Post-nappe brittle tectonics and kinematic evolution of the north-western Alps: an integrated...

Post-nappe brittle tectonics and kinematic evolution of thenorth-western Alps: an integrated approach

A. Bistacchi, M. Massironi*

Dipartimento di Geologia, Paleontologia e Geo®sica, UniversitaÁ di Padova, Via Giotto 1, 35137 Padova, Italy

Received 8 December 1999; accepted 15 August 2000

Abstract

Data from remote sensing, structural geology and thermochronology provide the basis for this integrated reconstruction of

the Oligocene to Present kinematic evolution of the north-western Alpine nappe stack. Two brittle tectonic phases post-date the

Cretaceous±Eocene ductile deformation. A NW±SE extension developed in the Oligocene (D1) along three main conjugate

fault systems arranged in orthorhombic symmetry (N-, NW- and SE-dipping). Cooling rate contour maps, from published

apatite and zircon ®ssion-track ages and Rb/Sr biotite ages, highlight the differential exhumation of large fault-bounded blocks

during this phase, whilst synkinematic hydrothermal veins and calc-alkaline dykes (29±32 Ma) help to constrain its age. From

the Miocene onwards, a general rearrangement of the strain pattern led to SW-directed lateral extrusion (D2) of the Pennine-

Graian Alps block, bounded by a network of seismogenic shear zones, the most important being the Ospizio Sottile, Simplon,

Rhone and Chamonix faults. The internal deformation of the Pennine-Graian Alps block is characterised by an overall more or

less homogeneous NE±SW extension. The approach undertaken, integrating remote sensing, structural analysis on different

scales, and thermochronology (with the cooling rate map representation), is therefore effective in reconstructing the late-

orogenic extensional tectonic evolution of metamorphic nappe stacks. q 2000 Elsevier Science B.V. All rights reserved.

Keywords: Western Alps; brittle deformation; remote sensing; normal faults; cooling; exhumation

1. Introduction

The Alpine belt results from the subduction-collision

history of Cretaceous±Cenozoic age, between the

European and Adria (African) plates (e.g. Coward and

Dietrich, 1989; Polino et al., 1990; Michard et al., 1996).

The axial domain of the north-western Alps is charac-

terised by polyphase metamorphic history and penetra-

tive ductile deformation, and also by brittle faulting

whose importance has been underestimated. This is

evidenced by the relatively small number of faults that

are reported on published geological maps (e.g. Spicher,

1980; Bigi et al., 1990; Steck and Bigioggero, 2000).

Owing to the limited knowledge of these faults,

several problems concerning the last stages (Oligo-

cene±Present) of Alpine tectonics are still open. The

Oligocene mode of faulting, the age of some back-

thrusts and back-folds, and the Miocene±Present

kinematics remain matters of debate.

To tackle these problems, an integrated analysis

combining satellite image interpretation, structural

analysis and thermochronology was carried out in

the axial sector of the NW Alps (Fig. 1), between

the Chamonix line (W of the Mont Blanc Massif)

and the Canavese line (SE border of Alpine meta-

morphic nappes).

Tectonophysics 327 (2000) 267±292

0040-1951/00/$ - see front matter q 2000 Elsevier Science B.V. All rights reserved.

PII: S0040-1951(00)00206-7

www.elsevier.com/locate/tecto

* Corresponding author. Tel.: 139-498272050; fax: 139-

498272070.

E-mail address: [email protected] (M. Massironi).

A.

Bista

cchi,

M.

Ma

ssironi

/T

ectonophysics

327

(2000)

267

±292

26

8

Fig. 1. Tectonic sketch-map of north-western Alps, showing principal Oligocene to Present brittle faults (thick lines) and other lineaments derived from satellite remote-sensing

interpretation (narrow lines). Inset (I): principal tectonic domains of Alps. European Foreland (EF), Jura (J), Helvetic-Dauphinois (HD), Penninic (PN), Austroalpine (AU),

Southern Alps (SA), Po Plain (Po), Appennines (AP).

In a previous paper (Bistacchi et al., 2000), we

focused on the Miocene±Present kinematics, with a

particular emphasis on the relationships between

structural and seismotectonic data. In the present

paper more details about methodologies are given,

followed by new data and discussion of the Oligocene

to Present polyphase evolution.

2. Geological outline

2.1. Regional setting

The Europe-vergent Austroalpine-Penninic wedge

is a subduction-related wedge devoid of welded litho-

spheric mantle, accreted before and during continental

collision (e.g. Nicolas et al., 1990; Polino et al., 1990;

Dal Piaz, 1999). This axial wedge consists of a

tectonic multi-layer composed of rootless basement

slices, deÂcollement cover units and minor ophiolitic

interleavings characterised by ductile deformation

and subduction-related low-T metamorphism. The

principal units of the orogenic wedge are, from top

to bottom, and from SE to NW (Fig. 1):

(1) the Adria-derived Austroalpine system,

composed of the Sesia-Lanzo inlier and of a number

of upper and lower Austroalpine outliers;

(2) the ophiolitic Piedmont zone, composed of the

Combin (upper) and Zermatt-Saas (lower) nappes

(details in Fig. 6);

(3) the Europe-derived continental nappes, includ-

ing:

(a) the inner/upper Penninic Monte Rosa, Arcesa-

Brusson and Gran Paradiso nappes;

(b) the middle Penninic Grand St. Bernard multi-

nappe system;

(c) the outer Penninic ¯ysch-dominated Sion-

Courmayeur zone;

(d) the lower Penninic Lepontine gneissic nappes

(Ossola-Tessin window), in places with minor

ophiolites.

In more detail (Fig. 1), the Monte Rosa nappe is

ultimately refolded in the Ossola Valley by the large

Vanzone back-fold (Laduron and Merlyn, 1974;

Klein, 1978; Milnes et al., 1981). The subsurface

continuity of the upper Penninic Monte Rosa and

Gran Paradiso nappes is suggested by basement

rocks of the same af®nity that crop out in the

Arcesa-Brusson window, south of the Aosta±Ranzola

fault (Fig. 1).

The Austroalpine-Penninic wedge is bounded by

two major crustal discontinuities: the Penninic frontal

thrust on its external border and the Canavese line on

the internal one (Fig. 1). The Penninic frontal thrust

separates the orogenic wedge from the Helvetic

(European) continental basement slices (Mont

Blanc, Aiguilles Rouges, Belledonne and Aar-

Gotthard external massifs) and deÂcollement cover

nappes. The Canavese line is the tectonic boundary

between the orogenic wedge and the antithetic (Adria-

vergent) Southalpine thrust belt, which developed in

the Adriatic hinterland from the Miocene onwards

(e.g. Schmid et al., 1989). Both the Helvetic and

Southalpine domains are free of ophiolites and

subduction-related metamorphic imprint.

2.2. Tectonic and metamorphic evolution of the nappe

stack

The Austroalpine-Penninic wedge developed under

a subduction-related low-T regime, during the subduc-

tion of the Mesozoic Piedmont ocean beneath Adria.

Different peak metamorphic conditions are

recorded in different groups of nappes, ranging from

blueschist to eclogite facies, and locally reaching

ultra-high-P conditions (Frey et al., 1999). Blueschist

facies climax assemblages have been found in the

upper Austroalpine outliers, Combin ophiolitic

nappe and mid-Penninic Grand St. Bernard system.

Instead, eclogite facies assemblages characterise the

Austroalpine Sesia-Lanzo inlier, lower Austroalpine

outliers, Zermatt-Saas ophiolitic nappe and upper

Penninic Monte Rosa and Gran Paradiso nappes

(BalleÁvre et al., 1986; Frey et al., 1999). Diachronous

ages of peak metamorphic assemblages in these

groups of nappes may indicate long-lasting subduc-

tion, from the Late Cretaceous to the Middle Eocene.

These nappes were independently dragged to different

depths and then exhumed within the accretionary

wedge (Michard et al., 1996; Cortiana et al., 1998;

Dal Piaz, 1999).

After collision, the subduction-related low-T

regime (ca. 208C/km) was replaced by a Late

Eocene±Early Oligocene higher-T regime (ca. 408C/

km). In this area of the Alps, the sudden change

in thermal conditions is indicated by regional

A. Bistacchi, M. Massironi / Tectonophysics 327 (2000) 267±292 269

metamorphic re-equilibration in greenschist (Aosta

Valley-Valais) to amphibolite facies (Ossola-Tessin)

conditions (Cannic et al., 1999; Desmons et al., 1999)

and by the Oligocene emplacement of the Biella and

Traversella plutons and swarms of calc-alkaline to

ultra-potassic dykes in the Austroalpine and ophiolitic

units (Dal Piaz et al., 1979).

2.3. Post-nappe brittle tectonics and pending

problems

According to the literature, the ®nal exhumation of

the Penninic-Austroalpine nappe stack was accom-

plished, from the Oligocene to the Present, through

the following steps:

(a) Oligocene extension coupled with the Periadria-

tic magmatism and emplacement of gold-bearing

quartz veins in the middle Aosta Valley (Dal Piaz et

al., 1979; Venturelli et al., 1984; Diamond, 1990; Dal

Piaz and Gosso, 1994).

(b) Differential exhumation at the Oligocene±

Miocene boundary, evidenced by thermochronologi-

cal data (Hurford et al., 1991; Hunziker et al., 1992).

This event is generally considered to be a conse-

quence of back-thrusting and back-folding, which

were in turn generated by the north-westward displa-

cement of the Adriatic indenter, a process which still

continues today (Schmid et al., 1989; Polino et al.,

1990; Giglia et al., 1996).

(c) Rapid exhumation of the western sector of the

Lepontine dome (ªSimplon sub-domeº of Merle et al.,

1989) from 18 to at least 3 Ma, accommodated by

normal displacement along the Simplon fault

(Mancktelow, 1985, 1990, 1992; Grasemann and

Mancktelow, 1993). Bistacchi et al. (2000) showed

that this still-active kinematics is coupled with the

SW-ward gravitational lateral extrusion of the

Pennine-Graian Alps block (PGA block), bounded

by the Chamonix line to the NW, the Rhone±Simplon

system to the N, and the Ospizio Sottile line to the SE

(Fig. 1).

Within this tectonic framework, many problems

still arise:

(1) No kinematic interpretation explaining the

mode of faulting during the Oligocene extension

(step ªaº) has been proposed.

(2) The late Oligocene±early Miocene NW±SE

compression (step ªbº) seems to be supported by:

(a) dextral transpression along the Canavese line

east of the Ossola Valley, and (b) the development

of the southern steep belt, which may be linked to

the Vanzone antiform (Schmid et al., 1989). In this

way, the late Oligocene±early Miocene transpression

along the Periadriatic lineament, well constrained in

the Central Alps, may extend to the SW segment of

the Canavese line (Schmid et al., 1989). However, in

this portion of the line, there is no clear ®eld evidence

for such kinematics at the Oligo-Miocene boundary

(Martin et al., 1999; Bistacchi et al., 2000). In addi-

tion, recent radiometric dating (Barnicoat et al., 1995;

Freeman et al., 1997) con®rm that, in the north-

western Alps, back-thrusting and back-folding devel-

oped in the late Eocene±earliest Oligocene, as

previously suggested by Milnes et al. (1981).

(3) The onset age of the lateral extrusion (step ªcº)

is debated (Grasemann and Mancktelow, 1993; Ring

and Merle, 1992; Seward and Mancktelow, 1994;

Schlunegger and Willett, 1999; Bistacchi et al.,

2000); according to different authors, it may be placed

between 22 and 18 Ma.

3. Methodology

3.1. Remote sensing

The brittle structural pattern of the NW-Alps was

comprehensively investigated using the regional

A. Bistacchi, M. Massironi / Tectonophysics 327 (2000) 267±292270

Fig. 2. Eastern end of Aosta±Ranzola fault system branching into Ospizio Sottile fault on Landsat TM image (band 4). (a) Spatial-edge

enhancement ®ltering emphasises high-frequency textural patterns, allowing detection of linear discontinuities evidenced by abrupt tonal

changes (arrows); sun illumination from lower right corner (SE). These changes are mainly linked to fault-related morphological features

(alignment of crests, valleys and passes, trenches, uphill- and downhill-facing scarps, ridge doubling), fault-related geological contacts, large

layers of fractured rocks. (b) Ospizio Sottile fault (white arrows) is evident throughout its extension as demonstrated by a 30-km-long array of

lineaments on this portion of Landsat TM scene; Aosta±Ranzola fault system (black arrows) is evidenced on satellite images by a 2-km-wide

system of east±west trending lineaments which do not extend east of Ospizio Sottile fault.

synoptic vision of satellite images. In an area of such

rugged topography and deep glacial erosion, remote

sensing is very useful for tracing large faults exten-

sively covered by Quaternary deposits.

One Landsat 5 TM geocoded scene and two ERS-1

SAR/GTC (Synthetic Aperture Radar/Geocoded

Terrain Correct) images were processed, analysed

and compared. Interpretation of the Landsat TM

product was carried out mainly on bands 4 and 7

and was then cross-checked with other bands (except

for band 6, due to its low spatial resolution) and with

the 3±2±1 multi-band composition.

As a consequence of side-looking radar acquisition,

SAR images of mountain ranges are characterised by

intense geometrical and radiometric distortions

(shadowing, layover, foreshortening), which princi-

pally affect sensor-looking slopes (Curlander and

Mcdonough, 1991). For this reason, the two highly

overlapping ERS-1 scenes, acquired in ascending

and descending orbits, were used in integrated analy-

sis of slopes facing W and E, respectively. The GTC

product is the result of the correction of the raw SAR

image with a Digital Terrain Model (ESA, 1992). In

this way, any kind of image distortion is geometrically

removed, whereas the simultaneous interpretation of

both SAR scenes partially avoids radiometric distor-

tions on sensor-looking slopes.

Initially, remote sensing was used as a preliminary

tool to detect potential brittle tectonic features

(Massironi et al., 1997). Lineament interpretation on

either radar or optical data was obtained by visual

inspection. Simple contrast histogram adjustments

(equalisation, linear-piecewise, logarithmic and expo-

nential stretch; see Richards (1986), Gupta (1991) and

Sabins (1997) for complete reviews) were frequently

implemented during interpretation on both optical and

radar data, in order to enhance differently illuminated

areas. Once a fault system was detected, not all linea-

ments that could represent its possible expression

were traced; only continuous arrays of lineaments

(.4 km long), with strong morphological and/or

spectral evidence, were drawn. This technique

identi®ed clear lineament patterns with good ®eld

correspondence. Nearly 2000 lineaments were recog-

nised, ranging from 400 m to 10 km in length: 429

Landsat TM lineaments (out of 1093) and 201

ERS-1 lineaments (out of 747) were checked through

systematic ®eld survey. When the correspondence of a

lineament with any brittle structure (either faults

related to regional tectonics or deep-seated gravita-

tional slope deformations) is considered as a positive

check, the reliability estimate ranged from 65.6 to

84.2% (depending on lineament length and different

satellite products; Massironi, 1998; Bistacchi et al.,

2000). Field analysis proved that most of the linea-

ments detected by image interpretation were high-

angle brittle features (dip . 658).In a more advanced analytical stage, continuous

feedback between satellite image interpretation and

®eld surveys turned out to be a robust correlation

tool, useful in reconstructing the geometrical details

of the structural model. During this stage, major faults

(thick lines in Fig. 1) were distinguished from minor

brittle features. It was also possible to better constrain

the trace, on the Landsat TM image, of some inter-

mediate angle faults (dip . 408 and ,65±708), which

were locally recognised by ®eld analysis. At this

stage, spatial frequency and directional edge enhance-

ment processing on optical images (Figs. 2 and 3;

Richards, 1986; Gupta, 1991; Sabins, 1997) were

often implemented in order to highlight any linear

discontinuities ill de®ned in the preliminary analysis,

but revealing important ®eld evidence (e.g. thick

pseudotachylite or cataclasite layers).

Satellite image interpretation, integrated with ®eld

surveys, allowed the brittle elements to be grouped

into three different sets, striking NE±SW, E±W and

NW±SE and pervasively dissecting the NW Alpine

nappe stack (Fig. 1).

3.2. Kinematic analysis

Before remote-sensing analysis was carried out,

only a few major faults were known in this portion


Fig. 3. M. GeleÂ and Trois Villes faults (D1 tectonic phase). Some major faults with dip # 65±708 are not clearly visible on satellite images and

were ®rst detected during ®eld surveys; later, their regional attitude and orientation were better constrained using remote-sensing. (a) Landsat

TM band 4: M. GeleÂ fault (black arrows) and, in particular, Trois Villes fault (white arrows) are hardly distinguishable on image. (b) NE±SW

directional ®ltering on Landsat TM band 4: NE±SW trending edges are enhanced and both faults become visible (arrows as in Fig. 3a).

(c) Structural sketch-map showing M. GeleÂ and Trois Villes faults (stars show outcrops where microtectonic data were collected).

of the Western Alps (Fig. 1): the Aosta±Ranzola line,

the Simplon±Rhone±Chamonix system and the

Canavese line (Spicher, 1980; Bigi et al., 1990;

Steck and Bigioggero, 2000). In addition, some data

have been published on brittle structures along the

NW and SE borders of the Dent Blanche nappe

(Wust and Silveberg, 1989; Ring, 1994). In the

present work, the complex lineament pattern detected

by remote sensing was taken as a guideline for

detailed ®eldwork. In addition, more low-angle

features were recognised during the ®eld survey.

Two principal brittle±ductile to brittle tectonic

phases (D1 and D2) were distinguished by means of

consistent cross-cutting relationships. They post-date

the Late Eocene±early Oligocene greenschist facies

regional fabrics. In areas where the strain was mainly

partitioned during D2, it was very dif®cult to ®nd

ªfossilº D1 kinematic indicators. Hence, much of

the ®eld analysis on D1 was carried out in the

Austroalpine, upper-Penninic and ophiolitic units of

the middle Aosta Valley, where the D2 strain is rela-

tively low and homogeneous (Fig. 6). In the D2 high-

strain areas, the relative timing is straightforward and

D2 kinematics could be studied in detail.

3D fault geometry was reconstructed with remote

sensing and ®eld data. Often, a remote-sensing linea-

ment was the ®rst clue to the existence of a fault and

its geometry was later constrained in the ®eld. Map-

traces of major faults that were active during D1 and

D2 are shown in Figs. 6 and 8, respectively.

The kinematics of these faults was determined by

means of microtectonic analysis carried out on selected

outcrops. This analysis was based on over 200 sites,

selected on the basis of remote sensing and ®eld

evidence. For each site, the average kinematics was

determined for different sets of fault planes

(Figs. 5±8). For example, Fig. 4 documents measure-

ments at an outcrop in the S. Bartelemy valley: two sets

of D1 mesoscopic faults and two sets of D2 mesoscopic

faults, plus R, R 0 and T joints linked to each phase (see

Petit (1987) and Hancock (1985) for terminology of

secondary joints). An average fault plane and slip

vector were calculated for each set, and average data

were rejected when they did not satisfy the following

conditions: (1) the aperture of the 95% con®dence cone

must be less than 158, and (2) the average slip vector

must be contained in the average fault plane with an

error of less than 58. Average fault planes for each set

were then compared with the large-scale 3D geometry

of faults, reconstructed from remote sensing and geol-

ogy. In the example, one set of mesoscopic faults for

each phase displays the same attitude as the corre-

sponding large-scale fault: hence, these are considered

to be ªprincipalº sets and representative of the kine-

matics of the large-scale ªmasterº fault. The other sets

are considered to be ªsecondaryº sets.

After averaging and selection, only average data

from the ªprincipalº sets were used to reconstruct

the regional kinematic model, assuming that large

faults best represent the regional kinematic pattern.

These data are shown in Figs. 5±8 for phases D1

and D2, respectively.

The result of this processing is to allocate to each

site the same weight in the reconstruction of the regio-

nal kinematic model, notwithstanding the fact that

datasets of each site were necessarily heterogeneous

owing to the high lithological variability. The occur-

rence of consistent kinematic indicators in different

rocks, found along high-angle faults cross-cutting

various levels of the almost ¯at-lying nappe stack

(Fig. 9), has been considered as important evidence

of the consistency of microtectonic data.

3.3. Cooling rates

Over 400 ®ssion-track ages on apatite (AFT) and

zircon (ZFT) and Rb/Sr biotite cooling ages

have been published on the study area. Hunziker

et al. (1992) provide a critical review of many

of these data and discuss their reliability. Seward


Fig. 4. Example of processing of microtectonic data in order to obtain average data representative of each site. Data from a homogeneous

outcrop of Piedmont Zone calcschists in S. BarteÂlemy Valley. Four sets of mesoscopic faults and related secondary joints were measured and

attributed to phase D1 (a) and D2 (b), thanks to consistent cross-cutting relationships (great circles: fault planes; arrows: sense of movement of

hanging-wall; small circles: poles of secondary fractures). Principal faults, representative of attitudes of two cross-cutting master faults

(reconstructed by remote sensing and ®eldwork) are shown in (c) and (d). Secondary conjugate faults shown in (e) and (f). Average fault-

slip data for four sets in (g) and (h); only ªprincipalº data are used in following regional kinematic reconstruction (see Figs. 5±8).


Fig. 5. Stereoplots of site-average kinematic data for phase D1: (a) reactivated foliation in calcschists and shear sense of hanging-wall, as

deduced from extensional crenulation cleavage (great circles and arrows, respectively); (b) low-angle fault planes in calcschists and sense of

movement of hanging-wall, as deduced from fault-gouge composite fabric (great circles and arrows, respectively); (c) high-angle fault planes in

all units and slip vectors, as deduced from slickenside lineations, Riedel-type shears, tension gashes, and pseudotachylite-®lled pull-aparts

(great circles and arrows, respectively); (d) poles of gold-bearing quartz veins of Brusson district (32±30 Ma, open circles) and calc-alkaline

dikes (31±29 Ma, triangles). Large arrows: regional NW±SE extension direction.

and Mancktelow (1994) and Balestrieri et al.

(1999) list more data on two key areas: the

Valais and the Mont Blanc massif, and the middle

Aosta Valley, respectively.

AFT ages clearly show that the principal fault-

bounded blocks underwent different exhumation

histories (Hurford et al., 1991). However, AFT ages

are not suf®cient to reconstruct the complex brittle

tectonic history of this part of the Alps. In fact,

when dealing with multi-phase tectonic history,

cooling age data are the discontinuous result of the

whole (continuous) cooling history, since they mark

cooling at a given closure temperature but do not give

information about the path that was followed at higher

and lower temperatures. In order to reconstruct the

kinematics and age of individual displacements, a

step-by-step instantaneous analysis is actually needed.

In particular, the kinematics of a fault at a given time can

only be reconstructed by comparing the instantaneous

exhumation rates of its hanging-wall and foot-wall.

The exhumation rate (dh/dt) can be computed only

from a set of P±T data, but at these shallow crustal

levels no P estimates are available. Alternatively, the

exhumation rate may be inferred from the cooling rate

(dT/dt) via thermal modelling (calculating T as a func-

tion of depth at given times), but only with a conco-

mitant degree of uncertainty re¯ecting assumptions in

the input data (such as thermal conductivity, radio-

metric heat production and mantle heat ¯ow). There-

fore, we prefer to show raw cooling rate data, which

are not in themselves biased by any uncertainty in

geophysical modelling. Cooling rates are not linearly

proportional to exhumation rates, since geotherms are

not straight lines in a T-depth space and may differ

signi®cantly in adjacent crustal blocks owing to

advection terms, which are different in the case of

differential exhumation of large blocks (see the

numerical example for the Simplon fault in Grase-

mann and Mancktelow, 1993). Nevertheless, there is

always a positive relationship between cooling and

exhumation rate (the faster the cooling, the faster

the exhumation), since on a crustal scale the geotherm

is a positive function of depth; therefore cooling rates

can be used, qualitatively, to infer the kinematics of

fault-bounded blocks.

Cooling rates (8C/Ma) were computed from

sequences of AFT, ZFT and biotite cooling ages

obtained on the same samples. Closure temperatures

were taken as stated by Hunziker et al. (1992) for

similar cooling rates (300 ^ 508C for Bt Rb/Sr,

240 ^ 408C for ZFT and 110 ^ 208C for AFT) and

a temperature of 108C was assumed for the present-

day surface. For each sample, a cooling curve was

reconstructed by linking the array of T/cooling-age

data with linear segments in a T/time diagram

(Fig. 10). Although the real cooling history of the

sample may be quite complicated, a simple linear

interpolation was undertaken because the actual

cooling path between data-points cannot be recon-

structed with better detail. The slope of varying

segments is the average cooling rate (DT/Dt) for the

given time interval.

Cooling rates resulting from this calculation range

between 1 and 608C/Ma. If we assume a linear geother-

mal gradient of 408C/km (Cannic et al., 1999), the exhu-

mation rate turns out to be of the order of 0.025±1.5 km/

Ma but, as mentioned before, the reliability of this esti-

mate is very low. Nevertheless, present-day uplift rates

reported by Khale et al. (1997) are of the same order of

magnitude in the Valais area.

Alternatively, exhumation rates could have been

reconstructed using the age-elevation method, gener-

ally undertaken in AFT studies (e.g. Fitzgerald and

Gleadow, 1988). However, using higher-T data

(ZFT and biotite ages), a wider age interval may be

covered, allowing analysis to be extended back in

time to the Oligocene. In addition, the age-elevation

method requires that considered samples come

from the same tectonic block and display a consistent

exhumation history. However, most suitable sampling

pro®les in the study area are affected by faulting,

which may post-date cooling through the apatite

partial annealing zone. This explains the general

lack of a convincing age-elevation correlation.

Cooling rates were gridded with kriging for ages

ranging from 32 to 0 Ma, and the resulting ªcooling

rate contour mapsº are characterised by well-de®ned

and long-lasting regional cooling patterns (Fig. 11),

which qualitatively correspond to exhumation

patterns with similar topology (fast cooling areas

should correspond to fast exhumation areas).

Cooling-rate contour maps show areas of relatively

constant or gently varying cooling rates (rigid blocks),

bounded by steep gradient belts developing along

principal faults that were active (with some verti-

cal slip) in the time increment considered. Purely


strike-slip or inactive faults are cross-cut by undis-

turbed cooling-rate contour lines.

Cooling curves, as interpolated here, using ªagesº

recorded at decreasing temperatures by single samples,

are not in themselves in¯uenced by topography (all

ªagesº in a curve come from the same sample, hence

from the same elevation). On the other hand, rugged

topography can induce lateral heterogeneities in the

geothermal ®eld, particularly at shallow depths. Thus,

interpretation of cooling rates in terms of exhumation

rates can be misleading, specially in the low-T ®eld

�T , 100±1508C; StuÈwe et al., 1994; Mancktelow and

Grasemann, 1997). However, we assume that the broad

tectonic reconstruction presented in this paper cannot be

signi®cantly affected by this error. In fact, our recon-

struction is based on the very high contrast in cooling

rates, which characterise different fault-bounded blocks

(e.g. 4 vs. 308C/km for the middle Aosta Valley and the

Lepontine dome at 12 Ma; Fig. 11c).

Grasemann and Mancktelow (1993) showed that

crustal blocks exhumed at different rates maintain a

different geothermal gradient for some million years

after the differential movement ceased, since conductive

thermal re-equilibration between adjacent blocks is

subject to delay. Therefore, the use of cooling rate

contour maps to infer the timing of movements along

major faults may be misleading; the onset or end of

a given activity along a fault is generally likely to

pre-date its thermal evidence. In any case, in the western

Alps, the contrast between exhumation rates in different

blocks should be ,2 km/Ma, and hence the horizontal

geothermal gradient across faults should be limited

(Grasemann and Mancktelow, 1993). We therefore

assume that this kind of error is of the same order as

that of thermochronological methods (about ^ 2 Ma).

4. Tectonic evolution

4.1. Oligocene extension: phase D1

In the middle and lower Aosta Valley, four sets of

extensional structures overprint the latest greenschist

facies foliations and are in turn overprinted by other

brittle structures. These systems, grouped according

to attitude, are:

(1) A system of low-angle NW- and SE-dipping

brittle±ductile detachment horizons, developing in

relatively weak rocks, such as the Piedmont

calcschists. These detachments consist of wide hori-

zons of penetrative extensional crenulation cleavage

and more localised shear zones characterised by 20- to

200-cm-thick clayey/chloritic fault-gouge layers (Fig.

9a).

(2) A system of low-angle N-dipping detachments

showing the same features as set 1. This system

mainly developed on the southern slope of the middle

Aosta Valley, locally reactivating the steep N-dipping

greenschist facies regional foliation.

(3) An intermediate to high-angle, NW- and SE-

dipping conjugate fault set, homogeneously devel-

oped all over the Aosta Valley region. These faults

show very different features, depending on the host-

rock. In carbonate rocks (e.g. Piedmont zone

calcschists), they are characterised by polished

slickensides (Fig. 9b), which always overprint the

low-angle structures, with calcite ®brous steps,

swarms of Riedel-type joints and calcite-®lled veins.

In harder rocks, such as gneisses (Austroalpine and

Penninic basement), the fault zones are characterised

by thick cataclastic layers (up to 500 m thick) with

abundant pseudotachylite-®lled joints (Fig. 9e). In

places, veining is strongly developed, indicating a

focus of hydrothermal ¯uid ¯ow owing to the

enhanced permeability of fault zones. The synkine-

matic gold-bearing quartz veins of the Brusson area

(Fig. 9c) result from this hydrothermal activity, as

well as a rather peculiar kind of hydrothermal meta-

somatism that took place where large faults cut

through serpentinites. The results of this process are

fault breccias called listvenites, characterised by

strong enrichment in Ca, K and OH, with respect to

the original composition, and by a high variability in

Si content (up to 70%; Richard, 1981; Dal Piaz and


Fig. 6. Sketch-maps of the area where structures linked to phase D1 can be continuously traced. Site-average kinematic data (attitude and vector

symbols) shown in their geographic location: (a) reactivated foliation in calcschists and shear sense of hanging-wall as deduced from

extensional crenulation cleavage (see Fig. 5a); (b) low-angle fault planes in calcschists and sense of movement of hanging-wall (see

Fig. 5b); (c) high-angle fault planes and slip vectors (see Fig. 5c). Principal high-angle fault traces are shown in (c) as thick grey lines.

Principal units as in Fig. 1.


Fig. 7. Stereoplots of site-average fault-slip data for phase D2, as deduced by slickenside lineations, Riedel-type shears, tension gashes and

pseudotachylite-®lled pull-aparts (great circles Ð fault planes; arrows Ð sense of movement of hanging-wall): (a) faults inside PGA block

(large white arrows Ð regional extension direction); (b) faults along NW border zone (black arrows Ð dextral strike-slip); (c) faults along SE

border zone (black arrows Ð left-lateral strike-slip).

Fig. 8. Sketch-map of area where structures linked to phase D2 can be continuously traced. Site-average kinematic data (see Fig. 7) are shown

in their geographic location. Dextral strike-slip prevails along the NW border-zone (Gourlay and Richou, 1983; Ramsay, 1989; Hubbard and

Mancktelow, 1992; Maurer et al., 1997); a left-lateral overall kinematics is inferred for the SE border zone. Ductile to brittle extension along

Simplon fault (NE border zone; Mancktelow, 1992) is parallel to extension within the PGA block (about 0508). Principal fault traces shown as

thick dark grey lines. Principal units as in Fig. 1.

A.

Bista

cchi,

M.

Ma

ssironi

/T

ectonophysics

327

(2000)

267

±292

281

Omenetto, 1978). The most important faults of this set

are the Ospizio Sottile, Mont GeleÂ and Trois Villes

lines (Figs. 1±3).

(4) An intermediate to high-angle N- and S-dipping

conjugate fault set, mainly running along the middle

Aosta Valley (Fig. 2), showing the same features as

set 3, but with a strong asymmetry. The master fault

(Aosta±Ranzola fault system) dips to the N and is

characterised by strong hydrothermal ¯uid ¯ow;

conversely, antithetic S-dipping faults always show

minor displacements and veining.

The kinematics of extensional crenulation clea-

vage and localised low-angle shear zones (Systems

1 and 2) is shown in Figs. 5a, 6a and 5b, 6b, respec-

tively. Microtectonic data of intermediate to high-

angle faults (Systems 3 and 4) are shown in Figs.

5c and 6c.

As we have seen, all these four sets show the same

deformation mechanisms and are associated with the

same type of hydrothermal ¯uid ¯ow; hence they may

be considered coeval. Given that thermal conditions

change continually during exhumation, similarity of

deformation mechanisms and hydrothermal mineral

assemblages may be considered evidence for

simultaneity.

No systematic cross-cutting relationships are asso-

ciated with the strike of the different sets. Structures

belonging to sets 1 and 2 are developed in different

areas, depending on the regional attitude of the

greenschist facies foliation. Faults belonging to sets

3 and 4 cut each other repeatedly, and therefore must

be considered coeval from a geometrical point of

view.

Steep faults (3 and 4) consistently cross-cut low-

angle detachments (1 and 2) in weak rocks such as

Piedmont calcschists, but the principal extension

direction, inferred from congruent kinematic indica-

tors, is everywhere horizontal NW±SE (Fig. 5a±d).

This is interpreted as the effect of the transition from

brittle±ductile to colder, truly brittle conditions

during continuing NW±SE extension.

Taking into account these considerations, the four

sets listed above developed in response to a single

deformation phase (D1), characterised by an overall

NW±SE extension (with vertical shortening) and

intense hydrothermal activity, focused along two

major faults: the Aosta±Ranzola and Ospizio Sottile

fault systems.

The Oligocene age of phase D1 is supported by:

(1) the occurrence of Oligocene (31±29 Ma) ande-

sitic-lamprophyric dikes in the inner Aosta and Sesia

valleys, synkinematically emplaced along sub-verti-

cal E- to NE-striking fractures (Dal Piaz et al.,

1979; Figs. 5d and 9d);

(2) the coeval (32±30 Ma) emplacement of


Fig. 10. Two hypothetical composite cooling curves, indicating

contrasted exhumation.

Fig. 9. (a) D1 low-angle detachments in calcschists of Piedmont zone. Shear sense is top-down to the left (NW; see Rutter et al., 1986, for

terminology of composite fabric in clayey gouge). (b) High-angle slickensides in calcschists of Piedmont zone. Same outcrop as (a) and same

extension direction (NW±SE). (c) Gold-bearing quartz vein injected into a NNW-dipping dilatant normal fault in Brusson gold district (eastern

end of the Aosta±Ranzola fault system). Sub-vertical offshoots branching from fault vein indicate NW±SE extension. (d) Lamprophyric dike

injected along an E±W fracture cross-cutting the greenschist facies foliation of Piedmont calcschists. (e) Pseudotachylite fault and injection

veins along Ospizio Sottile fault, (hand sample; see Swanson, 1992, for terminology of pseudotachylites). (f) Dextral reactivation along

Penninic frontal thrust, revealed by very penetrative en-echelon quartz±calcite vein arrays.


Fig. 11. Cooling rate contours at: (a) 30 Ma (D1), (b) 26 Ma (D1±D2 transition), (c) 12 Ma and (d) 0 Ma (D2). Principal lineaments as in Fig. 1.

Stars: sampling sites, corresponding to gridded data-points of contour maps. Contouring area limited by availability of reliable age determina-

tions.


Fig. 11. (continued)

synkinematic gold-bearing quartz veins along high-

angle E- to NE-striking normal faults and tension

gashes in the Brusson area (Diamond, 1990; Figs.

5d and 9c); ¯uids related to these veins are also

responsible for listvenitic metasomatism all along

the Aosta±Ranzola and Ospizio Sottile faults;

(3) the steep cooling rate gradient across the Aosta

Valley, con®rming at 30 Ma fast cooling of the blocks

to the south of the Aosta±Ranzola fault system and

slowest cooling in correspondence to the Dent

Blanche nappe (Fig. 11a);

(4) ZFT ages, which record cooling under 2508C at

about 32 Ma throughout the Aosta Valley region

(Hunziker et al., 1992; Seward and Mancktelow,

1994); this cooling age is consistent with the

ductile-to-brittle transition on NW- and SE-dipping

faults in calcschists of the Piedmont ophiolitic units,

since 2508C is an accepted estimate for the brittle±

ductile transition in carbonate rocks (e.g. Carter and

Tsenn, 1987).

If extensional activity along the sets listed

above can be ascribed to a single tectonic phase,


Fig. 12. Block diagram (a) showing proposed faulting model for phase D1 (Oligocene). Higher/lower elevations qualitatively indicate higher/

lower exhumation rates, deduced from cooling rates (elevation of topographic surface is not speci®ed). Orthorhombic symmetry of faults

evidenced in stereoplot (b).

synchronous displacements along these four conju-

gate sets (N-, S-, NW-, SE-dipping) can be interpreted

with the 3D faulting model of Aydin and Reches

(1982). According to this model, conjugate displace-

ments along four sets, arranged in orthorhombic

symmetry (Fig. 12), take place in response to a ªtrue

triaxialº strain ®eld, where all principal strains

�ex; ey; ez� are different from zero (Reches, 1978;

Aydin and Reches, 1982; Reches and Dieterich,

1983; Reches, 1983; Krantz, 1988). This model is

more general than the Anderson (1951) classical

model, which predicts displacement along two conju-

gate faults in the restrictive condition of plane strain

�ey � 0�; therefore, ®nding evidence of triaxial strain

is not surprising.

Applying the triaxial model to the middle Aosta

Valley, the maximum extension axis �ex� is inferred

to be NW±SE oriented, the intermediate axis �ey� is

positive (extensional) and NE±SW oriented, and

the shortening axis �ez� is vertical (ªodd axisº of

Krantz, 1988; Fig. 12). Therefore, the regional strain

ellipsoid is oblate, with a NW±SE-directed maximum

extension.

As previously noted, the fastest cooling domains

were located south of the Aosta±Ranzola fault system

(lower Champorcher Valley), whereas the area with

the slowest cooling was the Dent Blanche nappe,

bordered on both sides by antithetic NW- and SE-

dipping extensional structures. Steep cooling rate

gradients may be recognised along the Aosta±

Ranzola and Ospizio Sottile faults, which are also

characterised by hydrothermal ¯uid ¯ow; hence,

these two faults are considered the two most important

structures for the Oligocene phase D1. The block

diagram of Fig. 12 shows the proposed tectonic

model, which satis®es both the cooling-rate pattern

and the orthorhombic faulting model.

4.2. Age of back-thrusts and back-folds

As mentioned in the introduction, different authors

do not agree on the age of the last ductile deforma-

tions in this part of the NW Alps. The latest ductile

deformations are represented by open folds related to

ductile thrusts, with a transport direction opposite to

that of the overall vergence of the nappe stack (back-

folds and back-thrusts; Argand, 1911, 1916).

Hurford et al. (1991) suggest that the Miocene

differential exhumation, recorded by AFT age distri-

bution, is related to the evolution of these large-scale

ductile back-folds and therefore favour the develop-

ment of a compressional tectonic phase at the Oligo-

cene±Miocene boundary, as proposed by Argand

(1916).

However, since in this temperature range (110±

608C of the AFT partial annealing zone) no ductile

deformation would occur in basement rocks (Carter

and Tsenn, 1987), the back-thrust and back-fold phase

should be older than D1. This is con®rmed by cross-

cutting relationships and white mica Rb/Sr and Ar/Ar

ages (36±34 Ma) from the Entrelor back-thrust

(Freeman et al., 1997) and the Mischabel back-fold

(Barnicoat et al., 1995). The brittle±ductile transition

therefore probably took place in the Oligocene along

the middle Aosta Valley (at about 32 Ma), after the

greenschist facies back-thrusting and back-folding of

the middle Penninic Grand St. Bernard nappe.

Since the NE end of the Canavese line (E of the

Ossola Valley) acted as a dextral transpressive fault

during the Oligo-Miocene over-steepening of the

ªroot zoneº (caused by the NW motion of the South-

alpine indenter at 26±20 Ma; Schmid et al., 1989;

Zingg and Hunziker, 1990), this kinematic behaviour

may also be assumed for the SW part of the line.

Nevertheless, ®eld evidence shows that the Canavese

line mylonites between Ivrea and the Ossola valley are

overprinted by high-angle brittle normal faults asso-

ciated with large cataclastic layers. These are the only

structures that cut the Oligocene volcanoclastic cover

of the inner Sesia Lanzo zone. Thus, from the Oligo-

cene onwards, activity along the SW portion of the

Canavese line was extensional and, along this fault,

back-thrusting took place before the Oligocene.

In conclusion, all the ductile shortening events took

place, in the Aosta Valley region and surroundings,

before about 33 Ma and the extensional D1 phase.

This reconstruction is con®rmed by the gradual tran-

sition between cooling patterns characteristic of each

phase, shown by the cooling-rate contour map calcu-

lated between D1 and D2 (e.g. 26 Ma, Fig. 11b).

4.3. Miocene to Present lateral extrusion: phase D2

From the early Miocene to the Present, continuous

lateral extrusion has been active from the Simplon

fault to the SW, involving the whole Pennine-Graian


Alps nappe stack which, from then on, may be consid-

ered as a continuous and almost homogeneously

deforming fault-bounded block (PGA in Fig. 13;

Hubbard and Mancktelow, 1992; Ring and Merle,

1992; Bistacchi et al., 2000). The discontinuities

bordering the PGA block are (Figs. 8 and 13): (1) to

the NE, the Simplon fault; (2) to the NW, the Rhone

line (between Brig and Martigny), Chamonix line,

Penninic frontal thrust, BriancËonnais frontal thrust,

and some high-angle fault zones between them; and

(3) to the SE, the Ospizio Sottile fault.

The NE extensional edge of the SW-escaping PGA

block (Simplon fault) underwent continuous extension

from the Miocene onwards (Mancktelow, 1992). The

Rhone line has a well-documented dextral strike-slip

between Brig and Martigny (Hubbard and Manckte-

low, 1992; Maurer et al., 1997). More to the SW it

extends into the Chamonix line, with a 208 counter-

clockwise bend resulting in a dextral transpressive

movement (Gourlay and Richou, 1983; Ramsay,

1989; Hubbard and Mancktelow, 1992). South-west

of the Piccolo S. Bernardo pass area, the low-angle

Penninic and BriancËonnais frontal thrusts were reac-

tivated as extensional detachment horizons (AilleÁres

et al., 1995; Cannic et al., 1995, 1999) and some high-

angle normal faults between these two thrusts have

been active in the last few million years (FuÈgenschuh

et al., 1999). Perello et al. (1999) have recently docu-

mented post-metamorphic dextral transpressive activ-

ity of the Penninic frontal thrust in the Courmayeur

area. Our ®eldwork has revealed late right-lateral

reactivation of the steep E-dipping slaty cleavage of

Ultrahelvetic Jurassic slates in the Ferret and Veny

valleys (Fig. 9f), as well as the presence of some

high-angle brittle faults dissecting the outer Penninic

cover sheets (Figs. 7b and 8). Within this kinematic


Fig. 13. Block diagram showing faulting model proposed for Miocene to Present orogen-parallel extrusion: while PGA block underwent more

or less homogeneous extension, Simplon sub-dome and Mont Blanc massif were exhumed faster. The latter, in transpressive conditions, may be

interpreted as a positive ¯ower structure. Higher/lower elevations qualitatively indicate higher/lower exhumation rates, deduced from cooling

rates (elevation of topographic surface is not speci®ed).

framework, the Helvetic basement units of the Mont

Blanc massif may constitute a positive ¯ower struc-

ture developing under a Neogene transpression (Fig.

13).

The principal SE edge of the PGA block is the

Ospizio Sottile line, reactivated as a sinistral transcur-

rent fault after its normal activity during D1 (Figs. 7c

and 8). Kinematic indicators, such as calcite steps in

carbonatic rocks, pseudotachylite-®lled pull-apart

structures and asymmetric injection veins (in

Austroalpine gneisses; Fig. 9e) consistently show its

sinistral strike-slip activity (Bistacchi et al., 2000).

The internal deformation of the PGA block is

characterised by a regular array of NE- and SW-

dipping normal faults, indicating a NE±SW extension

direction from the Simplon fault to the Gran Paradiso

massif. A very penetrative set of sub-vertical exten-

sional joints, striking NW±SE and extending from the

Valais to at least the Piccolo S. Bernardo Pass, is also

consistent with this extension direction (Figs. 7a and

8). Some E±W striking structures, partly inherited

from D1, are also consistent with the internal defor-

mation of the PGA block. One of the most important

is the Aosta±Ranzola fault system, which, to the W,

extended as a slightly diverging splay of mainly

dextral strike-slip faults (Aosta±Piccolo S. Bernardo

fault system; Fig. 8).

The spatial arrangement of cooling-rate contours

shows that the Lepontine dome and the Mont Blanc

massif cooled faster than the neighbouring areas from

about 26 Ma onwards. Undisturbed cooling-rate

contour lines cross-cut the Ospizio Sottile fault,

showing no evidence for any along-dip displacement

(Fig. 11c and d). This cooling pattern may be a result

of differential exhumation of fault-bounded blocks, as

shown in the tectonic model of Fig. 13. The distribu-

tion of microseismicity and fault plane solutions

indicates that this orogen-parallel extrusion is still

active (Bistacchi et al., 2000).

5. Discussion

In many sectors of the Alps, a post-collisional

extension has been referred to an orogen-parallel

extrusion, as in the case of phase D2 (Selverstone,

1988; Ratschbacher et al., 1989; Hubbard and

Mancktelow, 1992; Ring and Merle, 1992; Meyre et

al., 1998; Bistacchi et al., 2000). However, the Oligo-

cene D1 NW±SE extension, developed along four

conjugate fault sets with orthorhombic symmetry

(Fig. 12), is characterised by an oblate strain ellipsoid,

indicating net extension in all horizontal directions.

During this phase, the main extension direction,

which is perpendicular to the belt axis, apparently

contradicts the continuing convergence of European

and Adriatic plates in the Oligocene, supported by

thrusting in the Southern Alps and Helvetic domains

and by global plate-tectonic reconstructions (e.g.

Laubscher, 1985; Coward and Dietrich, 1989; Platt

et al., 1989). However, the concomitant occurrence

of extension and shortening in axial and marginal

areas, respectively, is known in other collisional

belts and is generally ascribed either to gravitational

instability arising from over-thickening of the litho-

sphere (Molnar and Lyon-Caen, 1988) or to isostatic

rebound following a slab breakoff (Dal Piaz and

Gosso, 1994; Von Blanckenburg and Davies, 1995).

With regard to the timing of D1 and D2, we recall

that a 2-Ma delay is expected between the onset (or

end) of a given tectonic phase and its thermal effect.

Hence, the transition from D1 to D2 should be earlier

by the same amount with respect to the change in the

overall cooling pattern, evidenced by cooling-rate

contour maps, at about 26 Ma. Therefore the D1±D2

transition may be placed at about 28 Ma. Dating the

end of phase D2 may also be affected by a similar

delay, but Bistacchi et al. (2000) showed, on the basis

of seismotectonic data, that this phase is still active.

6. Conclusions

In this paper, different datasets are compared in

order to reconstruct the brittle tectonic evolution of

the axial part of the north-western Alps. Remote

sensing was found to be a useful tool for preliminary

tectonic analysis of such a large area. Microtectonic

analysis is con®rmed as the fundamental basis for any

tectonic reconstruction, and thermochronology is very

useful to constrain differential exhumation of large

fault-bounded blocks. The cooling-rate contour map

approach, although affected by several uncertain-

ties, provides a synthetic view of the over-time

evolution of the regional cooling pattern. Since it

®ts independent structural data (in the present


study, areas of steep cooling rate gradient correspond

to most important normal faults), the cooling rate

pattern is considered a good approximation of the

exhumation pattern.

All these data concur to constrain the multi-phase

kinematic model for the north-western Alps from the

Oligocene onwards, which developed through the

following stages:

(1) ductile back-thrusting and back-folding (before

32 Ma);

(2) Oligocene extension perpendicular to the belt

axis (32±28 Ma, D1);

(3) D1±D2 transition at about 28±26 Ma;

(4) Miocene±Present orogen-parallel extrusion

(26±0 Ma, D2).

Acknowledgements

The authors would like to thank Prof. Giorgio

Vittorio Dal Piaz and Drs. Antonio Germani, Bruno

Monopoli, Silvana Martin, Giorgio Pennacchioni,

Alessio Schiavo and Massimiliano Zattin for useful

discussions, encouragement and support during ®eld-

work. Profs Jean-Pierre Burg, Neil Mancktelow and

Lothar Ratschbacher are acknowledged for careful

revisions. Prof. Paolo Baggio kindly provided access

to computer facilities at LAT (University of Padova).

ERS-GTC data were studied at ASI-CGS (Matera)

with the support of I-PAF. Nicola Michelon and

Stefano Castelli are acknowledged for technical and

graphical support; Gabriel Walton revised the English

text.

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