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Sediment release and storage in early deglaciatedareas: Towards an application of the exhaustion modelfrom the case of Massif des Écrins (French Alps) sincethe Little Ice AgeEtienne Cossart; Monique Fort

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To cite this Article: Cossart, Etienne and Fort, Monique (2008) 'Sediment releaseand storage in early deglaciated areas: Towards an application of the exhaustionmodel from the case of Massif des Écrins (French Alps) since the Little Ice Age',

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Sediment release and storage in early deglaciated areas: Towards anapplication of the exhaustion model from the case of Massif des Ecrins(French Alps) since the Little Ice Age

ETIENNE COSSART & MONIQUE FORT

Cossart, E. & Fort, M. 2008. Sediment release and storage in early deglaciated areas: Towards an application of the exhaustion

model from the case of Massif des Ecrins (French Alps) since the Little Ice Age. Norsk Geografisk Tidsskrift�Norwegian Journal of

Geography Vol. 62, 115�131. Oslo. ISSN 0029-1951.

Deglaciation in high mountain areas triggers a series of adjustments affecting the various geomorphic components of formerly

glaciated catchments. Active processes, part of the ‘paraglacial’ open system, release and transport sediments from formerly

glacierized sources to catchment sinks. Such cascade sedimentary systems may be disrupted by temporary sediments traps. To

evaluate the consequences of such traps on sediment export, several test areas were selected, where former glacial geometries and

glacial retreat chronology were reconstructed from field data. Special attention is given to the link between the glacial margins and

the glacio-fluvial systems. During glacial retreat, the presence of morainic ridges may temporarily interrupt the sedimentary cascade

system, thus forcing local aggradation and change in the glacio-fluvial pattern. The volume of trapped sediments is controlled by the

volume and the position of morainic ridges, while the storage residence time is strongly dependent upon the rate of moraine erosion.

The ice surface elevation and its lowering control ice-contact storages, as well as sedimentary and water fluxes. The time lag between

the rate of sediment accumulation on the valley floor and the rate of glacial shrinking influences the pattern and behaviour of

drainage many decades after the peak in glacial melting.

Keywords: deglaciation, Little Ice Age, Massif des Ecrins, paraglacial, sediment budget

Etienne Cossart, UMR PRODIG � 8586 CNRS, and University Paris 1 � Pantheon Sorbonne, 2 rue Valette, F-75005 Paris, France.

E-mail: etienne.cossart@univ-paris1.fr; Monique Fort, UMR PRODIG � 8586 CNRS, and University Paris-Diderot (Paris 7), Case

7001, UFR GHSS, Case postale 7001 (Site Javelot), 105 rue de Tolbiac, F-75 205 PARIS cedex 13, France. E-mail: fort@univ-paris-

diderot.fr

Introduction

Establishing a sediment budget generally means quantifying

sediment sources, sinks and pathways within a catchment

cell. During the last decade, new methods (geophysical

soundings, morphometry, GIS framework) have made the

assessment of sediment budgets easier in mountain areas

(Evans 1997; Trimble 1999; Hinderer 2001; Schrott &

Adams 2002; Schrott et al. 2002). The working of the

cascade sedimentary system as defined by Caine (1976) has

thus become better understood. Today, concepts such as

geomorphic coupling of sediment sources and sinks and

coupling of intermediate sediment storages (i.e. talus sheets,

debris cones, rockfall deposits) appear as fundamental to

model the sediment delivery from an alpine catchment

(Harvey 2002; Schrott et al. 2003), more specifically within

a paraglacial context (Jones 2000; Ballantyne 2002a).

However, since the definition of the paraglacial concept

proposed by Ryder (1971), most of the quantifications of the

evolution of sediment yield within a paraglacial context have

been fitted at large time-scales, specially at Holocene time-

scale (Church & Ryder 1972; Jackson et al. 1982; Church &

Slaymaker 1989; Beaudouin & King 1994; Muller 1999;

Hinderer 2001). Recently, Ballantyne (2002b; 2003) synthe-

sized and discussed such results, pointing out that the

sediment delivery from a catchment subject to paraglacial

readjustment may be quantified from an ‘exhaustion model’,

in which sediment yield is related to the amount of

remaining ‘available’ sediments by a negative exponential

function (Cruden & Hu 1993). The curves associating the

sediment yield with the time elapsed since the deglaciation

are quite smoothed, and may not reveal the complexity and

the efficiency of processes acting at the beginning of the

paraglacial stage (Ballantyne & Benn 1996), due to the

resolution of Holocene time-scale studies. Various authors

have assumed the occurrence of a peak of sediment yield a

few decades or centuries after the beginning of the deglacia-

tion due to a specific combination of processes (Jackson

et al. 1982; Eyles et al. 1988; Evans & Clague 1994;

Ballantyne & Benn 1996), but such assumptions are still

to be demonstrated. A shift from Holocene timescales to

shorter, decennial or centennial, timescales is used here to

provide the arguments for this demonstration.

During the first centuries or decades of the deglaciation

the possible perturbations to the sediment yield curve may

be caused by glacier decay and its intermingling, somewhat

contradictory, effects. At such short time-scales, meltwater

supply from glaciers and their consequences for the fluvial

systems have been well documented (Johnson 1995; Jansson

et al. 2003). The rate of glacier thinning influences the

modalities of reworking, and hence the erosion rate of lateral

moraines (Eyles 1983; Meigs et al. 2006) and the deposition

pattern of sediments along its flanks (Hewitt 1989; de Graaf

1996; Turbek & Lowell 1999). However, the final conse-

quences of glacier lowering on sediments transfer and release

remain poorly documented. Our contribution aims at

measuring the impacts of glacier waning on the functioning

of the cascade sedimentary system by comparing the

disruptive signal (glacier variations) with the response-signal

(sediment transfers).

Norsk Geografisk Tidsskrift�Norwegian Journal of Geography Vol. 62, 115�131. Oslo. ISSN 0029-1951

DOI 10.1080/00291950802095145 # 2008 Taylor & Francis

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8 More specifically, we aimed to establish whether the rate

of sediment transfers is consistent with the rate of glacier

thinning or not. For this purpose, we selected two glacierized

catchments where glacial behaviour and shrinking patterns

had differed since the Little Ice Age. A fine temporal scale

approach was adopted to identify all stages of paraglacial

cascade-system evolution during the glacial retreat: the time

frame considered is the 20th century, a period which mainly

covers the post Little Ice Age (LIA) deglaciation.

In particular, we attempted to reconstruct the variations

of glacier tongues in three dimensions to document the

possible role of local base-level fluctuations on sediment

paths. We then surveyed the main storage units defined

within the paraglacial cascade sedimentary system (i.e.

deglaciated footslope and valley floor) to specify the

patterns of sediment transfers after glacier retreat.

Study area

Physical setting

The study area is situated within the eastern flank of the

Massif des Ecrins (Southern Alps of France), and corre-

sponds to the Vallouise valley (Fig. 1a). The current

glaciation of this catchment is spatially limited (c.24

km2), despite the high elevations (Barre des Ecrins peak

reaching 4102 m a.s.l.). This area is partly sheltered from

oceanic influences due to its eastward location: for

instance, precipitation is c.995 mm.yr�1 (mm per year) at

Pelvoux station (1280 m a.s.l.), which is significantly lower

than in the western flank of the Massif des Ecrins (1195

mm.yr�1 at Valjouffrey station, 1160 m a.s.l.). The

equilibrium line altitude ranges from 3000 m to 3200 m

(Cossart 2004; 2005), a value 200 m higher than in the

western part of Massif des Ecrins.

The main valleys became free of ice at the beginning of

the Holocene (Cossart 2005), and the Holocene period

was mainly characterized by cirque glaciation in this area.

During the Little Ice Age (LIA), only three small glacier

tongues (a few kilometres long) were occupying the upper

part of the subcatchments: Glacier Blanc, Glacier Noir,

Celse-Niere. The glaciated surfaces have decreased by

30%, compared to their extent in 1850 (Cossart et al.

2006) and many of the small glaciers tongues have more

or less disappeared (Fig. 1). Finally, this differentiated

glaciated pattern means the area is suitable for our

objective of measuring the consequences of the waning

Fig. 1. Location and physical settings of the Celse-Niere and Glacier Noir catchments in the upper section of the Vallouise valley, eastern flank of the Massif des

Ecrins.

116 E. Cossart & M. Fort NORSK GEOGRAFISK TIDSSKRIFT 62 (2008)

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8 and disappearance of the glacier tongues on sediment

transfers.

We selected the Glacier Noir and the Celse-Niere sub-

catchments, occupied by a debris-covered, low-sensitive

glacier and a ‘clean’, high-sensitive glacier respectively, based

on the assumption that they are characterized by different

rates of decay, and thus provide the opportunity to discuss

whether and how these characteristics exert a control on

sediment fluxes.

Glacier Noir area

The Glacier Noir catchment covers an area of 12 km2,

ranging in elevation from 1850 to 4102 m a.s.l. (Fig. 1b). It is

a classical glacially shaped alpine valley, characterized by

high (relief �1800 m) and steep (gradient �100%) moun-

tain slopes (Fig. 2a). Due to hillslope steepness, only a few

small glaciers (e.g. Momie glacier) are hanging above the

valley floor. As a consequence, an important surface with

ice-free faces constitutes a major source of sediments. Today,

the glacierized area is c.4 km2, but it was c.6 km2 during the

Little Ice Age. The glacier front was then situated at 1870 m

a.s.l. Both vertical decay and horizontal retreat of the glacier

tongue have been well documented and synthesized from

field data and archives (Allix 1924; Reynaud 1998).

Downward of the LIA glacier front, the catchment shows

a typical landform assemblage of a cascade system sub-

divided into three subsystems: free faces, slope deposits

(scree, avalanche deposits, rockfalls, etc.) and valley floor

(the glacio-fluvial plain Pre de Madame Carle) (Fig. 2a).

Upward of the glacier front, large lateral morainic ridges

disrupt the cascade system. Moraines are a source of

sediments but they intercept the slope deposits (Fig. 2b).

Celse-Niere area

The Celse-Niere catchment covers an area of 19 km2,

ranging in elevation from 1650 to 3954 m a.s.l. (Fig. 1b).

The topography is characterized by steep mountain slopes

(gradient �100%) with a relief of c.1000 to c.1500 m. The

glacierized area is c.6 km2, but it reached 9 km2 during the

Little Ice Age (Cossart et al. 2006). During that period, one

main glacier tongue (Sele glacier) occupied the valley, with

its front located at 2250 m a.s.l., whereas in the upper part of

the catchment some small tributaries (Ailefroide, Coup de

Sabre) were converging, with their front being blocked

against the Sele glacier valley tongue. Such small catchments

are today occupied by cirque glaciers that provided both

sediments and meltwater to the valley floor (Fig. 3a).

Within the catchment, the landform assemblages are

different upward and downward of the LIA glacier front.

Downward, the catchment shows a cascade system divided

into ice-free, rocky faces, slope deposits (scree, avalanche

deposits, rockfalls, etc.) and a narrow channel (Fig. 3b).

Upward the LIA glacier front, the cascade system is

disrupted by the presence of morainic ridges that locally

interrupt the continuity of slope deposits, while glacigenic

deposits are also intensely reworked.

Methodology

In both Glacier Noir and Celse-Niere cases, the presence of

morainic ridges creates temporary sediments storage cells,

yet what are their impacts on sediment transfers, and hence

on the functioning of the cascade sedimentary system?

To specify the potential impacts of moraines on sediment

path transfers in a recently deglaciated area, we intended to

compare the spatio-temporal combination of processes (i) in

the Celse Niere catchment, where many morainic ridges have

been deposited (high sensitive glacier) with (ii) the Glacier

Noir catchment, where only one or two generations of

moraines have been deposited after the LIA (low sensitive

glacier). We proceeded in three steps. We first surveyed the

three dimensional geometric variations of the glaciers in

relation to their position at the time of moraine deposition

and abandonment. Second, we reconstructed the evolution

of deglaciated footslope to assess the rate of sediment stored

or supplied down to the valley floor. We finally assessed the

impacts of sediment storage or supplies from footslopes on

the evolution of the fluvial system (valley floor) (Fig. 4).

Fig. 2. Glacier Noir catchment, looking downstream (A) and upstream (B) of the LIA glacier front. Note that the sediments stored at the footslopes are mainly

supplied from rocky, ice-free faces. The morainic ridges create a slope break within footslopes units. (Photos: M. Fort, June 2004)

NORSK GEOGRAFISK TIDSSKRIFT 62 (2008) Sediment release and storage in early deglaciated areas 117

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Three main units were thus surveyed: the glacier, the

deglaciated footslopes, and the valley floor.

Glacier variations

We carried out a geomorphic mapping to make an inventory

of glacial features. In order to reconstruct both glacial

horizontal extent and minimal altitude reached by the

glacier tongues, careful attention was given to former

morainic ridges. Their identification is based upon both

morphological (ridge-shaped, characterized by an internal

side steeper than the external one) and sedimentological

criteria (diamicton made up of a mixture of coarse and fine

material). In the present cases, the lithology also provides

useful pieces of evidence of glacial transport: granites

outcrops are predominant in the LIA glacier accumulation

area, while valley slopes consist of gneisses. Granitic

boulders identified in the lower sections of the valleys can

thus be considered as erratics.

The elevation of morainic ridges was carefully checked by

GPS within a GIS, in order to assess the minimal altitude

reached by the glacier ice (Fig. 4). The accuracy of the

estimation firstly depends on the vertical resolution of the

Fig. 4. Conceptual scheme of the methodology applied for assessing the amount of downwasting of a glacier tongue and its consequences for sediment fluxes.

Fig. 3. The Celse-Niere catchment, looking downstream (A) and upstream (B) of the LIA glacier front. Note that the debris supplied from supraglacial slopes are

issued from hanging glaciers (such as Ailefroide glaciers) or rocky ice-free faces. (Photos: E. Cossart, July 2002)

118 E. Cossart & M. Fort NORSK GEOGRAFISK TIDSSKRIFT 62 (2008)

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8 GPS measurements (c.10�20 m in this case). Secondly,

potential reworking of glacial remnants implies that the

altitude reached by glacier tongues may be underestimated,

especially in the case of the oldest moraines; hence, the

values obtained should therefore be considered as minimum

values.

Slope evolution

The deglaciated footslope evolution was reconstructed in

order to document how (process, rate of activity) the

sediments are supplied to the valley floor. A comprehensive

database was set up from geomorphic mapping, based upon

fieldwork (2002�2004) and photo-interpretation. We identi-

fied morphosedimentary units in the lower part of the slope,

such as rockfall deposits, avalanche or debris cones, alluvial

cones, moraines. The identification was based upon well-

known criteria, summarized in previous works (Francou

1988; Bertran et al. 2004): shape (concave or rectilinear

longitudinal profile), surface grain-size (increasing or de-

creasing boulder size downwards), orientation and inclina-

tion of the main axis of boulders, fabric of materials

(openwork or not, percentage of matrix) were particularly

surveyed. Vegetation cover was also estimated to document

the rate of geomorphic activity, and following Schrott et al.

(2003), we assumed that a vegetation cover higher than 20%

implies low geomorphic activity of sub-units.

Fieldwork was coupled with aerial photographs (covering

the study area since 1952), archives (topographical map of

1928) and dating methods (lichenometry, archives docu-

ment) to provide chronological benchmarks for reconstruct-

ing the stages of the building up of the footslope storage

units.

Particular attention was paid to the reconstruction of the

geometry of the morphosedimentary units assemblage,

based on basic measurements in the field (GPS, laser

telemeter, Leica level), with an error bar less than 10 m.

We first aimed to identify erosional and depositional land-

forms to examine where debris was eroded, transferred and

deposited in the footslope area. The identifications were

based upon the interpretation of ‘pre-erosional’ surfaces,

considered as a straight plane (Campbell & Church 2003)

(Fig. 5).

Fluvial system evolution

In the final stage, we studied the fluvial pattern because its

evolution may have recorded the rate of sediment release

from deglaciated areas. We particularly focused on active

and inherited valley-trains deposits (i.e. outwash plains

extending down a valley away from the ice front; Benn &

Evans 1998), to assess both the width of the active channel

and its braiding index. The active channel is considered as

the portion of the alluvial landscape composed of active,

unvegetated gravel bars and low-flow channels (Rundle

1985). The braiding index is assessed by the number of

channels per cross-section, according to Howard et al.’s

(1967) method. Both parameters may reflect the rate of

sediments transferred within the fluvial system. For instance,

an active channel widening associated with an increase of the

braiding index may be explained by an increase of solid

discharge in relation to the evolution of the glacial front or

to any exceptional input of debris from adjacent hillslopes.

Field observations, photo-interpretation (photos taken in

1952 and 1981) and digitalization of former topographical

maps were applied for systematic analysis of the effect of

sediment release from deglaciated areas on active channel

width. A total of 30 sites were selected at regular intervals

(approximately every 100 m) along the river courses. The

total length studied amounted to 4 km, which corresponds

to the upper part of Celse-Niere and Glacier-Noir proglacial

streams. Both temporal (since the beginning of the 20th

century) and spatial (from the glacier front to downward)

trends of variations of channel bed width were then

reconstructed.

Attempts at modelling

The final aim of this study is to interpret collectively the field

data we acquired in order to develop a dynamic approach to

changing geomorphic features since deglaciation. Following

Church & Ryder (1974), Ballantyne (2002a; 2002b; 2003)

developed a general steady-state model based upon the

exhaustion model, in which sediment yield decreases as an

exponential function through the time elapsed since the

beginning of deglaciation. The evolution of the volume of

paraglacial sediment storages can be estimated from equa-

tion 1:

Si�(1�e�lt): e�ki(t) (1)

Si is the volume of sediments within the storage units, l is

the rate of sediment supply from sediment source, ki is the

Fig. 5. Footslope landform assemblages (adapted from Campbell & Church

2003). Simple geometric units are reconstructed and interpreted from a ‘pre-

erosional’ or a ‘pre-depositional’ landform.

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8 rate of sediment release, and t is time. Here, it is assumed

that available sediments within the source equal 1 at t�0.

This model assumes that both l and k rates do not change

through time. More specifically, the exhaustion model

synthesizes a geomorphic evolution that can be subdivided

into specific sequences, each of them driven by a set of

geomorphic processes, the succession of which shifts

through time and space as soon as a new portion of land

becomes ice-free (Ballantyne 2002a; 2002b; 2003; Mercier

2002; Schrott et al. 2002; Jomelli et al. 2003; Cossart 2004;

Meigs et al. 2006). In the following, we aim to (i) recognizing

each distinctive sequence, (ii) establishing their chronological

succession, and (iii) describing the combination of active

processes distinctive of each sequence. Then we will discuss

how the succession of various geomorphic events can be

integrated to the equation 1.

Glacier Noir study case

Glacier retreat

The extent of the Glacier Noir during the LIA is easily

deduced from the large, latero-frontal, morainic ridges (c.100

m high). We observed a volume asymmetry between the

smaller right-flank moraine and the larger left-flank mor-

aine, caused by contrasted rock types and subsequent

uneven debris supply within the catchment (highly shattered

granites and gneisses on the left flank of the valley). The

surface of the Glacier Noir has reduced by c.35% since 1810.

This decay is associated with a horizontal retreat of the

glacier front by c.1.2 km. While the position of the glacier

front since the LIA has been recorded (data summarized in

Reynaud 1998) (Fig. 6), some chronological uncertainties

subsisted, especially for the period before 1930. We tried to

complete this data set by the identification and dating of

morainic ridges, and with information from archive docu-

ments.

The scenario of the horizontal retreat of the glacier

front can be reconstructed as follows. The front was

retreating during the second half of the 19th century;

more specifically, its horizontal recession was c.700 m

between 1850 and 1890 (Allix 1924) (Fig. 6A). This trend

was interrupted by a short period of stagnation of the

glacier front, as suggested by the deposition of a subdued

frontal moraine between 1890 and 1900 (former maps,

reproduced in Colas 2000). During the first half of the

20th century, the altitude of the glacier front remained

more or less constant. The deposition of a subdued

frontal moraine occurred between 1920 and 1930 (as

attested by former maps and old photographs of the

Fonds Dauphinois in Grenoble in 1939). In 1950, another

major period of retreat began, and is still ongoing. The

total retreat of the glacier front for the second half of the

20th century is estimated to be c.500 m in horizontal

dimension (Reynaud 1998), and c.50 m in altitude (data

based on field observation).

The rate of glacier downwastage has been surveyed in the

lower part of the tongue (Reynaud 1998), thus providing

evidence of two main stages (Fig. 6B). The first one occurred

before 1930 and is characterized by an increase of glacier

thickness of 17 m (�0.4 m.yr�1 (metres per year)). Since

1930, the tongue thinning has evolved at a constant rate

(�0.8 m.yr�1), and to date the total thickness lowering has

amounted to approximately 60 m in the lower part of the

glacial tongue.

This evolution pattern shows that the rate of glacier

front retreat was smooth and very progressive. The trend of

decay was not interrupted by major readvances but by a

few periods of stagnation of the glacier front. Only

subdued moraines have been deposited since the LIA.

This pattern is typical of a debris-covered glacier and

significantly differs from other patterns established within

the study area where major readvances of glaciers during

the 1920s, 1950s and 1970s have been identified (Cossart

et al. 2006).

Fig. 6. Geometric variations of the Glacier Noir. A. The horizontal retreat of

the front has been well documented by glaciologists and geographers since the

end of 19th century, and is summarized by Colas (2000). B. The vertical decay

has been surveyed by glaciologists and data are synthesized in Reynaud

(1998). The glacial downwasting results in the progressive lowering of the

valley floor elevation, and hence of local base level.

120 E. Cossart & M. Fort NORSK GEOGRAFISK TIDSSKRIFT 62 (2008)

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8 Slope evolution

Landform assemblages. As previously noted, the lateral

moraines are characterized by significant asymmetry in their

volume, creating two distinct assemblages of footslope

deposits respectively on the right flank and on the left flank

of the glacier. On the right flank, the morainic ridges are

smaller. They are buried by scree�avalanches deposits and by

torrential deposits. According to the landforms assemblages,

this footslope can be subdivided into two distinct parts:

upvalley (Bosses footslope) and downvalley (Momie foot-

slope) (Figs. 7 and 8).

Upvalley, scree and avalanche deposits are supplied from

the steep free faces (�458) and have built up a cone-shaped

landform that accumulated across the LIA right lateral

moraine of the Glacier Noir. Along the central radius of the

cone, the scree�avalanche deposits entirely cover the mor-

ainic ridge, and extend down to the valley floor; they

correspond to an active path for debris transfer, as suggested

by the lack of vegetation. On the external radiuses of the

cone, gravity-fed debris is intercepted by the morainic ridge

and accumulates behind it. This debris is now largely

Fig. 7. Momie and Bosses composite cone-shaped landforms. The left-flank

moraine of the Glacier Noir subdivides the cone assemblages, suggesting that

it intercepted the sediment fluxes issued from supraglacial slopes. (Photo: E.

Cossart, June 2002)

Fig. 8. Geomorphic map of the Glacier Noir foreland. A. In 1952, the right lateral Little Ice Age moraines were cut through by glacial meltwaters, and a

concomitant accumulation of meltwater deposits down to the valley floor took place (Momie cone) (sketch drawn from an aerial photograph of La Grave �Chorges, provided by Institut Geographique National). B. By 2002, the right-flank moraine ridge was buried by scree and avalanches deposits. Cones had

prograded down to the valley floor. The channel width of the main stream was narrower, but note the large active channel in the Glacier Noir proglacial area,

upstream of the 1950 moraine (drawn from field data).

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8 vegetated, thus suggesting its deposition is not currently

active.

Downvalley, the debris assemblage of the Momie cone-

shaped landform is a little more complex. Scree and

avalanche deposits are dammed upslope by the LIA

moraine. Their aggradation seems to be much less active

today (vegetation cover �20%); in fact, they are cut and

filled by the proglacial meltwaters issued from the Momie

hanging glacier. Also, the glacial meltwaters breached out

the LIA moraine. The reworked debris was partly redis-

tributed within a torrential cone which covers the inner side

of the moraine. According to the distribution of vegetation

cover, 40% of the area of this cone is now inactive.

On the left flank of the Glacier Noir, the LIA lateral

moraine is large and well preserved along a dominantly

concave, U-shaped, rocky slope. This massive and highly

compacted morainic ridge acts as an efficient barrier to large

rockfalls and scree taluses (several hundred metres high) that

are still supplied by the adjacent, steep rocky faces (vegeta-

tion cover less than 20%). In contrast, the inner (glacier

facing) slope of this morainic ridge is reworked by gullying,

the products of which accumulate below as small cones (c.10

m high) next to the glacio-fluvial stream and/or the debris-

covered glacier.

Interpretation. When glacial retreat occurred around the

second half of the 19th century, patterns of sediment

redistribution from the footslopes down to the valley bottom

were dependant upon both the characteristics of the

morainic ridges material (volume, cohesiveness) and the

presence (or not) of a continuous supply of glacial melt-

waters from the adjacent hillslopes (hanging glacier).

On the right flank of Glacier Noir, the small LIA moraine

created small sediment traps, forcing the aggradation of

scree and avalanche deposits:

. In the case of the Momie cone, the impoundment stage

was short because the waterfalls (collecting the melt-

waters of the Momie glacier) (Fig. 7) breached out the

morainic dam and gullied the scree�avalanche deposits

formerly built up below. This eroded debris started

accumulating within a torrential cone from the 1860s

(according to Jomelli et al. 2003), thus as soon as the base

level (i.e. the Glacier Noir surface) dropped. The

construction of the torrential cone became stabilized

during the second half of the 20th century, as evident

from the areal extent of the vegetation cover (from 8% in

1952 to 38% today, as assessed from photo-interpreta-

tion). Finally, large amounts of sediments were supplied

to the valley floor during the first half of the 20th century.

. In the case of the Bosses cone, the impoundment stage

lasted until 1950 (based on analysis of aerial photo-

graphs). Approximately between 1952 and 1960, the

sediment trap created by the lateral moraine was entirely

filled in, thus initiating a major period of progradation of

the debris over the moraine down to the valley floor.

Finally, the projected surface of the entire cone increased

by almost 100% after 1952, providing evidence of a

supply of debris directly to the valley floor during the

second half of the 20th century, many decades after the

Glacier Noir retreat.

Along the Glacier Noir left flank, the lateral moraines

formed a large and cohesive dam, thus generating upslope

and laterally an efficient sediment trap of a larger volume.

Nearly two centuries later, the trap is not yet full and the

damming effect by the morainic crest is still effective.

However, with the progressive lowering of the glacier

surface, the morainic ridge becomes more and more affected

by gullying on its internal face, thus ensuring the transfer of

some debris directly from the moraine flank to the valley

floor.

Fluvial system evolution

Along the valley floor, the evolution of the fluvial system can

be reconstructed by the distinction of three generations of

fluvio-glacial deposits.

A first generation of valley train (T1) was built up

downstream of the 1930 glacier front (Figs. 9 and 10),

encroaching over the entire valley floor. This aggradation,

currently vegetated by larch trees, was totally vegetation-free

before 1930, as attested by old postcards (1912) and maps

(drawn in 1928). The active channel bed was 160�190 m wide

Fig. 9. Evolution of the channel width (black bars) of the Glacier Noir proglacial stream since 1930. The narrowing of the channel bed occurred steadily between

1930 and 1981, in relation to the upward migration of sediment source subsequent to the glacial front retreat. The hydrosystem has stabilized since 1981, although

a channel widening upstream of the 1950 moraine was noted during the study.

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(Fig. 8); we noted it was wider upstream (190 m) than

downstream (170 m) of the 1895 moraine.

A second generation of glacio-fluvial (T2) deposits can be

individualized downstream of the 1950 front of the Glacier

Noir. In between the two 1930 and 1950 morainic ridges, five

different proglacial channels developed and cut through

the ablation-basal tills abandoned during that period by the

Glacier Noir during its decay. The cumulative width of these

former active channels was c.70 m. We noted these channels

to be converging downstream close to the 1930 moraine,

with their width reduced consequently (50 m). Downstream

of the 1930 moraine, the T2 deposits are cut-and-filled into

the T1 deposits, in relation to an incision of 2�3 m (Fig. 10).

In this section the active channel becomes narrower (60�70

m), yet it is significantly wider downstream of the Momie

cone (120 m wide).

The third generation of glacio-fluvial deposits is currently

active, extending downstream of the present glacier front.

Upstream of the 1950 moraine, active deposits are occupy-

ing the whole valley floor, reworking the recent till deposits.

Downwards, proglacial streams are flowing into two main

channels that are cutting through the T2 terrace and the

1930 frontal moraine (incision c.2 m). In this area, the

channel bed width is c.60 m. In the lower part of the study

section (downstream of the Momie cone), the current

channel bed width is c.100 m.

To summarize, two stages can be individualized in the

recent evolution of the Glacier Noir margins. In the first

stage, between 1930 and 1950, the glacier front receded, down

to an elevation of 2100 m a.s.l. in 1950. The difference in the

fluvial pattern observed between the upstream (many active

channels) and the downstream (single channel) reaches of the

1930 frontal moraine suggests that the 1930 moraine was at

that time acting as a local base level. The significant decrease

(c.100�120 m) of channel bed width since 1930 and the

incision of T1 deposits collectively reveal a reduction of

debris supply downstream of the 1930 frontal moraine. This

indirectly demonstrates that such moraine behaved as a trap

for debris supplied from the upper section of the meltwaters

streams. Downstream of the studied section, the channel was

particularly wide, probably because of the large amounts of

materials supplied from the Momie cone. In the second stage,

since 1950, this overall pattern shifted upwards; the glacier

front is currently at 2150 m a.s.l. The active channel bed is

very wide upstream of the 1950 frontal moraine, occupying

the whole valley floor, whereas downstream of the 1950

moraine, the T2 terrace and the 1930 moraine are incised,

and the active channels have become narrower (�10 to �15

m). This corroborates the existence of a relative depletion of

debris brought downstream from the glacier front, depletion

that can be explained by a damming effect exerted by the

1950 frontal moraine.

Finally, the fluvial pattern evolved and shifted upwards in

relation to the retreat of the glacier front since the beginning

of the 20th century. We noted a constant decrease of channel

bed width, in relation to a progressive incision of glacial

meltwaters into older valley-train terraces. This quite regular

trend can be explained by two factors: (1) the relatively rapid

degradation of subdued frontal moraines that were never

sharply shaped and (2) the limited amount of sediment

reworked from the adjacent hillslopes due to the large

quantities of sediments that are still stored against the

footslopes cells, especially along the longitudinal furrow

between the left lateral moraine crest and the rock walls.

Celse-Niere catchment

Glacial setting

The deglaciation pattern in Celse-Niere catchment differs

from the Glacier Noir catchment, as the ‘clean’, debris-free

surface of the Sele glacier is more sensitive to climate varia-

tions. Five generations of moraines were deposited in

approximately 1860, 1925, 1950, 1975, and 1993 (Fig. 11).

The dates generally correspond to pauses in an overall trend

dominated by glacial retreat (Reynaud 1998; Cossart et al.

2006).

In 1860, all glaciers (except for the Glacier du Coup de

Sabre) were converging into one single trunk glacier, the

front of which was located at 2200 m a.s.l. The relative

elevation of the lateral moraines suggests the Sele glacial

tongue was at least 130 metres thick. In 1925, the glaciers of

Ailefroide and Boeufs-Rouges were disconnected from the

main Sele valley glacier. Meanwhile, the thickness of the

latter glacier had decreased significantly (�40 m next to the

Coup de Sabre subcatchment outlet) while the glacier front

had shifted upvalley, to c.2400 m a.s.l. Between 1925 and

1950 a major stage of vertical decaying occurred, with a

thickness lowering estimated to be c.60�70 m. From 1950

onwards, both the retreat of the glacier front (shifting 500 m

upwards) and vertical decaying of the glacier tongue (30 m

lower) were effective, until the Sele glacier tongue definitely

disappeared and the outlet of the Coup de Sabre small

catchment became ice free.

In sum, the glacier surface decreased of approximately

�40% since 1860. The total disappearance of the glacier

from the valley resulted in a lowering of the local base level

of c.80�90 m.

Lateral margins

Footslope assemblages. In the upper part of the Celse-Niere

catchment, deglaciated footslopes are mainly composed of

Fig. 10. Vertical incision of the Glacier Noir proglacial stream since the late

19th century. Subdued frontal ridges were breached out just after their

deposition, generating an incision of the channel bed, and hence preventing

any persistent damming of the stream.

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scree taluses and cones. However, at the outlet of some

subcatchment tributaries (Coup de Sabre and Ailefroide

small catchments), particular, roughly cone-shaped assem-

blages are observed.

At the issue of the Coup de Sabre catchment, the

landform assemblage corresponds to three ice-contact fan

terraces and deposits (Fig. 12). The first one (C1) is perched

140 metres above the current valley floor. The second one

(C2) is cut-and-fill into C1, and is perched 50�60 metres

above the current valley floor. The third cone-terrace is cut-

and-fill within C2, and is encroaching over the valley floor.

This cone C3 is currently incised (1�2 m) by both the Coup

de Sabre and Sele proglacial streams.

Downslope the Ailefroide catchment, the landform as-

semblage is as follows (Fig. 13). Upstream the LIA lateral

moraine of Glacier du Sele, scree and avalanches deposits

are supplied from the ice-free faces and from the hanging

Ailefroide Glacier. The internal face of the moraine is

eroded by gullying. The eroded debris is carried away and

redistributed by streams, and is eventually deposited as

cone-shaped morphosedimentary units. Two main genera-

tions of such units are observed. The first one (Level 1�L1)

is inherited and lies at 2600�2650 m a.s.l., c.140 m above the

valley. This L1 unit is currently subject to erosion: reworked

debris is stored within small cones (Level 2�L2), built at the

contact with the valley floor or along the glacier.

Interpretation. The pattern of sediment redistribution was

developed during three main stages:

(1) During the first stage (1925�1950), the altitude reached

by the valley glacier surface had decreased of c.45�60

m. In the Coup de Sabre area, the lowering of the base-

level resulted in both breaching of the 1925 moraine

and the incision of the first generation of torrential

cone (C1). The incision is c.15 m. The eroded debris

was carried out and stored along the lateral moraine of

glacier du Sele, within a second generation of cone-

shaped morphosedimentary unit (C1). During this

period, the Ailefroide glaciers no longer converged

with the main glacial trunk (disconnection around

1920). The deposition of large lateral morainic ridges

fragmented the sediment fluxes issued from the Aile-

froide glacier foreland: the depression located upward

Fig. 11. Map of glacial geomorphic remnants of the upper section of Celse-Niere valley. Five generations of moraines can be observed, indicating an irregular

retreat of glaciers over time.

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of the moraines was then infilled with glacio-fluvial

deposits (Ailefroide torrent) and by avalanches�scree

deposits.

(2) During the 1950�1975 stage, the Sele glacial tongue

retreated sufficiently so that the glacio-fluvial debris

issued from the Coup de Sabre small catchment could

freely prograde over the main valley floor to form the

last generation of torrential cone (C3). As shown in the

aerial photographs, the 1950 moraine was already

breached in 1960 and the older deposits (C1 and C2)

deeply incised to a depth of 5 m. During the same

period, the rapid lowering of the Sele glacier surface,

while favouring gullying of the inner faces of its lateral

moraines, also facilitated their incision by the progla-

cial outlet issued from the Ailefroide subcatchment

(�50 m), at such a pace that these moraine ridges were

breached out around 1965 (95 yrs), according to aerial

photograph interpretation. Released sediments were

then stored along the glacier left flank and along the

1975 lateral moraine, hence forming a perched, kame-

like accumulation (Level 1�L1).

(3) From 1975 onwards, a period of stabilization occurred

at the outlet of Coup de Sabre small catchment. The

three generations of cone terraces became vegetated;

the youngest cone was now incised c.1�2 m by the Coup

de Sabre outlet, an evolution that suggests the geo-

morphic activity was currently low. At the confluence

with the Ailefroide subcatchment, the valley floor

became partly ice-free in 1981; meanwhile the 1975

lateral moraine was breached due to active gullying and

torrential stream incision. The sediments were then

partly redistributed in a complex patchwork of storage

units connected either with the valley floor or the

glacier surface (Level 2�L2).

Thus, the pattern of sediment release was characterized by

temporal discontinuities, linked with the successive stages of

glacier waning and melting out. Lateral margins adjusted

rapidly to the lowering of glacier surface: morainic materials

and sediments which had accumulated along the moraines

were reworked and redistributed according to the new local

base-level. At each stage of valley glacier recession, some

temporary sediments storage units, especially along the

valley glacier flanks, again became available for the sediment

cascade. Such units can supply huge quantities of sediments

as soon as the valley floor becomes ice-free: these conditions

Fig. 12. Geomorphic evolution of the proglacial area of the Coup de Sabre catchment since the early 20th century. A. Geomorphic mapping, showing the different

generations of sediment storages units developed along the right flank of the former glacier of the Sele tongue. B. Longitudinal profile, showing the significant rate

of incision driven by the downwasting of the Sele glacier. Sediment storage developed between the morainic dams and the valley walls. While glacial retreat began

in the late 19th century, sediments were progressively released after the breaching of the 1950 moraine.

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were fulfilled in the 1950s in the Coup de Sabre area, then in

1970s in the Ailefroide area.

Fluvial system evolution

To what extent was the fluvial system responding to the two

main periods of debris supply we identified? In the Celse-

Niere catchment, the valley floor is narrow, which limits the

storage potential of fluvial deposits, and hence field

evidence. Therefore, the evolution of the stream channel

planform and width was mainly deduced from archives

documents, especially from aerial photographs. Our obser-

vations were made in the upper Celse Niere area, 3 km

downstream of the 1880 morainic ridge (Fig. 11). Several

stages of evolution were distinguished (Fig. 14):

. Before 1950, the stream was flowing into a single

channel, constrained on both flanks by scree deposits:

its width was less than 20 m (Cossart 2004).

. A first episode of enlargement took place between 1950

and 1960 (Fig. 14). In 1952, the active channel was still

sinuous and narrow, with a width varying from less than

20 m to up to 30 m (upstream of the 1880 moraine). The

channel bed width increased in 1960, locally reaching 35

m next to the glacier front and upstream of the 1880 and

1920 frontal moraines; meanwhile, the channel planform

became braided (braiding index 3 to 4). These observa-

tions suggest a trend to aggradation, in relation to an

excessive sediment yield. This evolution may have been

influenced by the presence of the 1880 and 1920 moraines

which were acting as local base level for glacio-fluvial

waters. It may also be explained by additional debris

inputs from the tributaries, in particular from the Coup

de Sabre footslope.

. Between 1960 and 1981 the channel bed width became

more homogenous. The influences of the 1880 and 1920

frontal moraines on channel width were probably over,

due to the enlargement of their breach. This evolution

suggests that the downstream sediment transfer within

the fluvial system was more efficient at this time.

However, we also note that the channel bed width was

c.35 m and c.40 m in the upper part of the study section,

upstream of the 1950 and 1975 frontal moraines respec-

tively; the braiding pattern was about 4 on the braiding

index. These observations suggest an excessive sediment

yield. This can be mainly caused by two factors: firstly by

an increase of debris supply from the footslopes (for

instance, with the establishment of a connection between

the Ailefroide stream and the valley floor); secondly, the

formation of the 1975 frontal moraine may have forced

the temporary storage of fluvio-glacial deposits up-

stream. However, we cannot exclude the possibility of a

transient response to a major flood event, even if the

available meteorological records do not support direct

arguments for this.

. From 1981 to 2002, the formation of the 1993 moraine

ridge caused the enlargement and associated aggradation

locally observed in the upper part of the study section,

aggradation which was also probably exacerbated by

debris supply from the Ailefroide storage unit. Down-

stream, the channel width became narrower again, with

less contrast between the upstream and downstream

reaches of the 1975 frontal moraine.

. Additionally, for the entire period, the variation in the

channel width observed immediately upstream of the

1880 moraine is also related to specific meteorological

events which triggered many catastrophic debris flows on

the mountainslopes (1987 event) (Fanthou & Gambier

1991).

Fig. 13. Geomorphic evolution of the proglacial confluence area of the Ailefroide catchment. A. Three generations of Sele moraines interrupted the sediments

fluxes issued from the Ailefroide glacier forelands, forcing local aggradation of scree and avalanches deposits. B. With the decay of the Sele glacier tongue, the

morainic ridges are subject to gullying and are locally breached out. Sediments previously stored are released and accumulate along the glacier flanks (Level 1

noted L1). C. Meltwater streams are now connected to the valley floor, thus providing amounts of debris to the fluvial system (Level 2 noted L2). Note that some

sediments are still stored within the former storage sub-units.

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Thus, a complex evolution of the fluvial system can be

drawn. Firstly, the first two stages of active channel

widening are linked with the rate of sediment inputs from

the Coup de Sabre and Ailefroide footslopes. Secondly, the

frontal moraine ridges acted as barriers, locally controlling

the base levels, and forcing the temporary storage of

sediments into the valley floor many decades after moraine

deposition. Thirdly, the impact of exceptional meteorologi-

cal events cannot be ruled out of explanations relating to

channel width variations.

Generalization

The influence of morainic ridges on the cascadesedimentary system

The comparison between the Glacier Noir and Celse-Niere

study cases points out the role of lateral and frontal

moraines in causing intrinsic perturbations of the cascade

sedimentary systems. Along the lateral margins of the glacier

(Fig. 15), the lateral morainic ridges generate temporary

storage of sediments. Sediments carried out by meltwater

activity, gravity falls and avalanches, fill in the depression

created on the upstream reach of lateral moraine. The

redistribution of such stored sediments depends on two

parameters:

. The characteristics of the moraines are of prime im-

portance. The volume of the ridge influences the volume

of the trap, and thus the time necessary for the depression

to be filled up. The cohesiveness of the morainic dam also

influences the time necessary for the moraine to be

breached out, and thus the time at which the trapped

sediments would be released.

. The pattern of sediment release depends on the geo-

morphic processes which are acting on both sides of the

lateral moraines. On the one hand, if gullying encroach-

ing over the inner side of moraines acts in addition to

glacio-fluvial processes progressing from the outer side

of moraines, then the moraines would be incised rapidly

(B10 yrs after base-level lowering). The sediments

previously stored are then rapidly remobilized and

redistributed down to the valley floor. On the other

hand, if gravity-fed processes are very active on the outer

side of the lateral moraine, the evolution of the inner

reach of the moraine by gullying would rapidly cease and

even be prevented by the progressive accumulation of

scree and avalanches cones prograding over the morainic

ridge down to the valley floor.

On the valley floor, the sediment transfer is partly con-

strained by the frontal moraines. Just after the deposition of

the moraine, the sediment load exceeds the stream flow

capacity upstream of the frontal morainic ridge, as expressed

by the widening and high braiding pattern (Fig. 15). As the

frontal moraine becomes progressively eroded the local base-

level becomes lower, resulting in the incision of the channel

bed and the release of sediments stored upstream of the

moraine. This trend is well expressed by the progressive

calibration of the channel bed width along the river section

during the decade following the moraine deposition, hence

highlighting the restoration of a continuous sediment flux

from the deglaciated area to the downstream reaches of the

fluvial system.

Can the exhaustion model be applied to the Little IceAge time frame?

For both of the examples we studied, we have shown how the

presence of recessional moraines abandoned during glacial

retreat causes an intrinsic perturbation to the cascading

system. Fragmented temporary storage units can thus build

Fig. 14. Evolution of the channel width (black bars) of the Celse-Niere proglacial stream since 1952. The widening of the channel after 1952 suggests an increase

of debris supply from the Coup de Sabre catchment. However, between 1952 and 1960, the formation of new moraines (shown in grey) caused major

discontinuities along the fluvial system and forced sedimentation upstream of them. Since the 1980s, the channel width has tended to become more homogenous,

indicating that sediment transfer by meltwater streams is more efficient.

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up in the shadow of morainic dams, and sediment fluxes

generated by paraglacial adjustment may be locally slowed

or even interrupted. By contrast, huge quantities of sedi-

ments are released when moraines are breached out. Some

studies have suggested that sediment transfers cannot be

summarized by a simple function of time elapsed since

deglaciation because they not only depends on the rate of ice

melting but also on sediment supply from the slopes

(Warburton 1994; Cossart 2003; 2004) and on threshold

effects due to base-level fall (Eyles 1983; Jomelli et al. 2003;

Meigs et al. 2006). Such threshold effects may produce

secondary peaks in sedimentary fluxes that are not directly

expressed by the exhaustion model curve.

In detail, the influence of morainic dams results in a two-

stage evolution of storage units (Fig. 16). During the first

stage, or impoundment stage, the sediments carried out from

supraglacial slopes are stopped and stored along the

moraines. Equation (1), which utilizes the exhaustion model

to quantify the volume of stored sediments (Ballantyne

2003), can then be applied. In this equation, the rate l

Fig. 15. Sketch showing the geomorphic evolution of recently deglaciated glacial margins. The presence of moraines creates a damming effect, hence a

fragmentation of the cascade sedimentary system, as shown by local aggradation of sediments usptream of both lateral and frontal moraines. The duration of

such damming effects depends on (i) the number, volume and cohesion of moraines, and (ii) the erosion processes at work on the moraine (note the difference

between dammed meltwater cones and undammed scree cones).

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depends on the processes of sediment transfer from the

source; it may be considered as independent of time. k also

depends on the process of sediment release from storage unit

to valley floor, yet it is not independent of time in our case

studies. During the impoundment stage, as the sediments

transfer is low, k is close to 0.

As soon as the moraine is breached (stage 2), at time t]t?,amounts of materials are mobilized by erosion from morainic

sediments and supraglacial slopes. The sediment balance

within the store enters into deficit. As exemplified by storage

units such as Ailefroide, Coup de Sabre or Momie cones,

sediment release is mostly due to retrogressive erosion initially

caused by the incision along the main stream system, which

then propagates upward to the detriment of sediments stored

upstream of the moraine during the impoundment stage 1.

Considering the exhaustion model, the occurrence of a second

process for sediment redistribution implies that the storage

unit (of volume Si) is affected by a second rate of sediment

release (noted kj). The evolution of Sj, the volume of stored

sediments during the second stage, can be assessed from Si:

Sj�Si: e�kj(t�t?) (2)

kj is the rate of sediment release after the moraine breaching

out (at time t’), during the stage 2. Because ki and l are

supposed to be constant during stages 1 and 2, Si can be

quantified from equation 1:

Sj�[(1�e�lt): e�ki(t)]: e�kj(t�t?) (3)

The equation 3 is applied to plot curves showing the evolution

of sediment volume within storage units built up along

morainic dams (Fig. 17). These curves fit for different values

of ki and kj and show the threshold between stages 1 and 2,

which is marked by a major decrease of sediment volume

within the store, and hence by a huge transfer of materials

down to the valley floor. This result is more in accordance with

Fig. 16. Conceptual scheme of the hydrosystem evolution in frontal margins of glaciers. A. During the first stage the frontal moraine acts as a local barrier, forcing

the sediment aggradation upstream of it, hence causing channel widening. B. After the breaching out of moraine, sediments are progressively released

downstream, and a relative homogenization of the active channel width is observed on both sides of the moraine.

Fig. 17. Plotted curves in case of shifting rates of sediment release. The black

curve reflects the threshold effect due to the breaching out of the moraine: the

first stage corresponds to a forced aggradation upstream of the moraine;

during the second stage, sediments are rapidly released down to the valley

floor.

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the descriptions of massive sediment yield in early deglaciated

areas than curves plotted with a constant rate k.

Conclusions

Deglaciated alpine catchments are characterized by a peak of

sediment yield, which may be delayed from the beginning of

deglaciation. Such a lag may last for many decades, as we

have demonstrated, and can be explained by the complexity

of depositional landforms, in particular morainic ridges, that

affect the functioning of the cascade sedimentary system

(Fig. 18). Lateral moraines force the deposition of sediments

supplied by supraglacial slopes, whereas frontal moraines

force aggradation on the valley floor on their upstream reach:

in our examples, the effectiveness of lateral dams lasted from

0�15 year (debris carried out by meltwater streams), or even

more (at least 30 years) in the case of avalanches and scree

deposits; along the valley floor, the morainic ridges acted as a

dam for one decade. However, debris paths are even more

complex: debris is progressively transferred downstream

when the damming effects are over, usually many decades

after glacier retreat. Thus, the sedimentary cascading system

is not only linked to ice-melt rates, but also depends on the

rate of removal of sediments from storage units built up

against morainic ridges, therefore depends on external,

hydro-climatic conditions.

This complex pattern of paraglacial sediment release is

not clearly shown by the exhaustion model. However, one

promising aspect of the research is to break down the period

elapsed since glacial retreat in two stages within the equation

model. The first one corresponds to the impoundment stage

by morainic ridges, followed by the second stage of sediment

release once the morainic dams are no longer effective. By

applying different rates of sediment release from storage

units (rate k) for each stage, the massive sediment release due

to the moraine incision can be reflected by the plotted

curves.

Acknowledgements. � Financial support was provided by the Dynamique des

Milieux et Risques team (Paris 7) of UMR 8586 PRODIG (CNRS and Paris

1). We warmly acknowledge C.K. Ballantyne and A. A. Beylich for their

fruitful comments on a former draft of the manuscript. We are deeply

indebted to Catriona Turner who carefully edited the English manuscript.

Manuscript submitted 30 November 2007; accepted 5 December 2007

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