Slope instability in relation to glacial debuttressing in alpine areas (Upper Durance catchment,...

transcript

Available online at www.sciencedirect.com

(2008) 3–26www.elsevier.com/locate/geomorph

Geomorphology 95

Slope instability in relation to glacial debuttressing in alpine areas(Upper Durance catchment, southeastern France): Evidence from

field data and 10Be cosmic ray exposure ages

E. Cossart a,⁎, R. Braucher b, M. Fort a, D.L. Bourlès b, J. Carcaillet b

a Equipe DYNMIRIS, UMR PRODIG — 8586 CNRS, Universités Paris 1-7, 2 rue Valette, 75005 Paris, Franceb CEREGE, UMR 6635— CNRS, Université Aix-Marseille 3, Europôle Méditerranéen de l'Arbois, BP 80, 13545 Aix en Provence cedex 04, France

Received 10 June 2005; received in revised form 30 January 2006; accepted 16 December 2006Available online 29 April 2007

Abstract

The Upper Durance catchment is an area prone to rock-slope failures. Such failures reflect the combination of high relief, litho-structural controls and paraglacial stress release. The aim of this study is to determine the role of deglacial unloading and resultingparaglacial stress release in conditioning or triggering slope failure. Former dimensions of the Durance glacier are reconstructed,then combined with Digital Elevation Model data in a raster Geographic Information System to quantify the spatial pattern ofstresses associated with glacial loading at the Last Glacial Maximum. Preliminary calculations suggest that major rock falls androck avalanches are associated with areas subject to the highest decompression stresses. Focus on two case studies allows theconsequences of paraglacial stress release on slope instability to be evaluated. Description of slope failure runout deposits allowsreconstruction of the nature of slope failure. Surface exposure dates based on concentration of cosmogenic 10Be allows the timingof both deglaciation and that of post-glacial rock-slope failures to be established. It is shown that rock-slope failures areconcentrated on lower valley-side slopes within the area occupied by ice at the Last Glacial Maximum, and that their locationscoincide with zones of inferred high glacial loading stress, consistent with interpretation of both bedrock disruption and large-scalerock-slope failures as paraglacial phenomena induced by stress release following deglaciation. Timing of initial rock avalancherunout deposition at one site is consistent with this conclusion, though later instability episodes at the same site may have occurredindependent of the influence of paraglacial stress release.© 2007 Elsevier B.V. All rights reserved.

Keywords: Paraglacial; Rock-slope failure; Glacial debuttressing; Stress release; Cosmogenic nuclides; Late Glacial; French Alps

1. Introduction

The paraglacial concept was first conceived in analluvial context, referring to “non glacial processesconditioned by glaciation” (Ryder, 1971; Church and

⁎ Corresponding author. Tel.: +33 1 44 07 75 83.E-mail address: etienne.cossart@univ-paris1.fr (E. Cossart).

Ryder, 1972). However, over the last 30 years, a numberof investigations have extended this concept to othercontexts, such as hillslopes, glacier forelands and thecoastal zone (see Ballantyne, 2002a,b). Recently, glacialdebuttressing has been increasingly recognized asimportant in conditioning rock-slope instability (e.g.Bovis, 1990; Cruden and Hu, 1993; Shakesby andMatthews, 1996; Ballantyne et al., 1998; Ballantyne andStone, 2004; Holm et al., 2004; Matthews and

4 E. Cossart et al. / Geomorphology 95 (2008) 3–26

Shakesby, 2004), in part because slope failures appear tobe particularly common in deglaciated mountain areas.The relaxation of internal stresses after deglaciation(Evans and Clague, 1994; Augustinus, 1995) and theassociated propagation of internal joint networks(Ballantyne et al., 1998) are considered to result intime-dependent rock-slope weakening. However, rock-slope instability is also associated with tectonic stressesor river incision in tectonically-active mountains (Fort,2000, 2003; Korup, 2004), and rock-slope failure insuch environments is often triggered by seismic events.In tectonically-active alpine mountains, the role ofdeglacial unloading (debuttressing) in promoting rock-slope failure is difficult to determine, and the factorsresponsible for triggering failure are often unknown (DeJong et al., 1995; Korup, 2004; Seijmonsbergen et al.,2005). The role of glacial unloading on rock-slopeinstability thus requires further investigation: in partic-ular, what are the effects of glacial unloading on bedrockcohesion? And can glacial unloading itself constitute afailure trigger?

Previous studies have mainly documented the possiblerole of debuttressing and paraglacial rock-slope failure attwo different scales. At the smaller scale, Lewis (1954)noted that glacial retreat may be associatedwith shatteringor disruption of glacially-polished bedrock surfaces.Lewis interpreted displacement along joints as due toglacial unloading because (i) joint geometry was notcontrolled by the local structural pattern and (ii) manyjoints formed parallel to the former glacier bed, consistentwith stress release following deglacial unloading. Otherstudies have shown that a parallelism between fractureorientation and glacier valley topography is indicative offracture resulting from paraglacial stress release (e.g.Bovis, 1990; Ballantyne and Stone, 2004).

At a larger scale, major rockslides are thought to bethe result of debuttressing following glacial recession. Ifthis is the case, then the age of such rock-slope failuresshould be related to the timing of deglaciation: Crudenand Hu (1993) and Ballantyne (2002a,b) have suggestedthat paraglacial evolution of rock slopes is characterizedby decreasing frequency of rock-slope failures sincedeglaciation, and have proposed that the frequency ofrock-slope failures may be described by a negativeexponential exhaustion model. In contrast, slopeinstability driven by tectonic stress and the timing ofseismic events implies that the frequency of rock-slopefailure is independent of time elapsed since deglaciation.

Two main approaches have been applied to evaluatechanges in rock-slope stability since deglaciation.Firstly, some studies have estimated long-term rock-fall accretion rates on talus slopes in deglaciated areas

(Luckman, 1988; Luckman and Fiske, 1995; André,1997; Hétu and Gray, 2000). Such studies havecombined estimates of recent rock-fall rates or talusaccumulation rates with estimates of talus volume tocompare recent accretion rates (hundred-year scale) withformer rates (Holocene and post-glacial thousand-yearscale). The results suggest that accumulation of talusdeposits occurred rapidly after glacial retreat, and thatpresent rates of talus accumulation are much slower thanEarly Holocene rates. This pattern is consistent with theparaglacial exhaustion model. Secondly, the relationshipbetween the timing of deglaciation and that ofsubsequent rock-slope failures may also be investigatedto test the exhaustion model (Bovis and Jones, 1992;Abele, 1997; Flageollet et al., 1999; Corsini et al., 2001;Soldati et al., 2004). Such studies rely on establishing aninventory of mass-movements and their timing, basedupon numerous 14C dates. If the distribution ofradiocarbon ages is clustered soon after deglaciation(rather than being randomly scattered throughout theHolocene) then it indicates a probable link betweenslope failure occurrences and deglaciation.

However, because of the scarcity of organic matter inmany alpine areas, and the difficulty of retrieving organicmatter associated with rock-slope failures, surface expo-sure dating using cosmogenic isotopes offers a moresatisfactory and widely-applicable technique for establishthe ages of post-glacial rockslides (Ballantyne et al., 1998;Ivy-Ochs et al., 1998; Kubik et al., 1998; Ballantyne andStone, 2004; Kubik and Ivy-Ochs, 2004). Surfaceexposure dating using cosmogenic isotopes is particularlyuseful as it also offers a means of establishing the age ofdeglaciation (e.g. Brook, 1994; Gosse et al., 1995; Feiss,2000; Owen et al., 2001; Bourlès et al., 2004), thusproviding the opportunity to assess the time-lag betweendeglaciation and rock-slope failure (Ballantyne et al.,1998; Ballantyne and Stone, 2004). Cost constraints meanthat this approach has to be applied selectively, once theevidence of glacial debuttressing has been establishedusing geomorphic criteria.

Geomorphic evidence for a paraglacial (debuttressingand stress release) origin of rock-slope failure is indicatedby spatial coincidence between the location of failuresand areas where glacial debuttressing was at its maximum(Panizza, 1973; Holm et al., 2004). The use of GIS makesthis comparison easier, but such an approach must bebased upon an evaluation of glacial stress release, esti-mated from the basal stresses applied by a former glacieron adjacent bedrock. This latter parameter varies at threescales (Benn and Evans, 1998). First, at a regional scaleit is a function of ice thickness and former ice surfacegradient; second at the scale of individual slopes, basal

5E. Cossart et al. / Geomorphology 95 (2008) 3–26

stress is at its maximum close to the valley floor; andthird, at a more local scale, basal stress is higher on thestoss face of glacial rock-bars. Thus, coincidences be-tween glacial debuttressing and areas of rock displace-ment must be verified at all three scales.

This paper couples reconstruction of stresses at all threescales with a chronological framework for deglaciationand rock-slope failure for the Upper Durance catchment(Southern French Alps). We first document the debuttres-sing forces implied by former glacial extent and theapplication of glaciological laws within a GIS frame.Secondly, at a regional scale, we inventory landslides andcompare their locations with areas where both normaland longitudinal stresses (applied by the former glacier)were at maximum. Thirdly, we select sites (i) to considerwhether slope failures have been preferentially triggered

Fig. 1. The Upper Durance catchment. (A) General location in southeastern3 = main ridges; 4 = pass. Note that the highest altitudes are located in the w(NTF, Lambert III, units: meters).

on the lower part of the slope, where glacially-imposedstresses were highest, and (ii) to compare the jointingpatterns of bedrock on the stoss faces of rock-bars (highimposed stress) with those on lee faces (low imposedstress). Finally, some chronological benchmarks providethe opportunity to testwhether the timing of slope failure isconsistent with proposed paraglacial models of slopeevolution.

2. Physical setting

2.1. Controls on slope instability in the Upper Durancecatchment

The study area is located in the Upper Durancecatchment in the Southern Alps of France and comprises

France. (B) The study area covers 1011 km2. 1 = town; 2 = summit;estern part. Geographical coordinates refer to French geodesic norm

a 600-km2 block of rugged terrain (Fig. 1a). This areaexhibits many large rock-slope failures (Lahousse,1994; Breteaux, 1998; Colas, 2000) and was formerlyextensively glacierized.

The topography is characterized by a marked contrastbetween the western and eastern parts of the studyarea (Fig. 1b). The western part is higher with summitslocally reaching 4000 m (Barre des Ecrins, 4102 m a.s.l.)and very rugged terrain. The Vallouise valley is typical ofthis zone: slope gradients vary from 80 to 130% (over50 m long segments) and differences in elevationbetween summits and valley floor exceed 2500m aroundBarre des Ecrins or Mont Pelvoux. In contrast, theeastern zone has a lower and gentler relative relief: only afew summits reach 3000 m (e.g. Pic de Rochebrune,3320 m a.s.l., or Font Sancte, 3350 m a.s.l.). In Claréevalley, for example, only 6 summits are above 3000 m a.s.l., and slope gradients range from 50 to 80%.

There are also marked differences in geology(Fig. 2). The western part lies within the alpine outercrystalline zone where granites and gneisses are themost frequent bedrock lithology (Fig. 2b). These

Fig. 2. Geological setting. (A) Lithological setting of the Upper Durance catc(massive limestones), internal zone: 3 = limestones and schists; 4 = conglomedolomites; 7 = “Schistes-Lustrés”; 8 = ophiolite complex. Note the wide varieeasier: carboniferous sandstones (Upper Clarée), gneisses and granites (Up(Upper Guil and Cerveyrette) were essentially used to reconstruct palaeo1 = stream; 2 = main ridges; 3 = hypsometric curve; 4 = fault-line; 5 = limestoand structure of the upper Clarée catchment; the valley is mainly cut within

highly cohesive rocks, together with high uplift rates,have formed high relative relief, which may haveinfluenced slope instability. The eastern part belongsto the alpine inner geological zone, which includeslimestones, sandstones and schists (Fig. 2c). In thiszone, instability reflects lithological setting: weakbedrocks, such as “schistes lustrés” (highly-foliatedcalcschists) or poorly cohesive carboniferous sand-stones, are particularly prone to failure (Barféty et al.,1995).

Two case studies were selected to reflect probabledifferences in the effects of glacial debuttressing ineastern and western areas of the Upper Durance catch-ment. The Vallouise (western) case study is well suitedto assess the importance of glacial debuttressing in in-fluencing slope failure in an area of steep, rugged terrainunderlain by resistant lithologies. By contrast, theClarée area provides the opportunity to document theconsequences of glacial unloading on bedrock jointing;in particular, sandstones in this area have proved sen-sitive to shattering but are sufficiently massive to havewell recorded many stages of jointing.

hment. External zone: 1 = granites and gneisses; 2 = sedimentary unitsrates and sandstones; 5 = carboniferous sandstones; 6 = limestones andty of lithological units which made the identification of glacial erraticsper Gyr, Onde and Guisane), “green-rocks” from ophiolite complex-glacier fluxes. (B) Lithological setting of the Vallouise catchment.nes; 6 = sandstones; 7 = granites; 8 = diorites; 9 = gneiss. (C) Lithologyweak carboniferous sandstones.

2.2. Timing of deglaciation

Present ice cover is limited in extent, being locatedmainly in the western part of the area (Massif des Ecrins)and is approximately 25 km2 in extent. The Duranceglacier, one of the main glaciers within the French Alpsduring the Last Glacial Maximum, extended down toSisteron. Current knowledge of the pattern of deglaciationis deduced from geomorphic mapping of glacial remnantsand from radiocarbon dates. The Ubaye and Verdonvalleys (southern flank of the study area) were probablyfree of ice between 20 and 15 14C ka BP (Montjuvent,1973; Jorda, 1988, 1993; Jorda et al., 2000; Rosique,2004), while in the western part of the Massif des Ecrins(northwestern flank of the study area), some radiocarbondates imply that the glacier front stayed at about 1800 m a.s.l. until the Early Holocene (Edouard, 1979; Coûteaux,1983a,b,c; Coûteaux and Edouard, 1987; Chardon, 1991).Further dating evidence is therefore required to associate

Fig. 3. Methodology used to assess the glacial thickness and the debuttressingthe raster GIS software; 3 = results. (B) During the glaciation, the bedrockdeglaciation, the bedrock reaction is still sensitive, generating the deglacial s

the timing of deglaciation in the study area with one ofthese two conflicting chronologies.

3. Methods

3.1. Reconstruction of former glacial extent

Prior to any consideration of paraglacial effects, theformer extent and age of glaciation had to be establishedby mapping glacial landforms and deposits, especiallyerratics and roches moutonnées. It is assumed that anyglacial evidence pre-dating the Last Glacial Maximumwas not preserved on the rocky slopes of this highlygeodynamically active environment. The wide varietyof lithologies in the Upper Durance catchment made theidentification of erratics straightforward. For instance,sandstone boulders (carried from upper Clarée catch-ments), granite and gneiss boulders (carried from theupper Guisane and Vallouise catchments), and ophiolite

forces. (A) Geomatic processing: 1 = database; 2 = operator used withinis reacting against the glacial pressure: an equilibrium exists. Aftertress release.

boulders (carried from eastern parts of the study area)can be employed as sedimentological tracers for thereconstruction of directions of former ice movement.

The elevations of erratics and other glacial remnantswere carefully checked within a GIS framework to as-sess the minimal altitude reached by glacier ice at theLast Glacial Maximum (Fig. 3a). On the basis of thesedata, hypsometric curves representing the former icesurface level were computed every kilometer. To inter-polate the altitude reached by the former glacier surfaceacross the whole catchment we performed a Delaunaytriangulation (TIN Model) based on these curves(Appendix A). The results were then combined with aDEM (provided by Institut Géographique National),representing the present topography. The minimumthickness of the former glacier was calculated by sub-tracting the altitude of the present terrain surface fromthat of the reconstructed glacier (Fig. 3).

The accuracy of this calculation depends on threeconsiderations. The first is the vertical resolution ofthe DEM of present topography (b10 m). The second isthe amount of post-glacial sediment aggradation on thevalley floor; according to Barféty et al. (1995), thisamounts to 10–70 m, implying that immediately afterdeglaciation the valley floor was markedly lower thannow. The third is the accuracy of the glacier-surface

Table 1Estimations of beryllium production rates and exposure ages

Longitude Latitude Altitude(m)

Samplethicknesscorrection (cm)

CLA_04_01 06°35′16″ 45°01′14″ 1768 0 0CLA_04_02 06°35′4.5″ 45°01′14.2″ 1785 0 0CLA_04_03 06°35′4.5″ 45°01′14.2″ 1783 0 0CLA_04_04 06°32′46.8″ 45°01′45.3″ 1979 0 0CLA_04_05 06°31′36.1″ 45°01′28.9″ 2348 0 0CLA_04_06 06°31′36.1″ 45°01′28.9″ 2348 0 0CLA_04_07 06°31′41.7″ 45°03′33.7″ 2036 0 0CLA_04_08 06°31′41.7″ 45°03′33.7″ 2034 0 0CLA_04_09 06°31′07.6″ 45°03′32.5″ 2065 0 0CLA_04_10 06°31′07.6″ 45°03′32.5″ 2060 0 0CLA_04_11 06°33′57.0″ 45°01′25.4″ 1817 0 0CARL_04_01_inf 06°25′50.9″ 44°54′49.8″ 1913 504 0CARL_04_01_sup 06°25′50.9″ 44°54′49.8″ 1913 0 0CARL_04_02_inf 06°25′52.1″ 44°54′47.3″ 1900 822 0CARL_04_02_sup 06°25′52.1″ 44°54′47.3″ 1900 0 0CARL_04_03_sup 06°25′50.5″ 44°54′46″ 1895 0 0CARL_04_04_inf 06°25′49″ 44°54′43.8″ 1870 424 0CARL_04_04_sup 06°25′49″ 44°54′43.8″ 1870 0 0CARL_04_05_inf 06°25′48.7″ 44°54′43.3″ 1862 331 0CARL_04_05_sup 06°25′48.7″ 44°54′43.3″ 1862 0 0GYR_04_01 06°25′48.4″ 44°51′57.1″ 2175 0 0

For each site of the Pré de Madame Carle (CARL), two samples were extractethe bottom face (inf). Some dates were discarded because of bore pollution

reconstruction. Because of reworking of glacial evi-dence (moraines and erratics) it is possible that icesuface altitude at the Last Glacial Maximum has beenunderestimated. The calculated thickness of former gla-cier cover is thus a minimal value, but unlikely to under-represent true ice thickness by more than 100 m.

The gradient of glacier surface is another parameterto consider because it influences both basal shear stress(Paterson, 1981; Lliboutry, 2002) and former glacierdischarge (MacGregor et al., 2000). The calculation offormer ice surface gradient was deduced from thealtitude of the glacier surface within the raster GIS(Fig. 3b). Both normal (σ) and longitudinal (τ)? stresses(Paterson, 1981; Lliboutry, 2002) may be estimated asfollows within the raster GIS:

r ¼ qgh ð1Þ

s ¼ ðq g h sin aÞ f ð2Þwhere σ is the normal stress and τ the longitudinal stress(kPa), ρ is ice density (900 kg m−3), g is gravitationalacceleration (9.81 m s−2), h is former ice thickness (m)and α is former glacier surface gradient (°) and f is anempirical constant (Bindschadler et al., 1977). Thelongitudinal stress is derived from the basal shear stress,

creening Assumed rateof 10Be production(atom g−1 yr−1)

10Beconcentration(105 atom g−1)

Exposureage (yr)

Ageuncertainty(yr)

.89 20.23 1.090 5393 ±1081

.95 21.84 Bore pollution

.96 21.92 2.290 10,472 ±1257

.98 26.21 7.503 28,813 ±5604

.99 34.87 2.895 8319 ±1429

.99 34.87 2.753 7911 ±946

.98 27.30 2.752 10,106 ±2060

.98 28.01 2.371 8481 ±1218

.97 22.87 1.959 8579 ±1076

.95 23.98 1.12 1832 ±326

.98 24.44 1.51 10,635 ±4400

.98 24.44 2.59 33,940 ±11,540

.91 22.69 1.47 6489 ±1923

.90 22.04 3.11 1411 ±283

.92 22.73 2.51 9155 ±2713

.92 22.73 1.46 6481 ±1093

.92 28.29 2.10 7420 ±1079

d; the first one on the upper face of the boulder (sup), the second one onin the samples.

Fig. 4. Selected sampling sites for cosmic ray exposure dating ofglacial retreat. (A) Roches moutonnées in the Clarée valley. Samplingsite is at the highest point next to the steep, rocky lee face, to minimizethe possibility of till deposits preservation and subsequent shadingeffect following deglaciation. (B) Rock avalanche deposit (Pré deMadame Carle, 1930m). The surface (Sup) and the bottom (Inf) facesof large boulders (5 to 10-m long) were sampled to avoid influence byprior exposure to cosmic rays before slope failure (photograph by M.Fort, June 2004).

and accounts for the compressive effect due to glaciergeometry. The factor f is equal to 1 for an unconfined icesheet and decreases with increasing confinement; thevalue of 0.5 is applied in many studies of alpine glacialtongues (Kerschner et al., 2000).

If it is assumed that deglacial stress release is directlyproportional to the stress applied by the former glacier(Fig. 3b), a map of potential stress release can begenerated from normal-stress mapping. Another map ofthe longitudinal stress applied by a former glacier canalso be computed from the combination of ice thicknessmap with gradient of ice surface map. These maps werethen compared with the distribution of areas showingevidence for rock-slope failure.

3.2. Identification of post-glacial debuttressing impacts

3.2.1. Slope instability survey at a regional scaleSlope failures were surveyed in the field and mapped

at a regional scale, with the help of aerial photographs(1:17,000 scale) and analysis of geological maps (Barfétyet al., 1995). Three landslide types were differentiated:major rock falls or rock avalanches, translational rockslides and rock topples, and earthflows. Rock avalanchesand major rock falls are defined by large volumes (105

to 106 m3) of rock falling freely or bouncing down a cliff.Translational rock slides and rock topples typically sup-ply a volume of around 101 to 102 m3 of rock debris.Earthflows involve slow-to-rapid downslope movementof debris moving as a very viscous fluid. For each typeoccurrence, the extent of the failure zone and debris runoutzone were mapped and integrated in the data set.

At a regional scale, the identification of areassupporting a high density of slope failures is the primarycriterion for assessing whether these are essentiallyparaglacial (glacially-conditioned) in origin. We assumethat paraglacial slope failures were most likely to havebeen initiated where subglacial stresses and consequent-ly the resulting debuttressing (stress-release) forces wereat maximum, for instance at the confluence of majorglacial troughs (Panizza, 1973) or below trimlines(Bovis, 1990). The elevation of failure sites was there-fore plotted and compared with that of local trimlines.We then determined whether the areas affected by bothhighest normal and longitudinal buttressing stressescoincide with the distribution of landslides.

3.2.2. Smaller-scale effects of post-glacial stress releaseThe above regional-scale investigations were sup-

ported by local-scale studies of bedrock response todeglacial stress release. At more local scales, two dif-ferent settings were considered, depending on strength

of the underlying bedrock. First, outcrops of relatively-weak bedrock are assumed to record the most visibleimpacts of glacial unloading such as bedrock shattering.The study sites were chosen to minimize differences ingeology and inherent variability in joint pattern anddensity by selecting limited areas of homogeneousbedrock. In some cases an entire glacially-polishedvalley slope was investigated to compare the response ofglacially-polished surfaces located where ice loadingand consequent stress-release forces were high, withothers where ice loading and consequent stress-releaseforces were low. Comparison of joint development inglacially-polished rock surfaces was carried out at bothhillslope scale and through investigation of individualroches moutonnées. At the hillslope scale, we comparedthe pattern of bedrock disruption on lower slopes withthat on the upper slopes, close to the trimline. At the

roches-moutonnées scale, we compared the geometry ofjoints on the stoss face (highest basal stresses), with thaton the lee face (lowest basal stresses). Joint orientationswere measured in the field with a precision of 1° andtheir orientation frequency plotted to discriminatebetween joint orientations consistent with structuralcontrol and neo-joints produced by stress release. Wethen attempted to interpret the origin of non-structuralfractures. In particular, parallelism between dominantfracture orientation and glacier bed topography is animportant indication of joint formation through para-glacial stress release (Bovis, 1990). If such joints arelocated in areas where glacial loading or buttressingstresses were high, the hypothesis of rock-mass weak-ening through the development of paraglacial stress-release joints is supported.

Fig. 5. Glacial extent during the Late Glacial Maximum. (A) Minimal altimapping (erratics and polished surfaces were especially considered). Formeglacier surface from the present hypsometry (DEM provided by Institut Géogstress, proportional to ice thickness, can then be estimated. Note that the max1000 m. 1 = summit; 2 = main ridges; 3 = pass. (B) Quantification of longireached by the glacier. The calculation was performed within the raster GIS (5around Barre des Ecrins (Vallouise). Note that the Italian Doira Riparia catc

Second, to evaluate the role of paraglacial stressrelease in influencing slope evolution and slope failurein areas underlain by stronger bedrock, we focused on alarge, presently-stabilized mass-movement site with aview to determining whether paraglacial stress releasehad been directly responsible for triggering the finalfailure event at this site, or merely constituted one ofseveral factors (such as relief and lithology) predispos-ing the slope to failure. To identify possible successivestages of debris release and their source areas wesurveyed longitudinal transects across the site, thenreconstructed the morphometry of runout deposits andidentified constituent morpho-sedimentary units. Forthese, sediment characteristics were logged, and dip andsize of boulders were measured to differentiate betweena rock-fall origin (downslope-dipping clasts) and a rock-

tude reached by glacier ice is deduced from former glacial remnantsr glacial thickness is estimated by subtracting altitudes of the formerraphique National, 50-m grid) in a raster GIS (TNTmips©). The normalimal values are obtained around Barre des Ecrins (Guisane, Vallouise):tudinal stress. This map is derived from the map of minimal altitudes0-m grid). Maximal values of both thickness and gradient are estimatedhment, in the north east corner, is not considered in this study.

Fig. 5 (continued ).

avalanche origin (upslope-dipping clasts, particularly inthe runout toe where compressive forces were at max-imum; Hsü, 1975; Hewitt, 2002a,b). Lithological anal-ysis of boulders allowed the determination of whethersource areas correspond to zones where debuttressingforces were effective (i.e. in lower or median part of theslopes, below the trimline; Augustinus, 1995). The li-thology of 100 randomly-selected boulders was identi-fied within each morpho-sedimentary unit. Finally,absolute dating of runout deposits using cosmic rayexposure dating was carried out to test the likelihood ofrock-slope failure being triggered by paraglacial stressrelease, through comparison of timing of deglaciationwith that of slope failure.

3.3. Cosmic-ray exposure dating

3.3.1. Principles of cosmogenic radionuclideexposure dating

Cosmogenic nuclide production rate at a given sitedepends on location (latitude and altitude) and the expo-

sure geometry (relief shielding) of the sampled surface.We analysed concentrations of cosmogenic 10Be retainedwithin the quartz mineral fraction of exposed rocksurfaces (see Table 1). Cosmogenic 10Be is particularlyappropriate for this study as it does not suffer from loss bydiffusion from quartz minerals and its radiogenicformation in quartz minerals is negligible (Gosse andPhilips, 2001). In-situ 10Be production and accumulationwithin quartz minerals result from neutron- and muon-induced spallation reactions. The relationships between10Be concentration and both time elapsed since exposure(t) and depth (x) are well known (Braucher et al., 2003).

In calculating exposure ages we employed the 10Beproduction rate (21.8±0.6 atom g−1 yr−1) derived fromthe well-dated Koefels landslide (Kubik et al., 1998;Sorensen and Bauer, 2003; Kubik and Ivy-Ochs, 2004),as this covers the same latitudinal and altitudinal rangeas our study area and has been independently verifiedusing calibrated 14C ages. This production rate alsotakes account of any geomagnetic field variations (di-polar and multipolar variations). Production rates at

individual sampling sites were corrected for samplinglatitude using the scaling factors of Stone (2000), andcorrections for elevation were calculated using the dif-ference of air pressure from our sampling sites and theKoefels site using Standard Atmosphere Computations(http://aero.stanford.edu/StdAtm.html). Production rateswere also normalized to eliminate the effects by localtopography following the methods of Dunne et al.(1999); local shielding factors range from 0.658 to0.990. Analytical uncertainties (reported as ±1σ) arebased on counting statistics, an additional conservative3% uncertainty based on long-term measurements ofstandards, and a 50% uncertainty in the chemical andanalytical blank correction.

Quartz was isolated and purified from the samples andtargets prepared for AMS analysis of 10Be following theprocedure described by Brown et al. (1991, 1998). AMSanalyses were carried out at the Tandétron AcceleratorMass spectrometer at Gif-sur-Yvette, France (Raisbeck

Fig. 6. Deglaciation pattern in the Clarée catchment (A) and in the Vallouise caremnants: 3 = cirque glaciers; 4 = polished surfaces (the arrow indicates the orkame terrace; 9 = truncated scree. Glacial extent: 10 = Late Glacial Maximustagnation of glaciers during Early Holocene is evidenced by both morainic riwas already ice-free.

et al., 1994) using the National Institute of Standards andTechnology (NIST) 10Be standard (SRM 4325).

For comparison of exposure ageswithin the study area,ages are reported in terms of 10Be years (10Be yr) thatincorporate experimental uncertainties but not systematicuncertainties in production rate (Gosse et al., 1995; Brownet al., 1998). Comparison of these ages with absolute agesrequires incorporation of an additional 15% uncertaintyrepresenting production-rate uncertainty.

3.3.2. Sampling strategyIn calculating cosmic ray exposure ages from the

measured in situ-produced 10Be concentrations, twoassumptions are necessary: first, that 10Be concentrationat the initiation of the present surface exposure episodewas zero; second, that negligible erosion has affected thesampled surface since its exposure to cosmic radiation.Sampling strategy was designed to maximize the chan-ces of these assumptions being met. We sampled well-

tchment (B). Topography: 1 = summits; 2 = hypsometric curve. Glacialientation of glacial flows); 5 = cirque threshold; 6 = till; 7 = erratics; 8 =m extent; 11 = Holocene extent. In the upper Vallouise catchment, thedges and cosmogenic dates; at the same period, the upper Clarée valley

Fig. 6 (continued ).

preserved rochesmoutonnées (Fig. 4a), because (i) glacialerosion of such glacially-abraded surfaces reduces thepossibility of remnant 10Be inherited from a previousperiod of exposure, and (ii) negligible erosion has affectedthese surfaces since deglaciation. We sampled at thehighest point next to the steep lee side, to minimize thepossibility of sediment cover following deglaciation.

To date the timing of rock-slope failure in the Vallouisearea, we sampled large (N10 m3) boulders located onridges in the runout zone. Sampleswere collected from thesurface and bottom of each boulder (Fig. 4b) to ensure thatthe ages obtained are not influenced by prior exposure tocosmic rays on the pre-failure rock face. In Clarée Valley,we also sampled a rock surface exposed by rock-slopefailure.

4. Results

4.1. Deglaciation pattern

4.1.1. Glacial extent and quantification of basal buttressingThe Durance palaeo-glacier was one of the main

alpine glaciers during the Last Glacial Maximum. It

extended 200 km downvalley, and terminated close toSisteron. We estimate that the minimal thickness ofthe glacier in the Upper Durance area ranged from600 m to 900 m in the tributary valleys (Fig. 5a). Themaximum estimated values occur in the Briançonbasin and in the lower part of Vallouise and Guisanevalleys, where glacier thickness locally reached at least1000 m (Fig. 5b). The inferred gradient of the formerglacier surface was steepest in the Vallouise valley,where it locally reached 15%. As a consequence,estimated debuttressing (stress-release) values are thehighest in these western valleys. In both the Vallouiseand Guisane valleys, estimated normal stresses (7000–8000 kPa) and longitudinal stresses (200–300 kPa) dueto glacier overburden were greatest. In contrast, in theeastern part of the study area (Clarée, Cerveyrette andGuil valleys), the reconstructed ice thickness is lessthan 800 m and is associated with a relatively lowsurface gradient (less than 10%). For these valleys, thecalculated normal stress is 5000–7000 kPa and thelongitudinal stress 50–200 kPa. Thus, inferred debut-tressing values are significantly lower in these areasthan in the western valleys.

4.1.2. Chronological benchmarksIn the Upper Clarée Valley, the approximate thickness

of ice cover at the Last Glacial Maximum is constrainedby the difference between the 10Be age obtained for theupper part (CLA_04_04, 2170 m; 28.8±5.6 10Be ka) andthe lower part (CLA_04_03; CLA_04_11, 1880m; 10.5×1.3 10Be ka) of the valley slope. Although former icelimits cannot be determined, cosmic ray exposure datesfrom polished, roche-moutonnée surfaces in the UpperClarée Valley imply that the higher parts of the catchmentbecame progressively ice free during the Early Holocene,between 10.5±1.3 10Be ka and 7.8±0.9 10Be ka (Fig. 6a).

In the Vallouise valley, the last glacial advance prior tothe Little Ice Age advance is recorded by the Ailefroidemoraines (Cossart, 2005). Awell-preserved polished rocksurface that had been deglaciated before this stage yielded

Fig. 7. Slope failures in the Upper Durance area. 1 = summit; 2 = main ridges;This inventory is based on fieldwork, aerial photographs and geological mapsto 106m3). They are preferentially localized in the Vallouise Valley, East of

an exposure age (GYR_04_01) of 7.4±1.1 10Be ka,consistent with identification of a major stage of de-glaciation in the upper valleys during the Early Holocene(Fig. 6b).

Collectively, the dating evidence indicates that theupper valleys were occupied by glacier ice until the LateGlacial, and that glacier termini retreated to altitudesabove 2000 m in the Early Holocene. These resultsaccord with those of pedological and palynologicalanalyses carried out for sites within the study area, withclimatic warming during the Early Holocene, prior to6.0 ka BP (Talon et al., 1998; Muller et al., 2000). Thetiming of deglaciation implied by the exposure datingevidence is also consistent with the deglaciationchronology previously established for the western partof Massif des Ecrins (De Beaulieu, 1977).

3 = pass. 4 = rock topple; 5 = rock avalanche; 6 = landslide or earthflow.. Rock avalanches are supplying the largest volumes of sediments (105

the Barre des Ecrins.

Fig. 8. Slope failures database within Vallouise catchment. (A) Location of mass-wasting occurrences. 1 = main ridges; 2 = hypsometric curve;3 = stream; 4 = glaciers (present); 5 = glacial extent during the LGM; 6 = LGM trimline directly reconstructed from geomorphic data; 7 = rockfall (including rock avalanches); 8 = translationnal rock slides; 9 = earthflow; 10 = mass-movement reference. (B) Comparison betweenaltitudes of slope failure occurrences and local trimline. Note that most failures occurred below the trimline.

4.2. Slope failure at a regional scale

Slope failure sites are widespread in the UpperDurance catchment. In this study, 81 slope failures withvolumes N100 m3 have been identified, mapped andclassified according to the three types defined above. Theresulting map (Fig. 7) shows that different types of failureare associated with specific areas: rock avalanches arefound around the Massif des Ecrins, earthflows arefrequent in the Guil, Cerveyrette and Guisane valleys, androck topples mainly occur in the upper Clarée catchment.

The distribution of major rock falls is stronglyassociated with areas affected by the highest decom-pression values (σ=5000–9000 kPa and τ=200–300 kPa) and the steepest slope gradients (100 to150%). In the Vallouise area, 34 slope failures wereidentified and mapped (Fig. 8a): 6 major rock falls and/or rock avalanches, 9 earthflows and 19 translationalrockslides. Comparison of the altitudinal ranges of these

Fig. 9. Slope failures inventory in the Clarée catchment. Topography: 1 =3 = trimline; 4 = glaciated area. Instability features: 5 = rock topple; 6 =for details). Rock topples occurred in the upper catchment, most of them

slope failures with the reconstructed trimline altitude(Fig. 8b) shows that the slope failures mainly occurredin formerly glaciated areas. The largest slope failures arelocated where the former longitudinal stress was high(up to 250 kPa), such as in the Gyr valley (1 in Fig. 8a),the Pelvoux basin (12, 13), and the lower part of theOnde Valley (18, 19, 24, 25).

Conversely, the sites of rock topples and earthflowstend to be associated to moderate former basal stresses(σ=000–5000 kPa and τ=100–200 kPa), but thesefailures occur in specific lithological settings: rock top-ples on slopes underlain by weak Carboniferous sand-stones, and earthflows on slopes underlain by chloritic“lustrés” schists. In the Clarée valley the sites of 28slope failures were identified and mapped: 3 transla-tional rockslides and 25 rock topples (Fig. 9). Rocktopples in this area occur below the trimline, on rochesmoutonnées underlain by carboniferous sandstones inthe upper part of the catchment.

main ridges and summits; 2 = hypsometric curve. Glacial remnants:translational rock slide; 7 = area prone to rock toppling (see Fig. 10being located close to Pte de la Cassille.

Collectively, these data show that lithology hasconditioned the type of slope failure in most cases.However, the spatial coincidence of most rock-slopefailures with areas where glacial loading imposed highnormal and longitudinal stresses buttressing forces, to-gether with the preponderance of slope failures locatedbelow the trimline, clearly supports the notion thatparaglacial stress release played a major role in rock-slope weakening and failure.

4.3. Case studies

Two examples that document in more detail theeffects of glacial unloading on bedrock modification areoutlined below. In the first example, the effects of para-glacial stress release on bedrock disruption are discussedwith reference to the Upper Clarée valley, where rochesmoutonnées are often affected by toppling failure. In thisarea, relatively weak Carboniferous sandstones haveexperienced post-glacial bedrock deformation and jointdevelopment. In the second example, the impact ofparaglacial stress release on slope instability is illustratedby the Pré de Madame Carle failure, a large rock ava-lanche located in the upper Gyronde valley where boththe normal and longitudinal stresses imposed by glacialloading were high.

Fig. 10. The upper Clarée rock topples (see location in 7, Fig. 9): detailsandstones. The former glacier thickness is estimated to be around 400m“roches moutonnées” are identified, 15 of which are affected by rockthe valley slopes. Topography: 1 = hypsometric curve; 2 = stream. Lithomoutonnée; 6 = roche moutonnée affected by type A rock topple; 7 = ro

4.3.1. Post-glacial modification of roches moutonnéesin the Upper Clarée Valley

To assess the role of deglacial unloading on bedrockweakening, we examined a moderately rugged area of12.5 km2 (Fig. 10). Estimated normal and longitudinalstresses applied by former glacier loading weremoderate (σb3500 kPa, τb100 kPa). Within this area,21 roches moutonnées were identified, 15 of which havebeen affected by post-glacial disruption or blockdisplacement. Most of the latter occur on the lowerpart of the slope. On every roche moutonnée, the mainjoint sets are orientated at N-105° with a secondarymode at N-50° (Fig. 11). These joints are characteris-tically planar, and correspond to the recognized mainstructural alignments in the area (Tricart et al., 2001).

Two situations are encountered. Some roches mou-tonnées support only the rectilinear structural joints thatconform to the regional structural pattern. In some cases(6, 7, 12, 13, 17, 21 in Figs. 10 and 11), a type A rocktopple or displaced block occurs on the lee side of theoutcrop (Fig. 12). In such cases block displacementappears to reflect structural configuration, as the surfacegradient parallels the underlying bedding planes, whichhave acted as failure planes for joint-delimited blocks.Moreover, such topples or block displacements occurat various altitudes, even close to the trimline, where

of a slope section of 12.5 km2, carved out of weak carboniferous. The calculated values of deglacial stress release are moderate. 21topples. Note that instability is maximal in the lower part of

logy: 3 = limestones; 4 = sandstones. Glacial remnants: 5 = rocheche moutonnée affected by type B rock topple; 8 = trimline.

Fig. 11. Mirror diagrams of shattering patterns. 1 = stable glacially polished (roches moutonnées) surfaces; 2 = glacially polished surfaces affected bytype A rock topples, 3 = surfaces affected by type B rock topple. N-100 to N-105° orientations correspond to the regional, structural pattern. Note thatsurfaces affected by the type B rock topples are fragmented by two other joint sets of varying orientations parallel to the topography of the formerglacial valley floor.

Fig. 12. Roche moutonnée outcrop affected by type A rock topples (site no. 12, see Fig. 10). The unstable area is located on the lee side, shelteredfrom glacier flows. It is affected by structural joints only: they cause the failure of some boulders and their subsequent sliding along the bedding plane.

debuttressing values are low. It is therefore not necessaryto invoke paraglacial stress release as a cause of blockdisplacement in such cases. Conversely, other rochesmoutonnées (3, 4, 5, 8, 11, 14, 18, 19 in Figs. 10 and 11)are densely fragmented by joint sets oriented at N-60°and N-160°; their jointing pattern is thus significantlydifferent from the structurally-aligned joints. Such jointsare located on the stoss side of roches moutonnées whereglacial loading (and hence deglacial stress release) was atmaximum. On rock surfaces, N-60° and N-160° jointsconverge to create an irregular pattern of discontinuities(Fig. 13a, b). Rock segments oriented N-60° are typicallya few meters long and cross-cut by segments orientedN-160°. Such joints also differ from the rectilinear,structurally-controlled joints in form. A vertical section(Fig. 13c) shows that such stress-release joints are char-acteristically curved, thus forming a splay-shaped patternof discontinuities that is independent of the alignment ofstructural joint sets. Jointing is particularly dense near thesurfaces of the roches moutonnées, but only major stress-release joints penetrate deeper into the bedrock. Thesejoints are interconnected with bedding discontinuities,thus reducing overall bedrock cohesion and rendering the

Fig. 13. Roche moutonnée outcrop affected by type B rock topples. Site no.arrow), where deglacial stress release is at maximum. Some structural joints (Norientations N-60° and N-160°) disrupting the polished surface of the rochecaused vertical displacements of rock blocks since deglaciation. (D) The structfavouring toppling occurrence on the southwestern flank of the roche mouto

outcrops more liable to block displacement or type B rocktopples (Fig. 13d).

The alignment and geometry of the stress-releasejoints clearly distinguish them from structural joints:they are slightly curved and are characterized byorientations roughly parallel (N-160°) or perpendicular(N-60°) to the direction of former glacier movement.They are, moreover, associated with small grabens andantiscarps on the stoss faces of roches moutonnées (Fig.13c), indicating differential vertical displacements ofrock blocks since deglaciation: the central part of theroche-moutonnées are lower than the flanks, implyinglateral extensional dilation of the outcrops. Such de-formational patterns, together with intense near-surfacedisruption of joint-bound blocks, strongly supports thenotion that paraglacial stress release has contributed tobedrock breakdown and block displacement.

The above observations can be encapsulated in atypology of rock topples and block displacement (Fig. 14)that permits discrimination between structural features(type A) that occur across a wide range of altitudes, andparaglacial stress-release features (type B), that arerestricted to sites where glacial loading and hence stress

5 (see Fig. 10). (A) The unstable area is located on the stoss face (see-105°) affect the entire roche moutonnée. (B) Jagged neo-joints (mainmoutonnée. (C) Vertical cross section. The neo-joints are curved andural bedding is dipping valley ward and is acting as a sliding plane, thusnnée (photographs by E. Cossart, September 2002).

Fig. 14. Roches moutonnées evolutionary pattern in the Clarée catchment. Type A rock topples are due to structural jointing. Type B rock topples areconsidered as paraglacial features, in relation to the development of neo-joints and to their location (lower part of the valley slope; stoss face of rochesmoutonnées).

releasewas at amaximum (i.e. on lower valley-side slopesand on the stoss faces of roche moutonnées).

Attempts were made to date both types of feature.Exposure dating of a type A topple failed because of thehigh concentration of boron in the sample, but an ex-posure age obtained from the failure plane of a type Brock topple (CLA_04_08: 10.1±2.1 10Be ka) is con-sistent with block displacement very soon after degla-ciation of the site (Fig. 15). This result is consistent with

Fig. 15. Cosmic ray exposure ages dating glacial retreat and rock toppling in200m above the valley floor, and refers to a waning stage of the glacial tonguon the valley floor (CLA_04_02, 03, 09 and 11; (1880 m a.s.l.) or on a cirque fwas deglaciated during Early Holocene. Note that rock topple occurrences f

the hypothesis that bedrock displacement and associatedblock toppling occurred as a result of paraglacial stressrelease in the immediate aftermath of glacial unloading.

4.3.2. The Pré de Madame Carle case studyTo reconstruct the evolution of a valley-side slope

potentially influenced by high post-glacial debuttressingforces, we focused our investigation on Pré de MadameCarle rock-slope failure in the upper Vallouise catchment

the Upper Clarée Valley. Sample CLA_04_04 (2170m a.s.l.) is locatede, probably after the Late Glacial Maximum. Other samples are locatedloor (CLA_04_05 and 06; 2300 m a.s.l.): their dating implies the valleyit with the timing of glacial retreat (CLA_04_08).

Fig. 16. The Pré de Madame Carle rock avalanche. (A) Location, map extracted from Fig. 8. (B) Geomorphological setting. Lithology: 1 = granite;2 = anatexite; 3 = gneiss; 4 = limestone. Topography: 5 = hypsometric curve; 6 = summit; 7 = main ridges. Deposits: 8 = scree; 9 = debris-flowsdeposits; 10 = rock avalanche deposits; 11 = slope break; 12 = transect (see C). The source area of rubble can be deduced from lithology of the hillslope.Slope gradients are ranging from 100 to 150%. The altitude reached by the former LGM glacier is about 2450 m a.s.l., hence favouring post-glacialstress release. (C) Sedimentary fabric of the Pré de Madame Carle rock avalanche deposit. Three breaks of slope (cf. slope gradient in degrees) areassociated with the occurrence of the biggest blocks (Dmax: maximal clasts size; D50: median clasts size), thus suggesting several stages of debrissupply. Note that granites are preponderant in unit 1, limestones in unit 3.

next to Mont Pelvoux. At this site the valley is deeply cutinto granitic and gneissic bedrock, with slope gradientsranging from 100% to 130%. The volume of landslide

runout debris is estimated to lie between 1.0×106 and10.0×106 m3. The runout debris has accumulated as asteep, cone-shaped mass of large boulders, thus ruling

Fig. 17. Cosmic ray exposure ages of the Pré de Madame Carle rock avalanche deposit (CARL) plotted against ages of deglaciation (GYR). Theclustering of dates suggests the occurrence of two main stages of slope failure: the first one probably occurred shortly after deglaciation, while thesecond one occurred many millennia later. Because of 10B contamination, CARL_04_02_inf and CARL_04_05_inf samples (points drawn in darkgrey) did not provide accurate data (uncertainty higher than 30%).

out a possible morainic origin. Calculated former glacialloading stresses are high at this site (σ=∼5000 kPa,τ=∼300 kPa), suggesting that paraglacial stress releasemay have influenced rock-slope failure at this locality.

The runout deposit mainly consists of coarse angularboulders up to 60 m3 in volume, and finer matrixmaterial that accounts for around 20% of the deposit.The angularity of the boulders and the abundance offines are inferred to reflect inter-clast collisions duringrapid downslope transport of mobile debris. Upslope-dipping blocks at breaks of slope are consistent withinterpretation of the slope failure as a rock avalanche.Three main morpho-sedimentary sub-units differentiat-ed by breaks in slope, concentrations of the largest clastsand clast lithology were identified (Fig. 16a, b). In unit1, granites are preponderant (60% of boulders), whilelimestone and gneiss boulders are roughly equally rep-resented (about 20% each). In unit 2, the proportion of

Fig. 18. Pattern of the Pré de Madame Carle slope evolution before and after ththick). Stage 1: soon after deglaciation, the lower part of the slope became unswhich dammed the upper Gyr valley. Stage 2: materials are supplied from the uwhich are presently still active (rock falls, debris flows, snow and debris av

granitic blocks is lower and the representation of lime-stone boulders is greater. In unit 3, limestone bouldersare more abundant than either gneissic or graniticblocks. Above the runout deposit, limestones crop outon the upper part of the slope and granites underlie thelower part of the slope (Fig. 16a, b).

Two hypotheses arise from these field observations.Firstly, the lithological differences between the threeunits may represent a single failure event that affectedthe entire slope, so that the bedrock stratigraphic se-quence is represented by the different proportions ofeach lithology in each unit of the runout deposit.Alternatively, the three units can be interpreted as theresult of three successive stages of debris supply, withdifferences in the lithological composition of the unitsrepresenting retrogressive shifting of zone of rockfailure from the lower to the upper part of the slope.This scenario implies that the largest failure occurred

e deglaciation. Stage 0: maximal extent of former glacier (about 800-mtable, thus triggering a large, paraglacial rock avalanche, the deposit ofpper slope, the instability being sustained by classical, alpine processesalanches).

first, emplacing unit 1, and that more restricted failuresoccurred later, depositing unit 2 then unit 3.

Surface exposure dating of boulders at this sitesuggests that there were at least two stages of runoutdeposition (Fig. 17). Exposure ages of 6489±1923 10Beyr (CARL-04-03) and 6481±1093 10Be yr (CARL-04-05) for a boulder at the foot of unit 1 imply emplacementof the lower part of the runout deposit at ±6.5 10Be ka.Deglaciation of this site is dated by sample GYR-04-01(7420±1079 10Be yr). Taking into account the analyt-ical uncertainties associated with these dates, the initialfailure event appears to have occurred penecontempor-aneously with, or shortly after, deglaciation. Rock sam-ples obtained from a boulder located in the upper partof the runout deposit (unit 3) and from another locatedin the central part yielded exposure ages of 1887±326 10Be yr (CARL-04-01) and 1411±283 10Be yr(CARL-04-04) respectively. Together, these datessuggest that a second failure event occurred severalmillennia after the first, at ~1.5 10Be ka.

Collectively, the sedimentological, morphologicaland dating evidence suggests three stages in theevolution of the Pré de Madame Carle site (Fig. 18).At the Last Glacial Maximum, the Vallouise glacier wasat its maximum thickness (stage 0 in Fig. 18), impartingmaximum overburden stresses on the underlying bed-rock. During stage 1, immediately after deglaciation, theinitial failure appears to have occurred on lower slopesthat had experienced maximum glacial loading, andhence were subject to the greatest stress release. Debrisrunout associated with this event accumulated on thevalley floor, damming the river and causing the ag-gradation of the flat ʽPré' (meadow) immediately up-valley. Both the location of the failure plane in the zoneof maximum glacial loading and the timing of failureimply that this initial failure was, at least in part, in-fluenced by paraglacial stress release and joint devel-opment. At stage 2 (Fig. 18) several millennia afterdeglaciation, destabilization and failure propagated up-slope into a zone of steeper slope gradients. This laterfailure event was largely sourced from above the max-imum altitude of former glacier cover, and henceappears unlikely to reflect the direct influence of para-glacial stress release.

5. Discussion

5.1. Results of the spatial approach

Comparison within a GIS framework of the distri-bution of rock-slope failures with zones of inferred highstresses imposed by glacial loading provides a new

method for integrating the influence of paraglacial stressrelease into models of slope instability. It has beenshown that the main unstable areas are consistentlylocated in areas formerly occupied by glacier ice. Thispattern conforms with observations in other alpinemountain areas (e.g. Shreve, 1966; Porter and Orom-belli, 1981; Evans and Clague, 1994; Augustinus,1995). The influence of paraglacial stress release onrock-slope instability is here supported by two sets ofdata. At a regional scale, the location of landslides isassociated with high normal and longitudinal basalstresses imposed by glacial loading; the location of rockavalanches is particularly strongly associated with sitesof inferred high stress release. At a more localised scale,the hypothesis of a paraglacial stress-release originfor post-glacial deformation of roches moutonnées issupported by both the geometry and location of stress-release joints. In the Clarée valley, joint sets that havedeveloped independently of regional structural patternsact as planes along which bedrock disruption of rochesmoutonnées has occurred. The principal orientations ofsuch stress-release joints accord with former directionsof glacier movement and with valley-floor topography.Some vertical displacements and lateral block deforma-tions are associated with the development of smallgrabens and antiscarps that modify the original form ofroche-moutonnée surfaces. Such features resemblethose documented by Bovis (1990) at the margins ofthe downwasting Affliction Glacier in British Columbia,but at a much smaller scale. The possible operation offrost action in disrupting roche-moutonnée surfaces canprobably be ruled out, as the altitudes at which ob-servations were made (2040 m) lie well below the lowerregional limit of alpine permafrost (2700 m; Bodin et al.,2004). The locations of non-structural joint sets, on thelower part of valley-side slopes and on the stoss faces ofrock-bars, are consistent with areas where debuttressingstresses were at maximum at a local scale, and thussupport a paraglacial stress-release origin for both thejoints and the associated bedrock disruption.

5.2. Results of the chronological approach

Surface exposure dating of deglaciated rock surfacesusing cosmogenic 10Be has allowed the timing of de-glaciation in the study area to be refined, and the degla-ciation history to be related to that previously establishedin the northwestern part of Massif des Ecrins. Exposuredating of rock-slope failure events suggests that initialfailure occurred fairly soon after deglaciation, consistentwith a relationship between deglacial unloading, therapid development of paraglacial stress-release joints,

and rock-slope failure. In this case, the time elapsedbetween deglaciation and rock-slope failure is muchshorter than that documented using a similar approach inthe Highlands of Scotland (Ballantyne, 1997; Ballantyneand Stone, 2004), probably because of the higher andsteeper relief of the study area. The Pré deMadame Carlecase study suggests that alpine-scale slopes may con-tinue to exist in a state of critical conditional stability forseveral millennia after the initial triggering event, withprogressive upslope or retrogressive movement of thezone of potential rock-slope failure. This inferred patternappears consistent with the results of other studies inEuropean mountains that indicate two distinct stages ofmass-movement activity: an initial period of slope in-stability during the Early Holocene, associated with nu-merous post-glacial landslide events; and a later periodof slope instability during Subboreal and Subatlanticchronozones, during which slope failure has often beenexplained by non-paraglacial triggering involving, forexample, increased precipitation (e.g. Alexandrowicz,1993; Starkel, 1997; Corsini et al., 2001; Soldati et al.,2004). Further investigations and dating are needed todocument a correlation between our results and suchstudies.

6. Conclusion

Rock-slope failures are widespread in the UpperDurance catchment: a total of 81 failures have beenidentified, 62 of which are located in the Vallouise orClarée valleys. The concentration of failures at differentspatial scales within areas where stresses imposed byformer glacier loading were high implies that glacialdebuttressing and associated stress release have playedan important role in triggering rock-slope failures. Morespecifically, it explains the numerous occurrences offailure sites on lower valley-side slopes (i.e. below thetrimline marking the upper limit of glacial erosion) andon the stoss faces of rock-bars, where inferred sub-glacial stresses were particularly high.

Cosmic ray exposure dating methods are particularlywell suited to a chronological framework in such a para-glacial context as they can be applied to establish both thedeglaciation age of glacially-polished bedrock surfaces asthe age of rockslide runout debris, and thus the timing ofrock-slope failure. The dates obtained suggest that the firstrock-slope failures occurred soon after deglaciation of thefailure site. Subsequent episodes of slope failure, how-ever, may occur independently of the influence of para-glacial stress release. Systematic exposure dating of rock-slope failures provides the opportunity to assess theresponse time of paraglacial slope failures after deglaci-

ation. In the future, it may prove possible to couple suchchronological data with GPSmeasurements and geophys-ical prospecting to reconstruct sediment budgets in aparaglacial context.

Acknowledgments

Financial support was provided by the “Dynamiquedes Milieux et Risques” research team (UMR PRODIG8586 — CNRS) and “Traçeurs Isotopiques” researchteam (UMR 6635 — CNRS). Thanks are due to AlainMarais (CEREGE, Aix en Provence) for his assistanceon the field and G. Aumaître (CNRS, Gif sur Yvette) forhis invaluable assistance with 10Be AMSmeasurements.We thoroughly acknowledge with thanks Professor C.K.Ballantyne (University of St Andrews, Scotland) whocarefully edited and reviewed the English manuscript.The review comments of O. Slaymaker and B.H.Luckman greatly helped us in improving the manuscript.

Appendix A

The Delaunay triangulation of a set of points (herewe considered the points of hypsometric curves) inthe plane is a set of triangles connecting the pointssatisfying an “empty circle” property: the circumcircleof each triangle does not contain any of the points.According to the TIN model, the surface is representedby a plane within each triangle: the equation of the planeallows the calculation of the altitude for each pixel ofthe map.

References

Abele, G., 1997. Influence of glacier and climatic variation on rock-slide activity in the Alps. Paläoklimaforschung 19, 1–6.

Alexandrowicz, S.W., 1993. Late Quaternary landslides at the easternperiphery of the national park of the Pieniny Mountains,Carpatians, Poland. Studia Geologica Polonica 102, 209–225.

André, M.F., 1997. Holocene rockwall retreat in Svalbard: a triple rateof evolution. Earth Surface Processes and Landforms 22, 423–440.

Augustinus, P.C., 1995. Rock mass strength and stability of someglacial valley slopes. Zeitschrift für Geomorphologie 39, 55–68.

Ballantyne, C.K., 1997. Holocene rock-slope failures in the ScottishHighlands. Paläoklimaforschung 19, 197–205.

Ballantyne, C.K., 2002a. A general model of paraglacial landscaperesponse. The Holocene 12 (3), 371–376.

Ballantyne, C.K., 2002b. Paraglacial geomorphology. QuaternaryScience Reviews 21, 1935–2017.

Ballantyne, C.K., Stone, J.O., 2004. The Beinn Alligin rock avalanche,NW Scotland: cosmogenic Be-10 dating, interpretation and signif-icance. The Holocene 14, 461–466.

Ballantyne, C.K., Stone, J.O., Fifield, L.K., 1998. Cosmogenic Cl-36dating of postglacial landsliding at The Storr, Isle of Skye, Scotland.The Holocene 8, 347–351.

Barféty, J.C., Lemoine, M., De Gracianski, P.C., Tricart, P., Mercier,D., 1995. Briançon: Notice de la Carte Géologique. Editions duBRGM, Orléans. 179 p.

Benn, D.I., Evans, D.J.A., 1998. Glaciers and Glaciation. Arnold,London. 734 p.

Bindschadler, R.A., Harrison, W.D., Raymond, C.F., Crosson, R.S.,1977. Geometry and dynamics of a surge-type glacier. Journal ofGlaciology 18 (79), 181–194.

Bodin, X., Cossart, E., Fort, M., Schoeneich, P., 2004. The threat ofglobal change on periglacial belt in alpine areas. Proceedings ofEuropean Geosciences Union Assembly, Nice, 25–30/04/2004,vol. 6, p. 99.

Bourlès, D.L., Braucher, R., Engel, Z., Kalvoda, J., Mercier, J.L.,2004. Deglaciation of the Giant Mountains Indicated by 10BeDating. In: Drbohlav, D., Kalvoda, J., Vozenilek, V. (Eds.), CzechGeography at the Dawn of the Millenium. Olomouc, PalackyUniversity in Olomouc, pp. 25–39.

Bovis, M.J., 1990. Rock slope deformation at Affliction Creek, SouthernCoast Mountains, British Columbia. Canadian Journal of EarthSciences 27, 243–254.

Bovis, M.J., Jones, P., 1992. Holocene history of earth flow mass move-ment in south-central British Columbia: the influence of hydrocli-matic changes. Canadian Journal of Earth Sciences 29, 1746–1755.

Braucher, R., Brown, E.T, Bourlès, D.L., Colin, F., 2003. In situ-produced 10Be measurements at great depths: implications forproduction rates by fast muons. Earth and Planetary Science Letters211, 251–258.

Breteaux, L., 1998. Etude morphodynamique du versant méridional duChenaillet. Unpublished Maîtrise report, Denis Diderot — Paris 7University, 128 pp.

Brook E.J., 1994 — Surface Exposure Geochronology usingcosmogenic nuclides: applications in Antarctic Glacial Geology.Ph.D. Thesis, MIT/WHOI, WHOI-93-50, 230 pp.

Brown, E.T., Edmond, J.M., Raisbeck, G.M., Yiou, F., Kurz, M.D.,Brook, E.J., 1991. Examination of surface exposure ages ofmoraines in Arena Valley, Antarctica, using in situ produced 10Beand 26Al. Geochimica et Cosmochimica Acta 55, 2269–2283.

Brown, E.T., Bourlès, D.L., Burchfield, B., Deng, Q., Li, J., Molnar, P.,Raisbeck, G.M., Yiou, F., 1998. Estimation of slip rates in thesouthern Tien Shan using cosmic ray exposure dates of abandonedalluvial fans. Geological Society of America Bulletin 110, 377–386.

Chardon, M., 1991. L'évolution tardiglaciaire et holocène des glacierset de la végétation autour de l'Alpe d'Huez. Revue de GéographieAlpine 2, 39–53.

Church, M., Ryder, J.M., 1972. Paraglacial sedimentation: a consi-deration of fluvial processes conditioned by glaciation. GeologicalSociety of America Bulletin 83, 3059–3071.

Colas, A., 2000. Etude géomorphologique de la vallée de la Vallouise.Unpublished PhD Thesis, Lille 1 University, 458 p.

Corsini, A.,Marchetti,M., Soldati,M., 2001. Holocene slope dynamicsin the area of Corvara in Badia (Dolomites, Italy): chronology andpaleoclimatic significance of some landslides. Geografia Fisica eQuaternaria 24, 127–139.

Cossart, E., 2005. Evolution géomorphologique du haut bassin duranciendepuis la dernière glaciation — contribution à la compréhension dufonctionnement du système paraglaciaire. Unpublished PhD Thesis,Paris 7 University, 380 p.

Coûteaux, M., 1983a. Géomorphologie et évolution phytogéographi-que tardiglaciaires et holocènes aux Deux-Alpes. Revue deGéographie Alpine 71 (2), 143–162.

Coûteaux, M., 1983b. Fluctuations climatiques et fluctuationsglaciaires du Tardiglaciaire en Oisans (Isère, France). Final Proc.

Int. Conf. Climatic Changes and Related Problems. Stockholm,Sweden, pp. 24–28.

Coûteaux, M., 1983c. Fluctuations glaciaires de la fin du Würm dansles Alpes françaises, établies par des analyses polliniques. Boreas12, 35–56.

Coûteaux, M., Edouard, J.L., 1987. La déglaciation du site du lac desBèches (Massif des Ecrins), étude pollenanalytique et glacio-morphologique. Revue de Géographie Alpine 75 (1), 63–78.

Cruden, D.M., Hu, X.Q., 1993. Exhaustion and steady-state models forpredicting landslide hazards in the Canadian Rocky Mountains.Geomorphology 8, 279–285.

De Beaulieu, J.L., 1977. Contribution pollenanalytique à l'histoiretardiglaciaire et holocène de la végétation des Alpes méridionalesfrançaises. Unpublished PhD thesis, Aix-Marseille III University,358 p.

De Jong, M.G.G., De Graaff, L.W.S., Rupke, J., 1995. Der Eisabbaumvorderen Bregenzerwald und in den Nachbargebieten (Vorarlberg,Osterreich; Bayern, Deutschland) nach dem LetzeiszeitlichenEishochstand). Jahrbuch Der Geologischen Bundesanstalt 138 (1),27–54.

Dunne, A., Elmore, D., Muzikar, P., 1999. Scaling of cosmogenicnuclide production rates for geometric shielding at depth on slopedsurfaces. Geomorphology 27, 3–11.

Edouard, J.L., 1979. Les fluctuations glaciaires dans le haut bassin de laRomanche. Unpublished PhD Thesis, Grenoble I University, 685 p.

Evans, S.G., Clague, J.J., 1994. Recent climatic change and catas-trophic geomorphic processes in mountain environments. Geo-morphology 10, 107–128.

Feiss, C., 2000. Evolution des versants en conditions froides: exemple dela péninsule de Dingle. Unpublished PhD Thesis, Paris 8 University,346 p.

Flageollet, J.-C., Maquaire, O., Martin, B., Weber, D., 1999. Land-slides and climatic conditions in the Barcelonnette and Vars basins(Southern French Alps, France). Geomorphology 30 (1–2), 65–78.

Fort, M., 2000. Glaciers and mass wasting processes: their influenceon the shaping of the Kali Gandaki valley (higher Himalaya ofNepal). Quaternary International 65/66, 101–119.

Fort, M., 2003. Are high altitude, lava stream-like, debris mixtures allrock glaciers? A perspective from the Western Himalaya.Zeitschrift für Geomorphologie (Suppl.-Bd 130), 11–29.

Gosse, J.C., Philips, F.M., 2001. Terrestrial in situ cosmogenicnuclides: theory and application. Quaternary Science Reviews 20,1475–1560.

Gosse, J.C., Klein, J., Evenson, E.B., Lawn, B., Middleton, R., 1995.Beryllium-10 dating of the duration and retreat of the last Pinedaleglacial sequence. Science 268, 1329–1333.

Hétu, B., Gray, J.T., 2000. Effects of environmental change on screeslope development throughout the postglacial period in the Chic-Choc Mountains in the northern Gaspé peninsula, Québec.Geomorphology 32, 335–355.

Hewitt, K., 2002a. Styles of rock-avalanche depositional complexesconditioned by very rugged terrain, Karakoram Himalaya,Pakistan. Geological Society of America Bulletin 15, 345–377.

Hewitt, K., 2002b. Postglacial landform and sediment associationsin a landslide-fragmented river system: the Transhimalayan Indusstreams, Central Asia. In: Hewitt, et al. (Ed.), Landscapes oftransition. Kluwer Academic–Plenum Publishers, New York,pp. 63–91.

Holm, K., Bovis, M.J., Jakob, M., 2004. The landslide responseof alpine basins to post-Little Ice Age glacial thinning andretreat in southwestern British Columbia. Geomorphology 57,201–216.

Hsü, K.J., 1975. Catastrophic debris streams (Sturzstroms) generatedby rockfalls. Geological Society of America Bulletin 86, 129–140.

Ivy-Ochs, S., Heuberger, H., Kubik, P.W., Kerschner, H., Bonani, G.,Frank, M., Schlüchter, C., 1998. The age of the Koefels event-relative, 14C and cosmogenic isotope dating of an early Holocenelandslide in the Central alps (Tyrol, Austria). Zeitschrift fürGletscherkunde and Glazialgeologie 34, 57–70.

Jorda, M., 1988. Modalités paléoclimatiques et chronologiques de ladéglaciation würmienne dans les Alpes Françaises du Sud. BulletinAssociation Française pour l'Etude du Quaternaire 2–3, 111–122.

Jorda, M., 1993. Histoire des paléo-environnements Tardi- etPostglaciaires sud alpins de moyenne altitude, essai de reconsti-tution cinématique. Caen, Géomorphologie et Aménagement de laMontagne, Hommage à P. Gabert, CNRS, pp. 99–111.

Jorda, M., Rosique, T., Evin, J., 2000. Données nouvelles sur l'âge duDernier Maximum Glaciaire dans les Alpes Méridionales Fran-çaises. C. R. Acad. Sci., Paris, Sciences de la Terre et des Planètes,331 (3). Elsevier, pp. 187–193.

Kerschner, H., Kaser, G., Sailer, R., 2000. Alpine Younger Dryasglaciers as paleo-precipitation gauges. Annals of Glaciology 31,80–84.

Korup, O., 2004. Geomorphometric characteristics of New Zealandlandslide dams. Engineering Geology 73, 13–35.

Kubik, P.W., Ivy-Ochs, S., 2004. A re-evaluation of the 0–10 ka 10Beproduction rate for exposure dating obtained from the Köfels(Austria) landslide. Nuclear Instruments & Methods in PhysicsResearch. Section B, Beam Interactions with Materials and Atoms223–224, 618–622.

Kubik, P.W., Ivy-Ochs, S., Masarik, J., Frank, M., Schlüchter, C.,1998. 10Be and 26Al production rates deduced from an instanta-neous event within the dendro-calibration curve, the landslide ofKöfels, Ötz valley, Austria. Earth and Planetary Science Letters161, 231–241.

Lahousse, P., 1994. Recherches géomorphologiques et cartographiedes aléas naturels dans la vallée de la Guisane (Briançonnais,Hautes-Alpes). PhD thesis, Université Lille I, 431 p.

Lewis, W.V., 1954. Pressure release and glacial erosion. Journal ofGlaciology 2, 417–422.

Lliboutry, L., 2002. Extension of Glacier de Saint-Sorlin, French Alps,and equilibrium-line altitude during the Little Ice Age. Journal ofGlaciology 48 (160), 118–124.

Luckman, B.H., 1988. Debris accumulation patterns on talus slopes inSurprise Valley (Alberta). Quaternary Proceedings 42 (3),247–278.

Luckman, B.H., Fiske, C.J., 1995. Estimating long-term rockfallaccretion rates by lichenometry. In: Slaymaker, O. (Ed.), SteeplandGeomorphology. John Wiley & Sons, pp. 232–255.

MacGregor, K.R., Anderson, R.S., Anderson, S.P., Waddington, E.D.,2000. Numerical simulations of glacial-valley longitudinal profileevolution. Geology 28 (11), 1031–1034.

Matthews, J.A., Shakesby, R.A., 2004. A twentieth century neopar-aglacial rock topple on a glacier foreland, Ötzal Alps, Austria. TheHolocene 14 (3), 454–458.

Montjuvent, G., 1973. La transfluence Durance — Isère, essai desynthèse du Quaternaire du bassin du Drac (Alpes françaises).Revue de Géologie Alpine 49, 57–118.

Muller, S.D., David, F., Wicha, S., 2000. Impact de l'exposition desversants et de l'anthropisation sur la dynamique forestière dans lesAlpes du Sud (France). Géographie Physique et Quaternaire 54 (2),231–243.

Owen, L.A., Gualtieri, L., Finkel, R.C., Caffee, M.W., Benn, D.I.,Sharma, M.C., 2001. Cosmogenic radionuclide dating of glaciallandforms in the Lahul Himalaya: defining the timing of LateQuaternary glaciation. Journal of Quaternary Science 16 (6),555–563.

Panizza, M., 1973. Glacio pressure implications in the production oflandslides in the dolomitic area. Geologia Applicata e Idrogeologia8 (1), 289–297.

Paterson, W.S.B., 1981. The Physics of Glaciers, Second edition.Pergamon Press. 380 p.

Porter, S.C., Orombelli, G., 1981. Alpine rockfall hazards. AmericanScientist 69, 67–75.

Raisbeck, G.M., Yiou, F., Bourlès, D.L., Brown, E., Deboffle, D.,Jouhanneau, P., Lestringuez, J., Zhou, Z.Q., 1994. The AMSfacility at Gif-sur-Yvette: progress, perturbations and projects.Nuclear Instruments & Methods in Physics Research. Section B,Beam Interactions with Materials and Atoms 92, 43–46.

Rosique, T., 2004. La dernière phase glaciaire de la Moyenne Durance(région de Gap à Sisteron): bilan des recherches, dernièresconclusions chronologiques. Méditerranée 1–2, 25–36.

Ryder, J.M., 1971. The stratigraphy and morphology of para-glacialalluvial fans in the South-Central British Columbia. CanadianJournal of Earth Sciences 8, 279–298.

Seijmonsbergen, A.C., Woning, M.P., Verhoef, P.N.W., de Graaff,L.W.S., 2005. The failure mechanism of a Late Glacial Sturzstromin the Subalpine Molasse (Leckner Valley, Vorarlberg, Austria).Geomorphology 66, 277–286.

Shakesby, R.A., Matthews, J.A., 1996. Glacial activity and paraglaciallandsliding in the Devensian Lateglacial: evidence from CraigCerrig Gleisiad and Fan Dringarth, Forest Fawr (Brecon Beacons),South Wales. Geological Journal 31, 143–157.

Shreve, R.L., 1966. Sherman landslide, Alaska. Science 154, 1639–1643.Soldati, M., Corsini, A., Pasuto, A., 2004. Landslides and climate

change in the Italian Dolomites since the Late Glacial. Catena 55,141–161.

Sorensen, S.A., Bauer, B., 2003. On the dynamics of the KöfelsSturzstrom. Geomorphology 54, 11–19.

Starkel, L., 1997. Mass movements during the Holocene: the Carpatianexample and the European example. Paläoklimaforschung-Palaeo-climate Research 19, 385–400.

Stone, J.O., 2000. Air pressure and cosmogenic isotope production.Journal of Geophysical Research 105 (B10), 23753–23759.

Talon, B., Carcaillet, C., Thinon, M., 1998. Etudes pédoanthracolo-giques de variations de la limite supérieure des arbres au cours del'Holocène dans les Alpes françaises. Géographie Physique etQuaternaire 52 (2), 1–14.

Tricart, P., Schwartz, S., Sue, C., Poupeau, G., Lardeaux, J.M., 2001.La dénudation tectonique de la zone ultra-dauphinoise etl'inversion du front briançonnais au sud-est du Pelvoux (Alpesoccidentales): une dynamique miocène à actuelle. Bulletin de laSociété Géologique de France 72 (1), 49–58.

Slope instability in relation to glacial debuttressing in alpine areas (Upper Durance catchment,...

Documents