in Anatomy of an Orogen: The Apennines and the AdjacentMediterranean Basins, edited by G. B. Vai and J. P. Martini,pp. 177–195, Kluwer Academic, Norwell, Mass.

Catanzariti, R., A. Cerrina Feroni, P. Martinelli, and G. Ottria (1996),Le marne dell’Oligocene-Miocene inferiore al limite tra DominioSubligure e Dominio Toscano: dati biostratigrafici ed evoluzione spazio-temporale, Atti Soc. Tosc. Sc. Nat., Mem., 103, 1–30.

Catuneanu, O. (2004), Retroarc foreland system – evolution through time,J. Afr. Earth Sci., 38, 225–242.

Cederbom, C., H. Sinclair, F. Schlunegger, and M. Rahn (2004), Climate-induced rebound and exhumation of the European Alps, Geology,32, 709–712.

Champagnac, J. D., C. Sue, B. Delacou, P. Tricart, C. Allanic, andM. Burkhard (2006), Miocene orogen-parallel extension in theinner Western Alps revealed by dynamical fault analyses, Tectonics,25, TC3014, doi:10.1029/2004TC001779.

Ciarcia, S., S. Mazzoli, S. Vitale, and M. Zattin (2012), On the tectonicevolution of the Ligurian accretionary complex in southern Italy, Geol.Soc. Am. Bull., 124(3/4), 463–483, doi:10.1130/B30437.1.

Cibin, U., A. Di Giulio, and L. Martelli (2003), Oligocene–Early Miocenetectonic evolution of the northern Apennines (northwestern Italy)traced through provenance of piggy-back basin fill successions, inTracing Tectonic Deformation Using the Sedimentary Record, vol. 208,edited by T. S. McCann, pp. 269–287, Geol. Soc. Spec. Publ., London.

Cloetingh, S., and J. Negendank (2009), New Frontiers in Integrated SolidEarth Sciences, Springer, Berlin.

Collombet, M., J. C. Thomas, A. Chauvin, P. Tricart, J. P. Bouillin, andJ. P. Gratier (2002), Counterclockwise rotation of the western Alps sincethe Oligocene: New insights from paleomagnetic data, Tectonics, 21(4),1032, doi:10.1029/2001TC901016.

Corsini, M., G. Ruffet, and R. Caby (2004), Alpine and late-hercynian geo-chronological constraints in the Argenetera massif (Western Alps), Eclog.Geol. Helv., 97, 3–15.

Cortesogno, L., and D. Haccard (1984), Note illustrative alla carta geologicadella zona Sestri-Voltaggio, Mem. Soc. Geol. It., 28, 115–150.

Crispini, L., and G. Capponi (2001), Tectonic evolution of the Voltri Groupand Sestri Voltaggio Zone (southern limit of the NWAlps): A review,Ofioliti 26(2a), 161–164.

Crispini, L., L. Federico, G. Capponi, and C. Spagnolo (2009), Late orogenictranspressional tectonics in the “Ligurian Knot,” Ital. J. Geosci.128, 433–441.

Dallagiovanna, G. (1995), Rilevamento geologico e analisi strutturale dellazona compresa fra le valli Mongia e Corsaglia (Alpi Liguri), Atti Tic. Sc.Terra, 37, 133–154.

Dallagiovanna, G., S. Silvio, and M. Vanossi (1997), An example of theAlpine structural evolution of the Penninic zone in the Ligurian Alps:Tectonics of the Barbassiria area, Eclog. Geol. Helv., 90(2), 337–344.

Dallagiovanna, G., L. Gaggero, M. Maino, S. Seno, and M. Tiepolo (2009),U–Pb zircon ages for post-Variscan volcanism in the Ligurian Alps(Northern Italy), J. Geol. Soc., 166, 1–14.

D’Atri, A., F. Piana, S. Tallone, G. Bodrato, and M. Roz Gastaldi (1997),Tettonoica oligo-miocenica nell’alto Monferrato (Bacino TerziarioPiemontese) e nel settore nord-occidentale del gruppo di Voltri (Acquiterme-Cassinelle, Al), Atti Tic. Sc. Terra, 5, 85–100.

D’Atri, A., F. Delapierre, A. Festa, R. Gelati, M. Gnaccolini, F. Piana, P. Clari,and R. Polino (2002), Tettonica e sedimentazione nel retroforeland alpino,81� Riunione estiva, Soc. Geol., It., Torino, Italy.

Decarlis, A., and A. Lualdi (2008), A late Triassic–Early Jurassic paleokarstfrom the Ligurian Alps and its geological significance (Siderolitico Auct.,Ligurian Briançonnais domain), Swiss J. Geosc. 101, 579–593.

Decarlis, A., and A. Lualdi (2009), A sequence stratigraphy approach to t amiddle Triassic shelf-slope complex of the Ligurian Alps (LigurianBriançonnais, Monte Carmo-Rialto unit, Italy), Facies 55, 267–290.

Decarlis, A., and A. Lualdi (2011), Synrift sedimentation on the northernTethys margin: An example from the Ligurian Alps (Upper Triassicto Lower Cretaceous, Prepiedmont domain, Italy), Int. J. Earth Sci.100, 1589–1604.

DeCelles, P. G., and K. Giles (1996), Foreland basin systems, Basin Res., 8,105–123.

De Voogd, B., et al. (1991), First deep seismic reflection transect from theGulf of Lions to Sardinia (ECORS-CROP profiles in westernMediterranean), in Continental Lithosphere: Deep Seismic ReflectionsGeodyn. Ser., vol. 22, pp. 265–274, AGU, Washington, D. C.

Di Giulio, A., and B. Galbiati (1991), Le facies caotiche dei Flysch dellaLiguria Occidentale: un nuovo modello interpretativo, Atti Tic. Sc.Terra, 34, 155–160.

Di Giulio, A., and B. Galbiati (1995), Interaction between tectonics anddeposition into an episutural basin in the Alps-Apennine knot, in Atti delConvegno Rapporti Alpi-Appennino 14, pp. 113–128, Accad. Naz. delleSci. detta dei XL, Rome.

Di Giulio, A., B. Carrapa, R. Fantoni, L. Gorla, and L. Valdisturlo (2001),Middle Eocene to Early Miocene sedimentary evolution of the westernsegment of the South Alpine foredeep (Italy), Int. J. Earth Sci., 90,534–548.

Dumont, T., T. Simon-Labric, C. Authemayou, and T. Heymes (2011),Lateral termination of the north-directed Alpine orogeny and onset of west-ward escape in the Western Alpine arc: Structural and sedimentary evidencefrom the external zone, Tectonics, 30, TC5006, doi:10.1029/2010TC002836.

Elter, P., and P. C. Pertusati (1973), Considerazioni sul limite Alpi-Appennino e sulle sue relazioni con l’arco delle Alpi Occidentali, Mem.Soc. Geol. Ita., 12, 359–375.

Elter, P., R. Catanzariti, F. Ghiselli, M. Marroni, G. Molli, G. Ottria, andL. Pandolfi (1999), L’unità Aveto (Appennino settentrionale): caratteristichelitostratigrafiche, biostratigrafia, petrografia, Boll. Soc. Geol. It., 118, 41–73.

Ernst, W. G. (1981), Petrogenesis of eclogites and peridotites from theWestern and Ligurian Alps. Am. Mineral., 66, 443–472.

Faccenna, C., C. Piromallo, A. Crespo-Blanc, L. Jolivet, and F. Rossetti(2004), Lateral slab deformation and the origin of the westernMediterranean arcs, Tectonics, 23, TC1012, doi:10.1029/2002TC001488.

Federico, L., G. Capponi, L. Crispini, M. Scambelluri, and I. M. Villa(2005), 39Ar/ 40Ar dating of high-pressure rocks from the Ligurian Alps:Evidence for a continuous subduction–exhumation cycle, Earth Planet.Sci. Lett. 240, 668–680.

Federico, L., L. Crispini, M. Scambelluri, and G. Capponi (2007), Ophiolitemélange zone records exhumation in a fossil subduction channel,Geology,35, 499–502.

Federico, L., C. Spagnolo, L. Crispini, and G. Capponi (2009), Fault-slipanalysis in the metaophiolites of the Voltri Massif: Constraints for thetectonic evolution at the Alps/Apennines boundary, Geol. J., 44, 225–240.

Felletti, F. (2002), Complex bedding geometries and facies associations of theturbiditic fill of a confined basin in a transpressive setting (Castagnola Fm.,Tertiary Piedmont Basin, NW Italy), Sedimentology, 49 (4), 645–667.

Felletti, F. (2004a), Statistical modelling and validation of correlation inturbidites: An example from the Tertiary Piedmont Basin (CastagnolaFm., Northern Italy), Mar. Pet. Geol., 21 (1), 23–39.


25

Felletti, F. (2004b), Spatial variability of Hurst statistics in the CastagnolaFormation, Tertiary Piedmont Basin, NW Italy: Discrimination of sub-environments in a confined turbidite system, in, Confined TurbiditeSystems, vol. 222, edited by S. A. Lomas and P. Joseph, pp. 285–306,Geol. Soc. Spec. Publ., London.

Felletti, F., and R. Bersezio (2010a), Quantification of the degree of confine-ment of a turbidite-filled basin: A statistical approach based on bed thick-ness distribution, Mar. Pet. Geol., 27(2), 515–532.

Felletti, F., and R. Bersezio (2010b), Validation of Hurst statistics: A predic-tive tool to discriminate turbiditic sub-environments in a confined basin,Pet. Geosci., 16, 401–412.

Flodin, E. A., and A. Aydin (2004), Evolution of a strike-slip fault network,Valley of Fire, southern Nevada, Geol. Soc. Am. Bull., 116, 1/2, 42–59.

Forcella, F., R. Gelati, M. Gnaccolini, P. M. Rossi, and R. Bersezio (1999),Il Bacino Terziario Ligure-Piemontese tra il Monregalese e la Valle del T.Lemme: stato delle ricerche e prospettive future, in Studi geografici egeologici in onore di Severino Belloni, edited by G. Ombrelloni,pp. 339–365, Brigatti, Genova.

Ford,M., andW.H. Lickorish (2004), Foreland basin evolution around thewesternalpine arc, inDeepWater Sedimentation in the Alpine Basin of SE France: NewPerspectives on the Gres d’Annot and Related Systems, vol. 211, edited by P.Joseph, and S.A. Lomas, pp. 39–63, Geol. Soc. Spec. Publ., London.

Ford, M., W. H. Lickorish, and N. J. Kusznir (1999), Tertiary foreland sed-imentation in the Southern Subalpine Chains, SE France: A geodynamicappraisal, Basin Res., 11, 315–336.

Ford, M., S. Duchêne, D. Gasquet, and O. Vanderhaeghe (2006), Two-phaseorogenic convergence in the external and internal SW Alps, J. Geol. Soc.London, 163, 815–826.

Fossati, A., C. R. Perotti, and P. L. Vercesi (1988), Deformazioni tettonicheal margine sud-occidentale del bacino terziario Ligure-Piemontese, Rend.Soc. Geol. It., 11, 305–308.

Franceschetti, B. (1967), Studi geologici sulla regione ad Ovest di Ovada(AL), Mem. Soc. Geol. It., 6, 379–420.

Fravega, P., S. Giammarino, M. Piazza, A. Russo, and G. Vannucci (1987),Significato degli episodi coralgali a Nord di Sassello. Nuovi dati per unaricostruzione paleogeografico evolutiva del margine settentrionale delBacino Terziario del Piemonte, Atti Soc. Tosc. Sc. Nat., Mem., 94, 11–17.

Gattacceca, J., A. Deino, R. Rizzo, D. S. Jones, B. Henry, B. Beaudoin, andF. Vadeboin (2007), Miocene rotation of Sardinia: New paleomagneticand geochronological constraints and geodynamic implications, EarthPlanet. Sci. Lett., 258, 359–377.

Gelati, R., and M. Gnaccolini (1998), Synsedimentary tectonics and sedi-mentation in the Tertiary Piedmont Basin, Northwestern Italy, Riv. It.Paleont. Strat., 104(2), 193–214.

Gelati, R., and M. Gnaccolini (2003), Genesis and evolution of the Langhebasin, with emphasis on the latest Oligocene–Earliest Miocene andSerravallian, Atti Tic. Sc. Terra, 44, 3–18.

Gelati, R., et al. (2010), Dego - 228 Sheet, Geological Map of Italy, 1:50.000scale, ISPRA, Rome, Italy. (Available online at: http://www.isprambiente.gov.it/MEDIA/carg/211_DEGO/Foglio.html.

Gnaccolini, M. (1978), Osservazioni su una piccola scogliera organogenaoligocenica affiorante presso Cairo Montenotte (SV), Riv. It. Paleont.Strat., 84, 403–410.

Giglia, G., G. Capponi, L. Crispini, and M. Piazza (1996), Dynamics andseismotectonics of the western-Alpine arc, Tectonophysics, 267, 143–175.

Goffé, B., S.Schwartz, J. M. Lardeaux, and R. Bousquet (2004), Explanatorynotes to the map: metamorphic structure of the Alps, Western and LigurianAlps, Mitt. der Öst. Min. Ges., 149, 125–144.

Gradstein, F. M., J. G. Ogg, Smith, A. G., W. Bleeker, and L. Lourens(2004), A new geological time scale with special reference toPrecambrian and Neogene, Episodes, 27, 83–100.

Handy, M. R., S. M. Schmid, R. Bousquet, E. Kissling, and D. Bernoulli(2010), Reconciling platetectonic reconstructions of Alpine Tethys withthe geological–geophysical record of spreading and subduction in theAlps, Earth Sci. Rev. 102, 121–158.

Harris, L. B., and P. R. Cobbold (1984), Development of conjugate shearbands during bulk simple shearing, J. Struct. Geol., 7, 37–44.

Hoogerduijn Strating, E. H. (1994), Extensional faulting in an intraoceanicsubduction complex-working hypothesis for the Palaeogene of the Alps–Apennine system., Tectonophysics 238, 255–273.

Hoogerduijn Strating, E. H., W. A van Wamel, and R. L. M. Vissers (1991),Some constraints on the kinematics of the Tertiary Piemonte Basin (north-western Italy), Tectonophysics 198, 47–51.

Jolivet, L., and C. Faccenna (2000), Mediterranean extension and theAfrica–Eurasia collision, Tectonics, 19, 1095–1106.

Jolivet, L., C. Faccenna, B. Goffé, M. Mattei, F. Rossetti, C. Brunet, F. Storti,R. Funiciello, J. P. Cadet, and T. Paria (1998), Mid-crustal shear zones inpost-orogenic extension: The northern Tyrrhenian Sea case, J. Geophys.Res., 103, 12,123–12,161.

Lavier, L., and G. Manatschal (2006), A mechanism to thin the continentallithosphere at magma-poor margins, Nature, 440, 324–328, doi:10.1038/nature04608.

Kerckhove, C. (1969), La ‘zone du Flysch’ dans les nappes de l’Embrunais-Ubaye (Alpes occidentales), Géol. Alpine, 45, 5–204.

Kissling, E., R. Bousquet, and M. Ford (2012), Formation of the arc of theWestern Alps and Alps-Apennines transition in the light of new geophys-ical data on the lithospheric architecture around the Ligurian knot, EGUGeneral Assembly 2012, 8432.

Kohlstedt, D. L., B. Evans, and S. J. Mackwell (1995), Strength of the lith-osphere: Constraints imposed by laboratory experiments, J. Geophys.Res., 100(B9), 587–602.

Lacombe, O., and L. Jolivet (2005), Structural and kinematic relation-ships between Corsica and the Pyrenees-Provence domain at the timeof the Pyrenean orogeny, Tectonics, 24, TC1003, doi:10.1029/2004TC001673.

Lanteaume, M. (1968), Contributions à l’étude géologique des AlpesMaritimes franco-italiennes, Mem. Serv. Carte Géol. De la France, Paris.

Lanteaume, M., N. Radulescu, N. Gavos, and J. Feraud (1990), Noticeexplicative de la Carte Géol. De France (1/50.000), feuille Viève-Tendande(948), BRGM, Orleans.

Laubscher, H. P., G. C. Biella, R. Cassinis, R. Gelati, A. Lozej, S. Scarascia,and I. Tabacco (1992), The collisional knot in Liguria, Geol. Rund.,81, 275–289.

Lorenz, C. (1969), Contribution a l’ètude stratigraphique de l’Oligoceneinferieur des confins Liguro-Piémontais (Italie), Atti Ist. Geol. Univ.Genova, 6, 253–888.

Lorenz, C. (1984), Evolution stratigraphique et structurale des Alpes Liguresdepuis l’Eocène supérieur, Mem. Soc. Geol. It., 28, 211–228.

Maffione, M., F. Speranza, C. Faccenna, A. Cascella, G. Vignaroli, andL. Sagnotti (2008), A synchronous Alpine and Corsica–Sardinia rotation,J. Geophys. Res., 113, B03104, doi:10.1029/2007JB005214.

Maino, M., G. Dallagiovanna, K. Dobson, L. Gaggero, C. Persano, S. Seno,and F. M. Stuart (2012a), Testing models of orogen exhumation usingzircon (U-Th)/He thermochronology: Insight from the Ligurian Alps,Northern Italy, Tectonophysics, 560–561, 84–93, doi:10.1016/j.tecto.2012.06.045.

Maino, M., G. Dallagiovanna, L. Gaggero, S. Seno, andM. Tiepolo (2012b),U-Pb zircon geochronological and petrographic constraints on late to post-collisional Variscan magmatism and metamorphism in the Ligurian Alps,Italy, Geol. J., 47(6), 632–652, doi:10.1002/gj.2421.

Malavieille, J., R. Lacassin, and M. Mattauer (1984), Significationtectonique des linéations d’allongement dans les Alpes occidentales,Bull. Soc. Geol. Fr., 26, 895–906.

Malinverno, A., and W. B. F. Ryan (1986), Extension in the Tyrrhenian Seaand shortening in the Apennines as a result of arc migration driven by thesinking of the lithosphere, Tectonics, 5, 227–245.

Malusà, M. G., R. Polino, and M. Zattin (2009), Strain partitioning in the ax-ial NW Alps since the Oligocene, Tectonics, 28, TC3005, doi:10.1029/2008TC002370.

Marrett, R. A., and R. W. Allmendinger (1990), Kinematic analysis of fault-slip data, J. Struct. Geol., 12, 973–986.

Marroni, M., A. Cerrina Feroni, D. Di Biase, G. Ottria, L. Pandolfi, andA. Taini (2002), Polyphase folding at upper structural levels in theBorbera valley (Northern Apennines, Italy): Implications for the tectonicevolution of the linkage area between Alps and Apennines, C. R.Geosci., 334, 565–572.

Merizzi, G., and S. Seno (1991), Deformations and gravity-driven translationof the San Remo-M. Saccarello Nappe (Helminthoid Flysch, LigurianAlps), Boll. Soc. Geol. It., 110, 757–770.

Merle, O., and J. P. Brun (1984), The curved translation path of theParpaillon Nappe (French Alps), J. Struct. Geol., 6, 711–719.

Messiga, B., and M. Scambelluri (1991), Retrograde P–T-t path for theVoltri Massif eclogites (Ligurian Alps, Italy): Some tectonic implications,J. Metamorph. Geol., 9, 93–109.

Michard, A., A. Chalouan, H. Feinberg, B. Goffé, R. Montigny (2002), Howdoes the Alpine belt end between Spain andMorocco?, Bull. Soc. Geol. Fr.173, 3–15.

Mosca, P., R. Polino, S. Rogledi, and M. Rossi (2010), New data for thekinematic interpretation of the Alps-Apennines junction (NorthwesternItaly), Int. J. Earth Sci., 99, 833–849.

Mutti, E., L. Papani, D. Di Biase,. G. Davoli, S. Mora, S. Segadelli, andR. Tinterri (1995), Il Bacino Terziario Epimesoalpino e le sue implicazionisui rapporti tra Alpi ed Appennino, Mem. Soc. Geol. It., 47, 217–244.

Naylor, M., and H. D. Sinclair (2008), Pro- versus Retro-Foreland Basins,Basin Res., 20, 285–303.

Patacca, E., R. Sartori, and P. Scandone (1990), Tyrrhenian basin andApeninic arc: Kinematic relations since Late Tortonian times, Mem. Soc.Geol. It., 45, 425–451.


26

http://www.isprambiente.gov.it/MEDIA/carg/211_DEGO/Foglio.html

http://www.isprambiente.gov.it/MEDIA/carg/211_DEGO/Foglio.html

Perotti, C. R. (1985), Analisi mesostrutturale della Formazione delle Marnedi Rigoroso fra il T. Lemme e il T. Scrivia (margine sud-orientale delBacino Terziario Ligure-Piemontese), Rend. Soc. Geol. It., 8, 53–56.

Perrone, G., P. Cadoppi, S. Tallone, and G. Balestro (2010), Post-collisionaltectonics in the Northern Cottian Alps (Italian Western Alps), Int. J. EarthSci., 100, 1349–1373, doi:10.1007/s00531-010-0534-1.

Piana, F., S. Tallone, S. Cavagna, and A. Conti (2006), Thrusting andfaulting in metamorphic and sedimentary units of Ligurian Alps: Anexample of integrated field work and geochemical analyses, Int. J. EarthSci., 95, 413–430.

Pieri, M., and G. Groppi (1981), Subsurface geological structure of the PoPlain. Progetto Finalizzato Geodinamica, 414, CNR, Italy.

Platt, J. P., J. H. Behrmann, P. C. Cunningham, J. F. Dewey, M. Helman,M. Parrish, M. G. Shepley, S. Wallis, and P. J. Weston (1989a),Kinematics of the Alpine arc and the motion history of Adria,Nature, 337, 158–161, doi:10.1038/337158a0.

Platt, J. P., G. S. Lister, P. Cunningham, P. Weston, F. Peel, T. Baudin, andH. Dondey (1989b), Thrusting and back thrusting in the Briançonnais domainof the Western Alps, in Alpine Tectonics, vol. 45, edited by M. P. Coward, P.Dietrich, and R.G. Park, pp. 135–152, Geol. Soc. Spec. Publ., London.

Réhault, J. P., G. Boillot, and A. Mauffret (1984), The westernMediterranean Basin geological evolution, Mar. Geol. 55, 447–477.

Reiners, P. W. (2005), Zircon (U-Th)/He thermochronometry, in Low-Temperature Thermochronology: Techniques, Interpretations andApplication, vol. 58, edited by P. W. Reiners and T. A. Ehlers, Rev.Mineral. Geochem., pp. 151–179, Mineral. Soc. of Am., Washington, D. C.

Ricci Lucchi, F. (1986), Foreland basins of the northern Apennines, inForeland Basins, edited by P.A. Allen and P. Homewood, Spec. IASPubl., 8, 105–139.

Ricou, L. E. (1981), Glissement senestre des nappes penniques le long de labordure nord de l’Argentera: Son role dans le jeu de l’arc alpin,C.R. Acad. Sci.Paris, 292, 1305–1308.

Roca, E. (2001), The Northwest Mediterranean Basin (Valencia Trough,Gulf of Lions and Liguro-Provençal basins): Structure andgeodynamic evolution, in Peri-Tethys Memoir 6: Peri-Tethyan Rift/Wrench Basins and Passive Margins, vol. 186, edited by P. A. Ziegleret al., pp. 671–706, Mem. Mus. Hist. Nat., Paris.

Rollet, N., J. Déverchère, M. O. Beslier, P. Guennoc, J. P. Réhault,M. Sosson, and C. Truffert (2002), Back arc extension, tectonic inheri-tance, and volcanism in the Ligurian Sea, Western Mediterranean,Tectonics, 21(3), 1015, doi: 10.1029/2001TC900027.

Rosenbaum, G., and G. S. Lister (2005), TheWestern Alps from the Jurassicto Oligocene: Spatio-temporal constraints and evolutionary reconstruc-tions, Earth Sci. Rev., 69, 281–306, doi:10.1016/j.earscirev.2004.10.001.

Rosenbaum, G., G. S. Lister, and C. Duboz (2002), Relative motionof Africa, Iberia and Europe during the Alpine orogeny, Tectonophysics,359, 117–129, doi:10.1016/S0040-1951(02)00442-0.

Rossi, P., G. Oggiano, and A. Cocherie (2009), A restored section of the“southern Variscan realm” across the Corsica–Sardinia microcontinent,C. R. Geosci., 341, 224–238.

Rubatto, D., and M. Scambelluri (2003), U–Pb dating of magmatic zirconand metamorphic baddeleyte in the Ligurian eclogites (Voltri Massif,Western Alps), Contrib. Mineral. Petrol., 146, 341–355.

Rutter, E. H. (1986), On the nomenclature of the mode of failure transitionsin rocks, Tectonophysics, 122, 381–387.

Sanchez, G., Y. Rolland, M. Jolivet, S. Brichau, M. Corsini, and A. Carter(2011) Exhumation controlled by transcurrent tectonics: The Argentera–Mercantour massif (SW Alps), Terra Nova, 23(2), 116–126.

Schlunegger, F., (1999), Controls of surface erosion on the evolution of theAlps: Constraints from the stratigraphies of the adjacent foreland basins,Int. J. Earth Sci., 86, 285–304.

Schmid, S. M., and M. R. Handy (1991), Towards a genetic classificationof fault rocks. Geological usage and tectnophysical implications, inControversies in Modern Geology, edited by D. W. Müller, J. A.McKenzie, and H. Weissert, 339–361, Academic, London.

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

Schmid, S., A. Pfiffner, N. Froitzheim, G. Schönborn, and E. Kissling(1996), Geophysical-geological transect and tectonic evolution of theSwiss-Italian Alps, Tectonics, 15, 1036–1064, doi:10.1029/96TC00433.

Scholz, C. H. (2002), The Mechanics of Earthquakes and Faulting,Cambridge Univ. Press, Cambridge, U. K.

Schwartz, S., P. Tricart, J. M. Lardeaux, S. Guillot, and O. Vidal (2009), Latetectonic and metamorphic evolution of the Piedmont accretionary wedge(Queyras Schistes lustres, Western Alps): Evidences for tilting during Alpinecollision, Geol. Soc. Am. Bull., 121, 502–518, doi:10.1130/B26223.1.

Sibson, R. H. (1977), Fault rock and mechanisms, J. Geol. Soc. London,133, 191––213.

Sinclair, H. D. (1997), Tectonostratigraphic model for underfilledperipheral foreland basins: An Alpine perspective, Geol. Soc. Am.Bull., 109, 324–346.

Seno, S., G. Dallagiovanna, and M. Vanossi (2005a), A kinematic evolutionmodel for the Penninic sector of the central Ligurian Alps, Int. J. EarthSci., 94, 114–129.

Seno, S., G. Dallagiovanna, and M. Vanossi (2005b), Pre-Piedmont andPiedmont-Ligurian nappes in the central sector of the Ligurian Alps: Apossible pathway for their superposition on to the inner Briançonnaisunits, Boll. Soc. Geol. Ita., 124(2), 455–464.

Seno, S., G. Dallagiovanna, L. Gaggero, F. Felletti, L. Pellegrini (2010),Cairo Montenotte - 228 Sheet, Geological Map of Italy, 1:50.000 scale,ISPRA, Rome, Italy (Available online at http://www.isprambiente.gov.it/Media/carg/liguria.html).

Séranne, M., (1999), The Gulf of Lions continental margin (NWMediterranean) revisited by IBS: An overview, in On theMediterranean Basins: Tertiary Extension Within Alpine Orogenedited by B. Durand, L. Jolivet, F. Horváth and M. Séranne, Geol.Soc. Spec. Publ., 156, 15–36.

Stampfli, G. M. and R. H. Marchant (1997), Geodynamic evolution ofthe Tethyan margins of the western Alps, in Deep Structure ofthe Swiss Alps: Results of NRP20, edited by O.A. Pfiffner et al.,pp. 223–239, Birkhäuser, Basel, Switzerland.

Sue, C., and P. Tricart (2003), Neogene to ongoing normal faulting in theinner western Alps: A major evolution of the Late Alpine tectonics,Tectonics, 22(5), 1050, doi:10.1029/2002TC001426.

Tricart, P., S. Schwartz, C. Sue and J.M. Lardeaux (2004), Evidence ofsynextension tilting and doming during final exhumation from analysisof multistage faults (Queyras Schistes lustrés, Western Alps), J. Struct.Geol., 26, 1633–1645.

Twiss, R. J., and J. R. Unruh (1998), Analysis of fault-slip inversions: Dothey constrain stress or strain rate?, J. Geophys. Res., 103,12,205–12,222.

Vance, J. A. (1999), Zircon fission track evidence for a Jurassic(Tethyan) thermal event in the Western Alps, Mem. Soc. Geol.Padova, 51, 473–476.

Vanossi, M., L. Cortesogno, B. Galbiati, B. Messiga, G. B. Piccardo andR. Vannucci, (1986), Geologia delle Alpi Liguri: dati, problemi, ipotesi,Mem. Soc. Geol. It., 28, 5–75.

Vanossi, M., C. R. Perotti and S. Seno (1994), TheMarittimes Alps arc in theLigurian and Tyrrheninan system, Tectonophysics, 230, 75–89.

Vignaroli, G., F. Rossetti, M. Bouybaouene, H.-J. Massonne, T. Theye,C. Faccenna, and R. Funiciello (2005), A counter-clockwise P-T pathfor the Voltri Massif eclogites (Ligurian Alps, Italy), J. Metamorph.Geol., 23, 533–555, doi:10.1111/j.1525-1314.2005.00592.x.

Vignaroli, G. L., C. Faccenna, L. Jolivet, C. Piromallo, and F. Rossetti(2008), Subduction polarity reversal at the junction between the WesternAlps and the Northern Apennines, Italy, Tectonophysics, 450, 34–50.

Vignaroli, G., C. Faccenna, and F. Rossetti (2009), Retrogressive fabricdevelopment during exhumation of the Voltri Massif (Ligurian Alps,Italy): arguments for an extensional origin and implications for theAlps–Apennines linkage, Int. J. Earth Sci., 98, 1077–1093.

Vignaroli, G., F. Rossetti, D. Rubatto, T. Theye, F. Lisker, and D. Phillips(2010), Pressure-temperature-deformation-time (P-T-d-t) exhumationhistory of the Voltri Massif HP complex, Ligurian Alps, Italy, Tectonics,29, TC6009, doi:10.1029/2009TC002621.

Woodcock, N. H., and K. Mort (2008), Classification of fault breccias andrelated fault rocks, Geol. Mag., 145, 435–440, doi:10.1017/S0016756808004883.

�Zalohar, J., and M. Vrabec (2008), Combined (paleo)stress and kinematicanalysis of fault-slip data: The Multiple-slip method, J. Struct. Geol.,30, 1603–1613, doi:10.1016/j.jsg.2008.09.004).


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tectono-sedimentary evolution of the tertiary piedmont basin (nw italy) within the oligo-miocene...

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