an overview of the consequences of paraglacial

Quaternaire, 24, (1), 2013, p. 13-24
AN OVERVIEW OF THE CONSEQUENCES OF PARAGLACIAL
LANDSLIDING ON DEGLACIATED MOUNTAIN SLOPES:
TYPOLOGY, TIMING AND CONTRIBUTION
TO CASCADING FLUXES
n
Étienne COSSART1,4, Denis MERCIER2,3,4, Armelle DECAULNE2,4,
Thierry FEUILLET2,4
ABSTRACT
Three decades after the early definition of the «paraglacial» concept, a general model of paraglaciation was established, integrating the typical sediment paraglacial cascading system and the responses of associated sediment storages. The residuals of this
model should now be examined and explained through controlling factors that may act at both regional (e.g. tectonics) and local
(topography, lithology, etc.) scales. We compare here the patterns identified in a few mountain ranges of the northern hemisphere,
located in various seismotectonic settings (especially western Alps, northwestern Scotland, central Norway, Svalbard and Iceland).
By combining our field observations with a literature review, the effectiveness of the paraglacial response on mountain-slope erosion
and the contribution of paraglaciation to the sedimentary cascading system are discussed. In most cases, paraglaciation generates
sediment storage and aggradation in headwaters and glacial trough, as the evacuation rate into sinks remains low. Finally, paraglaciation appears to be a period of hillslope denudation, preparing reservoirs of sediments, which can only be effectively evacuated
during further glaciation periods.
Keywords: mountain slopes, paraglaciation, landsliding, cascading system, sediment fluxes, northern hemisphere
RÉSUMÉ
CONSÉQUENCES DES MOUVEMENTS DE MASSE PARAGLACIAIRES SUR LES VERSANTS DÉSENGLACÉS : TYPOLOGIE, TEMPORALITÉ, CONTRIBUTIONS AUX BUDGETS SÉDIMENTAIRES
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’Écosse, Norvège centrale, Svalbard et l’Islande). En combinant nos observations de terrain avec une recherche 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.
Mots-clés : versants, paraglaciaire, mouvement de terrain, cascade sédimentaire, montagnes, flux sédimentaires, hémisphère nord
1 - INTRODUCTION
More 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 and 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
Université Paris 1, Laboratoire PRODIG, 2 rue Valette, F-75005 PARIS. Courriel : [email protected]
Université de Nantes, Laboratoire LETG-Nantes-Géolittomer UMR 6554 CNRS, Campus du Tertre, BP 81227, F-44312 NANTES cedex 3.
Courriels : [email protected], [email protected], [email protected]
3
Institut Universitaire de France, 103 boulevard Saint-Michel, 75005 PARIS.
4
CNRS - GDR 6032 «Mutations polaires», 30 rue Mégevand, F-25030 BESANÇON cedex.
1
2
Manuscrit reçu le 24/05/2012, accepté le 05/01/2013
1301-066 - Mep 1-2013.indd 13
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14
Sediment
sources
Primary
sediment
stores
Rockwalls
Drift-mantled
slopes
Valley-floor
glacigenic deposits
Costal
glacigenic deposits
Rock-slope failure
Rockfall
Slope failure
Debris flow
Gullying
Snow avalanches
Debris flow
Fluvial reworking
Wave action
Nearshore currents
Alluvial fans
Valley fill
Barrier beaches
Spits
Baymouth bars
Barrier islands
Back-barrier deposits
Rockslide deposits
Talus
(Rock glaciers)
Debris-flow deposits
Debris cones
Avalanche tongues
Alluvial fans
River incision
Terrace formation
Fluvial reworking
Secondary
sediment
stores
Alluvial
Valley-fill deposits
Lacustrine deposits
(bottom sediments,
deltas)
Coastal deltas
Fjord deposits
Barrier structures
Fluvial reworking
Reworking by waves and currents
Sediment
sinks
Alluvial
Valley-fill deposits
Lacustrine
deposits
Coastal and
neashore deposits
Shelf and
offshore deposits
Fig. 1: Simplified paraglacial sediment cascade (after Ballantyne, 2002a,b).
Fig. 1 : Schéma simplifié du transfert sédimentaire en cascade en milieu paraglaciaire (d’après Ballantyne, 2002a,b).
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; Étienne
et al., 2008; Knight & Harrison, 2009).
In 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
1301-066 - Mep 1-2013.indd 14
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.
This 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).
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B
Beginning of the
deglaciation
1,0
S = (1- e-λt) e -κt
λ = 1.0 ka-1
0,9
Value of κ :
1
0
0,8
End of the
deglaciation
End of the
paraglacial period
Proglacial period
Paraglacial period
Time
Volume of sediments stored
(S/S0)
Volume of available sediments
Sediment yield
A
0,7
2
0,6
0,10
0,5
3
0,4
0,3
0,2
4
0,1
5
0,25
0,5
0,75
0
0
1
2
3
4
5
Time elapsed since the deglaciation (ka)
6
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.
Fig. 2 : La période paraglaciaire. (A) la période paraglaciaire définie par Church & Ryder (1972). (B) estimation, par le modèle de tarissement, de
l’évolution du volume de sédiments stockés dans un réservoir paraglaciaire (Ballantyne, 2003b). Les courbes 1 à 5 correspondent à différentes calibrations de l’équation de tarissement sédimentaire ; 1/ cas d’un stockage durable dans le réservoir sédimentaire ; 5/ cas d’un déstockage rapide du
réservoir après la déglaciation.
2 - PARAGLACIAL LANDSLIDING IN SPACE
2.1 - IMPACTS AT OUTCROP SCALE
At 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.
In 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.
In 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
1301-066 - Mep 1-2013.indd 15
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. 3C and D).
2.2 - IMPACTS AT HILLSLOPE SCALE
At 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.
Paraglacial 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
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A
C
B
D
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).
Fig. 3 : Indices de décohésion post-glaciaire. (A) Diaclases parallèles au paléo-glacier à Stuphallet, Svalbard (D. Mercier, juillet 2004). (B) Détachement de blocs à partir d’une paroi de cirque dans le Skagafjörður, lié à la combinaison d’une détente post-glaciaire et au litage des basaltes (E. Cossart,
Juin 2011). (C) Structure de diaclases en éventail au sein d’une roche-moutonnée (vallée de la Clarée, Alpes du Sud françaises) en grès carbonifère (E.
Cossart, Juin 2004). (D) Fissures et abaissements différentiels affectant le sommet d’une roche-moutonnée en grès (E. Cossart, Juin 2004).
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.
Such 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).
1301-066 - Mep 1-2013.indd 16
2.3 - IMPACTS AT REGIONAL SCALE
At 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
In 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 stee-
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Dip-slope
A
Palaeo-glacier
surface
re
Failu
Basal Quartzitze
ce
surfa
Debuttressing
Bedding 10-11°
Scarp zone
Displaced material
Basal Quartzitze
Former
topography
Bedding 10-11°
B
Failure
surface
Palaeo-glacier
surface
Counterdip
slope
Former
topography
Debuttressing
Joints
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. 4 : Conséquences du surcreusement de la partie basale de versants sur le déclenchement de glissements de terrain. (A) Cas d’un versant conforme
au pendage (Ecosse, adapté de Sellier, 2008). (B) Cas d’un versant à contre-pendage.
tion 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.
pening 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 excava-
Confluence
Narrow
trough
Joints
Joints
Roche-moutonnée
4
Palaeo-flux
Scar
5
3
1
2
φ
Landslide mass
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.
Fig. 5 : Typologie des emplacements de glissements de terrain paraglaciaires. 1/ Glissement rotationnel dans les pentes à contre-pendage, 2/ Basculement de blocs sur les protubérances polies par les glaces (notamment sur la pente faisant face aux anciens flux glaciaires), 3/ Glissements de terrain en
versant conforme au pendage, 4/ Glissements de terrain dans des vallées étroites et sur-creusées. 5/ Glissements de terrain dans des zones de confluence.
1301-066 - Mep 1-2013.indd 17
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2.3.2 - Active orogens
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).
Grenoble
St Marcellin
La Clarée
Drac Massif des
Ecrins
Briançon
D
N
s
yra
C
c
ran
Du
e
Gap
e
Qu
300
150
Fobs (Alt max) = 1,95
> Fseuil = 1,81
0
-150
Fobs (Alt min) = 1,84
> Fseuil = 1,81
-300
-450
-600
0
5
10
15
20
25
Mass-movement identifiant Data source : BD Mvt - BRGM
800
Sisteron
600
45°N
rance
Clarée
Du
6°45'E
Marseille
Verdon
0
10 km
Maximal pleistocene advance (MIS 6 or 4)
Maximal würmian advance (MIS 2)
Accumulation zone of glaciers
Relative elevation of
mass-movement (m)
(0 = trimline altitude)
Alps
The
1000
Lyon
Rhône
In both Scotland and active orogen areas, paraglacial
destabilization may act through debuttressing and oversteepening, 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,
C
Ubaye
A
2.3.3 - Iceland
450
Romanche
Relative elevation of
mass-movement (m)
(0 = trimline altitude)
B
Isè
re
In 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 (Buoncristiani 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 strikeslip 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).
Nevertheless, 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).
400
200
0
-200
-400
-600
-800
D
0
5
10
15
20
25
30
35
Mass-movement identifiant Data source : Fieldwork
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. 6 : Comparaison des principales localisations de glissements de terrain dans les parties supérieure et inférieure de vallées anciennement englacées.
(A) Carte de localisation. (B) Étendue des glaciers au cours du dernier maximum glaciaire. (C) Répartition des glissements de terrain (de part et d’autre
de la trimline) dans la partie inférieure de la vallée du Drac (ex-zone d’ablation), réalisé à partir de la base de données du BRGM. (D) Répartition des
glissements de terrain (de part et d’autre de la trimline) en Gyronde (ex-zone d’accumulation glaciaire), réalisé à partir d’un inventaire effectué sur le
terrain. Les altitudes obtenues dans les cas C et D sont statistiquement testées par un test de Fischer, le F observé est dans chaque cas supérieur au seuil
de significativité (avec une incertitude de 0,05), indiquant des différences significatives entre les deux groupes de glissement de terrain.
1301-066 - Mep 1-2013.indd 18
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19
Over-representation
15
10
A
5
0
-5
-10
Latitude (°)
-15
65,2
to
65,3
65,3
to
65,4
65,4
to
65,5
Under-representation
65,5
to
65,6
65,6
to
65,7
65,7
to
65,8
65,8
to
65,9
65,9
to
66
66
to
66,1
66,1
to
66,2
B
Deglaciation
completed
Cumulated
glacio-isostatic rebound
C
Glacio-isostatic rebound
100
D
Reykir
Saudarkrokur
Center of the
island
Latitude (°)
Hofdaholar
Vatn
50
Fjord sink
65.8
65.9
66.0
Altitude (m.asl)
During the
deglaciation
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).
Nevertheless, 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;
15 000
0
14 000
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 (Skagafjörður).
Fig. 7 : Relation entre la localisation des glissements de terrain et le
rebond glacio-isostatique en Islande (Skagafjörður). (A) Sur-représentation (évaluée en hectares) des glissements de terrain à l’embouchure du fjord (fjord orientée nord-sud), estimé à partir d’une analyse
du khi². (B) et (C) Croquis représentant le rebond glacio-isostatique
suite à la disparition de l’inlandsis. (D) Altitudes de plages soulevées
(Skagafjörður, Islande).
Bolling
Allerod
13 000
Younger Dryas
12 000
11 000
Kartell
10 000
9 000
Age 10Be
paraglaciation acts preferentially through a significant
post-glacial uplift, which induces both rock dilation and
seismotectonic activity.
Older Dryas
8 000
7 000
?
6 000
5 000
3 - PARAGLACIAL LANDSLIDING OVER TIME
Dating 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
Recent results acquired in both Scotland and the European Alps first highlight that post-glacial landslides may
be triggered immediately after glacier disappearance.
1301-066 - Mep 1-2013.indd 19
Suboreal
4 000
3 000
2 000
1 000
S
L
PMC
C
Cl
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).
Fig. 8 : Synthèse chronologique des glissements de terrains postglaciaires datés dans les Alpes françaises. S/ Séchilienne, L/ Lauvitel,
PMC/ Pré de Madame Carle, C/ Clarée, Cl/ La Clapière. D’après
Bigot-Cormier et al. (2005), Cossart et al. (2008), Delunel et al. (2010),
El Bedoui et al. (2011).
20/02/13 12:32
20
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.
In 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.
3.2 - ICELANDIC MODEL
In 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).
Finally, 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
1301-066 - Mep 1-2013.indd 20
(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
While 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 rockwalls 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).
In 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 RockSlope 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 glaciofluvial 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.
In 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
20/02/13 12:32
21
No connection with
valley floor
Landslide
mass
Stream
Stream
Bank erosion
Glacio-fluvial
terrace
+
+
B
0
A
0
Bank erosion
Accumulated sediments
Incision of the stream
Accumulated sediments
Stream
Landslide
mass
Landslide
mass
Landslide
mass
Stream
Stream
Landslide
mass
Stream
Reservoir
incision
+
+
Upstream/Downstream
re-connection
0
0
C
Aggradation
stage
D
Landslide-mass
incision
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).
Fig. 9 : Typologie des couplages glissements de terrain/fond de la vallée dans un contexte paraglaciaire : croquis géomorphologique et estimation approximative de la charge de sédiments à l’exutoire (graphiques). (A) Modèle du Svalbard et de l’Islande: pas de connexion / couplage avec le fond de la vallée.
Le graphique montre l’absence d’apports de sédiments provenant des glissements de terrain. (B) Connexion et remaniement de la partie aval des masses
glissées. L’érosion des berges fournit également des sédiments à la rivière. (C) Connexion et création d’une barrière permanente. Après une aggradation
complète en amont du barrage, une connexion amont / aval se met en place, permettant la fourniture de sédiments en aval. (D) Connexion et remaniement
de la masse glissée : les sédiments sont remaniés à partir de la masse du barrage pendant une première phase. Puis, suite à l’apparition d’une brèche, le
réservoir formé en amont du barrage est à son tour remanié, au cours d’une deuxième phase (la production de sédiments est alors au maximum).
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.
4.2 - TYPOLOGY OF LANDSLIDE DAM EVOLUTION
Three 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.
In the case of gradual river diversion, rivers can
partially rework the landslide mass. However, the evolu-
1301-066 - Mep 1-2013.indd 21
tion 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).
In Scotland, such a diversion occurs in the case of “subcataclysmic 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).
In narrow valleys and/or when the volume of the landslide is sufficient to block the river, landslide may act as
20/02/13 12:32
22
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 highmountainous 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).
Even 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.
If 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,
1301-066 - Mep 1-2013.indd 22
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).
Finally, 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
The 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.).
Even 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
20/02/13 12:32
23
provoke sediment transfers (for instance, material prone to
seepage, high uplift rate in very active orogens).
Finally, 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.
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