4 RÉSULTATS : CHRONOLOGIE ABSOLUE DES
FLUCTUATIONS DU LOBE LYONNAIS

Ce chapitre présente 3 articles publiés : “Geometry of glaciofluviatile deposits and dynamics of the Lyonnais lobe ice front during the last glacial period (France, Northern Alps)”, “Out-of-phase Late Pleistocene glacial maxima in the Western Alps reflect past changes in North Atlantic atmospheric circulation” et “Palaeogeographical reconstruction of the western French Alps foreland during the LGM using cosmogenic exposure dating”.

Dans ces articles, les méthodes de datation par exposition ¹⁰Be et par luminescence (OSL) ont été appliquées sur le piémont des Alpes occidentales françaises. Plus précisément, les datations par luminescence ont été réalisées sur les épandages fluvioglaciaires au front du complexe morainique du lobe lyonnais. Les datations par exposition ont été réalisées sur blocs erratiques et glaciaires localisés sur le piémont. Ces 3 études visent à établir un cadre chronologique absolu des fluctuations glaciaires du lobe lyonnais au cours du Pléistocène supérieur.
## 4.1 Geometry of glaciofluviatile deposits and dynamics of the Lyonnais lobe ice front during the last glacial period (France, Northern Alps).

Les épandages fluvioglaciaires au front du complexe morainique du lobe lyonnais constituent une opportunité de dater indirectement les extensions du lobe lyonnais. Avant toute datation, la géométrie, la dynamique et les corrélations entre les morphologies fluvioglaciaires et morainiques doivent être connues.

Les trois principaux couloirs fluvio-glaciaires ont été étudiés d'un point de vue géomorphologique à l'aide de MNT haute résolution afin d'analyser les relations entre les terrasses fluvioglaciaires et les différents ensembles de moraines. Des analyses sédimentologiques et des reconstitutions de la géométrie de ces épandages fluvioglaciaires à l'aide de forages ont permis de reconstituer leur dynamique au cours du dernier cycle glaciaire.

Ces travaux ont fait l'objet d'un article publié en anglais dans le Bulletin de la Société géologique de France en 2021.
# Geometry of glaciofluviatile deposits and dynamics of the Lyonnais lobe ice front during the last glacial period (France, Northern Alps).

ROATTINO Thibault: Université Savoie Mont-Blanc, Université Grenoble Alpes, CNRS, IRD, IFSTTAR, ISTERRE, Chambéry, France

Thibault.Roattino@univ-savoie.fr

CROUZET Christian: Université Savoie Mont-Blanc, Université Grenoble Alpes, CNRS, IRD, IFSTTAR, ISTERRE, Chambéry, France

BUONCRISTIANI Jean-François: Biogéosciences, UMR 6282 CNRS / Université Bourgogne Franche-Comté, Dijon, France

TISSOUX Hélène: BRGM, Unité Géologique de l'Aménagement des Territoires, Orléans, France

## 4.1.1 Abstract

Previous studies in the foreland of the French Western Alps, based on the analysis of geomorphological criteria for the internal moraine complex, show several stages of retreat or stagnation of the Lyonnais ice lobe during marine isotopic stages 4 and 2. Based on chronological data of the literature, the age of the Würmian maximum extension of the Lyon ice lobe must have occurred during MIS 4 but this result is still debated. At the Last Glacial Maximum, in the western part of the Lyonnais ice lobe, glaciofluviatile corridors were active during flash floods draining glacial meltwater. Today, these corridors are dead valleys and display a series of terraces. In this paper, we analyse the sedimentary geometries and dynamics of three glaciofluviatile corridors (Moidieu, Septeme and Heyrieux) located at the front of the internal moraine complex of the Lyonnais ice lobe. Upstream, the Moidieu corridor splits into three branches called North Moidieu, Central Moidieu and South Moidieu. Glaciofluviatile deposits in the corridors are composed of pebbles and gravels in a sandy matrix. Sedimentary structures show mass flow events and the migration of river bars in braided channels which is characteristic of proximal glaciofluviatile rivers in a proglacial environment. According to a new geomorphological map built using a high-resolution digital elevation model and an isopach map of the Quaternary deposits created from a compilation of the borehole data, we suggest that these corridors correspond to 'tunnel valleys' built during the most extensive Riss glaciation. Then during the Würm maximum glacial extension, these 'tunnel valleys' show complex infilling by glacio-fluviatile sediments during the Würmian maximal extent. In the three corridors, the number of river terraces can be better defined by using new geomorphological analyses. A total of three Würm terraces can be observed: two in the north and three in the south. This difference between the south and north is probably a result of climatic and tectonic forcing.

Key words: Lyonnais lobe, Würm, Geomorphology, Sedimentology, glaciofluvial, Moraines
4.1.2 **Introduction**

Quaternary glacial extent in the French Alps has been studied since the 19th century. These studies have primarily focused on the identification of erratic boulders and the mapping of moraine structures (Falsan and Chantre, 1879). Two moraine complexes have been identified in the marginal zone around the Alps and the Jura Mountains (Penck and Buckner, 1909). The external moraines complex is attributed to the Riss glaciation in the Northern French Alps, probably corresponding to marine isotopic stages (MIS) 8 to 6. A more internal moraines complex records the Würm glaciation corresponding to MIS 4 and 2 (Buoncristiani and Campy, 2011). Alpine glaciers flowed down to their foothills forming piedmont glaciers with several ice lobes during the Last Glacial Maximum (LGM). Very large ice lobes are located north of the Alps, which are from the west to east the Lyonnais, Rhône, Reuss, Rhine and Isar-Loisach ice lobes. These ice lobes are much smaller in the southern part of the Alps, e.g. the Ivrea, Ticino, Oglio, Garda and Tagliamento ice lobes (Ehlers and Gibbard, 2004) (Figure 4.1).

<figure><img src="image_36.png" /><figcaption>Figure 4.1: Map of the extension of the Alpine glaciers during the Last Glacial Maximum (Ehlers et al., 2011) showing the ice lobes on the main piedmonts.</figcaption></figure>

Several studies carried out in Switzerland, Austria and Italy have dated the age of the LGM as between 30 and 18 ka (Heiri *et al.*, 2014b; Wirsig *et al.*, 2016b). The age of the LGM is dated as between 25 and 22 ka in the Jura Massif (Buoncristiani and Campy, 2004). Radiocarbon ($^{14}$C) ages on wood fragments in the southern French Alps, more precisely in the Durance glacial deposit, show a maximal glacial extension close to 18 ka (Jorda *et al.*, 2000). Previous studies in the western Alps suggest two correlation hypotheses for the last maximal glacial extent with the Greenland isotopic record. First hypothesis based on the radiocarbon dating on plant remains and lacustrine paleomagnetic secular variations profiles suggests the beginning of the LGM between the isotopic stage 4 and 5. The second hypothesis rely on collagen of bones and tusks dating deem the beginning of the last maximal glacial extent between the isotopic stage 2 and 3 (Schoeneich, 1998). In this last hypothesis, the last maximal glacial extent in the western Alps would correspond to the global LGM.
Glaciofluviatile corridors exist at the front of the Lyonnais ice lobe, displaying several glaciofluviatile terraces built during the last maximal glacial extent and the different glacial retreats stages (Mandier, 1984). These palaeo-sandurs are currently dead valleys in which fresh outcrops can be accessed via several gravel pits. The goal of this study is to determine the geometry and to analyse the sedimentology of these three glaciofluviatile corridors which are geomorphologically associated with the moraines of the Lyonnais ice lobe. This research will provide new geomorphological and sedimentological data so as to better understand interactions between climate and uplift, in order to determine how these factors, control glaciofluviatile sedimentation and incision.

### 4.1.3 Geological setting

The Lyonnais ice lobe is located in the lower 'Dauphiné' region in the western part of the French Alps. The study area corresponds to the southern part of the Lyonnais ice lobe, and it integrates the Moidieu, Septeme and Heyrieux palaeo-fluviatile corridors (Figure 4.2). The pre-Quaternary substratum in this area corresponds mainly to Miocene sediment deposited within the Rhône Graben and in the western part to the Paleozoic basement (Sissingh, 2001). Two Miocene lithologies are described. The most abundant faciès contains silt and micaceous sand which are locally cemented in sandstone. The second faciès is a very discontinuous continental conglomerate with eastward increasing occurrence and thickness. The pre-Quaternary basement, close to the Rhône Valley, corresponds to metamorphic and plutonic Palaeozoic rocks. They are only exposed in areas downstream from the corridors (Chenevoy et al., 1969; Chenevoy et al., 1971; Elmi et al., 1986; Sissingh, 2001).

According to the petrography of the erratic blocks and the various paleogeographic reconstructions, the ice that feeds the Lyon ice lobe would have flowed from the Mont-Blanc and Beaufortain massifs (Mandier, 2003; Buoncristiani and Campy, 2011), but also probably from the Tarentaise and part of the Maurienne accumulation zone (Mandier, 2003; Coutterand et al., 2009; Coutterand, 2010). Quaternary deposits correspond to the LGM internal moraine complex and glaciofluviatile outwash deposits (Delafond, 1889; Gigout, 1960; David, 1967). The glaciofluviatile bodies are located in three corridors, Moidieu, Septeme and Heyrieux (Figure 4.2), where fluviatile terraces are geomorphologically connected to frontal moraines. Using both fluviatile terraces and frontal moraines, five stages of retreat are identified (labelled A to E) (Mandier, 1984). Stage A, called 'old Würm', is older than 35 ka (beyond the ¹⁴C analytic limit) (Vilain et al., 1988) and represents the maximal extension of the Lyonnais ice lobe during the Würm period (Mandier, 2003). Stages B and C result from the limited recession of the glacier. During stage D, the glacier has retreated enough eastward, and glaciofluviatile corridors are no longer supplied by glacial meltwater. Stage E corresponds to a decay stage of the glacier (Mandier, 1984) that left deposits on the downslope of the Alpine and Jura massifs. This last stage probably occurred at around 18,250 cal. BP (Mandier et al., 2003).
<figure><img src="image_37.png" /></figure>

Figure 4.2: Map of the external (Riss) and internal (Würm) moraine complexes of the Lyonnais ice lobe including the three studied corridors: the Moidieux, Septeme and Heyrieux corridors. The squares with black borders show the location of the studied gravel pits. A: Artas gravel pit; EP: Eyzin-Pinet gravel pit; SPC: St Pierre-de-Chandieu gravel pit; O: Oytier gravel pit. The longitudinal profiles of the terraces shown in Figure 4.6 are indicated by red dotted lines. MSP: Moidieu south profile; MCP: Moidieu central profile; MNP: Moidieu north profile; SP: Septeme profile; HP: Heyrieux profile

### 4.1.4 Methods

#### 4.1.4.1 Geomorphological map

A new geomorphological map of the glaciofluviatile corridors is established using a high-resolution digital elevation model (DEM) RGE ALTI® 5 m published by the French National Institute of Geographic and Forest Information (IGN®). The DEM have a resolution of 5 m, and the Lambert 93 projection system and RGF93 local geodesic system are used.

Würmian moraines and fluviatile terraces are identified either based on their morphology on the DEM and new field observations or by using geological maps and previous studies (Mandier, 1984). Moreover, fluviatile terraces and erosion channels are identified based on slope analysis, and more precisely two slope ranges defined with Qgis. The first span, between 1.5 and 5°, is used to show erosional channels and erosional bank. The second slope interval (between 4 and 16°) highlights the margins of the terraces. Terraces are correlated from one side of the valley to the other one using transverse topographic sections. Along a valley, longitudinal topographic sections are used to make correlations both between the terraces and with the moraines. Longitudinal profiles of the terraces are constructed for the Moidieu corridor (three profiles), Septeme corridor (one profile) and Heyrieux corridor (one profile). Longitudinal profiles are
calculated as a dimensionless curve in order to better compare the three corridors. The standardization is carried out from altitude *H* at the point of measurement normalized by the altitude from the outlet river to headwaters *Ho*. The length of a stream from the river head to a measured point *L* is normalized by the total river length *Lo* (Demoulin, 1998). The Langbein concavity index (*CI*) is calculated to quantify the differential concavity for each profile. It is defined as follows:

$$CI = \frac{2A}{H}$$

<figure><img src="image_38.png" /><figcaption>Figure 4.3: Variables A and H of the Langbein concavity index represented on a schematic river profile.</figcaption></figure>

where *A* is the difference of altitude measured between the profile at the middle of the corridor length and a straight line linking the two endpoints of the profile, and *H* is the altitude between the headwaters and the outflow (Figure 4.3) (Langbein, 1964).

#### 4.1.4.2 Field investigations

In the field area, the best outcrops for sedimentological investigations are gravel pits. Therefore, we studied four quarries located at Artas, Eyzin-Pinet, Oytier and St Pierre de Chandieu (Figure 4.2). Fresh outcrops in gravel pits provided the opportunity to identify the lithological, textural and structural characteristics of glaciofluviatile deposits with very good outcropping conditions. It was also possible to identify the underlying Miocene molassic sandstones and conglomerates at the bottom of the excavations in particular at St Pierre de Chandieu. By describing faciès in the field, it was possible to define the characteristics of the very upper part of the Miocene and of the Quaternary deposits. These characteristics were then used in this study when reinterpreting the borehole data.

A composite synthetic log (Figure 4.4) was built from two sections observed in two gravel pits. Basal parts (A, B and C) come from the St Pierre de Chandieu gravel pit and the upper parts (C and D) come from Eyzin-Pinet gravel pit. Contact between basal and upper part was also observed.
<figure><img src="image_39.png" /></figure>

<figure><img src="image_40.png" /></figure>

<figure><img src="image_41.png" /></figure>

<figure><img src="image_42.png" /></figure>

<figure><img src="image_43.png" /></figure>

<figure><img src="image_44.png" /></figure>

F1

F2

F3
Figure 4.4: Summary log representative of the Würmian terraces showing the three sedimentological faciès (F1, F2 and F3) found in four gravel pits. Photos A, B, C, D and E illustrate the lithologies. Notice that faciès F3 is a combination of massive gravel beds (D) and sandy lenses (C).

### 4.1.4.3 Borehole data

Stratigraphical and petrographical descriptions of the boreholes are available on Infoterre's website (http://infoterre.brgm.fr/viewer/MainTileForward.do) from the French geological survey (BRGM) database (Banque du Sous Sol). In the study area, 152 boreholes were used to construct geological sections. Borehole descriptions can be used to identify the Quaternary deposits. The depth of the Miocene molasse is generally determined from drilling data well as petrographical characteristics defined during field observations. However, in the database, some borehole descriptions are sometimes imprecise and it was possible to correct several misinterpretations through observations made on the quarry sections. Therefore, it is sometimes difficult to identify the various different quaternary faciès and to determine the depth of the top of the Miocene. Thus, boreholes that cannot provide obvious distinctions between quaternary faciès and the Miocene were not used. Finally, enough boreholes provided relevant data to construct geological sections in the Moidieu and Heyrieux corridors. There were not enough available boreholes in the Septeme corridor.

### 4.1.4.4 Isopach map

An isopach map of the Quaternary deposits in the studied corridors is obtained from the analysis and re-interpretation of the 110 boreholes in our database. To complete this map, 32 Miocene outcrop points are also added when they are identified from field observations and geological maps around the corridors. All of the outcrop points used to construct the isopach map are located in the area encompassing the glaciofluviatile corridors as well as on their hillside. Using the points displaying the depth of the Miocene, the interpolation surface, corresponding to the footwall of the Quaternary, is calculated by ordinary kriging using SAGA. When this surface is subtracted from the DEM, the isopach map of the Quaternary deposit is generated. An isopach map of the Septeme corridor was not created due to a lack of boreholes reaching the molassic bedrock in this area.

## 4.1.5 Results

### 4.1.5.1 Geomorphological map

Maximal length and width of the three corridors is 35 km long and 2.5 km wide for Moidieu corridor, 20 km long and 2 km wide for Septeme corridor and lastly 20 km long and 4.5 km wide for Heyrieux corridor.

In the Moidieu, Septeme and Heyrieux corridors, the geomorphological map shows the following four terraces (from high to low altitude): FgR, FgW1, FgW2 and FgW3, which are respectively geomorphologically connected to four different groups of moraines MR, MW1, MW2 and MW3. The other morphologies that can be observed are erosional channels and non-glacial deposits such as colluviums and post-glacial fluviatile deposits (Figure 4.5).

The FgR terraces are the highest of all the evidenced terraces in the studied area. They cannot be geomorphologically connected to any moraines of the internal morainic complex. They also cannot be linked with the external morainic complex as this latter is far to the west. These FgR terraces are probably built during retreat phase of the penultimate glaciation and therefore they lie above the erosional surface of the Russian glacier.

Terraces FgW1, FgW2 and FgW3 are correlated with moraines belonging to the internal moraine complex. Among them, the FgW1 terraces are the highest terraces and are associated with the most external MW1 moraines. The FgW2 terraces, found at mid-altitude, are connected to intermediate-position MW2 moraines. The FgW3 terraces are the lowest terraces correlated with the innermost MW3 moraines. The FgW1, FgW2 and FgW3 terraces gradually change from being well expressed and become unidentifiable toward the downstream part of the whole palaeosandur.
In the Moidieu corridor, the four terraces, FgR, FgW1, FgW2 and FgW3, have been identified. All of the moraines associated with the different branches of the Moidieu corridor are located at an altitude of approximately 500 m on the westward slope. In the downstream part, starting from the first confluence of the corridor, FgW1, FgW2 and FgW3 terraces are merged (Figure 4.5, Figure 4.6A, B and C).

The Septeme corridor show a small-sized FgW1 terrace upstream of the palaeosandur connected to a few poorly-preserved MW1 moraines located at an altitude of approximately 320 m. This terrace is surrounded by a FgW2 terrace connected to a set of MW2 moraine ridges at an altitude of approximately 250 m and is located on the edge of Verpillère trough (Figure 4.5, Figure 4.6D).

The Heyrieux corridor shows a broad FgW1 terrace connected to a NNE-SSW oriented MW1 moraine and a FgW2 terrace correlated with a NW-SW oriented MW2 moraine. FgW2 is delimited by a discontinuous and poorly preserved slope representing erosional gash. FgW1 and FgW2 merge roughly six kilometres downstream from the moraines. MW1 and MW2 are located at an altitude of approximately 280 m and directly border the Verpillère trough (Figure 4.5, Figure 4.6E).

The longitudinal profiles of the terraces connected to the moraines of the three corridors show distinct characteristics that are highlighted by their concavity index (CI). In the present work, this index was calculated from longitudinal profiles of FgW1 and FgW2 terraces that are present in all three corridors. The three profiles of the Moidieu corridor show a relatively straight shape (Figure 4.6A, B and C) which is confirmed by the low CI values (Figure 4.6F). The Septeme and Heyrieux profiles are more concave (Figure 4.6D and E) with relatively high CI values (Figure 4.6F). Downstream knickpoints (K) located on all profiles correspond to a lithological transition between the Miocene sands and the metamorphic and plutonic Palaeozoic basement rocks.
<figure><img src="image_45.png" /></figure>

Figure 4.5: Geomorphological map showing the fluvial terraces and erosional channels of the Moidieu, Septeme and Heyrieux corridors at the front of the internal moraine complex of the Lyonnais ice lobe and the main post-glacial fluvial and colluvium deposits found in the study area. 50 m isolines of rock uplift come from numerical modelling of Quaternary erosion-induced alpine isostasy from Champagnac et al. (2007). Geodesic system used: WGS84, Projection system used: Lambert 93, Graticule units: Degrees minutes seconds.
<figure><img src="image_46.png" /><figcaption>Figure 4.6: Dimensionless curves of the longitudinal profiles of the terraces and the CI values for the FgW1 and FgW2 terraces for the three studied glaciofluviatile corridors. The full line connects the points taken at the site of each terrace, and each terrace is linked to the others by dotted lines. The triangles, circles and cross represent the points of the FgW1, FgW2 and FgW3 terraces, respectively. K: Downstream knickpoint position. A: Longitudinal profile for the South Moidieu branch. B: Longitudinal profile of the Central Moidieu branch. C: Longitudinal profile of the North Moidieu branch. D: Longitudinal profile of the Septeme corridor. E: Longitudinal profile of the Heyrieux corridor. F: CI values for the FgW1 and FgW2 terraces in all corridors. The location of the profiles is given in Figure 4.2</figcaption></figure>
4.1.5.2 Isopach maps

The thickness of the Quaternary deposits in the Heyrieux corridor is greater in the upstream part (approximately 50 m) than in the downstream part (only a few meters; Figure 4.7). The same pattern is also displayed in the upstream part of the northern branch of Moidieu corridor but not in the southern branch. The Moidieu and Heyrieux corridors do not show comparable Quaternary depositional thicknesses.

Thickness displaying the highest absolute frequency for Moidieu corridor is approximately 25 m compared to 16 m for Heyrieux corridor. The longitudinal and transversal geological sections (Figure 4.8) show an irregular morphology of the top of the Miocene with slight overdeepening between 10 and 20 m in both corridors. The Quaternary erosion surface shows the existence of steeply sloping valleys, as shown by the south-eastern edge of section 2 (Figure 4.8), and the flanks of these valleys show slopes more than 20°. Finally, the width of each corridor is clearly smaller when the lithology of the pre-quaternary substratum changes, and thus corresponds to the Palaeozoic and Miocene boundary (Figure 4.8).

<figure><img src="image_47.png" /><figcaption>Figure 4.7: Isopach maps of the Moidieu and Heyrieux corridors. Geodesic system used: WGS84, Projection system used: Lambert 93, Graticule units: Degrees minutes seconds.</figcaption></figure>
<figure><img src="image_48.png" /><figcaption>Figure 4.8: Location of the geological cross-sections (black lines) and position of the Palaeozoic and Miocene boundary (red dashed line). B: Longitudinal and transversal geological sections of the Heyrieux corridor (geological sections 1 and 2) and</figcaption></figure>
### 4.1.5.3 Sedimentological analysis

The sedimentary successions are divided into three sedimentary faciès, each defined on the basis of their grain-size, sedimentary structures and geometries (Figure 4.4).

**F1: Micaceous sand with sigmoidal cross-strata**

F1 is primarily comprised of well-sorted medium sand with yellow-to-brown micaceous and mud clasts showing sigmoidal cross-stratification within one-meter thick sequences. Sandy sequences are intercalated by horizontal silty layers measuring several tens of centimetres thick, and by rare gravel levels. Mud clasts sometimes contain marine gastropod shells. This F1 faciès is locally consolidated into sandstone due to diagenetic processes.

**F2: Trough cross-bedded gravels to pebbles with small sandy lenses**

F2 is composed of a polygenic gravels to pebbles with a sandy matrix. The observed thickness varies between 3 and 35 m. The clasts are slightly spherical, sub-angular to well-rounded. This faciès is poorly sorted, the cross-stratification is clearly visible and there are abundant alternating clast-supported and matrix-supported beds. Trough cross-bedded sequences are fining upward, sometimes grading into thin sand with current ripples. There are abundant small sandy lenses measuring several metres in length and several decimetres in thickness including current ripples.

**F3: Massive gravels to cobbles including metric sandy lenses**

F3 is mainly composed of polygenic gravels, pebbles and cobbles with a sandy matrix. The observed thickness varies between 8 and 10 m. Rare striated pebbles are only observed at the Eyzin-Pinet gravel pit. The clasts are slightly spherical and angular to well-rounded. The faciès is poorly sorted and sequences starting with erosional surfaces show diffuse horizontal planar bedding or crude massive stratification, and clast-supported beds prevail. Some rare sequences exhibit a fining-upward trend in a large-scale trough shaped erosion structure with planar cross-bedding filling. Few sandy lenses in trough-shaped erosion based are described; they are approximately 10 to 5 m long and are several metres thick. These lenses show oblique tangential bedding to trough- bedding, with some current ripples sometimes showing climbing-ripple assemblages.

## 4.1.6 Interpretation and discussion

### 4.1.6.1 Sedimentary analysis

**F1: Micaceous sand with sigmoidal cross-strata**

The faciès is composed of sand dunes with sigmoidal cross-stratification interpreted as the result of migrating tidal dunes (Reynaud and Dalrymple, 2012). Silty horizontal beds result from ebb flow conditions and mud clasts result from the erosion of clay layers during tidal flood cycles (Allen and Homewood, 1984). Gravel beds represent drainage channels infilling. These deposits are interpreted as a signature of tidal environment. They represent the Serravalian faciès (Miocene Molasse) found in the Heyrieux area (Lesueur *et al.*, 1990).

**F2: Trough cross-bedded gravels to pebbles with small sandy lenses**

Clast-supported and matrix-supported alternations in the gravel to pebble bar deposit indicate migration bars during high flow periods (Rust, 1972; Carling, 1990). These faciès correspond to cut-and-fill deposits
interpreted as a filling of erosion troughs (Maizels, 1993). Sandy lenses occur as a result of a lower flow regime and records bar depositions in chutes minor channels (Siegenthaler and Huggenberger, 1993). Current ripples in sandy lenses result from unidirectional low current velocity. The fining-upward sequences indicate deposition through the waning stages of high flow (Miall, 1983). This faciès is typical of the deposits of superposed bars from unidirectional current in a proglacial braided channel river (Boothroyd and Ashley, 1975).

### F3: Massive gravels to cobbles including metric sandy lenses

Horizontal bedding result from the migration of gravel deposits with a high velocity current during flood periods (Brodzikowski and Van Loon, 1991). Flood events usually generate erosion of the top of the underlying sequences. The presence of clast-supported beds corresponds to the progressive deposition of the bedload during flood phases (Rust, 1978). Massive and poorly-sorted gravels with sequences including cobbles result from hyperconcentrated flows. The gravel to cobble deposit with trough structures and oblique planar bedding represent the migration of 3D and 2D dunes linked to unidirectional current in lower flow regime (Maizels, 1993). The fining upward of some sequences are deposit through the waning stages of high flow (Miall, 1983). Sandy lenses are usually interpreted as channel abandonment infills under a low velocity current. Sand dunes and ripples result from low current velocity deposition (Miall, 1978). Therefore, F3 shows hyperconcentrated flow deposits and some rare striated pebbles, characterizing a proximal braided fluviatile proglacial environment (Boothroyd and Ashley, 1975).

For F2 and F3, the polygenic nature and granulometry of the clasts, and the presence of few rare striated pebbles and sedimentary structures are characteristic of a Scott type sandur (Miall, 1978). The Scott type succession is named after the Scott outwash fan in Alaska studied by Boothroyd and Ashley (1975) and is typical of braided fluviatile dominated proximal proglacial river recording aggradation and migration of gravel bars and deposition of sand lenses in abandoned channel.

#### 4.1.6.2 Isopach map

The thickness of the glaciofluviatile deposits in the studied area is similar to the glaciofluviatile outwash of the Durance River beyond the Durance glacier, between 50 and 18 m thick (Gabert, 1984). The higher Quaternary thickness towards the upstream part corresponds to a proximal glaciofluviatile fan. The isopach map and geological sections highlight the absence of significant pre-LGM glacial erosion. Comparatively, many alpine valleys are overdeepened such as the 'Grésivaudan' Valley which has been eroded by the Würm glaciers up to 314 m below sea level (Nicoud et al., 2002). In the studied corridors, the morphology of the valley shows an abrupt beginning and termination and an irregular Quaternary erosion surface on top of the Miocene visible on the isopach map (Figure 4.7) and the geological sections (Figure 4.8). This typical morphology of the valley, may correspond to the tunnel valleys described under continental ice sheets (Van Der Vegt et al., 2012 and references therein) or under alpine glaciers (Moscariello et al., 1998; Dürst Stucki et al., 2010). Considering the dimensions of the three corridors, the tunnel valleys interpretation is favoured, instead of glacials overdeepenings. Tunnel valley are excavated in relatively soft sediments such as the Miocene strata and overdeepening take place rather in consolidated rocks (Magrani et al., 2020). Tunnel valleys are oriented parallel to the ice flow and have flanks with high slope values more than 20° (Cofaigh, 1996; Dürst Stucki et al., 2010; Van Der Vegt et al., 2012). This is in agreement with the high values measured for the slopes of the Miocene outcrops at the edges of the glaciofluviatile corridors. The formation of porewater pressure-driven tunnel valleys can occur in the Miocene Molasse basement. The alternation of sand and sandstone layers with different permeabilities enables the processes involved in the formation of the tunnel valley to take place. According to the porewater pressure-driven model, ball-structures and dykes constitute the main deformations within preglacial sediment (Ravier et al., 2015). Outcrops affording observation of such deformation are uncommon, but some dykes were observed in St Pierre de Chandieu gravel pit intersecting the Miocene basement and caped by glaciofluviatile sediment (Figure 4.4).
valley formation is inhibited on crystalline basements (Clark and Walder, 1994; Boulton et al., 2009). This can explain the narrower width of the valleys and the disappearance of the corridors downstream, near the contact between the crystalline and the sedimentary pre-quaternary basement. These tunnel valleys were probably formed during the most extensive glacier advance in the area (Riss glaciation).

<figure><img src="image_49.png" /><figcaption>Figure 4.9: Orthogonal schematic section of one studied corridor summarising the chronological relation between the Miocene basement, the Russian glacier erosion surface, the Russian terraces and the Würmian proglacial deposits.</figcaption></figure>

The FgR terraces are pre-LGM preserved deposits assumed to be contemporaneous from an older glaciation, probably the Riss glaciation as suggested by several authors (Chenevoy et al., 1969; Chenevoy et al., 1971; Mandier, 1984; Elmi et al., 1986). In the studied area, FgR terraces constitute probably remains of kame terrace deposits built by meltwater floods during the first stage of the retreat of the Russian glaciation. Thereafter during the Würm glaciation extent, tunnel valleys form the main proglacial corridor and are filled above the Riss glacial erosion surface by a proglacial sediment wedge showing now the different FgW1, FgW2 and FgW3 terraces (Figure 4.9).

### 4.1.6.3 Terrace formation

The terraces that are connected to the internal moraine complex were therefore formed during the Würm glaciation. In the studied corridors, two possible hypotheses can explain the observations and the formation of those terraces hereafter developed: a climatic hypothesis and a tectonic hypothesis.

**Climatic hypothesis:**

In present-day proglacial systems, such as those in Iceland, the formation of glaciofluviatile terraces occurs during glacial retreat (Marren, 2008; Marren and Toomath, 2013). Successive glacial retreats recede headwater position and river equilibrium profile. This shift results in the erosion of the upstream part of the glaciofluviatile outwash (Thompson and Jones, 1986). This model applied to the Lyonnais ice lobe retreat stage explains the formation of terraces in the upstream part of the three corridors. The difference in the number of terraces between the three corridors is explained by the frontal position of the paleo-glaciers and by topography. The moraines in the Moidieu corridor are located on a west-facing slope, thereby enabling the supply of meltwater during each glacial retreat phase, and the formation of the FgW1, FgW2 and FgW3 terraces. Moreover, narrow valleys in the upstream part of the Moidieu corridors allowed for the deep erosion of palaeo-river beds and good terrace expression.

The location of the MW1 and MW2 moraines in the Septeme and Heyrieux corridors involves meltwater supply only during two glacial retreat phases forming the FgW1 and FgW2 terraces. Later, the glacial front is located far below within the Verpillère trough and meltwater no longer feeds the glaciofluviatile corridors.

In the Heyrieux corridor, the large width of the valley does not allow for substantial erosion phases, thereby explaining the good conservation of the FgW1 terrace and thus the poor expression of the FgW2 terrace. In addition, the close proximity between MW1 and MW2 limits the shifting of the river equilibrium profile between the glacial retreat stages and inhibits the formation of terraces. The presence of two glacial retreats stages can only be justified by the different orientation of the MW1 and MW2 moraines.

Significant anthropization in the study area interferes with geomorphology of the three corridors. In addition, the position of the different moraines does not allow for good terrace formation, and it becomes difficult to identify the glacial retreat stages, as observed in the Heyrieux corridor.
### Tectonic hypothesis:

The proglacial palaeo-rivers of the three corridors evolved under similar conditions in terms of basement lithology, climate and probably sediment availability. The different CI values between the three corridors underline the different shapes of the longitudinal profiles between the Moidieu, Septeme and Heyrieux corridors. The low CI values for the Moidieu corridor could be evidence of a disequilibrium profile linked to a vertical shift of the river equilibrium profile (Strahler, 1952; Whipple, 2004). In the western French Alps, a differential uplift generating a tilting is related to glacioeustatic and erosional effects (Champagnac et al., 2007; Norton and Hampel, 2010; Nocquet et al., 2016; Sternai et al., 2019). Alternative geodynamic processes such as a mantellic upwelling (Lyon-Caen and Molnar, 1989), slab break-off (Kuhlemann, 2007) or mineralogical transformations in the lower crust (Kuhlemann et al., 2002) may also be involved in this possible tilting as deduced from the formation of the terraces in the three corridors. The modelled, Quaternary erosion-induced, alpine isostatic uplift deduced from Champagnac et al. (2007) may have generated differential response between the three corridors as illustrated by the isolines of rock uplift drawn on Figure 4.5. In addition, during the LGM period, an important glacioeustatic rebound could have occurred over a short period of time (Norton and Hampel, 2010). Tilting mainly raise the upstream part of the Moidieu corridor. In this corridor, palaeo-rivers dig their beds deeper and therefore the terraces are better expressed. This cut does not happen in the downstream part of the Moidieu corridor and in other corridors where there is not as much raised. Weak tilting affecting the Septeme and Heyrieux corridors explain the weak expression of their terraces (Figure 4.5). At time, further work is needed to estimate the proportion of both climatic and tectonic forcings for the terraces formation.

## 4.1.7 Conclusion

This study provides new data about glacial and proglacial features during the Middle to Late Pleistocene in the Lyonnais ice lobe at the front of the Western Alps. The thickness of the Quaternary deposits is approximately ten metres in the Heyrieux corridor and 25 m in the Moidieu corridor. The irregular shape of the longitudinal profiles of the valleys and their slight deepening in the Miocene is attributed to the genesis of the tunnel valleys prior to the last glacial maximum extent. In the three glaciofluviatile corridors, sedimentary processes were related to hyperconcentrated flow and cut-and-fill deposits, resulting in the superposition of bar-shaped deposits found in braided proglacial river systems (such as Scott-type channels), which are characteristic of a proximal proglacial environment. New geomorphological analyses can be used to better define the levels of the river terraces expressed mainly upstream of the palaeosandurs. The number of terraces differs from south to north. Three terraces are present in the Moidieu corridor, and two are found in the Septeme and Heyrieux corridors. The formation of terraces and their differences in number most likely results from both climatic and tectonic forcing.

## 4.1.8 Acknowledgements

This study was possible thanks to several grants provided by the Bureau de Recherches Géologiques et Minières (BRGM) program RGF Alpes, the Institut National des Sciences de l'Univers (INSU) and the Savoie Mont-Blanc University (USMB) project DIEGO. The companies Gachet SA (Artas), Cemex Granulats Rhône Méditerranée (Oytier and St Pierre de Chandieu), Nord Isère Materiaux (Eyzin-Pinet) are sincerely thanked for giving us access to active gravel pits.
## 4.2 Out-of-phase Late Pleistocene glacial maxima in the Western Alps reflect past changes in North Atlantic atmospheric circulation

Cette étude est dédiée à la datation par luminescence des systèmes fluvioglaciaires géomorphologiquement corrélés au front du complexe morainique du lobe lyonnais. Le but est de déterminer l’âge absolue de l’activité fluvioglaciaire et de l’extension maximum de ce lobe de piémont. Cette chronologie des extensions du lobe lyonnais doit également permettre de reconstituer la circulation atmosphérique dans l’Atlantique nord au Pléistocène supérieur.

Ces travaux ont fait l'objet d'un article publié en anglais en 2021, dans le journal Geology.
Out-of-phase Late Pleistocene glacial maxima in the Western Alps
reflect past changes in North Atlantic atmospheric circulation.

Natacha Gribenski¹, Pierre G. Valla²,¹, Frank Preusser³, Thibault Roattino², Christian Crouzet², Jean-François Buoncristiani⁴.

1: *Institute of Geological Sciences, University of Bern, Bern, Switzerland*

2: *ISTerre, Université Grenoble Alpes, Université Savoie Mont Blanc, CNRS, IRD, IFSTAR, Grenoble, France*

3: *Institute of Earth and Environmental Sciences, University of Freiburg, Freiburg, Germany*

4: *UMR 6282 Biogéosciences, CNRS/Univ Bourgogne Franche-Comté, Dijon, France*

**4.2.1 Abstract**

Paleoglacier reconstructions in the northern and southern forelands of the European Alps indicate a synchronous Late Pleistocene glacial maximum during Marine Isotope Stage (MIS) 2, in phase with global ice volume records. However, strong controversy remains in the western foreland, where scarce and indirect dating as well as modelling studies suggest out-of-phase glacial maxima with the rest of the Alps. New luminescence dating brings the first direct Late Pleistocene glacial chronology for the western Alpine foreland and reveals two major glacier advances of similar maximum extent, at ca. 75-60 and ca. 40-30 ka, coinciding with MIS 4 and late MIS 3. We propose that asynchrony in glacial maxima between the western and the northern/southern Alpine forelands results from a progressive spatial reorganization of the atmospheric circulation over the North Atlantic, in response to Northern Hemisphere ice-sheet fluctuations. While such feedback mechanism has emerged from general circulation models, our Late Pleistocene paleoglacial reconstruction permits to track the spatio-temporal evolution of moisture advection patterns over Western Europe.
## 4.2.2 Introduction

Ice-core and marine paleoclimate archives show two main periods of global temperature minima during the Late Pleistocene (LP; the last ca. 130 kyr; NGRIP, 2004), coinciding with Marine Isotope Stage (MIS) 4 and 2 (71-57 and 29-14 ka; Lisiecki and Raymo, 2005). Marine records indicate the main peak in global ice volume, also largely reported from terrestrial glacier archives, during the Last Glacial Maximum (LGM; ca. 27-19 ka; Clark et al., 2009). However, local glacial maxima well before the LGM are increasingly recorded in polar and mountain regions (Hughes et al., 2013; Batchelor et al., 2019). The emerging picture of variable ice expansion in both time and space likely reflects the regional effect of dynamic atmospheric circulations (Löfverström et al., 2014). Accurate glacial reconstructions across the globe are thus crucial to constrain past atmospheric circulation and paleo-precipitation patterns (e.g. Kuhlemann et al., 2008).

In the European Alps, major LP glaciations are recorded in prominent morainic lobes spreading over several tens of kilometers in the forelands (Ehlers et al., 2011; Figure 4.10). In the northern and southern forelands, synchronous LP glacial maxima at ca. 26-23 ka have been documented (Monegato et al., 2017). On the western side, the most extensive LP glaciation occurred in the Lyon area (Figure 4.10a), where an ice lobe draining the Arve (western Mont Blanc massif) and the Isère catchments developed ~50 km westward into the piedmont (Coutterand et al., 2009). In contrast to other Alpine forelands, scarce and indirect radiocarbon data in this area suggest older LP ice-maxima than the global LGM (Mandier et al., 2003). Such asynchronous paleoglacial activity is supported by numerical simulations suggesting transgressive glacier dynamics across the Alps, with the observed maximum ice extent in the Western Alps only reached when considering paleo-precipitation partitioning or older (i.e. MIS 4) glacier advance (Becker et al., 2016; Seguinot et al., 2018).

To further examine the hypothesis of asynchronous paleoglacial dynamics across the Alps, a robust chronology in the Western Alps is now critical. Radiocarbon and cosmogenic nuclide exposure dating are challenging in this area due to the scarcity of organic material incorporated within glacial deposits and in-place moraine boulders, respectively. Here, we present the first direct chronology of LP ice-maximum extent in the Lyon area, using optically stimulated luminescence (OSL) dating of glaciofluvial sediments. These results not only enlighten the debate around the paleoglacial history of the Western Alps, but also have major implications in terms of past atmospheric circulation over the North Atlantic and Western Europe.

## 4.2.3 Setting and methods

Two main moraine complexes, separated by ~25 km, define the Middle (≥MIS 6; External Moraine Complex, EMC) and Late (Internal Moraine Complex, IMC) Pleistocene ice-maximum extents in the Lyon area (Buoncristiani and Campy, 2011; Figure 4.10a). While the EMC presents discrete, highly-degraded and discontinuous moraine crests, the IMC is characterized by a series of well-defined frontal moraine ridges (Figure 4.10a).

Glaciofluvial valleys (i.e. inactive today) expand over several tens of kilometers westward from the IMC (Figure 4.10a), containing large sediment infills organized in step terraces. In the western and southern sectors, the IMC lays on a bedrock topographic step rising abruptly >200 m above the present-day lowland (Figure 4.10b). This topographic configuration allows to unambiguously relate glaciofluvial sediment deposition and terraces formation to major ice-extent periods (i.e. IMC ridges), when ice and meltwater could overtop this escarpment. Previous geomorphic and sedimentological studies identified at least three main terrace levels which, from high (T1/T2) to low (T4), directly connect to external (M2) to internal (M4) moraine ridges from the IMC (Mandier, 1988b; Roattino et al., 2021; Figure 4.10a), and thus were formed by aggradation during periods of glacial advances (Marren, 2005). Typical coarse braided-river faciès exposed along the terraces do not permit to argue for major discontinuity in terrace construction. Discrete outermost IMC ridges (M1) are visible only in the NW sector, suggesting an additional glacial stage hardly preserved through subsequent glacial periods (overridden moraine patterns; Fig. S8 in the Supplemental Material), and with possible remnants in the east and south sectors of the IMC indistinguishable from the
M2 deposits. In the absence of direct numerical dating, it remains challenging to quantitatively relate the observed glacial stages to distinct glaciations or short-term glacier oscillations.

We targeted the three main terrace levels (T1/T2 to T4; Figure 4.10a), with six samples collected from well-sorted sand lenses exposed in quarry sections (Table 4.1; see the Supplemental Material). Samples were taken as close as possible to the terrace surfaces to capture the last depositional period related to a glacial stage. Two additional samples were collected from a glaciofluvial unit underlying till from an M1 ridge in the NW sector (PIZ site; Figure 4.10a, Table 4.1), providing maximum age constraints of the LP outermost ice extent.

Luminescence signals and equivalent doses (D<sub>e</sub>) were measured from quartz (green OSL) and feldspar (infra-red stimulation at 50°C and 225°C, IR<sub>50</sub> and post-IRIR<sub>225</sub>) single grains (SG, 200-250 µm fraction). Methodological details are given in the Supplemental Material. Total burial dose since sample deposition was determined using the Central Age Model (CAM D<sub>e</sub>) or Finite Mixture statistical model (FMM D<sub>e</sub>) in case of partial bleaching diagnosis (Galbraith and Roberts, 2012). Feldspar burial doses were corrected for anomalous fading (20-30% increase) and final ages were calculated using sample-specific dose rate (2-3 Gy ka-1 range), derived from element concentrations measured by laboratory gamma spectrometry.
<figure><img src="image_50.png" /></figure>

Figure 4.10: (a) Glacial geomorphology of the study area (Mandier, 1988; Roattino et al., 2021). Inset: geomorphic limits of the LP ice maximum in the Alps (Ehlers et al., 2011) with main LGM ice flows (light blue arrows; Wirsig et al., 2016b), and location of the study area (black box) with main ice-drainage catchments (dark blue limit, AR: Arve, IS: Isère). (b) Longitudinal topographic profile (LP1) showing bedrock escarpment underlying the IMC ridges and downstream glaciofluvial sediments (undifferentiated), when the glacier reached its LP maximum extent. (c) Topographic cross-profile (ART site; CP1) showing the main three terrace levels. The bedrock/sediment contact is extrapolated from borehole data (Roattino et al., 2021).
4.2.4 **Results**

SG quartz OSL data could only be acquired for three samples (EYP1, OYT1, SPC5; Table 4.1), for which ~2% of measured grains emit a suitable luminescence signal. This is consistent with low luminescence sensitivity typically observed in quartz from crystalline bedrock and/or with limited transport (Sakawuchi et al., 2011). All eight samples reveal good feldspar luminescence characteristics (20-40% suitable grains).

SG quartz *D*<sub>e</sub> distributions exhibit normal-like to moderately-skewed shapes, while SG feldspar *D*<sub>e</sub> distributions are widely spread, asymmetric and multimodal, with a significant fraction of saturated grain for post-IRIR<sub>225</sub> (Figure 4.11). In addition, feldspar CAM *D*<sub>e</sub> are significantly larger (>50%) than for quartz. Both complex SG *D*<sub>e</sub> distributions and CAM *D*<sub>e</sub> differences between feldspar and quartz are strong indicators for limited pre-depositional exposure to light, typical for proximal glaciofluvial sediments (Duller, 2006). Partial bleaching, leading to age overestimate, was diagnosed for most of the samples and especially for feldspar signals. We thus applied statistical techniques (i.e. FMM model) to extract *D*<sub>e</sub> from the best-bleached grain sub-populations, giving more accurate age determination. For final age calculation, quartz OSL and feldspar IR<sub>50</sub> were selected, based on their best bleaching potential.

<figure><img src="image_51.png" /></figure>

Figure 4.11: Quartz normal-like and feldspar (no fading correction) complex single-grain burial age distributions for a representative sample (SPC5). Consistent quartz and feldspar sample burial ages are obtained using the CAM and FMM models respectively, based on the different partial bleaching diagnosis.

Quartz OSL and feldspar IR<sub>50</sub> ages obtained from the highest terrace in the central part of the study

area (T1/T2: SPC5; Figure 4.10a) and from the intermediate (T3: ART2, OYT1, EYP1) to lower (T4: ART1) terrace levels overlap between ca. 42 and 39 ka (Figure 4.12b; Table 4.1). Where both available, quartz OSL and feldspar IR<sub>50</sub> ages agree within uncertainties, bringing further

confidence in the statistical treatment applied to complex feldspar IR<sub>50</sub> *D*<sub>e</sub> distributions. The southern highest terrace level (PEN1) resulted in a distinctly older IR<sub>50</sub> age of 63.0±11.4 ka, coinciding with two IR<sub>50</sub> ages of 60.8±12.3 and 76.4±15.1 ka obtained for the glaciofluvial unit underlying the north-western M1 ridge (PIZ1 and PIZ2; Table 4.1).
<table><caption>Table 4.1: Sample locations and luminescence ages</caption><thead><tr><th>Sample ID</th><th>Latitude, longitude (°N/°E)</th><th>Elevation (m asl)</th><th>Depth (m below surface)</th><th>Geomorphic unit (Fig. 1)</th><th>Quartz (OSL) age (ka)</th><th>Feldspar (IR<sub>50</sub>) age (ka)</th></tr></thead><tbody><tr><td>EYP1</td><td>45.496, 5.009</td><td>262</td><td>5</td><td>T3</td><td>32.9 ± 2.3</td><td>29.4 ± 2.8</td></tr><tr><td>OYT1</td><td>45.575, 5.030</td><td>265</td><td>3.2</td><td>T3</td><td>36.4 ± 3.8</td><td>32.4 ± 4.3</td></tr><tr><td>SPC5</td><td>45.666, 5.015</td><td>250</td><td>5</td><td>T1-T2</td><td>38.6 ± 4.5</td><td>41.6 ± 7.9</td></tr><tr><td>ART1</td><td>45.537, 5.182</td><td>430</td><td>3</td><td>T4</td><td>—</td><td>33.8 ± 7.3</td></tr><tr><td>ART2</td><td>45.536, 5.182</td><td>420</td><td>12</td><td>T3 or T4</td><td>—</td><td>40.6 ± 4.2</td></tr><tr><td>PEN1</td><td>45.363, 5.186</td><td>313</td><td>12</td><td>T1-T2</td><td>—</td><td>63.0 ± 11.4</td></tr><tr><td>PIZ1</td><td>45.876, 5.088</td><td>250</td><td>13</td><td>Underlying M1</td><td>—</td><td>60.8 ± 12.3</td></tr><tr><td>PIZ2</td><td>45.876, 5.088</td><td>256</td><td>2</td><td>Underlying M1</td><td>—</td><td>76.4 ± 15.1</td></tr></tbody></table>

Note: asl—above sea level; OSL—optically stimulated luminescence; IR<sub>50</sub>—infrared stimulation at 50 °C.
Dashes indicate samples for which no Quartz OSL age could be obtained due to unsuitable luminescence signal.

### 4.2.5 Discussion

Luminescence dating indicates two periods of glaciofluvial aggradation beyond the IMC: between ca. 60-75 ka and ca. 30-40 ka, coinciding with MIS 4 and late MIS 3 respectively. This result strongly suggests that western Alpine glaciers reached the IMC at least twice during the LP, with similar maximum ice-extent configuration. The main terrace-moraine sequence (T2/T3/T4 - M2/M3/M4) observed in the area would hence result from glacier margin oscillations during a single glaciation period (i.e. late MIS 3), whereas scarce and outer remnant deposits (T1 - M1) were preserved from an earlier MIS 4 glacial maxima.

Our new MIS 4 chronology further corroborates earlier studies based on scarce chronological data in the Swiss Alpine foreland (Gaar et al., 2019 and references therein), and in the nearby Pyrenees (Delmas et al., 2011), as well as results from numerical simulations (Figure 4.12a; Seguinot et al., 2018), all of which proposed significant glaciations in the Western European mountain ranges at that time. The late MIS 3 glaciation clearly predates the last glacial maxima recorded in the northern and southern Alpine forelands during MIS 2 (Monegato et al., 2017). However, an extended MIS 3 ice advance has been suggested in the southwestern Alpine foreland (e.g. Ivy-Ochs et al., 2018), while valleys and forelands in the northern and central Alps were reported to be ice-free before 30 ka (Preusser et al., 2011; Barret et al., 2017). The absence of MIS 2 ages in our dataset, in particular for the lowest (T4) terrace level (ART 1, 33.8±7.3 ka), suggests that no glacier advance reached the IMC after late MIS 3. This is further supported by radiocarbon ages (Figure 4.12a; Mandier et al., 2003) of ca. 27 to >35 ka from a paleosol above till (i.e. minimum ages) in the NW sector of the IMC (Balan) and of 28-30 ka from glacio-lacustrine deposits 10-20 km upstream of the IMC (Moras and Malville), which suggest shorter glacier extent in the area during MIS 2. It is unlikely that the time lag (ca. 10 kyr) observed for the last major advance between the Lyon and northern/southern central Alpine forelands is caused by methodological uncertainties, as the applied dating approaches have shown age consistency on such timescales in similar glacial settings (e.g. Smedley et al., 2016; Gribenski et al., 2018).

In addition to multiple and similar LP glacial maxima in the western Alpine foreland predating the global LGM, our data indicates spatially and temporally variable paleo-glacier extent patterns across the European Alps, which likely reflect differential precipitation distribution combined with periods of global cooling. In contrast to the modern dominance of north-west atmospheric circulation, it has been proposed that south-dominated moisture advection prevailed across the Alps during the LGM (Florineth and Schlüchter, 2000; Luetscher et al., 2015; Monegato et al., 2017). The southward latitudinal shift of the North Atlantic storm track, as far as 40°N, was forced by the development of large ice masses and sea ice in the North Hemisphere (NH; Figure 4.12c), as indicated by atmospheric general circulation models (e.g. Hofer et al., 2012). The extensive MIS 4 glaciation in the western Alpine foreland coincides with a major cooling period observed
at the global and regional scales (NGRIP, 2004; Helmens, 2014; Moseley et al., 2020). Further elaborating on the above circulation model, we propose that the moderate expansion of the NH ice sheets during MIS 4, as indicated by deep-marine sediment records (Lisiecki and Raymo, 2005), could have similarly initiated a more discrete storm-track migration southward (Löfverström et al., 2014). Moisture was thus dominantly brought from the west and promoted high precipitation in the west-facing massifs of the Western Alps, enabling favourable conditions (i.e. wet and cold) for large piedmont glaciers. While cold MIS 4 conditions are expected across the entire Alps, limited precipitation in the Central and Eastern Alps may have resulted in more restricted ice extents in the main northern/southern central forelands, with glacial deposits likely eroded during the greater MIS 2 extent.

During MIS 3, global ice records indicate a decrease of temperatures already at 40 ka before reaching minima at 25-20 ka (NGRIP, 2004). Few existing long-term regional records suggest similar trends in SW Europe, even with a potentially earlier cooling onset at 45 ka (Moreno et al., 2014; Moseley et al., 2020). After significant shrinkage during early MIS 3, NH ice sheets started to expand again from ca. 40 ka (Batchelor et al., 2019). Similar to our proposed MIS 4 scenario, dominant western moisture advection would have been maintained throughout mid/late MIS 3 with a progressive southward shift of the polar storm track, again favoring glacier extent in the Western Alps. This configuration changed with the ongoing growth of NH ice sheets towards the LGM, which further pushed the polar storm track southward and lead shortly after 30 ka to the main southerly moisture advection over the European Alps (Luetscher et al., 2015). Interestingly, the lacustrine record at Les Echets (Figure 4.10a) also points towards a major change in lake productivity at ca. 28-30 ka, with a climatic transition from highly-oscillating to relatively stable, dry and cold conditions (Veres et al., 2009). We propose that this change in atmospheric circulation resulted in significant moisture decrease over the Western Alps, while enabling maximum glacier extent in the southern and northern Alpine forelands during MIS 2.
<figure><img src="image_52.png" /></figure>

Figure 4.12: (a) Late Pleistocene Northern Alpine speleothems δ18O (purple; Moseley et al., 2020) and global marine sediments δ18O (grey curve with MIS in shaded bars; Lisiecki and Raymo, 2005) records, with modelled Alpine ice volume (blue, GRIP paleoclimate forcing; Seguinot et al., 2018). (b) Luminescence ages showing two main groups around MIS 4 (green) and late MIS 3 (orange), consistent with existing radiocarbon ages from the NW sector (green) and upstream (red) of the IMC (Mandier et al., 2003). (c) Conceptual sketch illustrating the southward shift of the North Atlantic storm track forced by gradual NH ice sheet growth (LGM extent in blue), with impact on the paleo-atmospheric circulation (coloured arrows). Red dot indicates location of the study area.
4.2.6 **Conclusion**

Our luminescence chronology constitutes the first direct evidence of two major Late Pleistocene glaciations reaching similar maximum extents in the western Alpine foreland, during MIS 4 (76-60 ka) and late MIS 3 (42-29 ka). These multiple and early glacial maxima revealed in the western foreland contrast with MIS 2 maximum ice extent recorded in the northern and southern sides of the European Alps, in phase with the global LGM. Such spatial variability in glacier fluctuations has been observed elsewhere around the globe and likely translates swaying atmospheric circulation and precipitation partitioning during the last glacial cycle. We propose that such configuration may be related to the gradual migration of the polar storm track, forced by the progressive development of NH ice sheets. This study further highlights the importance of regional paleo-glacier reconstructions and numerical chronologies for quantifying past changes in atmospheric circulation patterns.

4.2.7 Acknowledgments

This work was supported by the SNSF (PP00P2_170559), INSU and USMB (AAP Montagne). We thank Gachet TP, Eiffage, Cemex Granulat Rhône Méditerranée, Nord Isère Matériaux and Vicat Granulat for providing access to gravel pits. We thank the four anonymous reviewers for their valuable suggestions.
# Supplemental Material

## Content:

### I. Geomorphic context and sedimentology

- General geomorphic classification and detailed information per sampling sites

- Table S1: sample location details

- Figures S1 to S7: Detailed geomorphic and sedimentological context per sampling site

- Figure S8: Geomorphic mapping of overridden moraine ridges

### II. Optically stimulated luminescence dating

- Sample preparation, measurement and data analysis

- Figure S9: dose recovery tests

- Figure S10, S11: fading measurements (multi and single grain scale)

- Figure S12: comparison of FMM burial equivalent doses for varying σb

- Figure S13, S14: single grain equivalent dose distributions (quartz and feldspar)

- Table S2: SAR measurement protocols

- Tables S3, S4: summary of luminescence dating (feldspar and quartz)

- Table S5: element concentrations for dose rate determination
# I. Geomorphic context and sedimentology

## Geomorphic terrace and moraine classification

Previous studies have been conducted in the Lyon area and distinguished several substages associated with the IMC, based on glaciofluvial and moraine ridges stratigraphic correlations.

Mandier (1988) identified four Late Pleistocene glacial stages, from A to D, within the IMC, based on geomorphology and morphostratigraphy relationships between moraine ridges and glaciofluvial terraces. Stage A corresponds to the maximum ice extent (outermost and highest moraine ridge connected with the highest terrace level), and stage D corresponds to the smallest ice extent (inner and lowermost moraine ridge connected with the lowest terrace level).

Roattino et al. (2021) identified three main glaciofluvial terraces, from FgW1(highest) to FgW3 (lowest), connected with moraine ridges from the IMC, based on remote sensing cross checked with borehole data and on field sedimentological and geomorphological investigations.

In the present study, we mainly followed the classification of Roattino et al. (2021), such as terrace levels FgW1, FgW2 and FgW3 correspond to the high, middle and low terrace levels (T1/T2, T3 and T4), connected with the external, middle and internal moraine ridges (M2, M3 and M4). We however do recognise the presence of a fourth moraine ridge (M1) clearly visible in the NW part of the study area, uncovered by Roattino et al. (2021). As a consequence, we propose two possible (and geomorphologically very similar) high terrace levels T1/T2 for the rest of the study area, connected to the most external moraine ridge (M2, possibly including some remnants of M1 deposits).

In total, eight samples were collected in six sampling sites distributed across the study area, and from glaciofluvial geomorphic units representative of the different substages identified in the area (Table S1).

## OSL sampling sites: detailed geomorphic and sedimentological context

For each sampling site, detailed information about the geomorphological and sedimentological context are provided in the Figures S2 to S7, as follow:

1. Geomorphological map (zoom locations as red boxes in the general geomorphological map below);

2. Representative longitudinal and transect profiles. These profiles do not directly intersect with the sampling site as they were first selected to best illustrate the general topographic/geomorphic context of the sampling area. However, the approximate location of the logged sedimentary section is indicated in the most relevant longitudinal/cross profile (note: the logged section symbol is not scaled and its absolute top elevation does not always fit within the elevation range of the topographic profile);

3. Detailed sedimentological log at the sampling site. Main sedimentary facies and sample location within the sedimentary section are indicated. The absolute elevation of the top of the log corresponds to the terrace surface at the sampling site.
<table><caption>Table S1: Sample locations and geomorphological information</caption><thead><tr><th>Sample ID</th><th>Sampling site locality</th><th>Latitude/ Longitude (°N/°E)</th><th>Elevation (m a.s.l.)</th><th>Depth (m below surface)</th><th>Geomorphic unit (Mandier, 1988)</th><th>Geomorphic unit Roattino et al. (2021)*</th><th>Geomorphic unit This study***</th></tr></thead><tbody><tr><td>EYP1</td><td>Eyzin Pinet (EYP)</td><td>45.496/5.009</td><td>262</td><td>5</td><td>Terrace level B</td><td>FgW2</td><td>T3</td></tr><tr><td>OYT1</td><td>Oytier (OYT)</td><td>45.575/5.030</td><td>265</td><td>3.2</td><td>Terrace level B/C</td><td>FgW2</td><td>T3</td></tr><tr><td>SPC5</td><td>St Pierre de Chandieux (SPC)</td><td>45.666/5.015</td><td>250</td><td>5</td><td>Terrace level B</td><td>FgW1</td><td>T1/T2</td></tr><tr><td>ART1</td><td>Artas (ART)</td><td>45.537/5.182</td><td>430</td><td>3</td><td>Terrace level D</td><td>FgW3</td><td>T4</td></tr><tr><td>ART2</td><td>Artas (ART)</td><td>45.536/5.182</td><td>420</td><td>12</td><td>Terrace level C</td><td>FgW2/3</td><td>T3 or T4</td></tr><tr><td>PEN1</td><td>Penol (PEN)</td><td>45.363/5.186</td><td>313</td><td>12</td><td>Terrace level A</td><td>FgW1</td><td>T1/T2</td></tr><tr><td>PIZ1</td><td>Pizay (PIZ)</td><td>45.876/5.088</td><td>250</td><td>13</td><td>Underlying till from the outermost moraine subset (stage A)</td><td>-</td><td>underlying M1</td></tr><tr><td>PIZ2</td><td>Pizay (PIZ)</td><td>45.876/5.088</td><td>256</td><td>2</td><td>Underlying till from the outermost moraine subset (stage A)</td><td>-</td><td>underlying M1</td></tr></tbody></table>

<footnote>*This study was conducted across the area covering EYP, OYT, SPC, ART and PEN sampling sites.</footnote>
Detailed geomorphic and sedimentological context per sampling site

<figure><img src="image_53.png" /></figure>

Main sedimentary facies identified in the sedimentary sections logged

<figure><img src="image_54.png" /></figure>

Sand-gravel, massive to coarsely stratified (horizontal bedding)

<figure><img src="image_55.png" /></figure>

Sand-gravel, coarsely stratified (trough/planar cross-bedding)

<figure><img src="image_56.png" /></figure>

Sand, trough/planar cross-bedded

<figure><img src="image_57.png" /></figure>

Sand, ripple cross-laminated

<figure><img src="image_58.png" /></figure>

Diamicton

Figure S1: Up: Geomorphological map of the study area. The red boxes correspond to the zoomed areas for each sampling site, and for which a detailed sedimentological log was also recorded in the field. Down: Legend of symbols used to record lithological and structural data (see sedimentological logs after; SE indicates observed erosion surfaces).
Eyzin-Pinet

<figure><img src="image_59.png" /><figcaption>Figure S2: Eyzin-Pinet (EYP) sampling site: detailed geomorphological map (up), longitudinal and cross profiles representative of the glacial/glacio-fluvial geomorphological context (down left), and sedimentological log (down right) with sample measured feldspar (IR<sub>50</sub>) and quartz (green OSL) single-grain ages.</figcaption></figure>
Oytier

<figure><img src="image_60.png" /></figure>

<figure><img src="image_61.png" /></figure>

<figure><img src="image_62.png" /></figure>

Figure S3: Oytier (OYT) sampling site: detailed geomorphological map (up), longitudinal profile representative of the glacial/glacio-fluvial geomorphological context (down left), and sedimentological log (down right) with sample measured feldspar (IR<sub>50</sub>) and quartz (green OSL) single-grain ages.
St Pierre de Chandieux

<figure><img src="image_63.png" /></figure>

Figure S4: Saint Pierre de Chandieux (SPC) sampling site: detailed geomorphological map (up), longitudinal profile representative of the glacial/glacio-fluvial geomorphological context (down left), and sedimentological log (down right) with sample measured feldspar (IR<sub>50</sub>) and quartz (green OSL) single-grain ages.
<figure><img src="image_64.png" /><figcaption>Figure S5: Artas (ART) sampling site: detailed geomorphological map (up), longitudinal and cross profiles representative of the glacial/glacio-fluvial geomorphological context (middle), and sedimentological logs (down) with sample measured feldspar (IR50) single-grain ages.</figcaption></figure>
Penol

Geomorphological map

<figure><img src="image_65.png" /></figure>

Topographic profiles

<figure><img src="image_66.png" /></figure>

<figure><img src="image_67.png" /></figure>

<figure><img src="image_68.png" /></figure>

Sedimentological log

<figure><img src="image_69.png" /></figure>

Figure S6: Penol (PEN) sampling site: detailed geomorphological map (up), longitudinal and cross profiles representative of the glacial/glacio-fluvial geomorphological context (down left), and sedimentological log (down right) with sample measured feldspar (IR<sub>50</sub>) single-grain age.
Pizay

Geomorphological map

<figure><img src="image_70.png" /></figure>

Topographic profile

<figure><img src="image_71.png" /></figure>

Sedimentological log

<figure><img src="image_72.png" /></figure>

Figure S7: Pizay (PIZ) sampling site: detailed geomorphological map (up), longitudinal profile representative of the glacial/glacio-fluvial geomorphological context (down left), and sedimentological log (down right) with sample measured feldspar (IR<sub>50</sub>) single-grain ages.
<figure><img src="image_73.png" /></figure>

Figure S8: DEM observations (5-m resolution, confirmed by field observations) of overridden moraines (right panel), in the NW of the study area (red box in the left panel), indicated by the presence of subtle lines (straight red lines on the right panel) cross-cutting perpendicularly sinuous positive topographic features (dark blue lines on the right panel) that are aligned with the general IMC ridges pattern (although more discrete). We interpret the latter as older moraine ridges (M1) which have been overridden by subsequent glacier advance.
## II. Optically stimulated luminescence dating

### Sample preparation, instrumentation and measurement protocols

Coarse-sand samples were collected in black tubes and prepared in the laboratory under red-light conditions following standard procedures (e.g. Lowick et al., 2015). The 200-250 µm fraction was sieved and treated with HCl and H₂O₂ to remove carbonates and organic components, respectively. Heavy-liquid density separation (sodium heteropolytungstate solution) was applied to isolate feldspar and quartz minerals, and the quartz fraction was etched using HF (40% for 40 min). Individual feldspar and quartz grains were placed in separated holes (100 holes of 300 µm diameter) drilled in a 9.8 mm diameter aluminium disk for single-grain luminescence measurements, or mounted on 4-mm aliquots for large-aliquot measurements.

All luminescence measurements were performed at the Institute of Geological Sciences of the University of Bern, Switzerland, on a Risø TL/OSL DA-20 reader equipped with an internal ⁹⁰Sr/⁹⁰Y beta-source, and with a single-grain attachment (Bøtter-Jensen et al., 2003). The OSL signal was stimulated with a green laser (single grain) or blue LEDs (large aliquot) and detected in the near-UV spectrum (Hoya U-340 filter). The infra-red stimulated luminescence (IRSL) signals were stimulated with at 140-mW with a 830-nm IR laser (90%) power (single-grain aliquot) or with IR LEDs (large aliquot) or and were detected in the blue wavelength (Schott BG-39 with a 410-nm interference filter and with an additional RG-780 filter for single-grain measurements).

Modified single-aliquot regenerative (SAR) protocols (Murray and Wintle, 2000) were applied to measure single-grain quartz (OSL; Table S2) and feldspar (post-IRIR₂₂₅; Buylaert et al., 2012; Table S2) equivalent doses (*D*ₑ). The OSL (quartz) signal was integrated over the first 0.1s minus a background averaged over the last 2s, and IR₅₀ and post-IRIR₂₂₅ (feldspar) signals over the first 0.5s minus a background averaged over the last 2s. Dose-response curves were reconstructed using an exponential fitting, including 2.5% instrumental uncertainty for single-grain *D*ₑ calculation. Single-grain *D*ₑ measurements were kept for further analysis with the following criteria: recycling ratio within 10%, recuperation ratio < 5%, and *D*ₑ < 2*D*₀. The suitability of each SAR protocol was checked by conducting dose-recovery and residual tests for each sample at the single-grain level for feldspar IRSL measurement, and for one representative sample using large aliquots for quartz OSL measurement (due to the very low proportion, i.e. ≤2%, of quartz grains, emitting suitable luminescence signal). For all samples, recovered dose after removal of the residual signal was within 10% of the given dose (within uncertainty), except for the post-IRIR₂₂₅ signal of one sample (PIZ2 within 19±5%; Fig. S9).

### Data analysis: partial bleaching evaluation and burial dose determination

Quartz OSL single-grain data could be collected for only three out of eight samples (EYP1, OYT1, SPC5), due to the very low luminescence sensitivity of quartz. Feldspar (IR₅₀ and post-IRIR₂₂₅) single-grain data were collected for all samples. However, analysis for burial dose determination were conducted using the IR₅₀ signal only, as the post-IRIR₂₂₅ natural signal was systematically saturated for a large grain proportion (>30-50%; Fig. S14).

Burial doses used for age determination were calculated from single grain *D*ₑ distributions using the Central Age Model (CAM; Galbraith et al., 1999), or the Finite Mixture Model (FMM; Galbraith and Green, 1990) in case of partial bleaching (i.e. incomplete reset of luminescence signal prior deposition) diagnosis. The FMM was run for multiple components (testing 2, 3, and 4 components), assuming a sigma-b (*σ*<sub>b</sub>) value of
0.25 (quartz) or 0.35 (feldspar-IR<sub>50</sub>) for well-bleached sediments in glaciofluvial environments (Gaar et al., 2014; 2019; Gribenski et al., 2018; Smedley et al., 2016; 2019). Final selection of the FMM component and associated burial dose (FMM D<sub>e</sub>) was based on statistical indicators (BIC and likelihood; Bailey and Arnold, 2006; Jacobs et al., 2008), the significance of the identified “youngest” D<sub>e</sub> subpopulation (p<sub>0</sub> > 0.05), and on visual inspection of the single-grain D<sub>e</sub> distributions. The robustness of the approach was investigated by testing the sensitivity of the FFM D<sub>e</sub> to varying σ<sub>b</sub> values (0.30-0.40; tests run on IR<sub>50</sub> distribution; Fig. S12) and by checking the consistency between burial doses obtained using the FMM and other sub-statistical models (e.g. Minimum Age Model - 4 parameters; Galbraith et al., 1999).

Samples collected in such proximal glaciofluvial environment (Fig. S1) are typically prone to partial bleaching issue, due to the characteristic sediment transport mode over short distances and in turbid waters, limiting the grain exposure to natural light (Duller, 2006; Fuchs and Owen, 2008). The significant increase in CAM D<sub>e</sub> values and/or in the saturated-grain proportion between the signals with different bleaching rates (i.e. decreasing resetting probability going from quartz OSL, IR<sub>50</sub>, to post-IRIR<sub>225</sub>; Buylaert et al., 2012; Murray et al., 2012; Colarossi et al., 2015) suggest limited light exposure of grains prior deposition. Final partial bleaching diagnostic was established for all the quartz and feldspar single-grain dataset but one (quartz OSL signal for EYP1 sample), based on the complex (asymmetrical and multimodal) and largely-spread (overdispersion OD>30% for quartz and OD>40% for feldspar) single-grain D<sub>e</sub> distributions.

Fading measurements (IR<sub>50</sub>) were conducted following two different protocols: at the single-grain level and using 4-mm aliquots. Both approaches provided similar outcomes when comparing single-grain (weighted average) and 4-mm aliquots (arithmetic mean) g-values within two standard errors, except for PIZ1 (Fig. S10). No relationship could be identified between single-grain D<sub>e</sub> and g-value (Fig. S11). Based on the lower inter-aliquot variability and the higher signal magnitude (better counting statistics), sample-individual g-values from 4-mm aliquot were used for correction of burial doses (following Huntley and Lamothe, 2001).

Sample-individual dose rates for age calculation were calculated using U, Th and K concentrations measured by high-resolution gamma spectrometry (Table S5) on bulk sediment using the online DRAC calculator (Durcan et al., 2015). A water content of 10±5% was assumed to take in account large variations over the considered time. For feldspar (IR<sub>50</sub>) dose rate calculation, an internal K-content of 12.5±0.5% (Huntley and Baril, 1997) and an alpha efficiency value of 0.15±0.05 (Balescu and Lamothe, 1994) were used.
<figure><img src="image_74.png" /></figure>

Figure S9: Dose recovery tests for feldspar IR<sub>50</sub> and post-IRIR<sub>225</sub> protocols (CAM recovery dose from single-grain measurements, 1 SG disc for each sample) with intrinsic overdispersion values (given in %).

<figure><img src="image_75.png" /></figure>

Figure S10: Individual sample g-values (two standard error) for IR<sub>50</sub> feldspar from single-grain measurements (1-2 SG discs for each sample; weighted average, in red) and from 4-mm aliquot measurements (3 discs for each sample; arithmetic mean, in purple).
<figure><img src="image_76.png" /><figcaption>Figure S11: Individual single-grain IR<sub>50</sub> g-values plotted against corresponding single-grain *D*<sub>e</sub> (black dots). Sample g-values from single-grain (weighted average; red line) and 4-mm aliquot (arithmetic mean; purple line) are also indicated. Some samples count significantly less single-grain g-values due to the non-measurable fading at an acceptable level of sensitivity and/or the absence of corresponding single-grain *D*<sub>e</sub>.</figcaption></figure>
<figure><img src="image_77.png" /><figcaption>Figure S12: IR<sub>50</sub> feldspar burial dose (one sigma error; non-fading corrected) obtained when applying the FMM model (FMM D<sub>e</sub>) using σ<sub>b</sub> values at 30, 35, 40%. In most cases, FMM D<sub>e</sub> agree within 1σ error using different σ<sub>b</sub> values. For ART2 and EYP1, FMM D<sub>e</sub> agree within 2σ error. Only for SPC5, a significant difference persists between 30-35% and 40% σ<sub>b</sub> values. However, IR<sub>50</sub> feldspar-age using 30-35% σ<sub>b</sub> value agrees within error with the associated OSL quartz age, hence we selected σ<sub>b</sub> of 35% for IR<sub>50</sub> feldspar age calculation. For PIZ1, the proportion of the minimum subpopulation extracted when applying the FMM model with 30% σ<sub>b</sub> is not significant (p<sub>0</sub> < 0.05), hence no FMM D<sub>e</sub> is provided.</figcaption></figure>
<figure><img src="image_78.png" /><figcaption>Figure S13: Single-grain OSL quartz D_c distribution for EYP1, OYT1 and SPC5 samples (suitable signals/statistics), and burial doses obtained using the CAM (top circle) and FMM (vertical line/envelope) models. No FMM model was run for EYP1 due to the limited overdispersion of this dataset (OD<25%).</figcaption></figure>
<figure><img src="image_79.png" /><figcaption>Figure S14: IR<sub>50</sub> and post-IRIR<sub>225</sub> single-grain *D*<sub>e</sub> distribution (non-fading corrected) for all samples. Probability distribution and burial doses (non-fading corrected) obtained using the CAM (top circle) and the FMM (vertical envelope) models are only provided for the IR<sub>50</sub> signal, as a large proportion of grains were saturated when considering the post-IRIR<sub>225</sub> signal.</figcaption></figure>
<table><caption>Table S2: SAR protocols used for feldspar and quartz single-grain equivalent doses</caption><thead><tr><th>Feldspar post-IRIR<sub>225</sub></th><th>Quartz OSL</th></tr></thead><tbody><tr><td>Dose</td><td>Dose</td></tr><tr><td>Preheat at 250°C for 60s</td><td>Preheat at 220°C for 10s</td></tr><tr><td>SG IRSL (laser) at 50°C for 5s</td><td>IRSL (diodes) at 125°C for 40s</td></tr><tr><td>SG IRSL (laser) at 225°C for 5s</td><td>SG Green OSL (laser) at 125°C for 5s</td></tr><tr><td>Test dose (500s)</td><td>Test dose (500s)</td></tr><tr><td>Preheat at 250°C for 60s</td><td>Preheat at 220°C for 10s</td></tr><tr><td>SG IRSL (laser) at 50°C for 5s</td><td>IRSL (diodes) at 125°C for 40s</td></tr><tr><td>SG IRSL (laser) at 225°C for 5s</td><td>SG Green OSL (laser) at 125°C for 5s</td></tr></tbody></table>

<table><caption>Table S3: Details of single-grain IR<sub>50</sub> feldspar measurements and age determination</caption><thead><tr><th>Sample ID</th><th>n grains (n/N, %<sup>1</sup>)</th><th>Total OD (%)</th><th>CAM D<sub>e</sub> (Gy)</th><th>skewness</th><th>kurtosis</th><th>FMM D<sub>e</sub><sup>2</sup> (Gy; σ<sub>b</sub> 0.35)</th><th>g-value<sup>3</sup> (% per decade)</th><th>Dose rate (Gy/ka)</th><th>Age (ka)</th></tr></thead><tbody><tr><td>EYP1</td><td>106 (35%)</td><td>78</td><td>96.0±7.3</td><td>1.44</td><td>5.33</td><td>49.9±3.7 (FMM 2)</td><td>2.4±0.1</td><td>2.14±0.12</td><td>29.4±2.8</td></tr><tr><td>OYT1</td><td>120 (40%)</td><td>73</td><td>164.8±11.1</td><td>0.65</td><td>2.52</td><td>57.5±6.8 (FMM 3)</td><td>2.5±0.2</td><td>2.27±0.12</td><td>32.4±4.3</td></tr><tr><td>SPC5</td><td>123 (41%)</td><td>68</td><td>244.8±15.1</td><td>0.75</td><td>2.72</td><td>79.7±14.0 (FMM 3)</td><td>2.8±0.3</td><td>2.53±0.13</td><td>41.6±7.9</td></tr><tr><td>ART1</td><td>105 (12%)</td><td>68</td><td>228.6±15.4</td><td>0.37</td><td>2.29</td><td>55.2±11.1 (FMM 3)</td><td>2.3±0.4</td><td>2.03±0.12</td><td>33.8±7.3</td></tr><tr><td>ART2</td><td>123 (41%)</td><td>75</td><td>182.8±12.6</td><td>1</td><td>4.37</td><td>70.2±5.3 (FMM 3)</td><td>2.5±0.4</td><td>2.20±0.12</td><td>40.6±4.2</td></tr><tr><td>PEN1</td><td>173 (29%)</td><td>43</td><td>328.2±11.0</td><td>0.9</td><td>6.42</td><td>107.9±18.1 (FMM 2)</td><td>2.0±0.1</td><td>2.09±0.12</td><td>63.0±11.4</td></tr><tr><td>PIZ1</td><td>133 (22%)</td><td>48</td><td>325.9±13.9</td><td>0.75</td><td>3.49</td><td>148.6±29.5 (FMM 2)</td><td>1.9±0.1</td><td>2.94±0.14</td><td>60.8±12.3</td></tr><tr><td>PIZ2</td><td>133 (22%)</td><td>52</td><td>326.4±15.3</td><td>1.45</td><td>6.83</td><td>181.2±34.4 (FMM 2)</td><td>2.5±0.2</td><td>3.05±0.13</td><td>76.4±15.1</td></tr></tbody></table>

<footnote n="1">¹Percentage of grains used for analysis (n) compared to the total number of measured grains (N). ²FMM D<sub>e</sub> used for age calculation. For each sample, the number of components of the selected FMM model is indicated; ³mean and</footnote>

<footnote>standard error g-value from fading measurements using 4-mm aliquots.</footnote>
<table><caption>Table S4: Details of single-grain OSL quartz measurements and age determination</caption><thead><tr><th>Sample ID</th><th>n grains (n/N,%)</th><th>Total OD (%)</th><th>CAM D<sub>e</sub> (Gy)</th><th>skewness</th><th>kurtosis</th><th>FMM D<sub>e</sub> (Gy; σ<sub>b</sub> 0.25)<sup>1</sup></th><th>Age model</th><th>Dose rate (Gy/ka)</th><th>Age (ka)</th></tr></thead><tbody><tr><td>EYP1</td><td>28 (2%)</td><td>24</td><td>39.4±2.3</td><td>0.11</td><td>2.48</td><td>-</td><td>CAM</td><td>1.20±0.05</td><td>32.9±2.3</td></tr><tr><td>OYT1</td><td>34 (2%)</td><td>36</td><td>60.2±4.1</td><td>0.67</td><td>2.45</td><td>48.7±4.6</td><td>FMM 2</td><td>1.34±0.06</td><td>36.4±3.8</td></tr><tr><td>SPC5</td><td>27 (2%)</td><td>39</td><td>74.6±6.3</td><td>1.38</td><td>4.77</td><td>60.2±6.4</td><td>FMM 2</td><td>1.56±0.07</td><td>38.6±4.5</td></tr></tbody></table>

<table><caption>Table S5: Element concentrations from high-resolution gamma spectrometry used for dose-rate calculation (see Tables S3 and S4 for output dose rates)</caption><thead><tr><th>Sample ID</th><th>U (ppm)</th><th>Th (ppm)</th><th><sup>40</sup>K (%)</th></tr></thead><tbody><tr><td>EYP1</td><td>0.89±0.11</td><td>3.97±0.08</td><td>0.79±0.01</td></tr><tr><td>OYT1</td><td>0.53±0.17</td><td>3.02±0.02</td><td>1.07±0.01</td></tr><tr><td>SPC5</td><td>1.26±0.31</td><td>5.91±0.13</td><td>0.98±0.01</td></tr><tr><td>ART1</td><td>0.79±0.1</td><td>3.73±0.11</td><td>0.90±0.01</td></tr><tr><td>ART2</td><td>0.85±0.23</td><td>2.63±0.08</td><td>0.78±0.01</td></tr><tr><td>PEN1</td><td>0.6±0.11</td><td>2.34±0.08</td><td>1.01±0.01</td></tr><tr><td>PIZ1</td><td>1.27±0.16</td><td>7.1±0.24</td><td>1.43±0.02</td></tr><tr><td>PIZ2</td><td>1.20±0.13</td><td>6.16±0.2</td><td>1.57±0.02</td></tr></tbody></table>

<footnote n="1">¹using a σ<sub>b</sub> of 0.3 results in FMM D<sub>e</sub> with less than 5% difference.</footnote>
## 4.3 Palaeogeographical reconstruction of the western French Alps foreland during the LGM using cosmogenic exposure dating.

Afin d'obtenir une chronologie directe du complexe morainique du lobe lyonnais, des datations par exposition ¹⁰Be de blocs erratiques et glaciaires ont été réalisées. Ces âges doivent permettre de contraindre l'âge du maximum d'extension du lobe lyonnais ainsi que de dater des étapes de recessions glaciaires. A partir de ce cadre chronologique, des reconstitutions paléogéographiques du glacier seront réalisées afin de reconstituer la géographie du maximum d'extension et des étapes de retrait glaciaire sur le piémont des Alpes occidentales françaises.

Ces travaux ont fait l'objet d'un article publié en anglais en 2022, dans le journal Quaternary Research.
# Palaeogeographical reconstruction of the western French Alps foreland during the LGM using cosmogenic exposure dating.

Thibault ROATTINO ¹, Christian CROUZET ¹, Riccardo VASSALLO ¹, Jean-François BUONCRISTIANI ², Julien CARCAILLET ³, Natacha GRIBENSKI ⁴, Pierre VALLA ³,⁴.

1: Univ. Grenoble Alpes, Univ. Savoie Mont Blanc, CNRS, IRD, Univ. Gustave Eiffel, ISTerre. Batiment Belledonne – Campus Scientifique, 73376 Le Bourget du Lac, France.

2: Biogéosciences, UMR 6282 CNRS/Bourgogne Franche-Comté University. 6 Boulevard Gabriel, 21000 Dijon, France.

3: Univ. Grenoble Alpes, Univ. Savoie Mont Blanc, CNRS, IRD, Univ. Gustave Eiffel, ISTerre. 1381 rue de la Piscine, 38400 Saint Martin d'Hères, France.

4: Bern University, Institute of Geological Sciences. Baltzerstrasse 3, 3012 Bern, Switzerland.

## 4.3.1 Abstract

The extent of glaciers in the western French Alps foreland during the Last Glacial Maximum (LGM, 26.5-19 ka) has not yet been determined, so understanding glacial palaeogeography during the LGM remains an open question. This study focuses on the glacial chronology in the western French Alps piedmont using ¹⁰Be surface exposure ages on 9 glacial boulders and 12 erratic boulders. Results indicate an LGM glacier advance between ca. 24 and 21 ka. During the Late LGM, a smaller glacier readvance or stabilization phase occurred at ca. 19 ka, and was then followed by a withdrawal phase between ca. 19 and 16.5 ka. Our outcomes show that the LGM extent in the western French Alps was similar or slightly less extensive than the pre-LGM ice extents during the last glacial period. Such palaeogeography has also been suggested in the western Italian Alps, which share the same accumulation zone with the western French Alps glaciers. The LGM dynamic of the western French Alps foreland glaciers highlighted by our exposure ages is consistent with the timing of the LGM glacier advances and deglaciation with the western Italian ice lobes.

Keywords: LGM, Glacier, Cosmogenic nuclides, western French Alps, Palaeogeography.
4.3.2 Introduction

For the Late Pleistocene (i.e. the last 129 ka), δ¹⁸O variations in planktonic fossils from oceanic cores have been used as a proxy for the continental ice volume (Emiliani, 1955; Shackleton, 1987), reflecting global palaeoclimatic variations and defining marine isotope stages (MIS). Ice cores from Greenland (Grootes et al., 1993) and Antarctica (Petit et al., 1999) have also provided high-resolution temperature reconstructions mainly from ice isotopic records. From these palaeoclimate proxies, it was shown that Late Pleistocene temperature minima occur during the MIS 4 and mainly during the MIS 2 (Bond et al., 1993; Lisiecki & Raymo, 2005). The last glacial maximum (LGM), which occurred during the MIS 2, between 26.5 and 19 ka (Clark et al., 2009) or between 30 and 16.5 ka (Lambeck et al., 2014), refers to the most recent cold period when the global ice sheet volume reached its maximum.

In the marginal areas of the European Alps, morainic complexes have been recognized and well mapped (Penck & Brückner, 1909; Ehlers et al., 2011 and references therein), indicating the large extent of the Alpine Late Pleistocene glaciers forming broad piedmont lobes. Two main glacier advance phases have been recorded during the LGM in the European Alps. The first one occurred between 24 and 22 ka in the Italian foreland (Monegato et al., 2007; Gianotti et al., 2008; Gianotti et al., 2015; Ravazzi et al., 2012a; Fontana et al., 2014; Scapozza et al., 2014; Braakhekke et al., 2020; Ivy-Ochs et al., 2018; Kamleitner et al., 2022) and in the northern Swiss foreland (Ivy-Ochs et al., 2004; Akçar et al., 2011; Reber et al., 2014; Graf et al., 2015; Ivy-Ochs, 2015) followed by a simultaneous retreat of the northern Swiss foreland lobes from their maximum positions (Ivy-Ochs et al., 2004; Akçar et al., 2011; Reber et al., 2014; Graf et al., 2015; Ivy-Ochs, 2015). The second glacier advance phase happened around 21-19 ka on the northern and southern sides of the Alps (Jorda, 1988; Lister, 1988; Monegato et al., 2007; Starnberger et al., 2011; Federici et al., 2012; Ravazzi et al., 2012a; Fontana et al., 2014; Reber et al., 2014; Ivy-Ochs, 2015; Federici et al., 2016; Ivy-Ochs et al., 2018; Braakhekke et al., 2020; Kamleitner et al., 2022). For the western French Alps, until recently, only scarce radiocarbon dates were available (Loebell, 1980; Evin et al., 1985; Mandier, 1988; Mandier & Piegay, 1991; Bossuet et al., 1996; Bintz & Evin, 2002; Mandier et al., 2003; Doyen et al., 2015) and two hypotheses were proposed for the chronology of the Late Pleistocene maximum glacier extent, mainly based on a correlation with isotopic data from Greenland. The first hypothesis proposes a glacial maximum around 27 ka and the second hypothesis suggests a glacial maximum around 22 ka (Schoeneich, 1998). Studies about the LGM palaeogeography of the western French Alps glaciers were then mainly based only on a morpho-stratigraphical approach (Enay, 1981; Mandier, 1984; Monjuvent, 1988; Kerrien & Monjuvent, 1990) and are not supported by quantitative chronology. More recently, an earlier timing for the glacial maximum has been proposed in the western French foreland (Gribenski et al., 2021). The dated deposits by Optically Stimulated Luminescence (OSL) techniques, are well-developed glacio-fluviatile sediments, still preserved today and geomorphologically connected to the morainic landforms associated with the Late Pleistocene maximum glacier advances (Mandier, 1984; Roattino et al., 2021). The results indicate two geometrically similar maximum advances during the MIS 4 and late MIS 3, whereas MIS 2 ages were not recorded (Gribenski et al., 2021). Finally, numerical simulations of the last glacial cycle (120-0 ka) in the European Alps indicated the absence of glacier advance on the western French Alps foreland during the LGM (Seguinot et al., 2018), or that the glacier may have spread to a limited extent without reaching the maximum MIS 4/late MIS 3 configurations (Mey et al., 2016; Becker et al., 2016; Visnjevic et al., 2020). The current lack of quantitative data for the LGM period raises the question of the glacier palaeogeography in the western French Alps foreland during this period. In the western French Alps, ¹⁰Be surface-exposure dating spanning the Late-glacial and the Holocene period has only been applied on the inner valleys (Coutterand & Nicoud, 2005; Cossart et al., 2008; Cossart et al., 2012; Perret, 2014; Schwartz et al., 2017; Wirsig et al., 2016a; Wirsig et al., 2016b; Le Roy et al., 2017; Protin, 2019; Protin et al., 2019; Prud'homme et al., 2020) and not in the foreland area.

In this context, this study aims to address two main questions: Did the glacier reached the foreland during the LGM? And if yes, was the LGM advance similar or significantly less extensive than the earlier advances recorded in the Lyon foreland during the last glacial period? Here we present new ¹⁰Be surface exposure dating results that in concert with geomorphological evidence, we used to define the LGM palaeogeography of the western French Alps foreland.
### 4.3.3 Geological setting

#### 4.3.3.1 Geography and bedrock geology

This study focuses on the western French Alps foreland (east of the city of Lyon, Figure II.1). The highest topographic reliefs in this area belong to the southern Jura (around 1100 m a.s.l.), the Chat Mountain (around 1400 m a.s.l.) and the subalpine Chartreuse massif (around 1800 m a.s.l.). The Rhône Valley is a wide valley that separates the southern Jura and the Chat Mountain before reaching the foreland. In the southern part of the foreland (Terres Froides, Figure II.1), the highest hills reach ~800 m a.s.l. are comprised of Miocene molasses. The rest of the foreland topography is a low-relief area with an average elevation of 300 m a.s.l., mainly formed by Miocene sandstone and conglomerate and by the monoclinal limestone series of the Crémieu plateau (Gigout et al., 1976; David et al., 1978; Monjuvent et al., 1980; Elmi et al., 1986; Kerrien & Monjuvent, 1990).

#### 4.3.3.2 Glacial geomorphology and palaeogeography

The last glacial cycle maximum extent in the study area is characterised by a range of moraine landforms known as the Internal Moraine Complex (IMC) (Buoncristiani & Campy, 2011), delimiting the extent of a wide piedmont lobe so called the Lyonnais ice lobe. At the scale of the Lyonnais ice lobe, three main morainic domains can be distinguished. The western morainic domain close to the Dombes plateau, the intermediate morainic domain mainly constituted by the IMC, and the eastern morainic domain located in the Crémieu Plateau and southern Jura (Figure II.1).

A new geomorphological map (Figure II.1) of the morainic domains was done using high-resolution digital elevation model (DEM) RGE ALTI® with 5m resolution from French National Institute of Geographic and Forest Information (IGN®), in southern Jura, Terres Froides, Southwestern foreland, Est Lyonnais (Steiner, 2020; Roattino et al., 2021).

According to the provenance of the pebbles, Lyonnais ice lobe is fed by glaciers flowing from Mont-Blanc, Beaufortain (Mandier et al., 2003; Buoncristiani & Campy, 2011), Tarentaise and part of the Maurienne accumulation zone (Mandier et al., 2003; Coutterand et al., 2009; Coutterand, 2010), (Figure II.1). Before reaching the piedmont, the ice flowed through the Rhône Valley between the southern Jura and the Chat mountain (Coutterand, 2010).

During its maximum extent, the Lyonnais ice lobe was approximately 55 km wide and stretched over 40 km into the piedmont. It was bounded to the north by the southern part of the Jura Mountains and to the south by the Terres Froides where the ice flowed in narrow valleys (Figure II.1). In the southern Jura, the moraine deposits are mainly arcuate with multiple ridges characteristic of ice marginal deposits. In this area, several longitudinal ridges correspond to lateral and frontal moraines (Kerrien & Monjuvent, 1990). In the Terres Froides, the morphology of the moraines is constrained by local topography (Figure II.1). Overall, several arcuate moraine ridges can be observed, reflecting several episodes of marginal glacial deposition (Steinfatt, 1975). In the foreland plain, the Lyonnais ice lobe has built well-preserved moraines in the Est Lyonnais area (Figure II.1). These glacial landforms are constituted by several sets of crests highlighting several glacial stages of the Lyonnais ice lobe identified and correlated from a morpho-stratigraphic and pedological point of view (Mandier et al., 2003; Steiner, 2020; Roattino et al., 2021).

In the southern Jura, glacial retreat moraines were recognized as frontal and arcuate frontal moraines in valleys behind or topographically below the morainic front attributed to the IMC (Kerrien & Monjuvent, 1990). During the retreat phases of the Lyonnais ice lobe, several series of morainic landforms, generally oriented NE-SW, with single, multiple, long-linear ridges were also deposited on the Crémieu plateau (Figure II.1) (Monjuvent, 1988; Steiner, 2020).
4.3.3.3 Available chronological data

Available ages, obtained in an attempt to date glacier advances in the study area, come from radiocarbon and OSL techniques. Sampling has focused only on glacio-fluviatile and postglacial deposits and no direct dating of the moraines exists (Loebell, 1980; Vilain et al., 1988; Mandier & Piegay, 1991; Bossuet et al., 1996; Bintz & Evin, 2002; Gribenski et al., 2021), mainly due to the general lack of in place boulders on moraine crests, especially along the IMC margins in the Lyon foreland.

OSL dating of glacio-fluviatile outwash terraces associated with the most external IMC moraines reveal two geometrically similar glacier extents. The earliest of these two glacier extents happened between ca. 70 and 60 ka (MIS 4) and is found locally in the Est Lyonnais area and southwestern foreland. The youngest of these two glacier advances occurred between ca. 42 and 29 ka (late MIS 3) (Gribenski et al., 2021), consistent with radiocarbon dating collected above the till near the Dombes plateau (Figure II.1) (Evin et al., 1985; Vilain et al., 1988).

Radiocarbon ages in the study area are found for the postglacial lacustrine, fluviatile and karst deposits (Table II.1). In Est Lyonnais, near the Dombes plateau (Figure II.1), sampled peat deposits above the outer IMC moraine provided a radiocarbon age of 30.41 -26.87 cal ka BP (Evin et al., 1985; Vilain et al., 1988). In the Eastern Crémieu plateau, radiocarbon dating was performed at a depth of 66 m within a lacustrine deposit (Mandier, 1988) yielding a radiocarbon age of 33.15-28.98 cal ka BP. Other organic samples (Figure II.1) below lacustrine deposit of Moras lake provide radiocarbon ages of 33.72-31.22 cal ka BP (Mandier et al., 2003).

A radiocarbon age for human occupation remains yielded an age of 16.30-15.26 cal ka BP (Bintz & Evin, 2002). In the north-east part of the Crémieu plateau, radiocarbon dating within a fluviatile terrace yielded an age of 17.43-16.77 cal ka BP (Mandier & Piegay, 1991). In the southern Jura, a lacustrine deposit yielded radiocarbon ages of 18.35-14.79 cal ka BP (Bossuet et al., 1996). Other lacustrine deposits in the Terre Froides and Crémieu areas provided radiocarbon ages of 14.04-13.80 and 13.31-13.16 cal ka BP respectively (Doyen et al., 2015). Lastly, in the very inner zone (Rhône Valley, Figure II.1), radiocarbon dating suggested the first human settlements at 18.33 -16.49 cal ka BP (Table II.1) (Loebell, 1980).
<figure><img src="image_80.png" /></figure>

Figure II.1 : Topographic map showing the main reliefs, the rivers and geomorphological sectors of the study area. The mapping of the western morainic domain (black ridges), intermediate morainic domain (brown ridges) and the eastern morainic domain (green ridges) is a synthesis of the different works carried out in this area. (Steinfatt, 1975; Mandier, 1984; Monjuvent, 1988; Monjuvent & Nicoud, 1988; Kerrien & Monjuvent, 1990; Steiner, 2020; Roattino et al., 2021). The locations of the available radiocarbon dates presented in Table II.1 are identified with white dots. The locations of the OSL dating (Gribenski et al., 2021) are marked by the red triangles for the Late MIS 3 ages and by the green triangles for the MIS 4 ages. The inset in the lower left corner shows the main European Alpine ice lobes during their maximal extent updated from Ehlers et al. (2011). The study area is delimited by a black square and the catchment of the Lyonnais ice lobe (Coutterand, 2010) is
demarcated by the red polygon with the main accumulation areas (MB: Mont Blanc; B: Beaufortain, T: Tarentaise, M: Maurienne).

<table><caption>Table II.1 : Summary of the radiocarbon dates available in the study area. Ages taken from previous studies are given in ka. These data are calibrated into ka Cal BP with IntCal 20 calibration curve (Reimer et al., 2020) along with the associated uncertainty. For each location, the oldest reliable age is presented.</caption><thead><tr><th>Location</th><th>Sample</th><th>Lat (°N)</th><th>Long (°E)</th><th>Elev (m)</th><th>Age (ka BP)</th><th>Uncertainty (ka BP)</th><th>Age (cal ka BP)</th><th>Geological Setting</th><th>Dated material</th><th>Source</th></tr></thead><tbody><tr><td>Cerin lake</td><td>Ly-5082</td><td>45°46,50'</td><td>5°33,87'</td><td>754</td><td>13.68</td><td>0.72</td><td>18.35-14.79</td><td>lacustrine deposit</td><td>Peat</td><td>Bossuet et al. 1996</td></tr><tr><td>Creys-Malville</td><td>Ly-3814</td><td>45°45,60'</td><td>5°28,36'</td><td>142</td><td>26.47</td><td>0.92</td><td>33.15-28.98</td><td>lacustrine deposit</td><td>Wood</td><td>Mandier, 1988</td></tr><tr><td>Garenne cave</td><td>Ly 702 / OxA 8163</td><td>45°40,27'</td><td>5°16,60'</td><td>336</td><td>13.15</td><td>0.18</td><td>16.30-15.26</td><td>Cave, Human occupation</td><td>Reindeer bone</td><td>Bintz et Evin, 2002</td></tr><tr><td>Moras Lake</td><td>Ly-1773</td><td>45°41,03'</td><td>5°15,93'</td><td>283</td><td>28.16</td><td>0.52</td><td>33.72-31.22</td><td>Lacustrine deposit</td><td>Pollen</td><td>Mandier et al, 2003</td></tr><tr><td>Moras Lake</td><td>SacA-12347</td><td>45°41,03'</td><td>5°15,93'</td><td>250</td><td>11.34</td><td>0.05</td><td>13.31-13.16</td><td>Lacustrine deposit</td><td>Charcoal</td><td>Doyen et al, 2015</td></tr><tr><td>Paladru Lake</td><td>SacA-20698</td><td>45°27,46'</td><td>5°32,38'</td><td>441</td><td>12.02</td><td>0.05</td><td>14.04-13.80</td><td>Lacustrine deposit</td><td>Charcoal</td><td>Doyen et al, 2015</td></tr><tr><td>Porcieu-Amblagnieu</td><td>Ly-4814</td><td>45°51,27'</td><td>5°24,02'</td><td>205</td><td>14.09</td><td>0.12</td><td>17.43-16.77</td><td>Fluviatile deposit</td><td>Horse bone</td><td>Mandier et Pigeay, 1991</td></tr><tr><td>Pugnieux</td><td>Ly-2318</td><td>45°51,17'</td><td>5°5,73'</td><td>200</td><td>24.11</td><td>0.90</td><td>30.41-26.87</td><td>Lacustrine deposit</td><td>Peat</td><td>Evin et al, 1985</td></tr><tr><td>Romains cave</td><td>Ly-16</td><td>45°42,56'</td><td>5°43,50'</td><td>280</td><td>14.38</td><td>0.38</td><td>18.33-16.49</td><td>Cave, Human occupation</td><td>Charcoal</td><td>Loebell, 1980</td></tr></tbody></table>

### 4.3.4 Materials and methods

#### 4.3.4.1 Sites and sampling strategy for the Terrestrial Cosmogenic Nuclide analyses

The timing of the glaciations was constrained by cosmogenic exposure dating on erratic and glacial boulders using Terrestrial Cosmogenic Nuclides ¹⁰Be. Nine glacial boulders and twelve erratic boulders were described and sampled with the aim to track the LGM glacial deposits. On this basis, the investigated areas are located on the intermediate and eastern morainic domains and also in more internal regions of the foreland (Figure II.1). However, on the southwestern foreland and Est Lyonnais sectors, the glacial and erratic boulders for ¹⁰Be exposure dating are very scarce. Indeed, the dense urbanization and agriculture on this area are probably responsible for the destruction or displacement of boulders that further reduces the opportunities for sampling. Only one large boulder (ART 18-01) is still preserved within the intermediate morainic domain and is located in the Southwestern foreland (Figure II.1, Figure II.2). Better preserved glacial morphologies and/or boulders were surveyed in less anthropized areas: the southern Jura, the Terres froides and the Rhône Valley (Figure II.3). Nevertheless, large boulders (> 1 m³) are sparse limiting the number of samples per location and requiring to select the best boulders. In the field, the biggest, well-anchored and quartz-rich boulders were in priority chosen to minimize potential late remobilization, post-abandonment displacement, as well as sustainable burial. The glacial boulders sampled are located on moraine ridges, while erratic boulders are located on bedrock or till deposits (Figure II.2, Figure II.3). A hammer, chisel, and angle grinder were used to sample rock chips measuring 4 cm thick. Sampling strategy targeted the top of the boulders to minimize the effect of the snow and/or the sediment burial through time. Topographic shielding values were measured using a clinometer and a compass.

#### 4.3.4.2 Analytical procedure

Rock samples were prepared at the GTC platform (ISTerre, Grenoble, France). Samples were crushed and sieved to keep the grain size between 200 and 500 µm. The magnetic fraction was first isolated
using a Frantz® magnetic separator, and the quartz was isolated from the other minerals by selective dissolution with several etchings by H₂SiF₆/HCl mixture (2/3-1/3). Quartz grains were etched three times with concentrated hydrofluoric acid (HF) in order to remove the atmospheric ¹⁰Be fraction. ~ 0,5 g of the Be standard solution was used to spike the quartz samples (See Table II.3 for details).

The ¹⁰Be target was prepared following the protocol of Brown et al. (1991) modified by Merchel and Herpers (1999). After total dissolution by concentrated HF, ¹⁰Be was extracted using chromatography techniques using anionic and cationic exchange resins (Dowex 1X8 and Dowex 50WX8). Before the conditioning for the AMS analysis, the Be target was oxidized to beryllium oxide (BeO) by heating to 900°C. The beryllium ratio (¹⁰Be/⁹Be) was measured using accelerator mass spectrometry at the ASTER AMS facility (Aix en Provence, France) (Arnold et al., 2010). The ¹⁰Be/⁹Be ratio was normalized with in-house CEREGE Standard Reference material called STD-11 (Braucher et al., 2015) with a theoretical ratio of 1.191.10⁻¹¹.

### 4.3.4.3 Age calculation

The online calculator Cosmic Ray Exposure Program (CREP, univ-lorraine.fr) (Martin et al., 2017) was used to calculate the surface ¹⁰Be exposure ages. The production rate scaling scheme used is the LSD model (Lifton et al., 2014), with the ERA-40 database (Uppala et al., 2005), used to calculate an atmospheric pressure and the VDM geomagnetic record described by Lifton (2016). Scaled production rate based on the Chironico landslide production rate calibration was used (Claude et al., 2014). These parameters were corrected for the sample thickness (4 cm) and the sample density (2.65 g.cm⁻³), and were scaled with the sample elevation, latitude, longitude and the topographic shielding correction (Table II.2). The surface-exposure ages were calculated with the null erosion rate, therefore the ages are the minimum estimates. No correction for snow effect was applied since no information about the snow cover record over time is available for the studied area. The present-day snow-covered period at the study area is short (1-2 weeks) and insignificant (< 20 cm). Therefore, the snow effect on the surface-exposure ages can be considered as negligible.

<table><caption>Table II.2: Geographic coordinates of the sampled boulders, site information and apparent size of the boulders.</caption><thead><tr><th>Sample</th><th>Areas</th><th>Latitude</th><th>Longitude</th><th>Elevation</th><th>Samplig year</th><th>Lithology</th><th>Géomorphological position</th><th>Sample informations</th></tr><tr><th></th><th></th><th>° N</th><th>° E</th><th>m a.s.l.</th><th></th><th></th><th></th><th>Apparent size (length*width*height, m)</th></tr></thead><tbody><tr><td>HIL 20-01</td><td></td><td>45°40,25'</td><td>5°16,63'</td><td>346</td><td>2020</td><td>Polygenic conglomerate</td><td>Boulder on morainic crest</td><td>4.2*1.1*2.7*</td></tr><tr><td>TRB 18-01</td><td rowspan="2">Crémieu plateau</td><td>45°41,69'</td><td>5°18,27'</td><td>339</td><td>2018</td><td>Polygenic conglomerate</td><td>Boulder on morainic crest</td><td>9.0*3.7*3.5</td></tr><tr><td>TRE 18-01</td><td>45°41,61'</td><td>5°18,26'</td><td>340</td><td>2018</td><td>Polygenic conglomerate</td><td>Boulder on morainic crest</td><td>6.5*4.0*3.0</td></tr><tr><td>BIL 19-01</td><td></td><td>45°42,76'</td><td>5°48,65'</td><td>410</td><td>2019</td><td>Micaschist</td><td>Boulder on till</td><td>2.5*1.5*0.5</td></tr><tr><td>CRV 20-01</td><td></td><td>45°42,11'</td><td>5°50,90'</td><td>824</td><td>2020</td><td>Granite</td><td>Boulder on till</td><td>2.5*1.2*0.7</td></tr><tr><td>CUL 19-01</td><td rowspan="2">Rhône Valley</td><td>45°50,71'</td><td>5°46,55'</td><td>300</td><td>2019</td><td>Schist</td><td>Boulder on the bedrock</td><td>4.0*2.0*2.0</td></tr><tr><td>LAT 19-01</td><td>45°36,73'</td><td>5°43,46'</td><td>856</td><td>2019</td><td>Granite</td><td>Boulder on till</td><td>1.5*0.8*0.7</td></tr><tr><td>YEN 19-01</td><td></td><td>45°42,16'</td><td>5°44,44'</td><td>350</td><td>2019</td><td>Schist</td><td>Boulder on the bedrock</td><td>2.0*1.0*0.5</td></tr><tr><td>INN 18-01</td><td></td><td>45°47,34'</td><td>5°35,36'</td><td>1113</td><td>2018</td><td>Schist</td><td>Boulder on till</td><td>1.0*0.6*0.4</td></tr><tr><td>INN 18-02</td><td rowspan="2">Southern Jura</td><td>45°47,29'</td><td>5°35,38'</td><td>1099</td><td>2018</td><td>Schist</td><td>Boulder on till</td><td>1.2*0.9*0.8</td></tr><tr><td>INN 18-03</td><td>45°47,24'</td><td>5°35,40'</td><td>1087</td><td>2018</td><td>Schist</td><td>Boulder on till</td><td>1.0*0.5*0.5</td></tr></tbody></table>
<table><tr><td>INN 18-04</td><td></td><td>45°47,18'</td><td>5°35,02'</td><td>972</td><td>2018</td><td>Quartzite</td><td>Boulder on morainic crest</td><td>5.0*2.4*2.0</td></tr><tr><td>ORD 18-02</td><td></td><td>45°51,83'</td><td>5°30,50'</td><td>826</td><td>2018</td><td>Schist</td><td>Boulder on morainic crest</td><td>1.3*1.2*0.8</td></tr><tr><td>ORD 18-03</td><td></td><td>45°51,90'</td><td>5°30,47'</td><td>817</td><td>2018</td><td>Quartzite</td><td>Boulder on morainic crest</td><td>1.5*1.2*0.9</td></tr><tr><td>ORD 18-04</td><td></td><td>45°52,12'</td><td>5°30,47'</td><td>807</td><td>2018</td><td>Schist</td><td>Boulder on morainic crest</td><td>1.3*1.2*0.5</td></tr><tr><td>THZ 18-01</td><td></td><td>45°52,84'</td><td>5°35,21'</td><td>864</td><td>2018</td><td>Flysch</td><td>Boulder on morainic crest</td><td>1.2*1.0*0.9</td></tr><tr><td>ART 18-01</td><td>Southwestern foreland</td><td>45°32,78'</td><td>5°9,12'</td><td>439</td><td>2018</td><td>Micaschist</td><td>Boulder on morainic crest</td><td>12.0*6.0*3.5</td></tr><tr><td>MIR 20-01</td><td></td><td>45°26,13'</td><td>5°42,27'</td><td>696</td><td>2020</td><td>Micaschist</td><td>Boulder on till</td><td>2.0*1.7*1.5</td></tr><tr><td>MIR 20-02</td><td>Terres Froides</td><td>45°27,30'</td><td>5°41,61'</td><td>788</td><td>2020</td><td>Granite</td><td>Boulder on till</td><td>4.1*3.4*3.1</td></tr><tr><td>PAL 20-01</td><td></td><td>45°26,11'</td><td>5°32,84'</td><td>795</td><td>2020</td><td>Amphibolite</td><td>Boulder on till</td><td>3.0*1.2*1.0</td></tr><tr><td>RAT 19-01</td><td></td><td>45°21,57'</td><td>5°39,66'</td><td>670</td><td>2019</td><td>Granite</td><td>Boulder on till</td><td>1.5*1.3*1.0</td></tr></table>
<figure><img src="image_81.png" /><figcaption>*Figure II.2 : Examples of sampled boulders on the southern Jura: INN 18-04 (A), Rhône Valley: CUL 20-01 (B) and CRV 20-01 (C), Crémieu plateau: TRE 18-01 (D), Southwestern foreland: ART 18-01 (E) and Terres Froides: MIR 20-02 (F).*</figcaption></figure>

### 4.3.5 Results

Exposures ages are grouped into five sectors; Southern Jura, Terres Froides, Southwestern foreland, Crémieu plateau and Rhône Valley (Figure II.4). We obtain 21 terrestrial cosmogenic nuclides dates, 20 of them are distributed between 23.81±1.12 and 2.91±0.20 ka with one much older sample at 496.91±19.20 (Table II.3).

Eight boulders were sampled in the southern Jura area, at altitudes between 800 and 1100 m a.s.l (Table II.2) and located on or close to well-preserved arcuate moraines. ORD 18-02, ORD 18-03 and ORD 18-04 glacial boulders are located on a frontal morainic crest (Figure II.4A) and give respectively 23.81±1.12, 14.08±1.01 and 14.83±0.64 ka (Table II.3 and Figure II.4). THZ 18-01 glacial boulder lies on a lateral morainic crest (Figure II.3) and provides 16.86±0.75 ka (Table II.3 and Figure II.3).

INN 18-01, INN 18-02 and INN 18-03 erratic boulders are located on a till deposit and provide surface-exposure ages of 16.56±0.68, 15.79±0.75 and 2.91±0.20 ka respectively. INN 18-04 glacial boulder lies
within a well-preserved frontal moraine (Figure II.4B) and yields 21.32±0.80 ka (Table II.3 and Figure II.3).

Four erratic boulders were sampled in the Terres Froides, between 670 and 800 a.s.l. Erratic boulders, MIR 20-01, MIR 20-02, PAL 20-01 and RAT 19-01 are located on till. They are not associated with a clear morphologic feature. Nevertheless, the area corresponds to the intermediate morainic domain (Figure II.4). These erratic boulders provide surface exposure ages of respectively 12.94±0.85, 22.72±0.96, 14.00±0.59 and 16.96±0.87 ka (Table II.3 and Figure II.4).

ART 18-01 is the only glacial boulder sampled in the Southwestern foreland. It lies further west, at 440 m a.s.l, on a hummocky moraine (Table II.2 and Figure II.4D), and yields an age of 20.71±0.80 ka (Table II.3 and Figure II.4).

Glacial boulders on the Crémieu plateau are located near the southern margin of this topographic high. Three glacial boulders TRB 18-01, TRE 18-01 and HIL 20-01 were sampled on a linear morainic ridge (Figure II.4E), at an altitude of 340 m a.s.l (Table II.2). These three samples yield ages respectively 17.08±0.84; 18.81±0.93 and 496.91±19.20 ka (Table II.3 and Figure II.4).

The last five sampled erratic boulders CUL 19-01, YEN 19-01, CRV 20-01, BIL 19-01 and LAT 19-01 are not in the foreland and located in a more internal zone to the east, within the Rhône Valley, located between 300 and 850 m a.s.l (Table II.2). Two boulders (CUL 19-01 and YEN 19-01) lie directly on the bedrock. The three other boulders (CRV 20-01, BIL 19-01, and LAT 19-01) are associated with glacial deposits, but without clear landforms (Table II.2, Figure II.3). They provide surface-exposure ages of respectively 16.50±1.08, 21.80±0.85, 15.28±0.65, 4.99±0.40 and 7.33 ± 0.43 ka (Table II.3 and Figure II.3).
<table><caption>Table II.3 : Cosmogenic-exposure data. The surface-exposure ages are calculated with a zero-erosion rate. The scale scheme used for the in-situ cosmogenic nuclide production rates is the LSD scaling scheme given by Lifton et al (2014). The atmospheric model considered is the ERA40 model (Uppala et al., 2005). The geomagnetic database used is the VDM provided by Lifton (2016). Last, production rates are scaled from the reference Chironico production rate (Claude et al., 2014). The thickness is ~4 cm for all the samples. Spike (a): Scharlau beryllium standard solution, 998 mg.l⁻¹; (b): LN2C, 3025mg.l⁻¹.</caption><thead><tr><th rowspan="2">Sample</th><th rowspan="2">Areas</th><th rowspan="2">Blank <sup>10</sup>Be/<sup>9</sup>Be</th><th rowspan="2">Blank <sup>10</sup>Be/<sup>9</sup>Be uncertainty</th><th rowspan="2"><sup>10</sup>Be/<sup>9</sup>Be (blank corrected)</th><th rowspan="2"><sup>10</sup>Be/<sup>9</sup>Be uncertainty</th><th rowspan="2">Mass dissolved quartz</th><th rowspan="2">Mass spike <sup>3</sup>Be</th><th rowspan="2">Spike concentration</th><th rowspan="2">Shielding factor</th><th colspan="2">Scaled production rates</th><th colspan="2"><sup>10</sup>Be concentration</th><th colspan="3">Ages with null erosion rate</th></tr><tr><th>Value</th><th>Uncertainty</th><th>Value</th><th>Uncertainty</th><th>Value</th><th>External uncertainty</th><th>Internal Uncertainty</th></tr><tr><th></th><th></th><th></th><th>%</th><th>at.g<sup>-1</sup></th><th>%</th><th>g</th><th>mg</th><th>mg.l<sup>-1</sup></th><th></th><th colspan="2">atoms/g/yr</th><th colspan="2">x10<sup>3</sup> atoms/gQtz</th><th colspan="3">ka</th></tr></thead><tbody><tr><td>HIL 20-01</td><td rowspan="3">Crémieu plateau</td><td>8.32E-15</td><td>9.80</td><td>3.567E-13</td><td>3.088</td><td>5.21</td><td>0.5240</td><td>998<sup>(a)</sup></td><td>1.000</td><td>5.4912</td><td>0.132</td><td>2333.5</td><td>74.43</td><td>496.91</td><td>19.20</td><td>15.07</td></tr><tr><td>TRB 18-01</td><td>2.64E-15</td><td>13.66</td><td>4.959E-14</td><td>4.200</td><td>17.69</td><td>0.1542</td><td>3025<sup>(b)</sup></td><td>1.000</td><td>5.3248</td><td>0.128</td><td>87.39</td><td>3.93</td><td>17.08</td><td>0.84</td><td>0.74</td></tr><tr><td>TRE 18-01</td><td>2.64E-15</td><td>13.66</td><td>6.329E-14</td><td>4.313</td><td>20.38</td><td>0.1538</td><td>3025<sup>(b)</sup></td><td>1.000</td><td>5.3248</td><td>0.128</td><td>96.56</td><td>4.38</td><td>18.81</td><td>0.93</td><td>0.82</td></tr><tr><td>BIL 19-01</td><td rowspan="5">Rhône Valley</td><td>8.32E-15</td><td>9.80</td><td>1.716E-14</td><td>4.240</td><td>37.42</td><td>0.5136</td><td>998<sup>(a)</sup></td><td>0.996</td><td>5.4496</td><td>0.131</td><td>26.16</td><td>2.07</td><td>4.99</td><td>0.40</td><td>0.38</td></tr><tr><td>CRV 20-01</td><td>8.32E-15</td><td>9.80</td><td>1.399E-13</td><td>3.221</td><td>64.19</td><td>0.5133</td><td>998<sup>(a)</sup></td><td>0.994</td><td>8.1952</td><td>0.197</td><td>124.99</td><td>4.32</td><td>15.82</td><td>0.65</td><td>0.54</td></tr><tr><td>CUL 19-01</td><td>8.32E-15</td><td>9.80</td><td>1.853E-14</td><td>4.330</td><td>19.74</td><td>0.5112</td><td>998<sup>(a)</sup></td><td>1.000</td><td>5.1584</td><td>0.124</td><td>81.68</td><td>5.14</td><td>16.50</td><td>1.08</td><td>1.01</td></tr><tr><td>LAT 19-01</td><td>8.32E-15</td><td>9.80</td><td>1.859E-14</td><td>4.945</td><td>25.04</td><td>0.5001</td><td>998<sup>(a)</sup></td><td>1.000</td><td>8.2784</td><td>0.199</td><td>58.60</td><td>3.43</td><td>7.33</td><td>0.43</td><td>0.39</td></tr><tr><td>YEN 19-01</td><td>8.32E-15</td><td>9.80</td><td>1.260E-13</td><td>3.098</td><td>58.66</td><td>0.5131</td><td>998<sup>(a)</sup></td><td>0.998</td><td>5.4080</td><td>0.130</td><td>113.44</td><td>3.84</td><td>21.80</td><td>0.85</td><td>0.69</td></tr><tr><td>INN 18-01</td><td rowspan="8">Southern Jura</td><td>2.64E-15</td><td>13.66</td><td>1.470E-13</td><td>3.400</td><td>27.39</td><td>0.1546</td><td>3025<sup>(b)</sup></td><td>1.000</td><td>10.5248</td><td>0.253</td><td>167.69</td><td>5.84</td><td>16.56</td><td>0.68</td><td>0.56</td></tr><tr><td>INN 18-02</td><td>2.64E-15</td><td>13.66</td><td>1.328E-13</td><td>4.156</td><td>26.30</td><td>0.1547</td><td>3025<sup>(b)</sup></td><td>1.000</td><td>10.4000</td><td>0.250</td><td>157.91</td><td>6.72</td><td>15.79</td><td>0.75</td><td>0.65</td></tr><tr><td>INN 18-03</td><td>2.64E-15</td><td>13.66</td><td>2.224E-14</td><td>5.565</td><td>25.19</td><td>0.1550</td><td>3025<sup>(b)</sup></td><td>1.000</td><td>9.8592</td><td>0.237</td><td>27.67</td><td>1.78</td><td>2.91</td><td>0.20</td><td>0.19</td></tr><tr><td>INN 18-04</td><td>2.64E-15</td><td>13.66</td><td>1.565E-13</td><td>3.092</td><td>25.40</td><td>0.1546</td><td>3025<sup>(b)</sup></td><td>0.997</td><td>9.4016</td><td>0.226</td><td>192.48</td><td>6.10</td><td>21.32</td><td>0.80</td><td>0.64</td></tr><tr><td>ORD 18-02</td><td>2.64E-15</td><td>13.66</td><td>6.604E-14</td><td>4.050</td><td>10.73</td><td>0.1535</td><td>3025<sup>(b)</sup></td><td>0.999</td><td>8.3616</td><td>0.201</td><td>190.93</td><td>8.13</td><td>23.81</td><td>1.12</td><td>0.97</td></tr><tr><td>ORD 18-03</td><td>2.64E-15</td><td>13.66</td><td>1.046E-13</td><td>6.848</td><td>29.60</td><td>0.1543</td><td>3025<sup>(b)</sup></td><td>0.999</td><td>8.1120</td><td>0.195</td><td>110.21</td><td>7.75</td><td>14.08</td><td>1.00</td><td>0.95</td></tr><tr><td>ORD 18-04</td><td>2.64E-15</td><td>13.66</td><td>7.159E-14</td><td>3.623</td><td>19.19</td><td>0.1529</td><td>3025<sup>(b)</sup></td><td>1.000</td><td>8.0704</td><td>0.194</td><td>115.28</td><td>4.38</td><td>14.83</td><td>0.64</td><td>0.54</td></tr><tr><td>THZ 18-01</td><td>2.64E-15</td><td>13.66</td><td>9.243E-14</td><td>3.757</td><td>20.97</td><td>0.1550</td><td>3025<sup>(b)</sup></td><td>0.999</td><td>8.5280</td><td>0.205</td><td>138.04</td><td>5.38</td><td>16.86</td><td>0.75</td><td>0.63</td></tr><tr><td>ART 18-01</td><td>Southwestern foreland</td><td>2.64E-15</td><td>13.66</td><td>1.090E-13</td><td>3.200</td><td>28.85</td><td>0.1525</td><td>3025<sup>(b)</sup></td><td>1.000</td><td>5.7810</td><td>0.141</td><td>116.52</td><td>3.85</td><td>20.71</td><td>0.80</td><td>0.65</td></tr><tr><td>MIR 20-01</td><td rowspan="4">Terres Froides</td><td>8.32E-15</td><td>9.80</td><td>1.050E-14</td><td>4.369</td><td>39.85</td><td>0.5067</td><td>998<sup>(a)</sup></td><td>0.998</td><td>7.2384</td><td>0.174</td><td>89.92</td><td>5.63</td><td>12.94</td><td>0.85</td><td>0.80</td></tr><tr><td>MIR 20-02</td><td>8.32E-15</td><td>9.80</td><td>1.090E-14</td><td>4.239</td><td>22.54</td><td>0.5045</td><td>998<sup>(a)</sup></td><td>0.994</td><td>8.0288</td><td>0.193</td><td>174.03</td><td>6.49</td><td>22.72</td><td>0.96</td><td>0.81</td></tr><tr><td>PAL 20-01</td><td>8.32E-15</td><td>9.80</td><td>1.403E-14</td><td>4.221</td><td>48.48</td><td>0.5128</td><td>998<sup>(a)</sup></td><td>1.000</td><td>7.9040</td><td>0.190</td><td>106.66</td><td>3.93</td><td>14.00</td><td>0.59</td><td>0.50</td></tr><tr><td>RAT 19-01</td><td>8.32E-15</td><td>9.80</td><td>1.653E-14</td><td>4.950</td><td>25.53</td><td>0.5140</td><td>998<sup>(a)</sup></td><td>0.993</td><td>7.1136</td><td>0.171</td><td>115.35</td><td>5.50</td><td>16.96</td><td>0.87</td><td>0.78</td></tr></tbody></table>
<figure><img src="image_82.png" /></figure>

Figure II.3: Simplified geological map with Quaternary units from the French geological survey (BRGM) (Gigout et al., 1976; David et al., 1978; Elmi et al., 1986; Kerrien & Monjuvent, 1990; Monjuvent et al., 1980) and location of the western morainic domain (black), intermediate morainic domain (brown) and eastern morainic domain (green) annotated with sampling locations (white circles, with their respective labels and ¹⁰Be surface-exposure ages). Topographical sections AA' and BB' showing the main sampling sectors and the location of the moraines.
<figure><img src="image_83.png" /></figure>

Figure II.4: Extract from the DEM showing morainic landforms (black arrows) of ORD (A) and INN (B) boulders sites in southern Jura, of MIR 20-02 (C) in Terres froides, of ART 18-01 (D) in southwestern foreland and of TRB 18-01, TRE 18-01, HIL 20-01 in Crémieu plateau. Location of each extract is given on the Figure II.3.
### 4.3.6 Discussion

#### 4.3.6.1 Terrestrial cosmogenic nuclides dating

Two types of geological processes may alter interpretation of the exposures ages on erratic and glacial boulders and induce subsequent erroneous chronological conclusions (Brown et al., 2005; Chevalier et al., 2005; Applegate et al., 2007; Barrows et al., 2007). The first process is the prior exposure to the cosmic rays inducing additional cosmogenic nuclides concentration, also called inheritance. This inherited concentration produces an apparent exposure age older than the true deposition age of the target moraine. The other process, is the incomplete post-depositional exposure because of shielding due to burial into a moraine or till. The subsequent exhumation yields reduced cosmogenic nuclides concentration with respect to a full exposure at the surface since deposition. This incomplete exposure induces apparent younger exposure age than the true deposition age. Boulders with significant pre-exposure inheritance are not predominant in glacial contexts because of significant glacier erosion and ice shielding. Incomplete exposure is much more frequent and is the principal source of error in age interpretation (Putkonen & Swanson, 2003; Briner et al., 2005; Heyman et al., 2011; Heyman et al., 2016). Considering this assumption, the surface-exposure age can be assessed as a minimum age for the deposition. This assumption is classically used in similar glacier contexts in the Alpine foreland (Ivy-Ochs et al., 2018; Kamleitner et al., 2022). For each site with multiple samples, the oldest exposure age is retained and considered as closest to the real age of the deposit. The preferred age is often derived from the tallest boulder in the most stable position (Putkonen & Swanson, 2003; Heyman et al., 2016; Tomkins et al., 2021). We consider the calculation of mean ages as not reliable in this case. For each site with a single sample, the exposure age is considered as a minimum age unless contradictory with the morpho-stratigraphical position of the boulder at the regional scale.

**Southern Jura:**

According with their geological and geomorphological setting, the three sites (INN, ORD and THZ) of this area belong to the same glacial stage (Monjuvent, 1988). For the multi sampled INN site (Figure II.3), the retained age of INN 18-04 at 21.32±0.80 ka corresponds to the oldest sample, even if it is not the highest in elevation (~140 m difference with highest boulder). INN 18-04 is the tallest erratic boulder sampled on the southern Jura (Figure II.2 and Table II.2), thus limiting incomplete exposure by burial. For ORD site (Figure II.3), the three boulders have a similar size, are located on the top of a well-defined morainic crest and not exceed 1 m height above the ground (Table II.2). The retained age corresponds to the oldest sample ORD 18-02 at 23.81±1.12 ka. The THZ site (Figure II.3) contains only one boulder. It is located on the internal slope of the moraine and does not protrude much from the ground (~1 m) (Table II.2), making it susceptible to an incomplete exposure by burial or gravitational processes. According to morphostratigraphical relationship of INN, ORD and THZ sampled sites, the age of THZ 18-01 at 16.86 ± 0.75 ka may be considered as underestimated. For the southern Jura area, the two dated boulders sites (ORD and INN) provide a similar oldest exposure ages of 23.81 ± 1.12 ka and 21.32 ± 0.80 ka respectively in accordance with the LGM. These dated glacial landforms are the outermost moraines of the southern Jura. No morainic landform attributable to older glacier advances is occurring beyond the dated moraines complex, only patches of undated till exist.

**Terres Froides**

The oldest exposure age of the Terres Froides (Figure II.3) is MIR 20-02 (Figure II.2F) at 22.72±0.96 ka and correspond to the tallest boulder (Figure II.2 and Table II.2), limiting incomplete exposure by burial. Similarly to the southern Jura area, this age correspond to the LGM. MIR 20-02 is not directly located on a morainic landform but a few hundred metres from a moraine of the same altitude (Figure II.4C). The age of the closest moraine should be comparable to the exposure age of the boulder. The elevation of this moraine is between 140 and 50 m below the highest and oldest moraines of the Terres
Froides (Steinfatt, 1975). This geomorphological evidence proves a preservation of prior older moraines located topographically above the LGM moraines dated by our exposure ages. According to Steinfatt (1975) and Mandier (1984), these higher moraines are correlated with glacio-fluviatile deposits dated MIS4 by Gribenski et al. (2021).

## Southwestern foreland

Sample ART 18-01 (Figure II.2E, Figure II.4D) is the only large boulder preserved on the IMC in the East Lyonnais and southwestern foreland (Figure II.3). It is also the largest (252 m³) and tallest (~3.5 m) sampled boulder in the study (Figure II.2 and Table II.2). It yields an exposure age of 20.71 ± 0.90 ka. It is however placed on the internal slope of the moraine rendering it vulnerable to incomplete exposure by partial burial or gravitational processes. Because of a lack of robustness due to the small amount of data, this age should be considered, with caution, as a minimum age of a glacier advance in southwestern foreland during the LGM. The glacio-fluviatile deposits dated as pre-LGM (Gribenski et al., 2021) in this area are correlated with IMC. The exposure age of ART 18-01 suggests that this glacial morphology is more complex and constituted by a stack of several glacier advances, the latest being LGM.

## Crémieu plateau

In this sector (Figure II.3), the three samples have the same lithology (ie. polygenic conglomerate, Table II.2). Therefore, they come from the same polygenic conglomerate source located in the internal Alps and probably from a same rockfall on the glacier surface. According to Coutterand et al., (2009) and Coutterand, (2010), this source corresponds to the "Flysch de Tarentaise". The oldest exposure age of the Crémieu plateau sector is HIL 20-01 at 497.48 ± 19.57 ka. This age is older than the oldest exposure age recorded from any sample issued from the Alpine glaciations at 144.0 ± 5.3 ka (Graf et al., 2015). The ¹⁰Be concentration measured for HIL 20-01 (~2330 x10³ atoms/g, Table II.3) is of the same magnitude order than a ¹⁰Be concentration found on rock wall scars in the Mont Blanc Massif at an elevation above 3400 m a.s.l (Gallach et al., 2018). We hence interpret that the ¹⁰Be concentration of HIL 20-01 is inherited from a prior exposure to cosmic rays on the original outcrop. TRB 18-01 and TRE 18-01 boulders protrude respectively 3.5 and 3.0 meters from the ground (Table II.2), and lie on a low elevated morainic landform (Figure II.4E), limiting incomplete exposure by burial or gravitational processes. These two samples yield similar exposure ages, and the oldest exposure age is TRE 18-01 at 18.81 ± 0.93 ka. This exposure age does not overlap within uncertainty with the other LGM exposure ages of the other sectors. It is therefore significantly younger and suggests a readvance or stabilization of the glaciers during the Late LGM.

## Rhône Valley

The oldest exposure age of the Rhône Valley sector (Figure II.3) is YEN 19-01 at 21.80±0.85 ka corresponding to the LGM period. YEN 19-01 sample is located on a bedrock ridge at 350 m a.s.l, similar to the elevation of the Crémieu plateau boulders. Considering YEN 19-01 internal position and low elevation, the apparent exposure age is too old in regard with the other samples giving an LGM exposure ages located at higher elevation and/or further in the foreland. Therefore, its overestimated age is probably explained by inheritance. Excluding YEN 19-01 sample, the oldest boulder on the Rhône Valley is CUL 19-01 at 16.50±1.08 ka. This erratic boulder is the tallest (2.0 meters high, Figure II.2, Table II.2) sampled in the Rhône Valley sector. Despite the absence of significant geomorphological evidence allowing to reconstitute the glacial front associated with CUL 19-01, inner position and very low elevation (300 m a.s.l) of this sample indicate that the lobe lyonnais glacier remained in the foreland at least until 16.50±1.08 ka during the Late-glacial period. This interpretation is supported by the ¹⁴C ages (Figure II.1 and Table II.1) of human settlement in the same area (Romains cave 18.33-16.49 cal ka BP) and the Cerin Lake sediments (18.35-14.79 cal ka BP).
4.3.6.2 Reconstruction of extent and timing of the western French Alps glacier

<figure><img src="image_84.png" /><figcaption>Figure II.5 : Time-relative distance diagram for the last glacial cycle glacier advances and retreat of the Lyonnais ice lobe according to OSL, ¹⁴C and our exposure ages. IMD; Intermediate Morainic Domain. EMD; Eastern Morainic Domain. Distance origin is located at the level of CUL 19-01 boulder.</figcaption></figure>

Previous studies on glacial landforms in the French Alps foreland were only based on relative chronology approach (Mandier, 1984; Roattino et al., 2021). Well-developed glacio-fluviatile sediments were dated by Optically Stimulated Luminescence (OSL) techniques (Gribenski et al., 2021), providing MIS 4 and late MIS 3 ages. Glacio-fluviatile deposit associated with the western morainic domain (Figure II.1) belong to the MIS 4. These moraines have never been directly dated because of the lack of boulders. In the same area, Vilain et al., (1988) studied paleosoil above the same till providing consistent pre-LGM ¹⁴C dates (Figure II.1 and Table II.1). These results evidence important pre-LGM glacier extent during the last glacial period (Figure II.5 and Figure II.6).

According to our results, dated moraines are at ca. 24-21 ka in Southern Jura, at ca. 23 ka in Terres Froides and at ca. 21 ka in Southwestern foreland. These exposure ages are considered as minimum ages. Theoretically, they could be much younger than the real glacier advance age. In this case, they would belong to pre-LGM glacier advances, and the LGM glacier may have not reached the front of the IMC.

However, the absence of older boulders in our dataset belonging to the last glacial period strongly suggests that these ages are representative of a major glacier advance in the region during the LGM. This glacial stage should therefore be associated with main morphological features at the Lyonnais ice lobe scale. This is the case of the moraines from which come the sampled glacial boulders. In the foreland, the intermediate morainic domain is a continuous geomorphic feature running over a length of several tens of kilometers from Terres Froides to Southwestern foreland and Est Lyonnais (Figure II.1). In the Southern Jura, the outermost frontal moraines of Ordonnaz and Innimond are the locally best-preserved geomorphic expression of a glacier advance (Figure II.4). The LGM exposure ages therefore coincide with the major glacier advance features both in the foreland and in the surrounding relief. The spatio-temporal correlation among these morainic domains appears as the most plausible hypothesis.

These results do not allow a direct correlation between these morainic landforms and the glacio-fluviatile deposits dated using OSL in the Southwestern area. However, they imply that at least in this area the two glacier extents (LGM and pre-LGM) are superimposed and confounded in the morphology. Several hypothesis may then be considered to explain why LGM ages were not recorded in the OSL based dating: 1) the LGM ice front was not thick enough to overcome the topographic step allowing glacio-fluviatile terrace aggradation observed beyond the intermediate morainic domain on southwestern foreland; 2) considering that the ice front was thick enough, because OSL sampling targets were limited to sand lenses embedded in the glacio-fluviatile sequences, a time delay between the levels sampled and the upper terrace surfaces could be possible; 3) potential partial bleaching, common in glacial environment, would result in overestimated OSL ages. However, the resolution (single grain) to which the measurements were undertaken as well as the use and concordance of multiple signals (quartz and feldspar) with varying bleaching properties suggest the reliability of the OSL ages.

On Crémieu plateau, dated moraines at ca. 19 ka are significantly younger than the dated moraines on Southern Jura, Terres Froides and Southwestern foreland. They are included into a more internal series
of frontal moraines located on the Crémieu plateau and associated with a glacier readvance or stabilisation phase (Enay, 1981; Steiner, 2020). This Late LGM age can be extended to the whole eastern morainic domain.

In the Rhône Valley sector, exposure ages indicate the retreat of the glacier from the foreland at ca. 16.5 ka. The age of the deglaciation of the western French Alps foreland is supported by the radiocarbon data in the study area associated with post-glacial lacustrine, fluviatile and karst deposits (Figure II.1 and Table II.1).

Deglaciation of the western French Alps foreland between ca 19 and ca. 16.5 ka is consistent with the initiation of ice thinning in the western French Alps accumulation zones, more precisely in the Mont Blanc area at ca. 18.5 ka (Wirsig et al., 2016b; Lehmann et al., 2020).

Our results are also in agreement with previous dating of the glaciers position at the entrance of the Romanche and Arc valleys at 17.5 and 15 ka, respectively (Schwartz et al., 2017; Prud'homme et al., 2020), and between 17.5 and 16 ka in the Arve valley (Coutterand & Nicoud, 2005; Perret, 2014).

### 4.3.6.3 The Lyonnais ice lobe dynamics in the Alpine context

Our results can be used to establish a palaeogeographic comparison between the Late Pleistocene advances of the Lyonnais ice lobe and the other Alpine foreland glaciers.

Pre-LGM glacier advances during the MIS 4 have been documented in the Swiss Alps from OSL dated fluviatile and glacio-fluviatile deposits (Preusser & Degering, 2007; Gaar & Preusser, 2019). The Italian foreland glaciers also shows an interesting geometric relationship between the LGM moraines and earlier morainic landforms. An apparent smaller LGM advance than the pre-LGM extent had been also suggested for the Ivrea glacier lobe and Rivoli-Avigliana lobe which shares the Mont Blanc/Tarentaise and Maurienne accumulation zones, with the Lyonnais ice lobe (Gianotti et al., 2008; Coutterand, 2010; Ivy-Ochs et al., 2018). Similarly, to the western morainic domain of the Lyonnais ice lobe, the pre-LGM during the last glacial period dates of the Rivoli-Avigliana glacier moraines are more external than the LGM moraines (Ivy-Ochs et al., 2018). Conversely, the other ice lobes in Italy and Switzerland reached their maximum extent during the LGM in the Alpine foreland (Wirsig et al., 2016b; Monegato et al., 2017; Braakhekke et al., 2020).

LGM moraines in the western French Alps foreland dated between ca. 24 and 21 ka (Figure II.5 and Figure II.6A) are consistent with the available chronology of the LGM foreland glaciers across the entire Alps. This result points a synchronous LGM glacier extent of the western French Alps glacier with respect to the other alpine glacier systems.

Glacier readvance or stabilization phase of the Late LGM highlighted by the Crémieu plateau exposure age (ca. 19 ka) (Figure II.5 and Figure II.6B) is also documented in the Italian foreland for the Rivoli-Avigliana glacier at 19.6 ± 0.9 ka (Ivy-Ochs et al., 2018), for the Ivrea glacier at 20.1 ± 3 (Gianotti et al., 2015) and for the Orta glacier at 19 ± 1 ka (Braakhekke et al., 2020). Innermost ice margin of the Verbano foreland glacier dated at 19.7 ± 1.1 ka, have been correlated to a short-lived glacier readvance (Kamleitner et al., 2022). A Late LGM advance has been evidenced between 22 and 17.5 ka on the Garda ice lobe (Monegato et al., 2017). In Switzerland, recessional moraines of the Reuss glacier have been dated at 18.6 ± 0.9 ka (Reber et al., 2014). These results show that the glacial landforms of the eastern morainic domain represent the impact of a major climatic stage at least at the scale of the western Alps.

Western French Alps foreland deglaciation between ca. 19 and 16.5 ka (Figure II.5 and Figure II.6C) is consistent with the initiation of ice thinning in the Alps accumulation zones between 18.5 and 14 ka (Dielforder & Hetzel, 2014; Hippe et al., 2014; Scapozza et al., 2014; Wirsig et al., 2016b; Lehmann et al., 2020). Parallel to this ice surface lowering in the alpine accumulation zones deglaciation of the Po plain foreland take place between 18 and 17 ka (Gianotti et al., 2008; Gianotti et al., 2015; Ravazzi et al., 2012a; Ravazzi et al., 2014; Lauterbach et al., 2012, Serra et al., 2022). In the north-eastern part of the Alps, the withdrawal seems to be slightly older between 20-18 ka (Lister, 1988; van Husen, 1997;
<figure><img src="image_85.png" /></figure>

Figure II.6 : Reconstructed palaeogeography of the Lyonnais ice lobe during the LGM. A: Reconstruction between ca. 24 and 21 ka, solid lines are a glacial front based on the sets of moraines recognized and considered synchronous according to morphostratigraphical, pedological and geological arguments by the previous studies and dated by our exposure ages. The gradient area delimited by thin solid lines represents the width of the intermediate morainic complex in the Southwestern foreland and Est Lyonnais where the position of the LGM front is not known precisely. Dotted line represents a supposed glacial front.

Reitner, 2007; Schmidt et al., 2012; Fontana et al., 2014). Finally, our results will improve the modelling of alpine glacier dynamics during the last glacial cycle. Contrarily to the results of previous models that lack datings in our study area (Seguinot et al., 2018), our findings show that the Lyonnais ice lobe returned to the piedmont during the LGM.

### 4.3.7 Conclusion

¹⁰Be surface-exposure dating of 9 glacial boulders and 12 erratic boulders allowed to investigate the glacial palaeogeography of the western French Alps foreland during the global LGM. Our results provide first evidences of the Lyonnais ice lobe extension in the piedmont area during the LGM. More precisely, our data suggest that the intermediate morainic domain corresponds to a glacier advance between ca. 24 and 21 ka. Palaeogeography highlighted by our dating and geomorphological mapping implies that the Lyonnais ice lobe extended over the foreland during the LGM, reaching the pre-LGM glacier extents in Southwestern foreland and Southern Jura areas. This palaeoglacial extension of the Lyonnais ice lobe during the LGM is similar to that of the Ivrea and Rivoli-Avigliana glaciers (Italian Alps), which share similar accumulation areas. At ca. 19 ka, the eastern morainic domain was deposited by a readvance or stabilization phase. This phase is synchronous of the western Italian and Switzerland forelands ice lobes. According to our exposure ages and radiocarbon datings, after this phase, the Lyonnais ice lobe starts a withdrawal phase and leaves the western French Alps foreland at ca.16.5 ka. This is in agreement with the timing of the deglaciation of the western Italian foreland ice lobes. This synchronicity of glacial phases at least at the scale of the Western Alps suggests a climatic forcing in glacier dynamics.
## 4.4 Synthèse

Dans ce chapitre, trois études ont permis de reconstituer la chronologie des différentes avancées glaciaires du lobe lyonnais sur le piémont des Alpes occidentales françaises au cours du Pléistocène supérieure et de répondre à la première problématique de ce projet de thèse: Quelle est la chronologie absolue des fluctuations du lobe lyonnais ?

Les terrasses retrouvées dans les couloirs fluvioglaciaires au front du complexe morainique du lobe lyonnais nous ont permis de déterminer à la fois la géométrie et l'âge des positions des fronts glaciaires des Alpes occidentales françaises au cours du dernier cycle glaciaire. Ces terrasses sont issues d'un environnement proglaciaire proximal et sont topographiquement corrélées à différents ensembles de moraines. Les datations par luminescence réalisées sur les dépôts qui constituent ces terrasses indiquent deux périodes majeures d'activité fluvioglaciaire : la plus ancienne associée aux terrasses corrélées aux moraines les plus externes du complexe des moraines internes (domaine morainique ouest) à 75-60 ka et la plus récente associée aux terrasses corrélées aux moraines intermédiaires de ce même complexe (domaine morainique intermédiaire) à 30-40 ka. Ces résultats suggèrent deux avancées géométriquement proches du lobe lyonnais au cours du MIS 4 et au MIS 3. Ces résultats permettent de mieux connaître la paléogéographie du lobe lyonnais au cours du dernier cycle glaciaire, mais ne nous renseignent pas sur les extensions relatives au LGM.

Pour étudier la période du LGM dans cette région nous avons utilisé une méthode chronologique qui vise directement les morphologies associées aux fluctuations glaciaires. 21 datations par âges d'exposition $^{10}$Be sur blocs erratiques et glaciaires ont été réalisées sur le complexe morainique du lobe lyonnais et sur le reste du piémont des Alpes occidentales françaises. Les résultats indiquent un abandon de blocs à 24-21 ka et suggèrent que la réavancée glaciaire au LGM a atteint à nouveau le domaine morainique intermédiaire. Cette avancée LGM est moins étendue que l'extension du MIS 4, et montre une extension très similaire à l'avancée glaciaire du MIS 3. Des âges d'exposition sur des blocs glaciaires localisés sur le domaine morainique est indiquent une réavancée/stabilisation du lobe lyonnais à ~19 ka. Enfin, le lobe lyonnais disparait définitivement à ~16,5 ka du piémont des Alpes occidentales françaises.
# 5 RÉSULTATS : DYNAMIQUE DE LA DÉGLACIATION TARDIGLACIAIRE DES GLACIERS SAVOISIENS
## 5.1 INTRODUCTION

Les fluctuations (avancées et retraits) des glaciers de vallées sont le témoin des variations climatiques. En effet, le bilan de masse des glaciers est influencé par les températures et précipitations qui jouent un rôle majeur dans l'ablation et l'accumulation de glace. Cependant, des facteurs locaux non climatiques tels que la topographie et l'orientation peuvent générer des fluctuations différentes pour des glaciers voisins (Oerlemans, 2001; Mackintosh, et al., 2017). Ainsi la reconstitution de paléoglaciers et la comparaison de leurs caractéristiques à l'échelle régionale constitue un proxy des conditions climatiques du passé (Kerschner & Ivy-Ochs, 2008). Ces fluctuations glaciaires peuvent être mises en évidence par l'étude de morphologies et de dépôts glaciaires couplée à des datations absolues.

Le tardiglaciaire (19,0-11,7 ka, Ivy-Ochs, et al., 2007; Schmidt, et al., 2012) est une période où le climat de l'hémisphère nord suit une tendance au réchauffement, entrecoupé par des événements froids brefs (Rasmussen, et al., 2014). Dans les Alpes européennes, la première détérioration climatique à lieu après le réchauffement de l'oscillation du Längsee à 19-18,5 ka (Schmidt, et al., 2012), et est associée à l'événement de Heinrich 1 (17,5-15,4 ka, Stanford, et al., 2011). Cet événement est caractérisé dans les Alpes par des températures estivales inférieures de 10°C par rapport à la valeur moyenne du 20ème siècle et par des précipitations inférieures de 30% par rapport aux valeurs actuelles (Kerschner & Ivy-Ochs, 2008). L'événement de Heinrich 1 est suivi d'un réchauffement pendant l'interstade du Bølling-Allerød à 14,6-12,8 ka (Heiri, et al., 2014b). Pendant cet interstade, l'augmentation des températures estivales (Juillet) estimée à partir de proxy biologiques varie entre 2 et 5°C sur le plateau suisse (Lotter, et al., 2012; Heiri, et al., 2014b) et entre 3 et 6°C en Suisse centrale (Grafenstein, et al., 2013; Bolland, et al., 2020). L'interstade du Bølling-Allerød est suivi du refroidissement du Dryas récent (YD ; 12,7-11,7 ka) (Ivy-Ochs, et al., 2007). Les températures moyennes annuelles sont alors comprises entre -5 et -10°C par rapport aux températures de la fin du 20ème siècle (1961-1990), et les paléoprécipitations sont très variables à l'échelle du massif alpin (Baroni, et al., 2021). Le climat du tardiglaciaire est donc très fluctuant avec des périodes de réchauffement sur de brèves périodes de temps, rappelant de manière moins extrême, le réchauffement climatique actuel. En effet, les Alpes européennes sont fortement impactées par le dérèglement climatique actuel et se sont réchauffées deux fois plus vites que la moyenne mondiale entre le 19ème siècle et le début du 21ème siècle (Auer, et al., 2005; Ceppi, et al., 2012).

Des études de morphologies et de dépôts glaciaires couplées à des datations absolues à travers les Alpes européennes ont mis en évidence une réponse des glaciers alpins à ces variations climatiques du Tardiglaciaire. En effet, après la disparition les grands lobes de piémont, plusieurs réavancées ou stagnations glaciaires au tardiglaciaires ont été corrélées avec des événements climatiques enregistrés dans les forages de glace de l'hémisphère nord. Le stade du Gschnitz défini entre ~17 et 16 ka (Ivy-Ochs, et al., 2006) correspond à la première avancée glaciaire après le LGM, corrélée à l'événement de Heinrich 1. Le stade de Daun (~14 ka ; Ivy-Ochs, 2015) correspond à une réponse au petit refroidissement pendant l'interstade Bølling-Allerød (Lotter, et al., 1992; Samartin, et al., 2012; Lih, et al., 2021). Enfin, le stade d'Egesen (~13,5-12,0 ka; Ivy-Ochs, 2015) est la dernière grande avancée du Tardiglaciaire mise en évidence sur l'ensemble du massif des Alpes en Autriche (Kerschner, 1980; Ivy-Ochs, et al., 2006; Ivy-Ochs, et al., 2009; Moran, et al., 2016), en Italie (Ivy-Ochs, et al., 2008; Ivy-Ochs, et al., 2009; Federici, et al., 2008; Federici, et al., 2017; Baroni, et al., 2014; Baroni, et al., 2017; Reitner, et al., 2016; Spagnolo & Ribolini, 2019), en Suisse (Ivy-Ochs, et al., 2009; Boxleitner, et al., 2019a) et en France (Darnault, et al., 2011; Cossart, et al., 2012; Chenet, et al., 2016; Protin, 2019).

Dans les Alpes occidentales françaises, la chronologie de la déglaciation au Tardiglaciaire est parcellaire. En outre, les moraines sont mal préservées ce qui ne permet pas d'avoir une vision exhaustive des fluctuations glaciaires. Ainsi, une analyse approfondie de dépôts proglaciaires mieux préservés est une opportunité de mettre en évidence ces fluctuations.

Dans ce travail nous allons analyser les modalités du retrait durant le Tardiglaciaire des glaciers des Alpes occidentales françaises depuis leur maximum d'extension (le lobe lyonnais) sur le piémont, jusque dans la vallée de la Maurienne. Parmi les grands bassins versants qui alimentent le lobe lyonnais, c'est dans celui de l'Arc que l'on retrouve le plus d'objets datables et les dépôts les mieux préservés nous
renseignant sur l'histoire du retrait glaciaire. Le retrait glaciaire sera reconstitué en définissant différentes étapes de retrait ou stades d'avancée, où la chronologie de la déglaciation sera contrainte en utilisant de nouvelles données de chronologie absolue (âges d'exposition $^{10}$Be et OSL) et des datations publiées. Des études sédimentologiques nouvellement acquises sur des dépôts proglaciaires bien préservés en Maurienne permettrons de mettre en évidence de nouvelles fluctuations glaciaires à haute résolution temporelle. Enfin une modélisation géométrique des glaciers pour les différentes fluctuations glaciaires définies en Maurienne permettra de calculer la variation du volume et du bilan de masse moyen d'un glacier de vallée pendant le réchauffement du Bølling-Allerød. Ces résultats seront ensuite comparés avec les bilans de masse des glaciers actuels en réponse au réchauffement climatique sous forçages anthropiques.

### 5.1.1 Contexte géographique

La zone d'étude de ce chapitre correspond à un secteur s'étendant du piémont des Alpes occidentales françaises jusqu'à la Maurienne (Figure 5.1). En amont du piémont, le relief est caractérisé par de larges vallées dont celle de l'Isère séparant le massif de Belledonne des massifs des Bauges (culminant à 2200 m) et de la Chartreuse (culminant à 2082 m). Plus à l'ouest, la Cluse de Chambéry sépare le massif des Bauges et de Chartreuse-Mont du Chat et passe au nord-ouest au Val du Bourget accueillant le lac du Bourget. Ce lac est le plus grand lac naturel de France avec une surface de 44,5 km². La zone d'étude passe ensuite au bassin versant de l'Arc d'une surface de 1975 km². L'Arc est une rivière d'environ 116 km de long s'écoulant de l'est vers l'ouest à partir du massif des Alpes graies pour confluer en aval avec l'Isère. Les principaux sommets culminent à plus de 3500 m, le plus haut étant la Grande Casse à 3855 m. Les glaciers actuels présents dans le bassin-versant de l'Arc couvrent une surface de 98 km² et sont principalement développés sur les sommets orientaux (BRGM, 2005).

<figure><img src="image_86.png" /></figure>

Figure 5.1: Carte topographique montrant les principaux reliefs, rivières et lacs de la zone d'étude. La localisation des âges d'exposition publiés sur blocs erratiques et glaciaires est indiquée par les diamants verts, les âges d'exposition publiés sur polis sont indiqués par les losanges verts. Enfin, les âges radiocarbone publiés sont indiqués par les ronds rouges. Ces âges sont introduits dans le chapitre état de l'art. Dans le piémont, les âges d'exposition sont présentés dans le chapitre 4.B. Les zooms A et B montrent la position détaillée des blocs de Bessans et de Bonneval-sur-Arc.
### 5.1.2 Géologie du Quaternaire

Le paysage de la zone d'étude a fortement été modelé par l'action des glaciers des Alpes occidentales françaises au cours des glaciations du pléistocène. Sur le piémont, le complexe morainique du lobe lyonnais est divisé en trois domaines datés dans le chapitre 4. Le domaine morainique ouest est daté au MIS 4, les extensions du MIS 3 et du LGM (24-21 ka) sont confondues au niveau du domaine morainique intermédiaire (DMI) et enfin le domaine morainique est (DME) est daté à 19 ka (Figure 5.1).

Après le LGM, les glaciers alimentant le lobe lyonnais occupent le Val du Bourget, la cluse de Chambéry et la vallée de l'Isère. Toutes ces vallées englacées montrent de grands surcreusements séparés par des seuils de substratum (Nicoud, et al., 1987; Monjuvent & Nicoud, 1988). Un surcreusement affecte une grande partie de la vallée de l'Isère et est rempli principalement de faciès lacustres déposés après le retrait glaciaire, et dans une moindre mesure de dépôts grossiers fluviaux (Fourneaux, 1976; Fourneaux, 1979; Nicoud, et al., 1987; Nicoud, et al., 2002). Des analyses géophysiques et des analyses de sondages ont montré une morphologie du substratum caractérisée par des sillons parallèles à la direction d'écoulement de la glace (Fourneaux, 1976; Fourneaux, 1979).

Les massifs présents dans le bassin versant de l'Arc forment une des zones d'accumulation alimentant le glacier de l'Arc constituant un tributaire du lobe lyonnais. Après le LGM, le glacier de l'Arc forme un glacier de vallée se retirant pendant tout le Tardiglaciaire (Coutterand, 2010). Des dépôts et morphologies hérités de l'action de ce glacier et de son retrait sont donc localement préservés dans le bassin versant de l'Arc.

Le till est plaqué sur l'ensemble du bassin versant et forme des dépôts atteignant localement plus de 50 m d'épaisseur (Hugonin, 1988; Marnezy, 1999). Les moraines sont présentes uniquement à partir de 1000 m d'altitude et principalement sur les versants de la vallée de l'Arc. Les moraines en fond de vallée sont très rares et sont localisées uniquement en partie très amont de la vallée à Bessans (Figure 5.2 et Figure 5.3) (Barbier, et al., 1973; Barféty, 1980; Hugonin, 1988; Debelmas, 1989; Fudral, et al., 1994; Crouzet, et al., 1999; Doudoux, et al., 1999; Barféty, et al., 2006).

Les mouvements et dépôts gravitaires sont abondants sur l'ensemble de la surface du bassin versant de l'Arc. Ils sont représentés par de grands glissements pouvant affecter les versants sur des surfaces de plusieurs dizaines de kilomètres carrés (Figure 5.2 et Figure 5.3). Ces glissements affectent principalement des versants constitués de schistes (Martin-Cochet, 1984; Hugonin, 1988; Fudral, et al., 1994; Fudral, 1998; Marnezy, 1999). De grands glissements sont encore actifs aujourd'hui (Martin-Cochet, 1984; Marnezy, 1999). Les mouvements et dépôts gravitaires sont également représentés par des écoulements parfois très important. Le volume de l'un de ces écoulements affectant des schistes lustrés et localisé en partie amont de la vallée de l'Arc au niveau du col de la Madeleine (Figure 5.3) a été estimé à $10^8 m^3$ (Couture, et al., 1997). Les écoulements sont parfois anciens, d'âge tardiglaciaire (Couture, et al., 1997) ou plus récent (Moyen âge) ou encore actuels comme certains ayant eu lieu en 1990 (Marnezy, 1999). Les éboulis sont plus rares et forment de petits cônes au pied des versants. Certains de ces éboulis sont encore actifs aujourd'hui (Marnezy, 1999).

La plupart des cônes alluviaux sont localisés en zone de confluence des vallées latérales avec la vallée principale de l'Arc (Barbier, et al., 1973; Barféty, 1980; Hugonin, 1988; Debelmas, 1989; Fudral, et al., 1994; Crouzet, et al., 1999; Doudoux, et al., 1999; Amelot, et al., 2003; Barféty, et al., 2006). Ces cônes sont constitués d'éléments mal triés de taille allant jusqu'aux blocs dans une matrice argileuse (Hugonin, 1988; Crouzet, et al., 1999) indiquant un dépôt par des coulées de débris (Mulder & Alexander, 2001). La majorité de ces cônes sont actuellement végétalisés et anthropisés, certains sont encore actifs aujourd'hui comme par exemple à Modane.

Le remplissage du fond de la vallée de l'Arc est assez mal connu. Des analyses de forages ont mis en évidence la présence de dépôts gravitaires, de coulées de débris, de cônes alluviaux, d'écoulements et enfin lacustres. En partie aval de la vallée, ces dépôts sont recouverts par des épandages fluviatiles, comme montré par l'analyse de forage montrant un remplissage lacustre et un cône alluvial recouvert d'un dépôt fluviatile issu de l'Arc (Coupe A, Figure 5.4) (Marnezy, 1999). Une succession montrant des
dépôts à lithologie hétérogène attribuable à un dépôt fluviatile à fluvio-glaciaire surmontés par des dépôts de coulées de débris, d'écroulements et issus de cônes alluviaux ont été reconnus par des forages (Coupe B, Figure 5.5) (Crouzet, et al., 1999). Les dépôts gravitaires, lacustres et de cônes alluviaux en Maurienne deviennent de plus en plus incisés par l'Arc et par les rivières issues des versants en direction de l'amont (Hugonin, 1988; Marnezy, 1999). Des bassins accueillant des dépôts lacustres et deltaïques ont été identifiés principalement en amont de la vallée au Verney, à Lanslebourg et à Bessans (Figure 5.3), dont l'origine a été attribuée à des événements gravitaires, à la présence de verrous ou encore à des barrages glaciaires. Ces dépôts sont recouverts de dépôts fluviaux, de cônes de déjection ou encore par des alluvions d'origine glaciolacustres ou encore fluvioglaciaires (Letourneur, et al., 1983; Hugonin, 1988; Debelmas, 1989; Fudral, et al., 1994).
<figure><img src="image_87.png" /></figure>

Figure 5.2: Carte des morphologies et dépôts quaternaires de la partie aval du bassin versant de l'Arc (BRGM, 2005). Seules les incisions, les cônes alluviaux, les dépôts gravitaires et lacustres en fond de vallée sont représentés. Les emprises A et B indiquent la localisation des coupes des Figure 5.4 et Figure 5.5.
<figure><img src="image_88.png" /></figure>

Figure 5.3: Carte des morphologies et dépôts quaternaires de la partie amont du bassin versant de l'Arc (BRGM, 2005). Seules les incisions, les cônes alluviaux, les dépôts gravitaires et lacustres en fond de vallée sont représentés.

<figure><img src="image_89.png" /></figure>

Figure 5.4: Coupe A d'après Marnezy 1999 représentant les relations entre des dépôts de cônes alluviaux, lacustres et fluviatiles dans le bassin versant de l'Arc. La localisation est donnée dans la Figure 5.2.
<figure><img src="image_90.png" /></figure>

(a): Cône de déjection, (b): Coulées boueuses, (c): Ecoulement, (d): Formation hétérogène, (e): Alluvions "anciennes" de l'Arc.

Figure 5.5: Coupe B d'après Crouzet et al. (1999) représentant les relations entre des dépôts de coulées de débris, d'écroulements, issus de cônes alluviaux et fluviatiles dans le bassin versant de l'Arc. La localisation de la coupe est donnée dans la Figure 5.2.
## 5.2 DATATION $^{10}$Be DE LA DÉGLACIATION POST-LGM ENTRE CHAMBÉRY ET LA VALLÉE DE LA MAURIENNE

Cette partie vise à reconstituer les différentes positions occupées par les glaciers des Alpes occidentales françaises depuis le LGM jusqu'au Dryas récent. Ces positions glaciaires seront reconstituées à partir de l'analyse de nouveaux âges d'exposition sur blocs erratiques et d'âges d'exposition publiés.

### 5.2.1 Échantillonnage et calcul des âges d'exposition

En plus des datations déjà existantes dans la littérature (§ 2.5.4.4), plusieurs blocs erratiques ont été échantillonnés. Trois blocs erratiques CHA 20-01, LCR-1, CRU-1 entre Chambéry et l'entrée de la Maurienne et trois blocs erratiques ROC 20-01, ROC 20-02 et ROC 20-03 localisés en Maurienne (Tableau 5.1) ont été décrits et échantillonnés dans le but d'obtenir de nouvelles contraintes chronologiques absolues. Les échantillons ont été traités et les âges d'exposition $^{10}$Be calculés en suivant le protocole décrit dans le chapitre 3 « Matériel et Méthode ». De même, les âges d'exposition publiés localisés à Montsapey, Aussois, Bessans et Bonneval-sur-Arc ont été recalculés suivant la méthode de calcul détaillée dans le chapitre 3 « Matériel et Méthode ».

<table><caption>Tableau 5.1: Coordonnées et informations sur les blocs erratiques échantillonnés sur les sites de Cruet, de la Croix-de-la-Rochette, de Challes les eaux et du pas du Roc.</caption><thead><tr><th>N° échantillon</th><th>Type d'échantillon</th><th>Localisation</th><th>Lat (°dec)</th><th>Long (°dec)</th><th>Altitude (m. snm)</th><th>Taille (m)</th></tr></thead><tbody><tr><td>CRU_1</td><td>Bloc erratique</td><td>Cruet</td><td>45,5346</td><td>6,0932</td><td>472</td><td>5,6*4,7*4,7</td></tr><tr><td>LCR-1</td><td>Bloc erratique</td><td>La Croix-de-la-Rochette</td><td>45,4739</td><td>6,1029</td><td>752</td><td>6,5*6,0*5,0</td></tr><tr><td>CHA-20-01</td><td>Bloc erratique</td><td>Challes les eaux</td><td>45,5400</td><td>5,9830</td><td>350</td><td>2,0*1,5*1,0</td></tr><tr><td>ROC-20-01</td><td>Bloc erratique</td><td>Pas du Roc</td><td>45,2240</td><td>6,4549</td><td>911</td><td>3,3*2,6*1,4</td></tr><tr><td>ROC-20-02</td><td>Bloc erratique</td><td>Pas du Roc</td><td>45,2260</td><td>6,4546</td><td>947</td><td>1,6*1,7*0,9</td></tr><tr><td>ROC-20-03</td><td>Bloc erratique</td><td>Pas du Roc</td><td>45,2305</td><td>6,4559</td><td>918</td><td>2,0*1,3*0,45</td></tr></tbody></table>

### 5.2.2 Résultats

CHA 20-01, LCR-1 (Figure 5.6) et CRU-1 sont trois blocs erratiques de schiste, de migmatite et de quartzite respectivement. CHA 20-01 est localisé à 350 m d'altitude sur le fond de vallée de Chambéry (Figure 5.7 et Tableau 5.1), sur un petit drumlin. LCR-1 est situé à 752 m d'altitude (Tableau 5.1), sur un plaquage de till posé sur un massif de substratum allongé surplombant la vallée de l'Isère. CRU-1 est localisé à 472 m d'altitude (Tableau 5.1), sur du till localisé sur le versant nord de la vallée de l'Isère. Les âges d'exposition de ces trois blocs indiquent respectivement 16,91 ± 1,10, 15,35 ± 0,58 et 16,31 ± 1,04 ka (Figure 5.7, Tableau 5.2).

ROC 20-01 (Figure 5.6), ROC 20-02 et ROC 20-03 sont trois blocs erratiques de très grand nombre de localisations. Ces blocs sont localisés sur un massif rocheux (Pas du Roc) surplombant d'environ 250 m un étroit verrou (Figure 5.7A). Les âges d'exposition de ces trois blocs erratiques donnent respectivement 15,06 ± 0,86 ; 14,47 ± 0,58 et 16,17 ± 0,70 ka (Figure 5.7, Tableau 5.2). Les âges d'exposition recalculés pour Bessans 05-01 (Figure 5.6), Bessans 05-02, Bessans 05-03 et Bessans 05-04 indiquent respectivement 13,67 ± 1,00 ; 12,20 ± 0,56 ; 12,37 ± 1,03 et 13,23 ± 0,64 ka (Figure 5.7, Tableau 5.2). Concernant les datations issues de Couette et Ghienne (in prep.), les âges d'exposition des polis donnent 11,84 ± 0,52 et 11,21 ± 0,51 ka. L'âge de la moraine la plus externe de Bessans est donc contraint par 6 âges d'exposition. En amont des moraines de Bessans à Bonneval-sur-Arc, 2 polis (Ev-6 et Ev-8) donnent respectivement 11,66 ± 0,45 et 12,19 ±
0,51 ka (Figure 5.7, Tableau 5.2). Ces polis sont localisés à environ 2635 m d'altitude, soit 965 m au-dessus de la moraine externe de Bessans. Ensuite, les âges d'exposition recalculés pour MSAP12_P1_01 et MSAP12_B1_06 donnent respectivement 12,65 ± 0,50 et 14,89 ± 1,76 ka. Enfin, après le recalcul de AUSS12_P1_01, AUSS12_B1_05, AUSS12_B2_06 et AUSS12_B3_07, les âges d'exposition sont respectivement 11,54 ± 1,65 ; 14,25 ± 0,64 ; 15,10 ± 1,03 et 13,86 ± 0,75 ka (Figure 5.7, Tableau 5.2).
<table><caption>Tableau 5.2: Données sur les âges d'exposition entre le piémont des Alpes occidentales françaises et la Maurienne. Les âges d'exposition publiés et issus de cette étude (CRU, LCR, CHA et ROC) ont été respectivement recalculés et calculés avec le calculateur en ligne CREp et avec les mêmes paramètres : Le taux d'érosion utilisé est nul. Le facteur d'échelle utilisé pour les taux de production des nucléides cosmogéniques in situ est le LSD issu de Lifton et al (2014). Le modèle atmosphérique considéré est le modèle ERA40 (Uppala et al., 2005). La base de données géomagnétique utilisée est la VDM fournie par Lifton (2016). Enfin, les taux de production sont échelonnés à partir du taux de production de référence du glissement de Chironico (Claude et al., 2014).</caption><thead><tr><th>N° échantillon</th><th>Epaisseur (cm)</th><th>Densité (g/cm3)</th><th>Ecrantage</th><th>Concentration <sup>10</sup>Be (at/g)</th><th>Incertitude <sup>10</sup>Be (at/g)</th><th><sup>10</sup>Be-std</th><th>Année d'échantillonnage</th><th>Age <sup>10</sup>Be publié (ans)</th><th>Incertitude publiée (ans)</th><th>Age <sup>10</sup>Be (ans) CREp</th><th>Incertitude externe (ans) CREp</th><th>Incertitude interne (ans) CREp</th></tr></thead><tbody><tr><td>CRU_1</td><td>4,0</td><td>2,65</td><td>0,991</td><td>93164</td><td>5695</td><td>NIST_27900</td><td>2018</td><td></td><td></td><td>16310</td><td>1040</td><td>970</td></tr><tr><td>LCR-1</td><td>4,0</td><td>2,65</td><td>1,000</td><td>113142</td><td>3626</td><td>NIST_27900</td><td>2018</td><td></td><td></td><td>15350</td><td>580</td><td>470</td></tr><tr><td>CHA-20-01</td><td>4,0</td><td>2,65</td><td>0,998</td><td>87051</td><td>5471</td><td>NIST_27900</td><td>2020</td><td></td><td></td><td>16910</td><td>1100</td><td>1030</td></tr><tr><td>ROC-20-01</td><td>4,0</td><td>2,65</td><td>0,992</td><td>125789</td><td>6842</td><td>NIST_27900</td><td>2020</td><td></td><td></td><td>15060</td><td>860</td><td>780</td></tr><tr><td>ROC-20-02</td><td>4,0</td><td>2,65</td><td>0,992</td><td>124446</td><td>4251</td><td>NIST_27900</td><td>2020</td><td></td><td></td><td>14470</td><td>580</td><td>470</td></tr><tr><td>ROC-20-03</td><td>4,0</td><td>2,65</td><td>0,992</td><td>136091</td><td>5094</td><td>NIST_27900</td><td>2020</td><td></td><td></td><td>16170</td><td>700</td><td>590</td></tr><tr><td>MSAP12_P1_01</td><td>2,0</td><td>2,70</td><td>0,997</td><td>120400</td><td>3920</td><td>NIST_27900</td><td>2012</td><td>12820</td><td>640</td><td>12650</td><td>500</td><td>410</td></tr><tr><td>MSAP12_B1_06</td><td>2,0</td><td>2,70</td><td>0,997</td><td>142450</td><td>17160</td><td>NIST_27900</td><td>2012</td><td>15160</td><td>1850</td><td>14890</td><td>1760</td><td>1730</td></tr><tr><td>AUSS12_P1_01</td><td>2,0</td><td>2,70</td><td>0,998</td><td>156160</td><td>22830</td><td>NIST_27900</td><td>2012</td><td>11520</td><td>1750</td><td>11540</td><td>1650</td><td>1630</td></tr><tr><td>AUSS12_B1_05</td><td>2,0</td><td>2,70</td><td>0,994</td><td>192730</td><td>7830</td><td>NIST_27900</td><td>2012</td><td>14310</td><td>800</td><td>14250</td><td>640</td><td>560</td></tr><tr><td>AUSS12_B2_06</td><td>2,0</td><td>2,70</td><td>0,994</td><td>204740</td><td>13670</td><td>NIST_27900</td><td>2012</td><td>15190</td><td>1100</td><td>15100</td><td>1030</td><td>970</td></tr><tr><td>AUSS12_B3_07</td><td>2,0</td><td>2,70</td><td>0,994</td><td>187370</td><td>9490</td><td>NIST_27900</td><td>2012</td><td>13870</td><td>880</td><td>13860</td><td>750</td><td>670</td></tr><tr><td>Bessans 05-01</td><td>4,0</td><td>2,65</td><td>0,972</td><td>219716</td><td>15997</td><td>NIST SRM 4325 2.79.10-11</td><td>2007</td><td>11327</td><td>1064</td><td>13670</td><td>1000</td><td>950</td></tr><tr><td>Bessans 05-02</td><td>4,0</td><td>2,65</td><td>0,972</td><td>196737</td><td>7658</td><td>NIST SRM 4325 2.79.10-11</td><td>2007</td><td>9798</td><td>691</td><td>12200</td><td>560</td><td>480</td></tr><tr><td>Bessans 05-03</td><td>4,0</td><td>2,65</td><td>0,972</td><td>198966</td><td>16150</td><td>NIST SRM 4325 2.79.10-11</td><td>2007</td><td>10066</td><td>1008</td><td>12370</td><td>1030</td><td>980</td></tr><tr><td>Bessans 05-04</td><td>4,0</td><td>2,65</td><td>0,972</td><td>213644</td><td>9443</td><td>NIST SRM 4325 2.79.10-11</td><td>2007</td><td>10794</td><td>792</td><td>13230</td><td>640</td><td>550</td></tr><tr><td>Ev-14a</td><td>2,0</td><td>2,70</td><td>0,950</td><td>192238</td><td>7163</td><td>STD11</td><td>2021</td><td>12440</td><td>670</td><td>11840</td><td>520</td><td>440</td></tr><tr><td>Ev-15</td><td>2,5</td><td>2,70</td><td>0,940</td><td>176498</td><td>7293</td><td>STD11</td><td>2021</td><td>11720</td><td>660</td><td>11210</td><td>510</td><td>430</td></tr><tr><td>Ev-6</td><td>2,0</td><td>2,70</td><td>0,993</td><td>382603</td><td>12176</td><td>STD11</td><td>2021</td><td>12860</td><td>660</td><td>11660</td><td>450</td><td>360</td></tr><tr><td>Ev-8</td><td>2,0</td><td>2,70</td><td>0,993</td><td>397595</td><td>13446</td><td>STD11</td><td>2021</td><td>13410</td><td>700</td><td>12190</td><td>510</td><td>420</td></tr></tbody></table>
<figure><img src="image_91.png" /></figure>

Figure 5.6: Photographies des blocs erratiques échantillonnés A : A Challes les eaux (CHA 20-01), B : A La Croix-de-la-Rochette (LCR-1), C : Au Pas du Roc (ROC 20-02) et D : A Bessans (Bessans 05-01).
<figure><img src="image_92.png" /></figure>

Figure 5.7: Carte de la zone d'étude du chapitre 5, montrant les âges d'exposition ¹⁰Be des blocs erratiques et glaciaires échantillonnés, ainsi que les âges d'exposition ¹⁰Be recalculés à partir de données sur blocs erratiques, glaciaires et polis glaciaires déjà publiés (Nicoud, et al., 2009; Prud'homme, et al., 2020, Couette et Ghienne, in prep.). Les diamants blancs indiquent la position des blocs erratiques et glaciaires. Les losanges blancs indiquent la localisation des polis. La carte A localise les blocs erratiques échantillonnés au Pas du Roc. La carte B localise les blocs erratiques par rapport aux moraines de Bessans. La carte C indique la localisation des polis de Bonneval-sur-Arc. Les noms d'échantillon inscrits en gras indiquent les blocs échantillonnés lors de cette étude.
### 5.2.3 Interprétation et discussion

Dans le cas des échantillons situés à Montsapey et Aussois, les âges d'exposition sur polis glaciaires (Tableau 5.2) sont systématiquement plus jeunes que les âges issus de blocs erratiques. Ces âges sur polis sont facilement soumis à un enfouissement par un sédiment ou par de la neige et donnent des âges minimum du retrait glaciaire (Prud'homme, et al., 2020). Comme l'hypothèse de l'enfouissement est difficile à mettre en évidence, nous avons choisi de garder uniquement les âges d'exposition sur blocs erratiques.

CHA 20-01, LCR-1 et CRU-1 sont des blocs erratiques isolés qui seront traités individuellement. A l'inverse, les blocs erratiques datés sur les sites du Pas du Roc et d'Aussois, sont localisés très proches les uns des autres et à des altitudes similaires, ce qui suggère que l'abandon sur chaque site a eu lieu sur un intervalle de temps très court. Les moyennes et écarts types des âges des échantillons de ces deux sites indiquent donc respectivement l'âge et l'incertitude associée de l'abandon des blocs lors du retrait glaciaire. Les abandons des blocs du Pas du Roc et d'Aussois sont donc datés respectivement à 15,26 ± 0,68 et à 14,49 ± 0,55 ka (Figure 5.8). La datation de l'unique bloc de Montsapey à 14,89 ± 1,76 ka est caractérisée par une grande incertitude et l'âge central est plus jeune que l'âge moyen de l'abandon des blocs du Pas du Roc. L'âge du bloc de Montsapey peut être réestimé entre le maximum analytique à 16,65 ka (14,89 + 1,76) et le minimum de l'âge du Pas du Roc à 14,58 ka (15,26 – 0,68). Les âges d'exposition à Montsapey à 16,65-14,58 ka, au Pas du Roc à 15,26 ± 0,68 ka et à Aussois à 14,49 ± 0,54 ka ne sont associés à aucune morphologie ou aucun dépôt attestant de réavancées ou de stabilisations glaciaires (Prud'homme, et al., 2020). Les blocs erratiques de ces trois localités témoignent ainsi d'un retrait du glacier de l'Arc entre 16,7 et 14,5 ka.

Les blocs de Bessans (Bessans 05-01, Bessans 05-02, Bessans 05-03, Bessans 05-0, Ev-14a, Ev-15) sont localisés sur un même cordon morainique et sont donc issus d'une même réavancée glaciaire. La moyenne et l'écart type des âges des blocs de Bessans, indique donc l'âge et l'incertitude associée de la moraine de Bessans à 12,58 ± 0,82 ka, correspondant au Dryas récent (Figure 5.8).

A Bonneval, les deux polis datés Ev-6 et Ev-8 donnent respectivement 11,66 ± 0,45 et 12,19 ± 0,51 ka. Compte tenu des incertitudes, les deux âges d'exposition de Ev-6 et Ev-8 recouvrent l'âge de la moraine de Bessans. Les âges des 2 polis de Bonneval et les âges d'exposition de la moraine de Bessans peuvent être considérés comme synchrones.
<figure><img src="image_93.png" /></figure>

Figure 5.8: Distributions de probabilités des âges d'exposition et de leur incertitude externe associée en gris et distributions de probabilités sommées en noir des sites du Pas du Roc (A), d'Aussois (B) et de Bessans (C). Les moyennes sont représentées par les lignes pleines rouges et les écarts types (incertitudes) par les lignes rouges pointillées. Les âges et incertitudes associées des stades glaciaires sont indiqués en haut à droite.

Après le LGM, le piémont des Alpes occidentales françaises est déglacé à ~16,5 ka, avec un front situé à Culoz (Roattino, et al., 2022). En amont de ce point, cinq âges d'exposition sont disponibles (CHA 20-01, CRU-1, LCR-1, MSAP12_B1_06 et ROC) et donnent des âges d'exposition très proches les uns des autres à 16,91 ± 1,10, 16,31 ± 1,04, 15,35 ± 0,58, 16,65-14,58 et 15,26 ± 0,68 ka respectivement et proches de l'âge de Culoz à 16,5 ± 1,1 ka. La cohérence des âges sur ~120 km entre Culoz et le Pas du Roc indique un recul rapide du glacier. La différence d'âge réelle entre ces blocs est du même ordre de grandeur que la résolution temporelle de la datation $^{10}$Be (entre 500 et 1000 ans), ce qui ne permet pas d'avoir une meilleure précision sur le taux de recul du glacier lors de cette phase (Figure 5.9). Le retrait d'un glacier peut être significativement accéléré lorsque son front est localisé dans un lac proglaciaire par circulation d'eau chauffée en surface (Kääb & Haeberli, 2001). Entre Culoz et l'entrée de la Maurienne le retrait rapide du glacier delphino-savoisien a donc pu être favorisé par son vêlage dans les paléolacs proglaciaires du Bourget et de l'Isère.

### 5.2.4 Conclusion

L'ensemble des blocs erratiques et glaciaires du piémont et ceux localisés en fond de vallée ou sur des verrous (CHA, ROC et Bessans) permettent de positionner des fronts glaciaires. Les autres blocs erratiques (CRU, LCR, MSAP et AUSS) localisés sur des versants ne permettent pas d'indiquer la position d'un front glaciaire au sens strict. En effet, pour ces blocs, le front glaciaire est localisé à quelques kilomètres en aval de la position du site daté (Figure 5.10). Ainsi, les âges d'exposition issus de la littérature recalculés et ceux inédits présentés dans cette étude (Figure 5.9) permettent de reconstituer la position et l'âge de 7 positions glaciaires entre le piémont des Alpes occidentales
françaises et la partie amont de la Maurienne (Figure 5.10). Plus précisément, 2 fronts du lobe lyonnais indiquent la position du glacier au LGM à 24-21 ka sur le domaine morainique intermédiaire et à 19 ka au niveau du domaine morainique est. Un front du glacier delphino-savoisien peut être positionné à Culoz à environ 16,5 ka et enfin 4 positions glaciaires indiquent le retrait du glacier de l'Arc entre environ 16,7-14,6 et 12,5 ka. Ces résultats montrent un retrait rapide entre Culoz et le site du Pas du Roc en Maurienne à environ 16,5 ka.
<figure><img src="image_94.png" /><figcaption>Figure 5.9: Classification des âges d'exposition du piémont des Alpes occidentales françaises jusqu'en Maurienne. Le tracé du profil topographique de la Figure 5.10 est indiqué en rouge.</figcaption></figure>
<figure><img src="image_95.png" /></figure>

Figure 5.10 : Coupe passant par le piémont des Alpes occidentales françaises jusqu'à la source de l'Arc, montrant la position des fronts glaciaires datés par des publications précédentes (Nicoud, et al., 2009; Prud'homme, et al., 2020; Roattino, et al., 2022) et par les âges d'exposition présentés dans cette étude. La position de la coupe est présentée dans la Figure 5.9. Les blocs erratiques et glaciaires sont indiqués par les diamants et les polis par les losanges. Les âges d'exposition sont classés par couleur dont l'échelle est donnée par la Figure 5.9. Les polygones gris foncé indiquent la position des dépôts lacustres post LGM des paléolacs de la Balme (pl. La Balme), du Bourget (pl. du Bourget) et de l'Isère (pl. de l'Isère). La géométrie et la profondeur du paléolac du Bourget a été reconnue par géophysique (Finckh, et al., 1984). La profondeur minimum des paléolacs de la Balme et de l'Isère est également indiquée (Nicoud, et al., 1987).
## 5.3 SÉDIMENTOLOGIE ET DATATIONS PAR LUMINESCENCE

Cette partie se consacre à l'étude sédimentologique détaillée d'affleurements proglaciaires localisés en partie amont de la vallée de l'Arc. Cette étude sera réalisée sur deux affleurements (Verney et Lanslebourg) dont les analyses réalisées dans des études antérieures, ont soulevé des questionnements sur leurs mises en place. En effet, ces dépôts montrent parfois des structures sédimentaires et des déformations particulières dont l'origine faisant intervenir des processus glaciaires, gravitaires et même tectoniques (Hugonin, 1988; Marnezy, 1999; Champagnac, et al., 2006), n'a pas été clairement identifiée.

Cette analyse sédimentologique va permettre de définir des fluctuations tardiglaciaires du glacier de l'Arc n'ayant pas été enregistrées par des morphologies morainiques. Une fois caractérisés, ces dépôts seront datés par luminescence afin de compléter le cadre chronologique de la déglaciation du glacier de l'Arc.

### 5.3.1 Matériels et méthodes

#### 5.3.1.1 Affleurements étudiés

<figure><img src="image_96.png" /></figure>

Figure 5.11 : Carte de localisation des affleurements du Verney, de Lanslebourg et de Bessans, des principaux verrous rocheux, de l'écroulement de la Madeleine et des moraines frontales de Bessans. La localisation de la Figure 5.13, de la Figure 5.14 et de la Figure 5.15 est indiquée.
<figure><img src="image_97.png" /></figure>

Figure 5.12 : Coupe schématique résumant la géologie des bassins du Verney, de Lanslebourg et de Bessans (Letourneur, et al., 1983; Hugonin, 1988; Debelmas, 1989; Fudral, et al., 1994).

Trois affleurements localisés respectivement dans les bassins du Verney, de Lanslebourg et de Bessans (Figure 5.11 et Figure 5.12) ont été étudiés afin de reconstituer les environnements de dépôts et pour prélever des échantillons dans le but de réaliser des datations par luminescence.

L'affleurement du Verney est localisé dans un bassin situé à environ 1200 m d'altitude et plus précisément, dans une ancienne carrière au nord du village du Verney, au pied du versant nord de la vallée de l'Arc. Cet affleurement est localisé dans une terrasse dont la surface est constituée d'alluvions anciennes (Debelmas, 1989) reposant sur des dépôts deltaïques (Hugonin, 1988). En aval de l'affleurement, on observe un cône alluvial et un grand dépôt gravitaire (glissement) ainsi que des gorges formant un verrou (Figure 5.11 et Figure 5.12) (Debelmas, 1989). Le fonctionnement de ce glissement associé au verrou est à l'origine du barrage de la vallée, ce qui a permis le développement des dépôts lacustres deltaïques. Ces dépôts lacustres sont recouverts d'alluvion (Figure 5.12). Ces dépôts lacustres et les alluvions ont plus tard été incisés par l'Arc et sont actuellement retrouvés sur des zones préservées de cette érosion fluviatile, formant des replats (Figure 5.13) (Hugonin, 1988). L'affleurement est localisé dans l'un de ces replats. Deux échantillons pour des datations par luminescence ont été prélevés (BRA-01, BRA-02, Figure 5.13).
<figure><img src="image_98.png" /><figcaption>Figure 5.13 : Carte présentant le contexte géologique et géomorphologique (Hugonin, 1988; Debelmas, 1989) de l'affleurement du Verney, redessiné avec le RGE Alti 1m de l'IGN à partir de la carte géologique de la France harmonisée (BRGM, 2005). L'emprise de la carte est donnée dans la Figure 5.11.</figcaption></figure>

Les affleurements étudiés à Lanslebourg sont localisés dans le bassin éponyme, situé à environ 1370 m d'altitude. Les affleurements se trouvent à l'ouest du village de Lanslebourg, au pied du versant nord de la vallée de l'Arc, au niveau d'une surface incisée par l'Arc. La partie basale de ces affleurements a initialement été attribuée à un dépôt deltaïque lié au barrage de la vallée par un verrou localisé en aval. Le glacier du Doron, s'écoulant dans la vallée de la rivière éponyme (Figure 5.11 et Figure 5.12), a également été désigné pour expliquer le barrage de la vallée. Le front du glacier du Doron devait arriver jusqu'au village de Sollières (Figure 5.11) (Hugonin, 1988). Au-dessus du dépôt deltaïque, des alluvions anciennes ont été décrites et attribuées à des environnements glaciolacustres ou fluvioglaciaires (Fudral, et al., 1994) ou encore à un dépôt de cône alluvial (Figure 5.12) (Hugonin, 1988). La zone contient des cônes alluviaux alimentés par des rivières venant du versant. Sur le versant nord, sont plaqués localement des dépôts de till. Un échantillon pour une datation par luminescence a été prélevé sur l'affleurement (LSB-01, Figure 5.14).
<figure><img src="image_99.png" /></figure>

Figure 5.14: Carte présentant le contexte géomorphologique et géologique (Fudral, et al., 1994; Hugonin, 1988) de l'affleurement de Lanslebourg, redessiné avec le RGE Alti 1m de l'IGN à partir de la carte géologique de la France harmonisée (BRGM, 2005). L'emprise de la carte est donnée dans la Figure 5.11.

L'affleurement de Bessans est localisé dans un bassin sédimentaire situé à environ 1710 m d'altitude et montrant des dépôts lacustres argileux et deltaïques issus du barrement de la vallée de l'Arc (Couture, et al., 1997) par un écroulement situé en aval, au niveau du col de la Madeleine (Letourneur, et al., 1983; Hugonin, 1988). Ce bassin est caractérisé par la présence de deux moraines frontales (Figure 5.11 et Figure 5.12). Dans ce secteur, on observe de grands cônes localisés au débouché des vallées où s'écoule le Ribon et l'Avérole, deux affluents de l'Arc (Figure 5.15). Des analyses de forages indiquent que les dépôts fluviaux reposent sur des dépôts lacustres. Ces dépôts fluviaux ont ainsi été interprétés comme des dépôts de comblement d'un lac. De la même manière, des analyses de forages et d'affleurements ont montré que les deux cônes alluviaux du Ribon et de l'Avérole reposent sur des dépôts deltaïques (Figure 5.12) (Letourneur, et al., 1983; Fudral, et al., 1994; Hugonin, 1988). L'affleurement est localisé sur le cône du Ribon, dans une zone incisée par la rivière (Figure 5.15). Ce cône est issu d'un dépôt deltaïque bien caractérisé recouvert d'un cône alluvial (Hugonin, 1988). Un échantillon de sable (BES-01) a été prélevé dans ce dépôt deltaïque pour obtenir une datation par luminescence (Figure 5.15).
<figure><img src="image_100.png" /></figure>

Figure 5.15: Carte présentant le contexte géomorphologique et géologique (Letourneur, et al., 1983; Fudral, et al., 1994; Hugonin, 1988) de l'affleurement de Bessans, redessiné avec le RGE Alti 1m de l'IGN à partir de la carte géologique de la France harmonisée (BRGM, 2005). L'emprise de la carte est donnée dans la Figure 5.11.

## 5.3.2 Résultats

### 5.3.2.1 Sédimentologie

#### 5.3.2.1.1 Le Verney

L'affleurement mesure environ 330 m de longueur pour une épaisseur maximum d'environ 20 mètres. En partie nord-est, les dépôts reposent directement sur le substratum. Les dépôts sont globalement granocroissants, avec à la base des unités sableuses et au sommet, des unités de graviers, galets et blocs. Des déformations sont visibles dans la plupart de ces unités. Le sommet de l'affleurement forme une surface plane visible dans le paysage (Figure 5.16). Cet affleurement a été interprété comme un delta et les déformations que l'on y retrouve, à des effondrements liés à la dissolution du substratum gypseux, à de la cryoturbation ou encore à de la déstabilisation gravitaire pendant le dépôt (Hugonin, 1988). L'histoire de cet affleurement n'est donc pas clairement établie et justifie une étude approfondie.
<figure><img src="image_101.png" /></figure>

<figure><img src="image_102.png" /></figure>

Figure 5.16: A : Photographie panorama et B : Photographie habillée montrant le substratum et les faciès Sb, Gm, St, Gp et Gl du Verney. La forme lenticulaire des faciès Gm et St est ici bien visible. Les clasts flottants de grande taille de Gp ont été représentés, à noter que ce faciès est dans cette partie de l'affleurement très déformé.
## Description et interprétation des faciès :

La succession sédimentologique et stratigraphique dans la carrière Verney est composée de 5 faciès sédimentaires, chacun étant défini sur la base de sa granulométrie, de ses structures sédimentaires et de ses géométries, ces faciès seront interprétés en processus de dépôts. Ils ont été regroupés en trois associations de faciès (AF) représentant des environnements de dépôts.

**Sable à stratifications sinusoïdales et « backsets » (Sb).** Ce faciès est localisé directement sur le substratum carbonaté (Figure 5.17) et est constitué d'éléments allant du sable fin jusqu'aux granules bien triés. Les structures sédimentaires sont principalement représentées par une stratification sinusoïdale de longueurs d'ondes comprises entre 1 et 2 m pour une amplitude de 5 à 30 cm. On observe également des « backsets » peu pentés. Ces structures remplissent des structures en auges de 1 à 2 m de long, quelques contacts basaux sont érosifs (Figure 5.17). Ce faciès contient des clastes de grandes tailles allant des graviers jusqu'aux galets. Ces clastes suivent la stratification ou sont isolés (Figure 5.19D).

**Interprétation.** La stratification sinusoïdale provient de l'aggradation d'antidunes stationnaires sous un courant supercritique permettant la préservation des faces sous le courant et des faces face au courant de la morphologie (Cheel, 1990; Brennand, 1994; Russell & Arnott, 2003; Ito & Saito, 2006; Duller, et al., 2008; Ito, 2010). Les contacts érosifs indiquent que l'aggradation n'est pas continue (Duller, et al., 2008). Les « backsets » montrent un dépôt sur le côté face au courant d'antidunes migrant à contre-courant (Kennedy, 1963; Alexander, et al., 2001; Cartigny, et al., 2014). Les lits grossiers reflètent un dépôt lors de conditions de fortes énergies (Lee, et al., 2015). Les galets striés isolés sont issus de la fonte de glace basale riche en débris.
<figure><img src="image_103.png" /><figcaption>Figure 5.17: A : Photographie et B : Photographie habillée montrant le substratum et les faciès Sb, Gm, St et Gp du Verney. A noter que les structures sédimentaires de Gp sont ici très peu visibles.</figcaption></figure>

**Sable fin et moyen à lamination croisée (St).** Ce faciès forme des unités lenticulaires de 20 cm à 3 m d'épaisseur, constituées de sable fin à moyen structuré en lamination entrecroisée (Figure 5.19A). Certains contacts basaux sont érosifs. Ce faciès montre des rides grimpantes de type A de 8 à 10 cm de longueur d'onde, pour 2 à 3 cm d'amplitude. On observe également dans ce faciès, des laminations
sinusoïdales correspondant à des rides grimpantes de type B entre 5 et 8 cm de longueur d'onde pour 1 à 2 cm de haut (Figure 5.19A) définis par Jopling et Walter (1968). Ce faciès est interstratifié avec des lits de 5 à 15 cm d'épaisseur à granoclassement normal, montrant à la base des silts ou des sables fins à lamination parallèle et se terminant vers le haut par des niveaux silto-argileux massifs ou à lamination parallèle et pouvant montrer des convolutes (Figure 5.19B). Des galets isolés sont visibles dans ce faciès.

**Interprétation.** La lamination entrecroisée à petite échelle est issue d'un dépôt par la migration de rides 3D sous un courant subcritique. La présence de rides grimpantes implique une sédimentation à partir d'un courant perdant sa capacité de transport. Ainsi les rides grimpantes de type A résultent d'un faible taux d'aggradation par rapport au taux de migration (Jopling & Walker, 1968; Ashley, et al., 1982). Inversement les rides grimpantes de type B sont produites par un taux d'aggradation important par rapport au taux de migration (Jopling & Walker, 1968; Ashley, et al., 1982; Allen, 1973; Fielding, 2006). Ces structures sont issues d'un courant turbiditique de faible densité (Ashley, 1995; Mulder & Alexander, 2001). L'épaisseur parfois importante de ces unités de rides et de rides grimpantes implique des courants turbiditiques relativement longs dans le temps (Mulder & Alexander, 2001). Les surfaces d'érosions correspondent à des chenaux ayant accueilli les courants turbiditiques. Les lits de matériaux fins granodécroissants sont issus d'un dépôt lors de la dissipation de l'énergie d'un écoulement turbiditique de faible densité (Mulder & Alexander, 2001). Les directions de paléocourant enregistrées indiquent un transport depuis le nord. Les galets isolés sont issus de chutes de débris à travers une colonne d'eau résultant de la fonte d'iceberg.

**Graviers et galets massif (Gm).** Ce faciès forme des lentilles de 10 à 100 cm d'épaisseur pour environ 10 m de large, remplies de graviers et galets polygéniques à matrice composée de sable fin. Le granoclassement est normal et faiblement marqué (Figure 5.19C).

**Interprétation.** Ce faciès est issu d'écoulements hyperconcentrés (Mulder & Alexander, 2001).

**Graviers et galets à stratification plane parallèle (Gp).** Ce faciès est constitué de graviers et galets polygéniques montrant une stratification plane parallèle avec un pendage d'environ 15° vers l'aval et incluant des galets imbriqués (Figure 5.18 et Figure 5.19F). On observe des lits massifs avec matrice, qui montrent un granoclassement inverse ou normal (Figure 5.19F) et contenant des clasts flottants de grandes tailles (50-100 cm long) alignés parallèlement à la stratification (Figure 5.16). Les lits à granoclassement normal se terminent fréquemment par des niveaux sableux de quelques dizaines de centimètres d'épaisseur (Figure 5.19F). Localement, des structures en auges sont remplies par des graviers et granules suivant une stratification croisée (Figure 5.19E). Dans ces remplissages, des lits sans matrice sont présents.

**Interprétation.** La stratification plane parallèle oblique et les lits à granoclassement normal à inverse suggèrent un transport sédimentaire contrôlé par la gravité sur la pente d'un delta (Postma & Cruickshank, 1988; Nemec, 1990). Les lits massifs ou à granoclassement normal ou inverse à matrice et à clastes flottants résultent d'un dépôt après un transport par des coulées de débris (Shanmugam, 2000). Les niveaux sableux indiquent un dépôt lors de la dissipation de l'énergie du courant par des écoulements concentrés ou turbiditiques (Mulder & Alexander, 2001). Le remplissage des structures en auges par une stratification croisée, indique un comblement de dépressions par la migration de dune 2D (Allen, 1982). La présence de lits sans matrices côtoyant des niveaux à matrice indique une séparation des sables et des graviers par des processus hydrauliques (Carling, 1990), probablement liée à un courant turbulent de forte énergie (Allen, 1982; Carling & Glaister, 1987).

**Graviers et galets à lentilles sableuses (Gl).** Ce faciès est constitué par des graviers et galets polygéniques, mal triés, parfois imbriqués et à matrice composée d'éléments allant de la taille d'un sable grossier à des granules (Figure 5.19G). Les lits sont massifs ou montrent un granoclassement normal et se terminent parfois par des niveaux de sable grossier (Figure 5.19H). Les niveaux de sable grossier sont structurés suivant une lamination croisée issue de rides grimpantes. Les limites basales sont subhorizontales et érosives (Figure 5.16 et Figure 5.18).

**Interprétation.** Ce faciès est issu d'un dépôt par des écoulements hyperconcentrés (Mulder & Alexander, 2001). Les niveaux sableux reflètent un dépôt pendant la phase de dissipation de l'énergie
du courant (Miall, 1983). Les rides grimpantes sont issues d'une forte sédimentation lors de la perte de la capacité de transport d'un courant turbiditique (Ashley, et al., 1982; Jopling & Walker, 1968; Mulder & Alexander, 2001).

<figure><img src="image_104.png" /><figcaption>Figure 5.18: A : Photographie panorama et B : Photographie habillée montrant les faciès Gp et Gl du Verney. La stratification parallèle et pentée de Gp est ici bien visible.</figcaption></figure>
<table><caption>Tableau 5.3 : Classification des faciès sédimentaires et associations de faciès (AF) du Verney</caption><thead><tr><th>AF</th><th>Faciès</th><th>Description</th><th>Interprétation</th></tr></thead><tbody><tr><td rowspan="2">AF1</td><td>Sb, Sable à stratification sinusoïdale et "backsets"</td><td>Sables, granules à stratification sinusoïdale et à "backsets" remplissant des auges. Clasts striés de grandes tailles</td><td>Aggradation d'antidunes stationnaires ou migrant à contre-courant en milieu sous-glaciaire, proche de la marge du glacier</td></tr><tr><td>Gm, graviers et galets massif</td><td>Lentilles de graviers et galets à matrice sableuse, granoclassement normal ou faiblement marqué</td><td>Faciès issus d'écoulements hyperconcentrés</td></tr><tr><td rowspan="2">AF2</td><td>St, Sable fin et moyen à lamination croisée</td><td>Sables fins à moyen structurés en lamination croisée, rides grimpantes de types A et B. Présence de lits silto-argileux à granoclassement normal, terminé vers le haut par des convolutes. "Dropstones"</td><td>Migration de ride de courant sous un courant turbiditique perdant sa capacité de transport en milieu lacustre</td></tr><tr><td>Gp, Graviers et galets à stratification plane parallèle</td><td>Blocs, graviers et galets parfois imbriqués polygéniques à matrice sableuse et à stratification plane parallèle, penté vers l'aval à 15°. Lits massifs ou à granoclassement inverse ou normal se terminant par un niveau sableux. Présence d'auges remplis de graviers et granules parfois sans matrice à stratification croisé</td><td>Dépôts de foreset après un transport de coulées de débris, écoulements hyperconcentrés, turbiditiques et par migration de dune 2D</td></tr><tr><td>AF3</td><td>Gl, Graviers et galets à lentilles sableuses</td><td>Graviers et galets polygéniques mal triés à matrice de sable grossier et granules. Lits massifs ou à granoclassement normal se terminant par un niveau sableux. Limites basales subhorizontales et érosives</td><td>Dépôt après un transport par écoulements hyperconcentrés puis pendant une phase de dissipation de l'énergie du courant</td></tr></tbody></table>
<figure><img src="image_105.png" /></figure>

Figure 5.19: Photographies des faciès retrouvés à l'affleurement du Verney. A : Photographie des rides grimpantes de type B du faciès St. B : Photographie des lits de matériaux fins granodécroissants du faciès St. C : Photographie du faciès Gm. D : Photographie du substratum rocheux recouvert du faciès Sb dont la base contient des galets striés. Le faciès Sb est ici recouvert par le faciès Gm. E : Photographie du faciès Gp montrant des niveaux à stratification croisée remplissant une structure en auge. F : Photographie du faciès Gp montrant des lits plans parallèles massifs ou à granoclassement normal ou inverse. Certains niveaux granodécroissants se terminent par des niveaux sableux. G : Photographie du faciès Gl montrant un sédiment mal trié à matrice sableuse. H : Photographie du faciès Gl montrant des niveaux massifs et subhorizontaux granodécroissants parfois terminés par des lentilles sableuses. Les triangles symbolisent le granoclassement.
Déformations :

**Description.** Les déformations affectent les faciès St et Gp. St est injecté dans le faciès Gp (Figure 5.16 et Figure 5.20B), surtout au nord-est de l'affleurement où il peut former des intrusions de 5 à 10 m de haut. Les intrusions contiennent également des figures d'échappement de fluides millimétriques à centimétriques (Figure 5.20B). St est parfois injecté vers le sud-ouest dans Gp sous forme de coins de compression d'environ 1 m de longueur pour environ 30 cm d'épaisseur (Figure 5.20C). Ce faciès montre des failles inverses associées à des antiformes et synformes. Le pendage des failles à environ 45° est orienté vers le nord-est (Figure 5.20B). Au nord-est de l'affleurement, le faciès Gp montre des plis déjetés de vergence orientée vers le sud-ouest et localisés au-dessus des intrusions de St (Figure 5.16). Enfin, en partie aval, le faciès Gp montre des failles inverses à pendage vers le nord-est (Figure 5.20D).

**Interprétations.** Les coins se produisent lors d'un cisaillement simple, permettant à un sédiment de s'injecter dans un autre dépôt sus-jacent. Ce type de structures peuvent être le résultat d'une avancée glaciaire créant des conditions de cisaillement. Le pendage du coin vers le nord-est indique une avancée glaciaire en direction de l'aval (Rijsdijk, et al., 2010). Les synformes et antiformes associées aux failles inverses retrouvés dans le faciès Gp indiquent une déformation ductile liée à une contrainte en compression (Michaël, et al., 2009). Les figures d'échappement de fluides indiquent une expulsion d'eau contenue dans la porosité d'un sédiment. En milieu sous-glaciaire, les échappements de fluides sont le résultat d'une mise sous pression de l'eau pendant une avancée glaciaire ou une surcharge liée à la masse de la glace (Denis, et al., 2009). Les intrusions du faciès Sc dans le faciès Gp sont liées à une différence de densité et de viscosité entre les deux sédiments (Hindmarsh & Rijsdijk, 2000). La surpression de fluides issues d'une surcharge par de la glace génère une dilatation et une diminution de la densité du sédiment sous-jacent (Boulton & Caban, 1995). Dans le cas d'un sédiment sableux chargé en eau et de faible viscosité, cette diminution de densité peut générer une intrusion sableuse dans un sédiment sus-jacent de densité supérieure (Owen, 2003; Rijsdijk, et al., 2010). Les plis déjetés résultent d'une déformation ductile d'un sédiment chargé d'eau sous pression. La localisation des plis déjetés au-dessus des intrusions sableuses indiquent que le mouvement vertical ascendant à l'origine de ces intrusions est responsable de l'apparition des plis. La vergence des plis déjetés orientées vers le sud-ouest implique l'application d'une contrainte (McCarroll & Rijsdijk, 2003) venant du nord-est.