Navigation – Plan du site

An overview of the consequences of paraglacial landsliding on deglaciated mountain slopes: typology, timing and contribution to cascading fluxes

Conséquences des mouvements de masse paraglaciaires sur les versants désenglacés : typologie, temporalité, contributions aux budgets sédimentaires
Etienne Cossart, Denis Mercier, Armelle Decaulne et Thierry Feuillet
p. 13-24

Résumés

Trois décennies après la définition initiale du concept de "paraglaciaire", un modèle général de la paraglaciation a été formalisé, intégrant la cascade sédimentaire paraglaciaire et les types de réponses des réservoirs de sédiments associés. Les écarts à ce modèle doivent maintenant être examinés et expliqués par des facteurs de contrôle agissant aussi bien à l’échelle régionale (tectonique, notamment) que locale (topographie, lithologie, etc.). Nous comparons ici des schémas établis dans quelques massifs montagneux de l'hémisphère nord, localisés dans des contextes séismo-tectoniques variés (notamment Alpes occidentales, nord-ouest de l'Ecosse, Norvège centrale, Svalbard et l'Islande). En combinant nos observations de terrain avec une recension bibliographique, nous discutons de l'efficacité de la réponse paraglaciaire sur l'érosion des versants et de la contribution de cette érosion à la cascade sédimentaire. Dans la plupart des cas, la paraglaciation engendre un stockage de sédiments dans les têtes de bassin et le fond des auges glaciaires, le taux d'évacuation sédimentaire vers les exutoires restant faible. La paraglaciation apparaît donc comme une période de dénudation des versants, créant des réservoirs de sédiments dont l'évacuation ne peut être réalisée que lors des phases de glaciation ultérieures.

Haut de page

Texte intégral

1 - Introduction

1More than thirty years after the first definition of paraglaciation as “non-glacial processes influenced by glaciation” (Ryder, 1971a,b; Church & Ryder, 1972), the first model of paraglacial geomorphic processes integrating both space and time was constructed (Ballantyne, 2002a,b, 2003a,b). This model highlights potential paths for sediment transfer (from sediment sources to some specific stores and then to sinks) and the associated rates of sediment evacuation (fig. 1 & 2). In the latter case, sediment yield is considered to be related to the amount of remaining available sediments by a negative exponential function (known as the exhaustion model in Cruden and Hu (1993)). Hence, this model seems to fit satisfactorily with most physical settings (i.e. alpine and high latitude environments); we should however examine its residuals to highlight any local or regional specificities in paraglaciation. In fact, several authors have pointed out particular connectivity patterns in the paraglacial cascading system due to local settings: for instance, the creation of obstacles (moraines) or threshold effects in dam dismantlement can create residuals from the classic exhaustion model (Cossart, 2008; Cossart & Fort, 2008a,b; Etienne et al., 2008; Knight & Harrison, 2009).

2In this work, sediment paths structured by post-glacial dismantlement of mountain slopes are particularly studied in three steps. First, the processes involved in rock failure are identified and their possible influence on mass-movement locations at different spatial scales in various places is discussed. This comparison exhibits various patterns of paraglacial landslide distribution, and leads to identifying the local/regional parameters that explain these differences. Second, the rate of triggering of mass-movement over time is roughly assessed in various settings based on a review of recently published data. This comparison aims to typify models of slope evolution through the time elapsed since deglaciation. Once again, parameters leading to a possible differentiation are identified and discussed. Third, the contribution of landsliding to the whole paraglacial cascading system is debated. On the one hand, some authors highlight a high sediment yield at catchment sinks in relation to paraglacial landsliding (Church & Ryder, 1972; Ritter & Ten Brink, 1986). On the other hand, some long-lived sediment dams can occur after the deposition of a landslide mass, so that no sediment exportation can take place (Cossart & Fort, 2008a; Korup, 2009). From field observations and a review of the published data, a typology of geomorphic coupling between paraglacial landslides and other geomorphic processes is defined to contribute to this debate.

3This paper is thus based upon a comparative approach, carried out in various mountainous areas located in the northern hemisphere, in various seismotectonic settings. Thus both high latitudes (Iceland, Norway, Svalbard) and high altitudes (Western Alps) are compared here, as it encompasses active mountain areas (Western Alps, Iceland) and more stable ones (Scotland, Norway).

Fig. 1: Simplified paraglacial sediment cascade (after Ballantyne, 2002a,b).

Fig. 1: Simplified paraglacial sediment cascade (after Ballantyne, 2002a,b).

Fig. 2: The paraglacial period.

Fig. 2: The paraglacial period.

(A) paraglacial period defined by Church & Ryder (1972). (B) application of the exhaustion model to assess the evolution of the volume of sediments within a paraglacial store (Ballantyne, 2003b). Curves noted from 1 to 5 correspond to various calibrations of the exhaustion model; 1/ case of a durable storage within the store, 5/ case of a rapid degradation after deglaciation

2 - Paraglacial landsliding in space

2.1 - Impacts at outcrop scale

4At a fine (i.e. bedrock outcrop) scale, paraglaciation may act through decohesion processes: compression due to the glacier, followed by a consequent debuttressing, may shatter bedrocks (Lewis, 1954). In more detail, various patterns of paraglacial shatters are identified, the geometry of which is related to former glacier fluxes (transverse or parallel joints) and not to structural joints or foliation. Yet paraglacial shattering is not widespread; its location is highly dependent on the geological setting and palaeo-glacier characteristics.

5In basement areas (Svalbard, Norway, Scotland), paraglacial shattering is more efficient at the base of mountain slopes subject to a high lithostatic pressure, i.e. at the base of deep troughs (500 to 1000 meters-deep in Peulvast (1985)). In other cases, neo-joints are observed in massive but fragile outcrops such as sandstones or basalts (fig. 3A), where joints parallel to former glacial fluxes are identified: larger joints are close to the trimline. In Scotland, Sellier (2008) suggests a concept of "paraglacial differential erosion", where paraglacial jointing and weakening of bedrocks is more efficient in quartzites (Sellier & Lawson, 1998). Even though such paraglacial joints are clearly generated in accordance to the lithology (cohesion of bedrock), some debuttressing evidence is locally identified in unexpected areas (i.e. highly cohesive quartzophyllade outcrops and in limestone series of Svalbard; in André (1993, 1997) and Mercier (2002, 2011)). This reinforces the idea that paraglacial decohesion is not a myth and is effective at creating neo-joints.

6In active areas, most joints are related to the relief, geological structure and seismotectonic activity, so that no clear evidence of glacial debuttressing can be easily found (Bois et al., 2012; Bouissou et al., 2012). In the Western Alps, however, paraglacial joints have been identified in the upper part of formerly glaciated watersheds, i.e. where both the glacier surface slope and glacier thickness were high (Cossart et al., 2008; Darnault et al., 2011): it corresponds to neo-joints roughly parallel to former glacier-fluxes which density decreases in depth (fig. 3B). In addition, paraglacial neo-joints are prominent in carboniferous sandstones, even in quartzites (Monnier, 2006; Cossart et al., 2008), in which they can draw a splay-shaped pattern (fig. 3B).

2.2 - Impacts at hillslope scale

7At hillslope scale, most inventories point out the classic role of the lower part of hillslopes in generating failures in otherwise stable areas (Peulvast, 1985; Jarman, 2006; Ballantyne, 2006; Sellier, 2008; Jarman, 2009). Sliding processes are driven by a combination of debuttressing (higher at the base of hillslopes due to former ice-thickness) and steepening of the lower part of hillslopes (due to glacial erosion), whatever the structure. In the case of dip slopes, the slope steepening observed in the lower part of englaciated hillslopes may cut the structure, so that hillslopes become unstable (fig. 4A): translational slides may then occur due to the foliation pattern or the bedding pattern of the outcrops (Sellier, 2008). In the case of counterdip slopes, the main identified mechanism involves the development of neo-joints just above the former trimline (fig. 4B). These neo-joints are quite vertical and can become deeper in relation to debuttressing and vacuum due to glacier disappearance. This shattering may generate some rock-falls, or may evolve into a rotational landslide.

8Paraglacial jointing may also be coupled with slow movements that can reflect a lateral spread of mountain ranges. If they are of a large magnitude, such movements may lead to sackung features and then provoke deep-seated gravitational deformation of the entire hillslope, in relation to the development of normal faults (Gutiérrez-Santollala et al., 2005; Mège & Bourgeois, 2011). Sackungs are revealed by the following impacts on hillslopes: uphill-facing (antislope) scarps, tension cracks, grabens, and anomalous ridge-top depressions.

9Such spatial patterns of paraglacial failures impacts can be more complex in active areas. First, retrogressive movement of the zone of potential rock-slope failure can indeed be observed, as many non-paraglacial triggering can act after the initial failure. These triggers involve, for example, seisms, rejuvenation by river incision or increased precipitation (e.g. Soldati et al., 2004). Second, it is known that a temporal succession from sackung to landsliding may occur (Dikau et al., 1996; Cossart & Fort, 2008a). More generally, both jointing (at local scale) and faulting (at slope scale) weaken the internal cohesion of bedrock, favor water seepage, then make the displacement of material easier and keep the area prone to landsliding during millennia (Hippolyte et al., 2006, 2009).

Fig. 3: Evidence of post-glacial decohesion.

Fig. 3: Evidence of post-glacial decohesion.

(A) Joints parallel to palaeo-glacier fluxes at Stuphallet in Svalbard (D. Mercier, July 2004). (B) Block detachment from a cirque free-face in the Skagafjordur area due to a combination of post-glacial neo-joint and basalt bedding (E. Cossart, June 2011). (C) Splay-shaped joints in a rochemoutonnée (Clarée valley, southern French Alps) made of carboniferous sandstone (E. Cossart, June 2004). (D) Cracks and consequent differential lowering affecting the top of a roche-moutonnée made of sandstone (E. Cossart, June 2004).

2.3 - Impacts at regional scale

10At regional scale, inventories of landslides are drawn up in order to decipher the potential influence of paraglaciation on the location of mountain-slope instabilities. Many authors have tried to identify a relationship between an over-representation of landslides and glacial debuttressing or glacial deepening in basement areas, in active orogens, and in Iceland.

2.3.1 - Basement regions

11In basement areas such as Scotland, Jarman (2006) and Ballantyne (2003c, 2008) typified the main locations of rockslope failures. They identified two factors that particularly favor landsliding. First, all areas where glacial over-burdening reaches its maximum are prone to landsliding, such as narrow troughs associated with particularly constrained glacier fluxes, and areas of flux convergence (coalescent cirques, confluence of glacial valleys). Second, landslides also occur in over-deepened basins, where steepening of the lower part of the slope provokes an extended destabilization of mountain slopes. Hence, paraglaciation is here combined with geological structure, which also influences the geometry and the assemblages of glacial landforms (fig. 5). Glacial landform patterns are indeed predominantly derived from pre-glacial and structural features: quaternary evolution prolongs the late Pliocene incision phase (due to the efficient coupling of weathering and transportation), which occurred in response to tectonic movements (Le Cœur, 1999). Yet, over-deepening of basins, trough steepening and cirque enlargement are even more efficient in highly-shattered bedrocks and are driven by tertiary landforms; more precisely, pre-glacial excavation developed in relation to bedrock weathering or structural joint patterns (Godard, 1961). In western Norway (between 67°50’N and 69°50’N), Saintot et al. (2011) demonstrated that parameters leading to 72 rock slope instabilities were: (1) weak rocks; (2) foliation towards the fjords or the valley or steep foliation striking roughly slope-parallel; (3) folds and interference folds; (4) Caledonian thrusts cutting the slope; and (5) regional brecciated/cataclastic faults close to the slope. Collectively, these data in basement areas highlight that geological parameters are pre-conditioning factors (i.e. that fix inherent strength of a slope), which with paraglacial preparatory factors and triggers can be coupled to generate landsliding.

Fig. 4: Consequences of over-deepening of hillslope basal parts on landslide triggering.

Fig. 4: Consequences of over-deepening of hillslope basal parts on landslide triggering.

A/ Case of a dip slope (Scotland, adapted from Sellier, 2008). B/ Case of a counterdip slope

Fig. 5: Typology of paraglacial landslide location.

Fig. 5: Typology of paraglacial landslide location.

1/ Rotational slides in counterdip slopes, 2/ Rock-topples on glacially polished knobs (slope facing former glacier-fluxes), 3/ Translational slides in dip slopes, 4/ Landslides in narrow, over-deepened valleys (over-deepening in relation to fault emplacement), 5/ Landslides in zones of confluence.

2.3.2 - Active orogens

12In active orogens, the complexity of the joint patterns and the relief morphometry make the role of paraglaciation more difficult to decipher. Landslides are indeed common features in non-glaciated areas, such as in Prealps (Buonchristiani et al., 2002; Bravard et al., 2003; Fort et al., 2009), where landslides are geologically-driven features. Furthermore, fractures and joints patterns and both type and conditions of the rock layers are often pointed out as factors of prime importance in explaining the location of many investigated landslides located in formerly-glaciated valleys. For instance, in the case of Granier rock-failure, the superposition of Urgonian limestone on weak Hauterivian marls (100 m), coupled with the development of many faults and strike-slip faults favored the collapse (Bozonnat, 1980; Gidon, 1990). Dip of bedding planes, rock-weakening due to faults are other classical factors often considered (von Poschinger et al., 2006; Delunel et al., 2010).

13Nevertheless, in the French Alps, the distribution of landslides is quite different in the upper parts of formerly glaciated watersheds (high glacial erosion rates) and in the lower parts of formerly glaciated watersheds (low glacial erosion rates). An inventory is carried out in two areas of similar bedrocks (granites, gneisses, sandstones): the Gyronde catchment (former accumulation zone of the Durance glacier) and the Drac catchment (former ablation zone of the Isère glacier) (fig. 6). Landslides are over-represented below the trimline in the Gyronde (fig. 6C), a trend that is statistically not significant in the Drac (fig. 6D). Paraglacial landslides may thus occur where calculated normal and longitudinal ice loading stresses were higher (i.e. in upper catchments), thus modifying the overall spatial distribution of landslides. At this scale, the relief of stresses damages the rock after the unloading of the ice (i.e. "stress-release" in McColl (2012)). Post-glacial stress release can also explain some very specific locations at the confluence of former glaciers (Panizza, 1973) or onto the upstream side of some knobs, facing palaeo-glacier fluxes (Cossart et al., 2008), where the over-burdening effects of former glaciers were at maximum (fig. 5).

2.3.3 - Iceland

14In both Scotland and active orogen areas, paraglacial destabilization may act through debuttressing and over-steepening, a pattern that is slightly different in Iceland. There, landslides are mostly located at the margins of the island, in fjord areas, partly in relation to relief (s.s.) patterns. However, at fjord scales, landslides are concentrated at the mouth of fjords, while both structure and lithology are constant; such a pattern is also observed in Svalbard (Mercier, 2007). A statistical study highlights a direct relationship between landslide density and the value of glacio-isostatic rebound (Cossart et al., in press; fig. 7), well recorded by raised-beaches. In this case, paraglaciation acts preferentially through a significant post-glacial uplift, which induces both rock dilation and seismotectonic activity.

Fig. 6: Comparison of landslide location patterns in the upper and lower parts of formerly glaciated valleys.

Fig. 6: Comparison of landslide location patterns in the upper and lower parts of formerly glaciated valleys.

(A) Location map. (B) Extent of glaciers during the Last Glacial Maximum. (C) distribution of landslides (above vs. below the trimline) in the lower part of the Drac valley (former ablation zone), realized from the BRGM database. (D) Distribution of landslides (above vs. below the trimline) in the Gyronde area (former accumulation zone), realized from field investigations. The comparison of altitudes, both landslide scars and toe deposits, in cases C and D is statistically tested by a Fischer test; in each case the observed F is higher than the significance threshold (with an uncertainty of 0.05).

Fig. 7: Relationship between landslide location and glacio-isostatic rebound in Iceland (Skagafjörður).

Fig. 7: Relationship between landslide location and glacio-isostatic rebound in Iceland (Skagafjörður).

(A) Over-representation of landslide at the mouth of the fjord (fjord oriented north-south) estimated from a chi-square analysis in comparison with randomly distributed landslides (assessed in hectares). (B) and (C) Sketch of glacio-isostatic rebound following inlandsis disappearance. (D) Altitudes of raised-beaches in Iceland (Skagafjordur).

3 - Paraglacial landsliding over time

15Dating landslides remains difficult, in spite of the emergence of new techniques (OSL, Cosmogenic Nuclides, etc.; Lang et al., 1999). This hampers the establishment of statistically reliable trends, to ensure that the exhaustion model can be applied to landslide frequency (exponential decrease over time). Nevertheless, dates acquired during the last two decades can be sumarized to define temporal patterns.

3.1 - Scottish and alpine models

16Recent results acquired in both Scotland and the European Alps first highlight that post-glacial landslides may be triggered immediately after glacier disappearance. The oldest dated landslides thus occurred during the Lateglacial in Scotland (Cairngorm Mountains, 10Be, 11.5 ka - Ballantyne, 2008) and in the Alps (Straneggtal in Upper Austria, 36Cl, at 18.8 ± 0.9 ka - Van Husen et al., 2007; La Clapière in Maritime Alps, 10Be, 10.3 ± 0.5 ka - Bigot-Cormier et al., 2005). Furthermore, in many cases (Maol Cheann-Dearg in Scotland, the Clarée valley in the French Alps) the deposits of such large landslides were redistributed by glacier ice, indicating that landslides occurred before complete glacier disappearance (i.e. before the Younger Dryas ending in published studies).

17Nevertheless, periods of landsliding can also last for millennia after deglaciation, especially during the first half of the Holocene. In the French Alps, the examples of Fontfroide (Pré de Madame Carle area, French Alps) and other deep-seated landslides in the Tinée Valley exhibit successive periods of gravitational instabilities: the first occurred shortly after the deglaciation event (12-13 ka), the second at 7-9 ka and the third at 2.5-5.5 ka (Cossart et al., 2008; Darnault et al., 2011; El Bedoui et al., 2011) (fig. 8). Although incomplete, this scenario is suggested in case of Lauvitel failure, where an old rock-avalanche is covered by a recent landslide deposit which age is 4.7 ± 0.4 ka (10Be; Delunel et al., 2010). A similar pattern is also observed around the Alm and Straneggtal, in Upper Austria (Calcareous Alps): the initial event (i.e. a rock-avalanche) was followed by at least four millennia of slope instability (Van Husen et al., 2007). Alpine sequences finally suggest that initial paraglacial events occurred shortly after deglaciation through failures below the trimline (Cossart et al., 2008) or deep-seated gravitational deformations (DSGD; Hippolyte et al., 2006, 2009). Following these events, the mountain slope remained unstable during the main part of the Holocene in relation to topographic and structural parameters. In case of DSGD, sackungs may also evolve into earthflow or rotational landslides because of water infiltration in neo-joints (Darnault et al., 2011; El Bedoui et al., 2011). In case of landslides located below the trimline, the unstable area was progressively extended above the trimline, following weakened outcrops.

18In Scotland, the recent occurrence of landslides cannot be ruled out (Ballantyne, 2008), so no precise temporal pattern of Holocene landslide activity can be drawn. However, the Storr landslide on Skye (36Cl, 6.5 ka - Ballantyne et al., 1998) shows evidence of a paraglacial origin, followed by various stages of instability.

Fig. 8: Chronological synthesis of post-glacial landslides in the French Alps.

Fig. 8: Chronological synthesis of post-glacial landslides in the French Alps.

S/ Séchilienne, L/ Lauvitel, PMC/ Pré de Madame Carle, C/ Clarée, Cl/ La Clapière. After Bigot-Cormier et al. (2005), Cossart et al. (2008), Delunel et al. (2010), El Bedoui et al. (2011).

3.2 - Icelandic model

19In Iceland, tephrochronology helps dating landslides. In the Skagafjordur area (Northern Iceland), extensive fieldwork (Decaulne et al., 2010; Mercier et al., 2013) was carried out to identify periods prone to landsliding. Over one hundred landslides were identified and mapped (Jónsson, 1957; Pétursson & Saemundsson, 2008; Cossart et al., in press), and ten of them were dated. In all cases, these landslides are later than the emplacement of raised-beaches, and occurred prior to the development of peat deposits, identified on the landslide deposits. Raised-beaches are common features in Iceland and have already been studied and dated in the area (9.6 to 12 ka 14C BP according to Rundgren et al. (1997)). Bogs were systematically cored, and a model of sequence was defined (Mercier et al., 2013): all landslides are older than 4.2 ± 0.1 ka cal. BP (H4 tephra layer) and, in three cases, vegetal remnants are observed (Betula sp.) and are all dated from 7.8 to 8.0 ka cal. BP. Thus, the ages of all landslides are well constrained: these features occurred during the first half of the Holocene, and probably during the Early Birch period, identified in Iceland as a period of vegetation growth (Einarsson, 1991; Ingólfsson, 1991; Óladóttir et al., 2001; Langdon et al., 2010). This period is also known to correspond to the time at which the glacio-isostatic rebound was at its maximum, with a rate of 2.1-9.2 cm yr-1 between 10 ± 0.3 and 8.15 ±0.35 ka cal. BP (Biessy et al., 2008; Le Breton et al., 2010).

20Finally, a main stage of landsliding was identified, which clearly occurred at the beginning of the Holocene in Iceland. In this case, the evolution of the volume of supplied sediment fits well with a rapid exhaustion model, so that no significant dismantlement of mountain slopes occurred during the second half of the Holocene, and probably after the end of the Early Birch period (8.0 ka). This timing further reinforces the idea that paraglaciation mostly acts through the effects of glacio-isostatic rebound in this area.

4 - The contribution of landsliding within the paraglacial sedimentary cascade

4.1 - Landslide connectivity

21While landsliding appears to be a symptom of paraglaciation, a discussion on the ability of paraglaciated basins to deliver sediments is needed, linked to the classic sediment delivery dilemma identified by Walling (1983). Many landforms, and especially landslides, may act as barriers (Fryirs et al., 2007), which affect landscape connectivity at various scales (Meade, 1982). For this reason, the regulators that directly influence how subsystems are connected to each other through geomorphic processes should be studied, following Chorley & Kennedy (1971). These authors show that a basin represents a typical landform assemblage of a sediment cascade. In mountainous areas, this assemblage is subdivided by Schrott et al. (2003) into three subsystems. Subsystem I corresponds to sediment sources and bedrock outcrops, and mostly involves rock-walls in mountainous areas. Sediments delivered from such sources are then stored within subsystems II (slope) and III (valley bottom), while vegetation cover, slope and size material are regulators that can influence the connectivity between these subsystems (Otto & Dikau, 2004; Schrott et al., 2006).

22In the case of paraglacial landsliding, the actions of regulators are first related to the mode of emplacement of landslides. Landslide masses can indeed be deposited entirely within subsystem II (slopes), and thus not in connection with the main streams that could rework them. Reworking is furthermore difficult because these deposits are often uncoupled from streams by gentle slope areas; for instance, fluvial or marine terraces that may act as buffers between hillslopes and the valley bottom (Jarman, 2006; Ballantyne, 2008) (fig. 9A). This pattern is particularly common in Scotland (Arrested Rock-Slope Failure in Jarman (2006)) and high latitude environments: landslides mostly occur at the mouth of valleys and fjords, where valley bottoms are predominantly wide. In Iceland, landslide masses are mostly stored on glacio-fluvial terraces or raised-beaches (101 cases on the 105 recorded in the Skagafjordur area, Northern Iceland). In any case, landslides are disconnected from any geomorphic process, so that they are still preserved and do not contribute to the cascading system.

23In many mountainous alpine areas or in active orogens, paraglacial landslides are located in narrow troughs and, more generally, in the upper part of watersheds. In such cases, landslide masses are mostly stored within subsystem III, but their influence on the cascading system can vary. In many cases, the occurrence of landslides generates persistent dams, which efficiently interrupt sediment delivery. Large sediment traps (aggradational plains) illustrate the fragmentation of such valleys (Hewitt, 2006; Cossart & Fort, 2008a; Fort et al., 2009). If incision of the landslide mass occurs, it is controlled and hampered by river bed armoring. The grain-size material that constitutes the landslide mass then acts as the main regulator, which is of great influence because the stream power of these upper alpine rivers is rather low and cannot remove large boulders. Finally, even if they are in connection with the valley floor and streams, landslide dams may be persistent features, whose evolution and dismantlement patterns control the sediment yield during the entire Holocene period.

Fig. 9: Typology of landslide/valley bottom coupling in a paraglacial setting: geomorphic sketch and rough estimation of sediment delivery (graphs).

Fig. 9: Typology of landslide/valley bottom coupling in a paraglacial setting: geomorphic sketch and rough estimation of sediment delivery (graphs).

(A) Iceland and Svalbard model: no connection/coupling with the valley bottom. Graph shows no input of sediment from landslide to the cascade sedimentary system. (B) Connection and creation of a buffer: reworking of the landslide mass toe, coupled with bank erosion, provides sediments to the river, before a gradual stabilization. (C) Connection and creation of a permanent barrier. After a complete aggradation upstream of the dam, the upstream/downstream connection is made, providing sediments downstream. (D) Connection and reworking of the landslide mass: sediments are provided downstream from dam breaching during the first phase, and then by dam breaching coupled with sediment reservoir erosion during the second phase (sediment yield is then at a maximum).

4.2 - Typology of landslide dam evolution

24Three different situations of landslide/river coupling have been typified (Fort, 2011): partial blockage and diversion of the river, complete damming and impeding of sediment flux, reworking (possibly catastrophically) of the landslide dam by the river.

25In the case of gradual river diversion, rivers can partially rework the landslide mass. However, the evolution of sediment fluxes over time depends on the nature of the opposite bank. When it consists of soft material (slope or alluvial deposits), bank erosion may occur and provide large amounts of sediment (fig. 9B), while sediment transfer from the upper catchment is only slightly slowed. If the opposite bank is cohesive, aggradation predominates upstream and favors the generation of alluvial pockets of floodplain (Phillips, 1992).

26In Scotland, such a diversion occurs in the case of “sub-cataclysmic translational” slides (Jarman, 2006). Yet, the remobilization of sediments at the contact between rivers and hillslopes provoked a net aggradation downstream within Scottish valley floors until 4.0 to 2.0 ka, followed by a net incision (Maizels & Aitken, 1991; Ballantyne & Whittington, 1999). As this period was not significantly affected by climate or anthropogenic changes, the trend from net aggradation to net erosion may be interpreted as an intrinsic self-organization of the cascading system, leading to an exhaustion of sediment supplied from the area affected by this landslide/river interaction (Ballantyne, 2008).

27In narrow valleys and/or when the volume of the landslide is sufficient to block the river, landslide may act as a barrier, i.e. it can disrupt sediments moving along the channel (Fryirs et al., 2007). The creation of such a local base-level blocks sediment conveyance and enhances retention (Hewitt, 2006), with this retention of sediments depending on the size of the reservoir created upstream and the sediment yield of the river. However, if the dam remains stable (bed armoring, lack of seepage, continuous supply of sediment from sources to landslide dam, etc.) sediment transfer is entirely blocked until the reservoir is totally filled. This pattern is well-documented in high-mountainous areas (the French Alps in Cossart & Fort, 2008a) in narrow headwaters. In fact, catastrophic landslides deliver huge debris that cannot be removed by small rivers. Furthermore, landslide masses continuously store some sediments delivered from adjacent slopes (instability maintained by seismotectonic activity and relief in these active orogens), which hampers the reworking of debris masses. In such high-mountain cases, sediment yield is thus very low after the landsliding stages. When reservoirs located upstream of the dams are filled, the uphill aggradational plains created may act as a transfer zone. A partial sediment transfer is then re-established from the upper part of the catchment. Sediment yield is nevertheless affected by pulsation as short-lived dams may temporarily act in accordance with sediment supply from still unstable mountain slopes (fig. 9C).

28Even if a dam occurs, some breaches may appear and create an exportation of sediments (possibly catastrophically). The amount of sediment carried out depends on the nature of the breach (i.e. catastrophic vs. progressive). Of course, catastrophic breaching (incision by overflowing or collapse due to seepage) involves the release of a large amount of sediment and water, the relative proportions of which depend on the sediment infill upstream (and thus the duration of the dam). Extreme peak discharges and high velocities enable the transportation of sediments derived from the dam and from the upstream lacustrine reservoir, but the duration of the peak remains ephemeral. Furthermore, most of the coarse debris supplied forms a wedge immediately downstream of the former dam, so that the effect on the whole cascading system is probably lower than suggested by the violence of the event, as examplified in Rhine valley (Schneider et al., 2004). Some examples (Benito et al., 1998; Brooks & Lawrence, 1999) highlight that the extreme discharge rapidly evolves in both time and space into hyper-saturated flows and ‘normal’ high flows. Once again, in spite of the (potentially dramatic and severe) violence of the event, the final contribution to the cascading system appears somewhat limited.

29If the incision of the dam is progressive, a sediment store can be created upstream and entirely filled in (fig. 9D). The adjustment of the longitudinal profile by retrogressive erosion provokes a progressive re-connection of the sediment cascade, initially by exporting the sediments of the landslide mass, and then by exporting sediments of the former aggradational plain. This second phase is generally associated with a significant rise in the sediment yield as sediments deposited upstream of the dam (silts in many cases) can be easily removed (Cossart & Fort, 2008a). For instance, in the case of the early-Holocene Chenaillet landslide (Cerveyrette valley, Southern French Alps), the second stage has just begun and the present situation corresponds to an amount of sediment export of only 1/60 of the total debris of the reservoir (Cossart & Fort, 2008a). Nevertheless, in highly active orogens river incision is higher in response to the uplift rate. The second stage may thus occur more rapidly (2.0 to 3.0 ka after the creation of the dam) and 50 to 75 % of the total amount of debris may then be evacuated (Hewitt, 2006).

30Finally, paraglaciation stores large amounts of debris in troughs and basins, especially through landsliding, rather than contributing to the sedimentary cascading system. Sediments are thus accumulated prior to further glaciation: glaciers are indeed the only agents able to remove and evacuate such debris (except in active orogens). Therefore, paraglacial landsliding is probably of prime importance in the enlargement and deepening of classic glacial landforms such as troughs and cirques, especially at high latitudes (Bentley & Dugmore, 1998; Mercier, 2011). If so, the glacial processes of excavation and ablation would not entirely explain valley or fjord development during the short Pleistocene time scale.

5 - Conclusion

31The comparative approach presents three main results concerning the role of paraglaciation in landsliding. Firstly, the effectiveness of paraglaciation in mountain-slope destabilization can be confirmed, while the processes that predispose or trigger instability are more varied than expected. Although post-glacial decohesion is identified in various settings, it is often coupled with the over-deepening of valleys and slope-steepening to generate instability. In Iceland, the role of glacio-isostatic rebound is demonstrated; such relationship between landsliding and rebound can probably be applied in other areas covered by inlandsis, such as Greenland, Svalbard, Scandinavia, Canada, etc. However, further research is needed. Secondly, paraglaciation appears to influence strongly the location of landslides: over-deepened and/or narrow valleys in mountainous areas, mouths of fjords in high-latitude areas. However, the exception of active orogens is noticeable: seismotectonic activity is of prime importance here and should be coupled with paraglaciation to explain landslide distribution. Thirdly, a paraglacial period prone to landsliding is probable but not statistically proven (except in the case of glacio-isostatic-triggered landslides, as in Iceland) because instability is probably maintained during the whole Holocene by classic factors of instability (relief, lithology, structure, etc.).

32Even if landsliding appears to be the main process leading to the dismantlement of sediment sources in formerly glaciated areas, its contribution to the cascade sedimentary system is probably lower than expected for two main reasons. First, most landslides (especially in high latitudes) are uncoupled from other geomorphic processes; second, if they reach the valley bottom, landslides often act as persistent dams: only particular local settings may provoke sediment transfers (for instance, material prone to seepage, high uplift rate in very active orogens).

33Finally, paraglacial landslides are efficient to dismantle mountain slopes after the deglaciation, providing large amounts of debris. Nevertheless, these sediments are mostly stored within reservoirs located in troughs and basins, and the final evacuation of paraglacial sediments remains rather low. It is thus suggested that paraglaciation is involved within troughs, cirques and fjords enlargement/deepening. Further sediment budgets would certainly specify and quantify this pattern.

Haut de page

Bibliographie

ANDRÉ M.-F., 1993 - Les versants du Spitsberg, approche géographique des paysages polaires. Presses Universitaires de Nancy, Nancy, 361 p.

ANDRÉ M.-F., 1997 - Holocene rockwall retreat in Svalbard: a triple-rate evolution. Earth Surface Processes and Landforms, 22 (5), 423-440.

BALLANTYNE C.K., STONE J.O. & FIFIELD L.K., 1998 - Cosmogenic Cl-36 dating of postglacial landsliding at The Storr, Isle of Skye, Scotland. The Holocene, 8 (3), 347-351.

BALLANTYNE C.K. & WHITTINGTON G., 1999 - Late Holocene floodplain incision and alluvial fan formation in the central Grampian Highlands, Scotland: chronology, environment and implications. Journal of Quaternary Science, 14 (7), 651-671.

BALLANTYNE C.K., 2002a - Paraglacial geomorphology. Quaternary Science Reviews, 21 (18), 1935-2017.

BALLANTYNE C.K., 2002b - A general model of paraglacial landscape response. The Holocene, 12 (3), 371-376.

BALLANTYNE C.K., 2003a - Paraglacial landsystems. In D.J. Evans (ed.), Glacial landsystems. E. Arnold, London, 432-461.

BALLANTYNE C.K., 2003b - Paraglacial landform succession and sediment storage in deglaciated mountain valleys: theory and approaches to calibration. Zeitschrift für Geomorphologie. Supplementband, 32, 1-18.

BALLANTYNE C.K., 2003c - A Scottish sturzstrom: the Beinn Alligin rock avalanche, Wester Ross. Scottish Geographical Journal, 119 (2), 159-167.

BALLANTYNE C.K., 2008 - After the Ice: Holocene geomorphic activity in the Scottish Highlands. Scottish Geographical Journal, 124 (1), 8-52.

BENITO G., GRODEK T. & ENZEL Y., 1998 - The geomorphic and hydrologic impacts of the catastrophic failure of flood-control dams during the 1996-Biescas flood (central Pyrenees, Spain). Zeitschrift fur Geomorphologie, 42 (4), 417-437.

BENTLEY M.J. & DUGMORE A., 1998 - Landslides and the rate of glacial through formation in Iceland. Journal of Quaternary Science, 13 (6), 11-15.

BIESSY G., DAUTEUIL O., VAN VLIET-LANOË B. & WAYOLLE A., 2008 - Fast and partitioned postglacial rebound of southwestern Iceland. Tectonics, 27 (3), TC3002, 1-18, doi: 10.1029/2007TC002177.

BIGOT-CORMIER F., BRAUCHER R., BOURLÈS D., GUGLIELMI Y., DUBAR M. & STÉPHAN J.-F., 2005 - Chronological constraints on processes leading to large active landslides. Earth and Planetary Science Letters, 235 (1-2), 141–150.

Bois T., Bouissou S. & Jaboyedoff M., 2012 - Influence of structural heterogeneities and of large scale topography on imbricate gravitational rock slope failures: New insights from 3-D physical modeling and geomorphological analysis. Tectonophysics, 526-529, 147-156.

Bouissou S., Darnault R., Chemenda A. & Rolland Y., 2011 - Evolution of gravity-driven rock slope failure and associated fracturing: Geological analysis and numerical modeling. Tectonophysics, 526-529, 157-166.

Bozonnat J.-P., 1980 - Infiltration des eaux dans les calcaires fissurés : hydrogéologie et bilan hydrique du secteur septentrional du Massif de la Chartreuse. Thèse de Doctorat, Université Scientifique et Médicale de Grenoble, Grenoble, 360 p.

Bravard J.-P., Barthélémy L., Brochier J., Petiot R., Landon N., Joly-Saad M.-C., Evin J., Astrade L., Thiebault S. & Roberts M., 2003 - Mouvements de masse et paléoenvironnement quaternaire : les paléo-lacs de Boulc (Haut-Diois, Alpes, France). Revue de Géographie Alpine, 91 (1), 9-27.

BROOKS G.R. & LAWRENCE D.E., 1999 - The drainage of the Lake Ha!Ha! reservoir and downstream geomorphic impacts along Ha!Ha! River, Saguenay area, Quebec, Canada. Geomorphology, 28 (1-2), 141-168.

Buoncristiani J.-F., Petit C., Campy M., Richard H. & Bossuet G. 2002 - Quantification de l’ablation d’un bassin versant marno-calcaire alpin durant le Petit Age Glaciaire par l’étude d’un système lacustre (cas du « Claps » de Luc-en-Diois). Geodinamica Acta, 15 (2), 103-111.

CHORLEY R.J. & KENNEDY, B.A., 1971 - Physical geography: a systems approach. Prentice-Hall International, London, 370 p.

CHURCH M. & RYDER J.M., 1972 - Paraglacial sedimentation: a consideration of fluvial processes conditioned by glaciation. Geological Society of America Bulletin, 83 (10), 3059-3072.

COSSART E., 2008 - Landform connectivity and waves of negative feedbacks during the paraglacial period, a case study: the Tabuc subcatchment since the end of the Little Ice Age (Massif des Ecrins, France). Géomorphologie: Relief, Processus, Environnement, 2008 (4), 249-260.

COSSART E., BRAUCHER R., FORT M., BOURLÈS D.L. & CARCAILLET J., 2008 - 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. Geomorphology, 95 (1-2), 3-26.

COSSART E. & FORT M., 2008a - Consequences of landslide dams on alpine river valleys: examples and typology from the French Southern Alps. Norsk Geografisk Tidsskrift - Norwegian Journal of Geography, 62 (2), 75-88.

COSSART E. & FORT M., 2008b - Sediment release and storage in early deglaciated areas: Towards an application of the exhaustion model from the case of Massif des Écrins (French Alps) since the Little Ice Age. Norsk Geografisk Tidsskrift - Norwegian Journal of Geography, 62 (2), 115-131.

COSSART E., MERCIER D., DECAULNE A., FEUILLET T., JÓNSSON H.P. & SÆMUNDSSON Þ, in press - Impacts of paraglaciation on landslides spatial distribution at a regional scale in Northern Iceland (Skagafjörður). Earth Surface Processes and Landforms.

CRUDEN D.M. & HU X.Q., 1993 - Exhaustion and steady state models for predicting landslide hazards in the Canadian Rocky Mountains. Geomorphology, 8 (4), 279-285.

Darnault R., Rolland Y., Braucher R., Bourlès D., Revel-Rolland M., Sanchez G., Bouissou S., 2011 - Timing of the last deglaciation revealed by receding glaciers at the Alpine-scale: impact on mountain geomorphology. Quaternary Science Reviews, 31, 127-142.

DECAULNE A., MERCIER D., COSSART E., FEUILLET T., SÆMUNDSSON Þ. & JÓNSSON H.P., 2010 - The Höfðahólar rock avalanche in Skagafjörður, Northern Iceland: geomorphological characteristics & relative dating. In Þ. Sæmundsson, A. Decaulne & A.A. Beylich (eds.), 5th I.A.G./A.I.G. SEDIBUD Workshop: Sediment Budgets in Cold Environments “Qualitative and Quantitative Analysis of Sedimentary Fluxes and Budgets in Changing Cold Climate Environments - Field-Based Approaches and Monitoring”. Sauðárkrókur, Iceland: September 19th – 25th, 2010, Extended Abstracts Contributions. Náttúrustofa Norðulands vestra, Sauðárkróki, 58 p.

Delunel R., Hantz D., Braucher R., BourlÈs D. L., Schoeneich P. & Deparis J., 2010 - Surface exposure dating and geophysical prospecting of the Holocene Lauvitel rock slide (French Alps). Landslides, 7 (4), 393-400.

Dikau R., Brunsden D., Schrott L. & Ibsen M.L., 1996 - Landslide recognition. Wiley, Chichester, 251 p.

EINARSSON Þ., 1991 - Geology of Iceland, rocks and landscape. Mál og Menning, Reykjavík, 309 p.

EL BEDOUI S., BOIS T., JOMARD H., SANCHEZ G., LEBOURG T., TRICS E., GUGLIELMI Y., BOUISSOU S., CHEMENDA A., ROLLAND Y., CORSINI M. & PÉREZ J.L., 2011 - Paraglacial gravitational deformations in the SW Alps: a review of field investigations, 10Be cosmogenic dating and physical modelling. Geological Society Special Publications, 351, 11-25.

ETIENNE S., MERCIER D. & VOLDOIRE O., 2008 - Temporal scales and deglaciation rhythms in a polar glacier margin, Baronbreen, Svalbard. Norsk Geografisk Tiddskrift - Norwegian Journal of Geography, 62 (2), 102-114.

FORT M., COSSART E., DELINE P., DZIKOWSKI M., NICOUD G., RAVANEL L., SCHOENEICH P. & WASSMER P., 2009 - Geomorphic impacts of large and rapid mass movements: a review . Géomorphologie : Relief, Processus, Environnement, 2009 (1), 47-64.

FORT M., 2011 - Two large late Quaternary rock slope failures and their geomorphic significance, Annapurna Himalayas (Nepal). Geografia Fisica e Dinamica Quaternaria, 34 (1), 5-14.

FRYIRS K., BRIERLEY G. J., PRESTON N.J. & KASAI M., 2007 - The (dis)connectivity of catchment-scale sediment cascades. Catena, 70 (1), 49-67.

Gidon M., 1990 - Les décrochements et leur place dans la structuration du Massif de la Chartreuse (Alpes occidentales françaises). Géologie Alpine, 66, 39- 55.

GODARD A., 1961 - L'efficacité de l'érosion glaciaire en Ecosse du Nord. Revue de Géomorphologie Dynamique, 12 (1), 32-42.

GUTIÉRREZ-SANTOLALLA F., ACOSTA E., RÍOS S., GUERRERO J. & LUCHA P., 2005 - Geomorphology and geochronology of sackung features (uphill-facing scarps) in the Central Spanish Pyrenees. Geomorphology, 69 (1-4), 298-314.

HEWITT K., 2006 - Disturbance regime landscapes: mountain drainage systems interrupted by large rockslides. Progress in Physical Geography, 30 (3), 365-393.

Hippolyte J.-C., Brocard G., Tardy M., Nicoud G., Bourlès D., Braucher R., Ménard G. & Souffaché B., 2006 - The recent fault scarps of the Western Alps (France): tectonic surface ruptures or gravitational sackung scarps? A combined mapping, geomorphic, levelling, and 10Be dating approach. Tectonophysics, 418 (3-4), 255-276.

Hippolyte J.-C., Bourlès D., Braucher R., Carcaillet J., Léanni L., Arnold M. & Aumaitre G., 2009 - Cosmogenic 10Be dating of a sackung and its faulted rock glaciers, in the Alps of Savoy (France). Geomorphology, 108 (3), 312-320.

INGÓLFSSON Ó., 1991 - A review of the Late Weichselian and early Holocene glacial and environmental history of Iceland. In J.K. Maizels & C.J. Caseldine (eds), Environmental Changes in Iceland: Past and Present. Kluwer, Dordrecht, 13-29.

JARMAN D., 2006 - Large rock slope failures in the Highlands of Scotland: characterisation, causes and spatial distribution. Engineering Geology, 83 (1-3), 161-182.

JARMAN D., 2009 - Paraglacial rock slope failure as an agent of glacial trough widening. Geological Society Special publications, 320, 102-131.

JÓNSSON Ó., 1957 - Skriðuföll og snjóflóð. Tome 1. Bókaútgáfan norðri, Akureyri, 141 p.

KNIGHT J. & HARRISON S. (eds.), 2009 - Periglacial and paraglacial processes and environments. Geological Society Special publications, 320, 272 p.

KORUP O., 2009 - Linking landslides, hillslope erosion, and landscape evolution. Earth Surface Processes and Landforms, 34 (9), 1315-1317.

LANG A., MOVA J., COROMINAS J., SCHROTT L. & DIKAU R. 1999 - Classic and new methods for assessing the temporal occurrence of mass movements. Geomorphology, 30, 33–52.

LANGDON P.G., LENG M.J., HOLMES N. & CASELDINE C.J., 2010 - Lacustrine evidence of early-Holocene environmental change in northern Iceland: a multiproxy palaeoecology and stable isotope study. The Holocene, 20 (2), 205-214.

LE BRETON E., DAUTEUIL O. & BIESSY G., 2010 - Post-glacial rebound of Iceland during the Holocene. Journal of the Geological Society London, 167 (2), 417-432.

LE CŒUR C., 1999 - Rythmes de dénudation tertiaire et quaternaire en Ecosse occidentale/Tertiary denudation and Quaternary erosion rates in Western Scotland. Géomorphologie: Relief, Processus, Environnement, 5 (4), 291-303.

LEWIS W.V., 1954 - Pressure release and glacial erosion. Journal of Glaciology, 16, 417-422.

MCCOLL S.T., 2012 - Paraglacial rock-slope stability. Geomorphology, 153-154, 1-16.

MAIZELS J.K. & AITKEN J.F., 1991 - Palaeohydrological change during deglaciation in upland Britain: a case study from northeast Scotland. In L. Starkel, K.J. Gregory & J.B. Thornes (eds.), Temperate Palaeohydrology: fluvial processes in the temperate zone during the last 15000 years. John Wiley and Sons, Chichester, 105-145.

MEADE R.H., 1982 - Sources, sinks and storage of river sediment in the Atlantic drainage of the United States. The Journal of Geology, 90 (3), 235-252.

Mège D. & Bourgeois O., 2011 - Equatorial glaciations on Mars revealed by gravitational collapse of Valles Marineris wallslopes. Earth and Planetary Science Letters, 310 (3-4), 182-191.

MERCIER D., 2002 - La dynamique paraglaciaire des versants du Svalbard. Zeitschrift für Geomorphologie, 46 (2), 203-222.

MERCIER D., 2007 - Le paraglaciaire: évolution d’un concept. In M.-F. André, S. Etienne, Y. Lageat, C. Le Cœur & D. Mercier (eds.), Du continent au bassin versant : Théories et pratiques en géographie physique : Hommage au Professeur Alain Godard. Presses universitaires Blaise Pascal, Clermont-Ferrand, 341-353.

MERCIER D., 2011 - La géomorphologie paraglaciaire : Changements climatiques, fonte des glaciers et crises érosives associées. Editions universitaires européennes, Saarbrücken, 256 p.

MERCIER D., COSSART E., DECAULNE A., FEUILLET T., SÆMUNDSSON Þ. & Jónsson H.P., 2013 - The Höfðahólar rock avalanche (sturzström): chronological constraint of a paraglacial landsliding on Icelandic hillslope. The Holocene, doi: 10.1177/0959683612463104.

MONNIER S., 2006 - Les glaciers-rocheux, objets géographiques: analyse spatiale multiscalaire et investigations environnementales. Thèse de Doctorat, Université Paris-Est Créteil Val de Marne, Paris, 678 p.

ÓLADÓTTIR R., SCHLYTER P. & HARALDSSON H., 2001 - Simulating Icelandic vegetation cover during the Holocene, implication for long-term land degradation. Geografiska Annaler, 83 (4), 203-215.

OTTO J.C. & DIKAU R., 2004 - Geomorphologic system analysis of a high mountain valley in the Swiss Alps. Zeitschrift für Geomorphologie, 48 (3), 323-341.

PANIZZA M., 1973 - Glacio pressure implications in the production of landslides in the dolomitic area. Geologia Applicata e Idrogeologia, 8 (1), 289-297.

PÉTURSSON H.G. & SAEMUNDSSON Þ., 2008 - Skriðuföll í Skagafirði. In Þ. Sæmundsson, A. Decaulne & H.P. Jónsson (eds.), Skagfirsk náttúra 2008: Málþing um náttúru Skagafjarðar: Sauðárkrókur, 12. apríl 2008. Náttúrustofa Norðulands vestra, Sauðárkróki, 25-30.

PEULVAST J.-P., 1985 - Relief, érosion différentielle et morphogenèse dans un bourrelet montagneux de haute latitude: Lofoten-vesteralen et Sogn-Jotun (Norvège). Thèse d’Etat, Université Paris 1 Panthéon-Sorbonne, Paris, 1642 p.

PHILLIPS J.D., 1992 - Nonlinear dynamical systems in geomorphology: revolution or evolution? Geomorphology, 5 (3-5), 219-229.

RITTER D.F. & TEN BRINK N.W., 1986 - Alluvial fan development and the glacial-glaciofluvial cycle. Nenana Valley, Alaska. Journal of Geology, 94 (4), 613-615.

RUNDGREN M., INGOLFSSON O., BJÖRCK S., JIANG H. & HAFLIDASON H., 1997 - Dynamic sea-level change during the last deglaciation of northern Iceland. Boreas, 26, 201-215.

RYDER J.M., 1971a - The stratigraphy and morphology of para-glacial alluvial fans in the south-central British Columbia. Canadian Journal of Earth Sciences, 8 (2), 279-298.

RYDER J.M., 1971b - Some aspects of the morphometry of paraglacial alluvial fans in south-central British Columbia. Canadian Journal of Earth Sciences, 8 (2), 1252-1264.

SAINTOT A., HENDERSON I.H.C. & DERRON M.-H., 2011 - Inheritance of ductile and brittle structures in the development of large rock slope instabilities: examples from western Norway. Geological Society Special Publications, 351, 27-78.

Schneider J.-L., Pollet N., Chapron E., Wessels M. & Wassmer P., 2004 - Signature of Rhine Valley sturzstrom dam failures in Holocene sediments of Lake Constance, Germany. Sedimentary Geology, 169, 75-91.

SCHROTT L., HUFSCHMIDT G., HANKAMMER M., HOFFMANN T. & DIKAU R., 2003 - Spatial distribution of sediment storage types and quantification of valley fill deposits in an alpine basin, Reintal, Bavarian Alps, Germany. Geomorphology, 55 (1-4), 45-63.

SCHROTT L., GÖTZ J., GEILHAUSEN M. & MORCHE D., 2006 - Spatial and temporal variability of sediment transfer and storage in an Alpine basin (Bavarian Alps, Germany). Geographica Helvetica, 3, 191-201.

SELLIER D. & LAWSON T.J., 1998 - A complex slope failure on Beinn nan Cnaimhseag, Assynt, Sutherland. Scottish Geographical Magazine, 114 (2), 85-93.

SELLIER D., 2008 - Les glissements translationnels paraglaciaires et l’évolution des grands reliefs monoclinaux du Sutherland occidental (Highlands d’Ecosse). Bulletin de l’Association des Géographes Français, 2, 141-152.

SOLDATI M., CORSINI A. & PASUTO A., 2004 - Landslides and climate change in the Italian Dolomites since the Late glacial. Catena, 55 (2), 141-161.

VAN HUSEN D., IVY-OCHS S. & ALFIMOV V., 2007 - Mechanism and age of late glacial landslides in the Calcareous Alps; The Almtal, Upper Austria. Austrian Journal of Earth Sciences, 100, 114-126.

VON POSCHINGER A., WASSMER P. & MAISCH M., 2006 - The Flims rockslide: history of interpretation and new insights. In S.G. Evans, G. Scarascia Mugnossa, A. Strom & R.L. Hermanns (eds.), Landslides from Massive Rock Slope Failure. NATO Science Series. Series IV, Earth and Environmental Sciences, 49, 329-356.

WALLING D.E., 1983 - The sediment delivery problem. Journal Hydrology, 65 (1-3), 209-237.

Haut de page

Table des illustrations

Titre Fig. 1: Simplified paraglacial sediment cascade (after Ballantyne, 2002a,b).
URL http://quaternaire.revues.org/docannexe/image/6444/img-1.png
Fichier image/png, 233k
Titre Fig. 2: The paraglacial period.
Légende (A) paraglacial period defined by Church & Ryder (1972). (B) application of the exhaustion model to assess the evolution of the volume of sediments within a paraglacial store (Ballantyne, 2003b). Curves noted from 1 to 5 correspond to various calibrations of the exhaustion model; 1/ case of a durable storage within the store, 5/ case of a rapid degradation after deglaciation
URL http://quaternaire.revues.org/docannexe/image/6444/img-2.png
Fichier image/png, 123k
Titre Fig. 3: Evidence of post-glacial decohesion.
Légende (A) Joints parallel to palaeo-glacier fluxes at Stuphallet in Svalbard (D. Mercier, July 2004). (B) Block detachment from a cirque free-face in the Skagafjordur area due to a combination of post-glacial neo-joint and basalt bedding (E. Cossart, June 2011). (C) Splay-shaped joints in a rochemoutonnée (Clarée valley, southern French Alps) made of carboniferous sandstone (E. Cossart, June 2004). (D) Cracks and consequent differential lowering affecting the top of a roche-moutonnée made of sandstone (E. Cossart, June 2004).
URL http://quaternaire.revues.org/docannexe/image/6444/img-3.png
Fichier image/png, 5,7M
Titre Fig. 4: Consequences of over-deepening of hillslope basal parts on landslide triggering.
Légende A/ Case of a dip slope (Scotland, adapted from Sellier, 2008). B/ Case of a counterdip slope
URL http://quaternaire.revues.org/docannexe/image/6444/img-4.png
Fichier image/png, 25k
Titre Fig. 5: Typology of paraglacial landslide location.
Légende 1/ Rotational slides in counterdip slopes, 2/ Rock-topples on glacially polished knobs (slope facing former glacier-fluxes), 3/ Translational slides in dip slopes, 4/ Landslides in narrow, over-deepened valleys (over-deepening in relation to fault emplacement), 5/ Landslides in zones of confluence.
URL http://quaternaire.revues.org/docannexe/image/6444/img-5.png
Fichier image/png, 228k
Titre Fig. 6: Comparison of landslide location patterns in the upper and lower parts of formerly glaciated valleys.
Légende (A) Location map. (B) Extent of glaciers during the Last Glacial Maximum. (C) distribution of landslides (above vs. below the trimline) in the lower part of the Drac valley (former ablation zone), realized from the BRGM database. (D) Distribution of landslides (above vs. below the trimline) in the Gyronde area (former accumulation zone), realized from field investigations. The comparison of altitudes, both landslide scars and toe deposits, in cases C and D is statistically tested by a Fischer test; in each case the observed F is higher than the significance threshold (with an uncertainty of 0.05).
URL http://quaternaire.revues.org/docannexe/image/6444/img-6.png
Fichier image/png, 709k
Titre Fig. 7: Relationship between landslide location and glacio-isostatic rebound in Iceland (Skagafjörður).
Légende (A) Over-representation of landslide at the mouth of the fjord (fjord oriented north-south) estimated from a chi-square analysis in comparison with randomly distributed landslides (assessed in hectares). (B) and (C) Sketch of glacio-isostatic rebound following inlandsis disappearance. (D) Altitudes of raised-beaches in Iceland (Skagafjordur).
URL http://quaternaire.revues.org/docannexe/image/6444/img-7.png
Fichier image/png, 132k
Titre Fig. 8: Chronological synthesis of post-glacial landslides in the French Alps.
Légende S/ Séchilienne, L/ Lauvitel, PMC/ Pré de Madame Carle, C/ Clarée, Cl/ La Clapière. After Bigot-Cormier et al. (2005), Cossart et al. (2008), Delunel et al. (2010), El Bedoui et al. (2011).
URL http://quaternaire.revues.org/docannexe/image/6444/img-8.png
Fichier image/png, 102k
Titre Fig. 9: Typology of landslide/valley bottom coupling in a paraglacial setting: geomorphic sketch and rough estimation of sediment delivery (graphs).
Légende (A) Iceland and Svalbard model: no connection/coupling with the valley bottom. Graph shows no input of sediment from landslide to the cascade sedimentary system. (B) Connection and creation of a buffer: reworking of the landslide mass toe, coupled with bank erosion, provides sediments to the river, before a gradual stabilization. (C) Connection and creation of a permanent barrier. After a complete aggradation upstream of the dam, the upstream/downstream connection is made, providing sediments downstream. (D) Connection and reworking of the landslide mass: sediments are provided downstream from dam breaching during the first phase, and then by dam breaching coupled with sediment reservoir erosion during the second phase (sediment yield is then at a maximum).
URL http://quaternaire.revues.org/docannexe/image/6444/img-9.png
Fichier image/png, 398k
Haut de page

Pour citer cet article

Référence papier

Etienne Cossart, Denis Mercier, Armelle Decaulne et Thierry Feuillet, « An overview of the consequences of paraglacial landsliding on deglaciated mountain slopes: typology, timing and contribution to cascading fluxes  », Quaternaire, vol. 24/1 | 2013, 13-24.

Référence électronique

Etienne Cossart, Denis Mercier, Armelle Decaulne et Thierry Feuillet, « An overview of the consequences of paraglacial landsliding on deglaciated mountain slopes: typology, timing and contribution to cascading fluxes  », Quaternaire [En ligne], vol. 24/1 | 2013, mis en ligne le 01 mars 2016, consulté le 23 novembre 2017. URL : http://quaternaire.revues.org/6444 ; DOI : 10.4000/quaternaire.6444

Haut de page

Auteurs

Etienne Cossart

Université Paris 1, Laboratoire PRODIG, 2 rue Valette, F-75005 PARIS. Courriel : etienne.cossart@univ-paris1.fr; CNRS - GDR 6032 "Mutations polaires", 30 rue Mégevand, F-25030 BESANÇON cedex.

Articles du même auteur

Denis Mercier

 Université de Nantes, Laboratoire LETG-Nantes-Géolittomer UMR 6554 CNRS, Campus du Tertre, BP 81227, F-44312 NANTES cedex 3. Courriel : denis.mercier@univ-nantes.fr;  Institut Universitaire de France, 103 boulevard Saint-Michel, 75005 PARIS;CNRS - GDR 6032 "Mutations polaires", 30 rue Mégevand, F-25030 BESANÇON cedex.

Armelle Decaulne

Université de Nantes, Laboratoire LETG-Nantes-Géolittomer UMR 6554 CNRS, Campus du Tertre, BP 81227, F-44312 NANTES cedex 3. Courriel : armelle.decaulne@univ-nantes.fr;CNRS - GDR 6032 "Mutations polaires", 30 rue Mégevand, F-25030 BESANÇON cedex.

Thierry Feuillet

Université de Nantes, Laboratoire LETG-Nantes-Géolittomer UMR 6554 CNRS, Campus du Tertre, BP 81227, F-44312 NANTES cedex 3. Courriel : thierry.feuillet@univ‑nantes.fr;  CNRS - GDR 6032 "Mutations polaires", 30 rue Mégevand, F-25030 BESANÇON cedex.

Haut de page

Droits d’auteur

© Tous droits réservés

Haut de page
  • Logo Association française pour l’étude du Quaternaire
  • Revues.org