# New evidence of mid-Pleistocene glaciations in the French Jura Mountains using debris flow deposits and U/Th dating of speleothems (Verneau karst network)

Margot Vivier, Stéphane Jaillet, Eglantine Husson, Edwige Pons-Branchu, Jean-Baptiste Charlier

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Margot Vivier, Stéphane Jaillet, Eglantine Husson, Edwige Pons-Branchu, Jean-Baptiste Charlier. New evidence of mid-Pleistocene glaciations in the French Jura Mountains using debris flow deposits and U/Th dating of speleothems (Verneau karst network). International Journal of Speleology, 2025, 54 (2), 10.5038/1827-806x.ijs2558. hal-05372974

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# New evidence of mid-Pleistocene glaciations in the French Jura Mountains using debris-flow deposits and U/Th-dating of speleothems (Verneau karst network)

Margot Vivier ID <fnref n="1" />,<fnref n="2" />,<fnref n="3" />*, Stéphane Jaillet ID <fnref n="3" />, Eglantine Husson ID <fnref n="4" />, Edwige Pons-Branchu ID <fnref n="5" />, and Jean-Baptiste Charlier ID <fnref n="1" />,<fnref n="2" />

**Abstract:** We redefine the extent of former glaciers in the northern Jura Mountains through the study of endo-karst deposits. Sedimentary deposits (pebbly debris flows, laminated clays, and allochthonous limestone blocks) and chemical deposits (speleothems) were identified and studied using three complementary approaches: (i) detailed geomorphological mapping, (ii) petrographic analysis, and (iii) U/Th-dating. In addition, a surface geomorphological analysis made it possible to calculate the minimum erosion time required for the formation of the sinkhole linked to the underground study site. The study site is the Dolois Chamber, an underground site upstream of the Verneau karst network (Doubs Department, France). The combined results aided in reconstructing the chronology of the various events affecting the study site, including 1) formation of the Baume-des-Crêtes paleo-sinkhole a few hundred thousand years (during Marine Isotope Stage (MIS) 8); 2) invasion by a pebbly debris flow, probably linked to a mid-Pleistocene glacial periods (MIS 8 or 6); 3) hydrological dysfunction, including significant flooding and deposition of laminated clays; 4) capture of the paleo-sinkhole by the Biefs-Bousset active sinkhole approximately 10 kyr (MIS 2); and 5) a recent return to water level variations of approximately 10 m during flooding. Furthermore, our results show that the extent of the ice cap in the Jura Mountains during the mid-Pleistocene (MIS 8-6) probably did not reach beyond the topographic high of the Salinois Fault Zone at the plateau margin.

**Keywords:** Karst deposit, speleothem dating, glaciation, mid-Pleistocene, Jura Mountains

Received 7 May 2025; Revised 13 October 2025; Accepted 21 October 2025

**Citation:** Vivier, M., Jaillet, S., Husson, E., Pons-Branchu, E., Charlier, J.-B., 2025. New evidence of mid-Pleistocene glaciations in the French Jura Mountains using debris-flow deposits and U/Th-dating of speleothems (Verneau karst network). International Journal of Speleology, 54(2), ijs2558. https://doi.org/10.5038/1827-806X.ijs2558

## INTRODUCTION

The question of the extent of former ice ages in the Alps was raised as early as the 19th century following the discovery of erratic boulders far away from ice-covered massifs (Venetz, 1821; Geikie, 1910). Since then, the amount (Agassiz, 1840; Bourdier & Lautridou, 1974; Campy, 1982; Coutterand, 2010; Kamleitner et al., 2022) and the diversity (Laverdière et al., 1985; Pons et al., 1989; Masson-Delmotte & Chappellaz, 2002; Spötl & Mangini, 2007) of field data and improved numerical models (Jouvet et al., 2023; Leger et al., 2025) have continued to add to our knowledge of the former glacial extent.

Over the last few decades, specific forms (sinkholes linked to glacier (Bini et al., 1998)) and specific deposits (fluvio-glacial deposits, laminated clays, etc.) in underground karst have made it possible to identify and date the extent of former ice ages. Indeed, during the advance glaciers, the bedrock underwent severe erosion by "abrasion, crushing, and stripping" (Boulton, 1974; Laverdière et al., 1985). Then, during glacier retreat, i.e. in paraglacial phase (Ballantyne, 2002; Mercier, 2008), sediments are transported and deposited with no sorting. This led to glacial sedimentary deposits (till) composed of loose material and boulders up to several tens of centimetres in size (Campy, 1982). Some access to underground (sinkhole or cave entrance) nearby can trapped these

<footnote n="1">¹BRGM, University of Montpellier, 1039 rue de Pinville, 34000 Montpellier, France</footnote>

<footnote n="2">²G-EAU, University of Montpellier, AgroParisTech, BRGM, CIRAD, INRAE, Institut Agro, IRD, Montpellier, France</footnote>

<footnote n="3">³EDYTEM, UMR 5204 CNRS, University of Savoie Mont Blanc, Pôle Montagne, 5 boulevard de la mer Caspienne, 73390 Le Bourget du Lac, France</footnote>

<footnote n="4">⁴BRGM, F-45060 Orléans, France</footnote>

<footnote n="5">⁵Laboratoire des Sciences du Climat et de l'Environnement, LSCE-IPSL (CEA-CNRS-UVSQ), University of Paris-Saclay, 91191 Gif-sur-Yvette, France</footnote>
sedimentary deposits. Localisation and dating of them
make it possible to constrain the extent of past ice
ages.

Endo-karst forms and deposits resulting from
glacial activity have been studied successfully in
Europe: in Italy (Bini et al., 1998), Switzerland
(Jeannin, 1991; Häuselmann, 2002), Austria (Spötl
& Mangini, 2007), England (Farrant et al., 2014),
Norway (Lauritzen & Skoglund, 2013; Skoglund et
al., 2024), the French Pyrenees (Sorriaux et al., 2016)
and Spain (Ballesteros et al., 2017; Bartolomé et al.,
2021). They have been also studied in Chile (Jaillet et
al., 2008) and New Zealand (Williams, 1996). Dating
methods vary in each case (Hajna et al., 2010; Jaillet
et al., 2025). In the Vallier cave (Vercors, France),
paleomagnetically-dated deposits revealed an ancient
glacial phase (Audra & Rochette, 1993). In the Bauges
(Savoie, France), U/Th-dating of speleothems made
it possible to constrain the age of associated clastic
deposits (Jaillet et al., 2018, 2023). In Spain, in the
Picos de Europa, optical stimulated luminescence
(OSL) and faunal dating have provided additional age
constraints (Ballesteros et al., 2017).

Similar studies have been used in the Jura
Mountains (Fig. 1) to constrain the glacial extent (MIS
4 and 2 - late-Pleistocene - and MIS 8 and 6 - mid-
Pleistocene) in the southern area, and also on the

Swiss side. In the southern Jura (Bugey), the "Sous les Sangles" cave has been used to constrain glacier flow in the Cluse des Hôpitaux (Sbai et al., 1995; Lignier & Desmet, 2002; Jaillet & Pons-Branchu, 2022). In the Swiss Jura, work in progress uses terrestrial cosmogenic nuclide (TCN) dating to constrain ancient glacial advances (Mathes et al., 2025; Val et al., 2025). Note the unique case of a cave entirely excavated in glacial till (Durlet et al., 2024). But not many studies of this type have been conducted in the northern Jura, where the extent of ice during pre-MIS 2 glacial periods remains poorly understood.

In this article, we propose a new framework for the
spatio-temporal extent of glaciers at a topographic
high in the northern Jura (Salinois Fault Zone) that
separates the Ornans and Levier plateaus, using
data collected in the Dolois Chamber in the Verneau
karstic network (Doubs Department, Eastern France).
According to current knowledge, this site is located
at the geographical limit of the glacial extent of the
mid-Pleistocene glaciers (MIS 8 and 6; Fig. 1). The
first section of this article documents the sedimentary
deposits (clastic and chemical) in the Dolois
Chamber, including U-Th-dating of speleothems. The
identification of a pebbly debris flow and chronological
constraints made it possible to reexamine the model
of glacial spatio-temporal extent at the site.

<figure><img src="image_3.png" /><figcaption>Fig. 1. Unity and structural context the Jura Mountains between the Alpine chain to the east and the Bresse Graben to the west (Sommaruga, 1997). The ice extent in the Alps and the Jura Mountains during the mid-Pleistocene (MIS 8-6, External Moraine Complex, EMC) and the late Pleistocene (MIS 4-2, Internal Moraine Complex, IMC) is indicated (Campy, 1982). The study area marked by the red rectangle is located at the boundary between the two ice domains of the mid-Pleistocene (MIS 8-6, EMC with or without crystalline rocks from the Alps) (modified from Bichet & Campy, 2008).</figcaption></figure>

SITE AND METHODS

The Jura Mountains and their Pleistocene ice cap

The Jura Mountains are a sub-alpine mountain
range in Western Europe on the border between
France and Switzerland, a folded foreland of the Alps.
The plateaus and overlapping anticlines (fault zone)
of the crescent-shaped Jura Mountains extend over

360 km, following the curve of the Alpine arc. The
Mara Mountains are Jurassic-age limestone with
marly intercalations (Fig. 1).

As early as 1909, Penck & Brückner, identified and
relatively dated four moraine complexes in the foothills
of the Alps. Subsequent studies focusing on the Jura
Mountains (Bourdier, 1958; Campy, 1982) described
the origin and extent of glaciers that deposited two
moraine complexes. The External Moraine Complex (EMC) is associated with MIS 6 and 8, also known as the mid-Pleistocene glacial periods (Campy, 1986). The Internal Moraine Complex (IMC), located 20-30 km further east, is associated with MIS 2 and 4 (late Pleistocene) (Fig. 1). The hypothesis of an ice cap on the Jura Mountains independent of that of the Alps was initially contested, but later won unanimous support on the basis of new evidence (Campy, 1982). Moreover, it is now recognized that the Jura Mountains ice cap was a warm-based glacier. Studies of Pleistocene Scandinavian ice cap, which was also warm-based (Kleman et al., 2008), have shown that most glacial deposits are concentrated on the periphery of the ice cap. Glacial sedimentary deposits were not transported long distances from the edge of the ice cap.

In the Salinois Fault Zone area, the abundance and preservation of glacial deposits differ greatly between valleys and plateaus. The Levier plateau, south of the Salinois Fault Zone, records a few lenses of mid-Pleistocene glacial deposits (labelled Gx) on the 1:50,000 Pontarlier and Salins-les-Bains geological maps (Glangeaud et al., 1960; Caire, 1967). These deposits are (1) pebbles and gravels of crystalline rocks, with iron enrichment through weathering processes and sometimes indurated, and (2) erratic boulders of Alpine origin. The authors of the geologic maps point out that these deposits are often reworked and barely recognizable. These deposits have been attributed to the mid-Pleistocene glacial periods due to their weathering and their generally poor state of preservation. Campy (1982) used geological surveys specifying that "they [the deposits] are remnants

that tell a varied and complex story to reconstruct" (Campy, 1982). With current data, it is difficult to define the ice extent in the northern Jura Mountains, particularly regarding ice thickness and exact spatio-temporal extent. The Jura Mountains are home to karst caves that may have recorded such glacial activity (Fig. 1). In fact, the carbonate rocks, tectonic deformation, and the high relief of the Jura Mountains provide conditions that have been highly suitable for karstification since the end of the Miocene (Bienfait, 1991). A large number of underground networks are known in this region (G.I.P.E.K., 1988, 1991, 1996, 2004, 2012). Among these, the Verneau cave is one of the major underground networks in the Jura Mountains (Fig. 2). However, despite the relatively exhaustive number of studies conducted (Chorvot, 1984; Aucant et al., 1985; Villegas et al., 2017), none has provided a chronological framework of past glacier influence on these caves. Sedimentary deposits in the upstream part of the network, in the Baume-des-Crêtes collapse pit, in the Dolois Chamber, enable us to specify the glacial extent of the mid-Pleistocene ice cap at the topographic high of the Salinois Fault Zone.

## The Verneau underground network

The Verneau underground network, explored since the 1960s, includes 32 km of underground conduits spread over an area 6.9 km long and 390 m in vertical extent (G.I.P.E.K., 1996). The Verneau has the distinction of being almost completely penetrable and topographically recognized, from the upstream end to the Verneau spring (Chorvot, 1984; Aucant et al., 1985) (Fig. 2).

<figure><img src="image_4.png" /><figcaption>Fig. 2. Map of the underground Verneau network and its geomorphological context. To the south, the Salinois Fault Zone (864 m NGF for the summit ridge) forms a barrier which may have blocked the Jura Mountain ice cap of the mid-Pleistocene (MIS 8-6). Below is a north-south geological cross-section through the Salinois Fault Zone showing the location of the Dolois Chamber within the folded units.</figcaption></figure>
The Verneau network lies on the northern flank of the Salinois Fault Zone, a thrusted anticline oriented N065. This large folded tectonic feature forms a topographic high that brings Liassic marls into surface contact with Upper Jurassic limestones (Tissot & Tresse, 1978). The Salinois Fault Zone topographic high separates the Levier plateau to the south from the Ornans plateau to the north. The Verneau underground network is controlled by this structural framework (Foltete, 1985) with an overall network orientation of N060, similar to that of the Salinois Fault Zone (Fig. 2) (Vivier, 2021; Vivier et al., 2022). With these marl-limestone contacts in a folded and faulted environment, the Verneau karst system shows a binary hydrological behaviour. Its catchment area covers 14 km², with localized infiltration through five sinkholes in the runoff zone on the Liassic marl watersheds to the south and diffuse infiltration into limestone to the north (Fig. 2).

Although the Verneau underground network was not covered by ice during the last glacial period (MIS 4 and 2, Fig. 1) (Campy, 1982), it appears that during MIS 8 and 6 this underground network may have been covered by ice originating partly from Alpine overflows (from Switzerland) and partly from a local Jura Mountains ice cap (Campy, 1982). The dynamics of the advance and retreat of these allochthonous and autochthonous ice caps are not known. It is likely that these ice caps did not function synchronously. The Verneau network is located precisely at the boundary between the Jurassic and Alpine ice caps,

at the location of a topographic high (Salinois Fault Zone) (Fig. 1). However, due to the scarcity of glacial deposits of MIS 6 and 8 at the surface, the glacial extent map (Campy, 1982) was produced at the scale of the Jura Mountains and not at the level of detail investigated here. This fact makes it all the more important to focus on eglacial deposits present in an underground network that lies at the geographical limit of the glacial extent of MIS 8 and 6.

The study site, the Dolois Chamber, is located upstream of the Verneau network (7 km upstream of the Verneau spring), 130 m below the Baume-des-Crêtes collapse pit, whose entrance is at 789 m a.s.l. (Fig. 3). The Dolois Chamber is situated in the non-active part of the network, which limits remobilization of deposits. The current entrance to the Baume-des-Crêtes collapse pit shows no traces of water flow (Vivier, 2022). However, the deposition of sediments studied in the Dolois Chamber required a substantial inflow of water from the surface. The Baume-des-Crêtes collapse pit would therefore have been a paleo-sinkhole, collecting water from old surface sinkholes. A topographically flat area at 774 m a.s.l. suggests that the Baume-des-Crêtes paleo-sinkhole was also located at ~775 m a.s.l. Today, the active sinkhole in the marly outcrop of the hydrographic network is the Biefs-Bousset Sinkhole, which infiltrates and joins the Verneau network main drain, 200 m to the southeast and at 765 m a.s.l. (Fig. 3). Surface water is now captured by the Biefs-Bousset Sinkhole rather than the Baume-des-Crêtes paleo-sinkhole.

<figure><img src="image_5.png" /><figcaption>Fig. 3. Location of the study site at the foot of the Salinois Fault Zone. The Baume-des-Crêtes collapse pit is at a higher elevation than the active Biefs-Bousset Sinkhole. The exact location of the Baume-des-Crêtes paleo-sinkhole is unknown. The elevation of the Jura Mountains ice cap of the mid-Pleistocene (MIS 6-8) is hypothetical.</figcaption></figure>

## Methods

Four methods were used in this study: (i) detailed geomorphological mapping, (ii) petrographic analysis of deposits, (iii) U/Th-dating of speleothems, and (iv) determination of an order-of-magnitude age of the capture of the Baume-des-Crêtes paleo-sinkhole.

### Detailed geomorphological map of the Dolois Chamber

To obtain a high-resolution topographic map of the selected area, we acquired a 3D model using terrestrial laser scanning (Mohammed Oludare & Pradhan, 2016)

of the Dolois Chamber in 36 scan scenes, each with 7,000,000 points. After cleaning, 25,442,921 points made up the selected point cloud (Fig. 4). A Digital Terrain Model (DTM) extracted from the 3D model ground/land surface with 1cm resolution was used as a geomorphological background map for the study site. However, it was not possible to acquire the northern portion of the study area, which includes a very narrow passage. A topographic top-up was then done on this area, which was not surveyed in 3D, and was then integrated into the background map (Fig. 4).
<figure><img src="image_6.png" /><figcaption>Fig. 4. 3D model from which the DTM for the detailed geomorphological map of the Dolois Chamber was extracted. Dated speleothems and sediment logs are indicated. The abbreviation LOG refers to the sedimentary logs collected in the clastic deposits of the Dolois Chamber, while the abbreviation DOL refers to coring conducted in speleothems of the same chamber.</figcaption></figure>

## Sedimentary study of clastic deposits

A sedimentological study was done to determine the organization and order of clastic deposits of the various deposits in the cave. The analysis was done in the field using four sedimenta logs. Spot samples were taken to better characterize the petrography of these deposits. We assume that they originated from the surface and were transported into caves, where they remained for long periods.

## U-Th-dating of speleothems

Speleothems are relevant indicators that offer potential for paleoclimatic reconstruction (Quinif et al., 1994; Couchoud, 2008; Pons-Branchu & Quinif, 2018), particularly for interglacials (Luetscher et al., 2021). We selected five relatively unweathered speleothems (Fig. 4) for coring and dating. Three of them are covered by clay and other fine sediments and the other two seal these clastic deposits. The ²³⁰Th/²³⁴U method was applied to these five samples. This method makes it possible to date samples back about 500,000 years (Pons-Branchu & Quinif, 2018). The samples were analyzed at the Laboratoire des Sciences du Climat et de l'Environnement at Gif-sur-Yvette, France. The samples were first dissolved in beakers containing known quantities of a ²²⁹Th-²³³U-²³⁶U spike. This step was followed by co-precipitation with FeOH. U-TEVA® resin was used to performed U and Th purification. The U and Th fractions obtained were measured with a Thermo Scientific Neptune Plus Multi Collection Inductively Coupled Plasma Mass Spectrometer fitted with a jet pump interface and an Aridus II desolvating system. The procedure is described in Pons-Branchu et al. (2022). After blank and mass fractionation correction, iterative age estimation was used to calculate ages from measured atomic ratios (Ludwig & Titterington, 1994). The ages were corrected for non-authigenic ²³⁰Th, using the assumption of a (²³⁰Th/²³²Th) ratio of 1.5 ± 0.75 in the detrital fraction.

## Order of magnitude age of the sinkhole and the paleo-sinkhole

It is possible to estimate the minimum erosion time required to export a certain volume of marl from a watershed that is stable over time (Vivier, 2022). The first step is to convert the eroded marl volume (in m³) into eroded weight (in tons, t) using the density of sedimentary rock (2.7 t/m³), then divide it by the watershed surface area (in km²), and finally divide it by the average erosion rate of marl under humid temperate conditions (45 t/km²/yr) (Jaillet, 2000). The average erosion rate is obtained by measuring the sediment transport (dissolved and solid load at the outlet of the river system in Lorraine (Corbonnois, 1998) and more generally in the Paris basin (Bravard & Petit, 1997). This method assumes linear erosion over time, a low watershed gradient, and does not take environmental variations into account (e.g., changes in vegetation cover, transition from periglacial to present-day conditions). The goal is to obtain an order of magnitude estimate. The watershed of the Biefs-Bousset Sinkhole is generally flat, with an average slope of less than 15°, and is primarily composed of Liassic marls (Fig. 5A). Because the Baume-des-Crètes paleo-sinkhole is close by (100 to 200 m to the north), we can assume that its watershed was characterized by a similar lithology. It is also assumed that the entire volume of eroded marl was exported through the associated sinkhole. Thus, calculation of the minimum erosion time required to export the volume of marl corresponds to the calculation of the minimum erosion time required for the formation of the sinkhole itself, i.e. its most recent age of formation. The volume of marl eroded from the Biefs-Bousset Sinkhole is related to the talweg that the river incised, some 10 m deep, clearly visible on the DTM (Fig. 5B).

The volume of marl eroded from the Baume-des-Crètes paleo-sinkhole is more difficult to characterize.
Indeed, capture of the Baume-des-Crêtes paleo-sinkhole river by the Biefs-Bousset Sinkhole has altered the landscape, and the uncertain location of this paleo-sinkhole further complicates estimate. However, two slope breaks consistent with the topography enabled us to make the calculation

(Fig. 5C). Moreover, the Biefs-Bousset Sinkhole captured surface water to the detriment of the Baume-des-Crêtes paleo-sinkhole. We therefore assume that the age of formation of the Biefs-Bousset Sinkhole corresponds to the age of the capture of the Baume-des-Crêtes paleo-sinkhole.

<figure><img src="image_7.png" /><figcaption>Fig. 5. Eroded volume of marl exported by the Baume-des-Crêtes paleo-sinkhole (C) and the active Biefs-Bousset Sinkhole (B). We assume that the two sinkholes have the same watershed, mainly composed of Liassic marl (A). The slope break is highlighted. The exact location of the Baume-des-Crêtes paleo-sinkhole is unknown, but a topographical flat (774 m a.s.l.) suggests that it was also around 775 m a.s.l.</figcaption></figure>

RESULTS

Sediment data

The Dolois Chamber in the Baume-des-Crêtes
collapse pit is 25 m wide and 20 m high. Four types
of sediments were identified: i) allochthonous fluvial
pebbles, ii) clays and other fine sediments, iii) autoch-
thonous limestone blocks, and iv) speleothems.

Allochthonous fluvial pebbles show two similar
groups. Both are composed of weathered pebbles
averaging 1 to 5 cm in size, with no sorting, and
with an unconsolidated matrix. However, the first
group includes angular pebbles and a silty matrix,
whereas the second group includes round pebbles
and a clay matrix. We measured a thickness of 1
m for the allochthonous fluvial pebble deposit. The
bedrock of the Dolois Chamber is not visible, so it
is difficult to assess the actual thickness of these
deposits, which in any case must be greater than 1 m.
Petrographic analysis of the pebbles shows them to
be carbonate-bearing, more precisely biocalcareous,
with a ferruginous matrix and massive dark and
white limestones (Fig. 6). This observation does
not allow us to constrain the spatial origin of these
deposits, since this type of limestone is found at all
scales in the Verneau network catchment area (Tissot

& Tresse, 1978; Vivier, 2022), in the Jura Mountains (Bichet & Campy, 2008), as well as in the Alps. The nature and lack of classification of those sediments point to an allochthonous fluvio-glacial deposit, i.e. a debris flow containing pebbles. This pebbly debris flow originated from the upstream watershed of the Baume-des-Crêtes paleo-sinkhole (Fig. 3).

Clays and other fine sediments are composed of
brown laminae. The laminae have no particular
structure but seem to follow the geometry of the deposit
and allochthonous fluvial pebbles. We estimated
their total thickness at 2 m, but this also remains an
minumum estimate without knowing the depth to the
Dolois Chamber bedrock. The nature of this deposit
and its lamination suggests a pond-type depositional
environment. In addition, evidence of re-flooding (wall
karren and connecting conduits) (Vivier, 2021) suggests
flooding of the underground main drain 60 m below
on a regular basis, which gave rise to a now-energy
environment and the deposition of laminated clays.

Autochthonous limestone blocks contain large
angular blocks of autochthonous limestone (2 to 3 m³)
scattered on the ground, some of which are covered
by clay and other fine sediments described above.
Traces of collapse of the ceiling and walls of the Dolois
Chamber indicate the local origin of these blocks.
<figure><img src="image_8.png" /><figcaption>Fig. 6. Four sediment logs of the clastic deposits. Refer to Figure 5 for the precise location of those logs in the Dolos Chamber. The coarse part is composed of weathered pebbles averaging 1 to 5 cm in size, with no clear sorting, suggesting a pebbly debris flow. The entire surface is topped by fine brown clayey and laminated sediments.</figcaption></figure>

## Chronological data

The allochthonous fluvial pebbles are sealed by the clays and other fine sediments, i.e., they are older than the clay. Three speleothems are covered by clays

and other fine sediments, whereas two speleothems overlie the fine clastic sediments (Fig. 7). Collapsed limestone blocks from the ceiling are partly covered by clays and other fine sediments.

<figure><img src="image_9.png" /><figcaption>Fig. 7. Example of speleothems above (A) and below (B) clay and other fine sediments. The scale bar is 20 cm.</figcaption></figure>

The ages of the speleothems framing the fine sediments fall into MIS 1, 6, and 8 (Table 1). From the oldest to the most recent, the dates obtained are as follows: DOL 3: 272.0 ± 5.7 ka; DOL 5: 255.4 ± 11.6 ka; DOL 4: 171.0 ± 1.6 ka, belonging to MIS 8 and 6. The three

Holocene speleothems are: DOL-1B-B: 3.96 ± 0.49 ka; DOL2-B: 3.4 ± 2.5 ka; DOL-1-B-A: 3.3 ± 1.1 ka. Three generations of speleothems can thus be distinguished, MIS 8, MIS 6, and a Holocene generation (Table 1). The first two generations were covered by clays and
other fine sediments, whereas the current generation of speleothems grew on top of these fine sediments. We can therefore affirm that the latter were deposited after 171.0 ± 1.6 ka (i.e., during MIS 6 or later). Speleothems and more specifically the stalagmites we studied grow in the vadose zone, implying that the study site has been part of the vadose zone since MIS 8.

The incised talweg of the Biefs-Bousset Sinkhole has a volume of 6.8 × 10⁴ m³ (Fig. 5B). Therefore, approximately 10 ka appears sufficient to export this volume of marl via the sinkhole itself. For the Baume-des-Crêtes paleo-sinkhole, the eroded volume is 1.5 × 10⁶ m³ (Fig. 5C), which corresponds to a minimum erosion time of a few hundred thousand years.

<table><caption>Table 1. U and Th content, isotopic activity ratios, and ages of the six speleothems sampled in the Dolois Chamber. δ²³U_M = (²³U/²³U)measured / (²³U/²³U)equilibrium – 1 × 1000, with ²³U/²³Uequilibrium = 54.89 × 10⁻⁶ (atomic ratio after Cheng et al., 2013). δ²³U_T is the initial δ²³U calculated using measured value and calculated age. Ages were corrected for non-authigenic ²³UTh assuming a (²³UTh/²³Th) of the detrital fraction of 1.5 ± 0.75.</caption><thead><tr><th>Sample name</th><th>²³U (ppm)</th><th>²³Th (ppb)</th><th>δ²³U_M (%)</th><th>²³Th/²³U</th><th>²³UTh/²³Th</th><th>Age (kyr before 2022)</th><th>δ²³U_T (%)</th><th>Corrected age (kyr before 2022)</th></tr></thead><tbody><tr><td>DOL 1-B-B</td><td>0.126 ± 0.001</td><td>2.57 ± 0.02</td><td>338.7 ± 1.2</td><td>0.057 ± 0.001</td><td>8.49 ± 0.13</td><td>4.78 ± 0.08</td><td>342.5 ± 1.3</td><td>3.96 ± 0.49</td></tr><tr><td>DOL 1-B-A</td><td>0.140 ± 0.001</td><td>6.90 ± 0.06</td><td>344.9 ± 2.5</td><td>0.063 ± 0.001</td><td>3.84 ± 0.05</td><td>5.23 ± 0.09</td><td>348.1 ± 2.8</td><td>3.3 ± 1.0</td></tr><tr><td>DOL 2-B</td><td>0.157 ± 0.001</td><td>18.11 ± 0.15</td><td>242.1 ± 1.5</td><td>0.091 ± 0.001</td><td>2.37 ± 0.02</td><td>8.28 ± 0.09</td><td>244.5 ± 2.3</td><td>3.4 ± 2.5</td></tr><tr><td>DOL 3</td><td>0.168 ± 0.001</td><td>11.32 ± 0.09</td><td>67.9 ± 1.2</td><td>1.000 ± 0.002</td><td>44.62 ± 0.10</td><td>275.11 ± 4.24</td><td>146.5 ± 3.5</td><td>272.0 ± 5.7</td></tr><tr><td>DOL 4</td><td>0.238 ± 0.002</td><td>5.80 ± 0.05</td><td>127.7 ± 1.2</td><td>0.916 ± 0.001</td><td>112.66 ± 0.15</td><td>172.09 ± 1.01</td><td>207.1 ± 2.1</td><td>171.0 ± 1.6</td></tr><tr><td>DOL 5</td><td>0.151 ± 0.001</td><td>58.23 ± 0.47</td><td>59.3 ± 1.4</td><td>0.988 ± 0.002</td><td>7.66 ± 0.01</td><td>272.15 ± 4.14</td><td>122.1 ± 4.9</td><td>255.4 ± 11.6</td></tr></tbody></table>

## DISCUSSION

The aim of this paper was to propose a new framework for the spatio-temporal extent of paleo-glaciers at a topographical high in the northern Jura Mountains based on subsurface sediment data. Our results are of three types: morphological, sedimentological, and chronological. The study area is made up of separate sites, i.e. (i) the summit ridge of the Salinois Fault Zone at 864 m a.s.l. (Fig. 3), (ii) the area of proglacial sinkholes (775 m a.s.l. for the Baume-des-Crêtes paleo-sinkhole (today unknown) and 765 m a.s.l. for the present-day Biefs-Bousset Sinkhole), (iii) the Dolois Chamber at approximately 670 to 650 m a.s.l., and finally (iv) the active underground main drain at 608 m a.s.l.

### Reconstruction of cave deposits

Four types of sediment deposits are present in the Dolois Chamber. The first is a collection of allochthonous weathered carbonate pebbles, averaging 1 to 5 cm in size, without sorting and within an unconsolidated silty or clayey matrix (Fig. 8, Unit 1). The pebbles are interpreted as a fluvial and pro-glacial debris flow that entered the chamber via the Baume-des-Crêtes paleo-sinkhole (location unknown). This pebbly debris flow is linked to a major glacial advance located upstream of the Baume-des-Crêtes paleo-sinkhole watershed (Fig. 3).

Horizontally laminated clays and other fine sediments overlie and cover both the pebbly debris deposit and the old (MIS 8 and 6) speleothems. We suggested that this deposit may be linked to significant flooding of the underground main drain (50 to 70 m of hydraulic head variation). Even today, the main drain downstream of the Dolois Chamber is underwater and has some narrow gallery section, indicating that it may have been locally obstructed. It is known that during glacier retreat, i.e. in paraglacial phases, the amount of sediment transported is significant (Ballantyne, 2002) and can obstruct karst conduits (Jaillet et al., 2023). These obstructions cause temporary rises in the karst water level and result in flooding and the deposition of laminated clays such as those identified at the Dolois Chamber (Fig. 8, Unit 5). Finally, the laminated clays show significant suffosion. Therefore, it has been eroded and transported towards the Verneau main drain. Limestone blocks have also been identified in this chamber, either resting on the laminated clays or encased by them (Fig. 8, Unit 2). These are autochthonous blocks that testify to the destabilization of the ceiling and walls during periods of low water levels.

### Chronological model

U-Th dating suggests a least three phases of speleothem deposition, one related to MIS 8 (Fig. 8, Unit 3), a second related to MIS 6 (Fig. 8, Unit 4), and the third during the Holocene (Fig. 8, Unit 7). The two old speleothem generations are overlain by laminated clays (Fig. 8, Unit 5), while the most recent generation seals these fine clastics.
No clear stratigraphic relationship exists between the two old speleothem generations (Fig. 8, Units 3 and 4) and the pebbly debris flow deposit (Fig. 8, Unit 1), but it is obvious that the pebbly debris flow is older than the laminated clay. Only a great quantity of meltwater could transport the pebbly debris flow underground, and a large quantity of meltwater is common in paraglacial phases (Eyles & Kocsis, 1988; Church & Ryder, 1989; Ballantyne, 2002; Mercier, 2008). Glacial sedimentary deposits are not transported long distances from ice cap margins (Kleman et al., 2008). The only known glacial period with a spatial extent that reached the study site and possibly produced proglacial flows such as pebbly debris flows are glacials during the mid-Pleistocene. We therefore suggest that the pebbly debris deposit

formed during a mid-Pleistocene glacial period.

The old (DOL 4) and very old (DOL 3-5) speleothems formed during cold climate periods (Fig. 8) which may be surprising, but recent research has shown the existence of subglacial speleothems, which developed in a meltwater-protected zone of a cave under a warm-based glacier in impure carbonate rocks containing disseminated sulphide minerals (Spötl et al., 2024). The Jura Mountains ice cap was likely also a warm-based glacier, the Verneau underground network develops at the contact between Liassic marls and Jurassic limestones, and the Dolois Chamber was protected from meltwater intrusions. Hence it is conceivable that the two old speleothem generations formed during cold climates when the cave was overlain by a temperate glacier.

<figure><img src="image_10.png" /><figcaption>Fig. 8. Chronological reconstruction of the sediment deposits in the Dolois Chamber and the origin of water fluxes associated with these deposits. Speleothem growth periods are compared to the ice core record from Vostok, Antarctica (Petit et al., 1999).</figcaption></figure>

We also attempted to constrain the timing of the formation of the Baume-des-Crêtes paleo-sinkhole and the Biefs-Bousset Sinkhole. A few hundred thousand years are required to export the volume of marl eroded from the watershed through the Baume-des-Crêtes paleo-sinkhole. Thus, according to this estimate, the Baume-des-Crêtes paleo-sinkhole may have formed at the latest during the mid-Pleistocene. This is consistent with the age of the oldest dated generation of speleothems (MIS 8). Similarly, approximately 10 kyr are sufficient to export the volume of marl eroded from the watershed of the active Biefs-Bousset Sinkhole. Its formation, i.e. the capture of the Baume-des-Crêtes paleo-sinkhole, may have occurred at the latest during the early-Holocene.

Based on these observations, we propose a model for the evolution of the study site. This evolution is presented in four stages (Fig. 9). For each stage, we mention the stages presented in the reconstruction of the sediments in the Dolois Chamber (Fig. 8, Units

1 to 7) as well as the origin of the water intrusions associated with these deposits (Fig. 8 A–C).

**Stage 1:** The mid-Pleistocene ice cap (MIS 6-8) was at its maximum extent, reaching the top of the Salinois topographic high (Fig. 9) between the Ornans and Levier plateaus. The Baume-des-Crêtes paleo-sinkhole formed at the latest in mid-Pleistocene and was part of the vadose zone as shown by speleothem deposition during this time. We propose that a proglacial debris flow carried Jurassic limestone pebbles into the Baume-des-Crêtes paleo-sinkhole (location unknown) and deposited allochthonous fluvial pebbles in the Dolois Chamber.

**Stage 2:** The mid-Pleistocene ice cap (MIS 6-8) retreated (Fig. 9). Influx of pebbly sediments into the caves terminated and the pebbly deposit was partially eroded. We tentatively link this stage to the deposition of vadose speleothems (DOL 4), but the stratigraphic relationship between the pebbly debris flow and the two old speleothem generations remains unknown.
**Stage 3:** The glacial retreat was completed. The watershed of the Baume-des-Crêtes paleo-sinkhole was incised, leading to the Biefs-Bousset Sinkhole capturing the river of the Baume-des-Crêtes paleo-sinkhole at the latest during the early-Holocene (approximately 10 kyr) (Fig. 9). Water no longer flowed from the surface into the Dolois Chamber. Within the Verneau underground main drain, water flow was limited by internal obstructions (clastic sediments clogging the main drain) and external obstructions (moraine deposits, valley outlet blockage by ice). This situation led to flooding of the cave to depths of of several tens of meters and the deposition of laminated clays. These deposits rest on pebbly debris flow deposits (Fig. 8, Unit 1), on some of the limestone blocks (Fig. 8, Unit 2) and cover the two old speleothem generations (Fig. 8, Units 3 and 4). Thus, the effect

of the glacier on the hydrology of the karst network (flooding) and on sediments in the cave (laminated clays) remains significant even after its retreat.

**Stage 4:** The removal of cave sediments had reached a stage where the hydrological functioning of the karst network returned to water level variations of approximately 10 m during flooding (Fig. 9). In the Dolois Chamber, the laminated clays were eroded and transported into the Verneau main drain (Fig. 8, Unit 6) and Holocene speleothems started to grow (Fig. 8, Unit 7).

For each stage, we indicated the origin of the water intrusions (Fig. 8 and Fig. 9, A–C). This shows that the Dolois Chamber, and by extension also the Verneau underground network, are at the crossroads of (i) hydrological and glacial sediment input from upstream and (ii) flooding and deposition of laminated clays from downstream.

<figure><img src="image_11.png" /><figcaption>Fig. 9. Model of the development and evolution of the Verneau karst network in the Baume des Crêtes sector above and below the surface from the mid-Pleistocene to the Holocene. Water origin refers to Figure 8.</figcaption></figure>

## Spatial extent of the Jura Mountains ice cap

Given the glacial nature and age of the pebbly debris flow deposit, we propose that only during the mid-Pleistocene the local glacier extent was large enough to give rise to sufficient proglacial sediment transport.

The pebbly deposit contains only carbonate pebbles. This does not allow us to constrain the spatial origin of the glacier that gave rise to this deposit. On the other hand, Campy (1982) attributed glacial deposits lacking rocks from the Alps to the Jura Mountains ice cap. Can we then follow the same hypothesis and attribute the pebbly debris flow to the Jura Mountains ice cap? Further study of the lithology of a larger number of pebbles is needed to address this question. In any case, the mid-Pleistocene ice cap (MIS 8-6) near the study area was constrained by the topographic high of the Salinois Fault Zone between the Ornans and Levier plateaus. The limits of the mid-Pleistocene glacial extent can now be locally refined in line with the proposals made by Campy (1982).

## CONCLUSIONS

This study provides new insights into the past functioning of the Verneau karst network. First, a pebbly debris deposit is linked to proglacial meltwater intrusions that reached the Dolois Chamber via a paleo-sinkhole (location unknown) that we call the Baume-des-Crêtes paleo-sinkhole. It would be located 100 to 200 m to the north, and some ten metres above the current Biefs-Bousset Sinkhole. The Baume-des-Crêtes paleo-sinkhole likely formed at the latest during the mid-Pleistocene. Laminated clays that cover most of the Dolois Chamber floor is linked to several tens of meters of flooding in the Verneau main drain. The capture of the Baume-des-Crêtes paleo-sinkhole by the Biefs-Bousset Sinkhole implies a local flow reorganization and the abandonment of flows towards the Dolois Chamber. This may have occurred at the latest during the early-Holocene. Dating of speleothems in the Biefs-Bousset Sinkhole would be required to constrain the age of this capture.
ACKNOWLEDGEMENTS

The work was funded by the French Geological Survey (BRGM), the Rhône-Mediterranée-Corse Water Agency, and by a French government grant managed by the Agence Nationale de la Recherche under the France 2030 program, reference ANR-22-PEXO-0012, as part of the OneWater - Eau Bien Commun national research program. The authors thank the Panoply technical platform for access to MC ICP MS and A. Dapoigny for assistance during MC-ICP MS measurement. The authors thank the reviewers for their relevant feedback.

**Authorship statement:** MV, SJ, EH, and JBC designed the study and conducted the observations, measurements, and field sampling. EPB carried the U/Th analyses at the LSCE facility. MV and SJ analyzed the data. MV, SJ, EH, EPB, and JBC contributed to writing and manuscript design.

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