At total of 129 Sr dating results are compiled in this study (Table S2); 12 samples display evidence for diagenetic alteration with δ13C values <-5 ‰ (Table S2) and cathodoluminescence imagery exhibiting recrystallization patterns. Since their stable isotope data exhibit no evidence of diagenetic bias (δ13C > -5 ‰) but the obtained ages are older than the age of the stratigraphic level, four samples were further interpreted as reworked samples and considered to be outliers (Table S2). Excluding those 16 recrystallized or outlier samples, the 87Sr/86Sr ratios range between 0.708338 and 0.708914, corresponding to stratigraphic ages between 21.45 ± 0.30 Ma (late Aquitanian) and 8.58 ± 1.78 Ma (Tortonian), thereby allowing for the dating of 113 stratigraphic levels.

For the Forezan and Grésy-sur-Aix sections (sections 5 and 4, respectively; Fig. 2 – locations in Fig. 3), the NRM intensity ranges between 1.53×10-5 and 8.43×10-3 Am-1 with an average of 1.12×10-3 Am-1 and between 3.19×10-5 and 3.95×10-3 Am-1 with an average of 8.83×10-4 Am-4. The Forezan section (enriched in clayey facies) presents slightly higher NRM intensities and better paleomagnetic results, probably due to finer and more abundant ferromagnetic grains. Orthogonal projection of paleomagnetic results exhibits different patterns: (i) a stable reverse polarity slightly overprinted by a normal polarity (Fig. S2A) in many samples of the Forezan sedimentary section, (ii) a large normal overprint with a path toward reverse polarity (Fig. S2B) for many samples of the Grésy sedimentary section, (iii) rarely stable normal polarity (Fig. S2C) attributed to full normal remagnetization or eventually a subchron, and (iv) in the 868–870 m interval of the Grésy section (for two neighboring samples), a stable normal polarity affected by a post-lock tilting (Fig. S2D). Declination and inclination of the stable magnetization, or the direction displayed in the very last steps of demagnetization, were plotted together with the reversal angle (Fig. S3A, B). Due to large normal magnetic overprints, the reversal angle, which is the angle between the calculated direction and the expected direction for the normal geomagnetic polarity, helps to determine geomagnetic polarity scale, especially in carbonates affected by partial but intense remagnetizations (e.g., Demory et al., 2011). The entire Forezan sedimentary section displays very high reversal angles (Fig. S3A) attributed to the record of the reverse geomagnetic Chron C5Cr according to the age frame inferred from Sr dating (Fig. 2). Most of the Grésy-sur-Aix section is characterized by reverse polarities except the normal polarity recorded at 772 m, which may be attributed to Subchron C5Dr.1n., and tilted normal polarities (Figs. S2D and S3B) located at 868 m – just below a major unconformity – that may be related to the record of Chron C5Dn (Fig. 2). In the case of post-lock tilting, the reversal angle is higher than expected for normal polarity.

A total of 34 out of 74 analyzed samples enabled the biostratigraphical dating of sedimentary sections based on nannofossil assemblages (Table S3). Species that allowed precise stratigraphical calibrations are Coccolithus miopelagicus (NN5–NN8; 14.91–10.55 Ma), Helicosphaera ampliaperta (middle part of NN2–top NN4; 20.4–14.91 Ma), Helicosphaera scissura (upper NN2–NN5; 20.1–14.91 Ma), Helicosphaera stalis (NN6–NN11; 13.53–5.59 Ma), Sphenolithus belemnos (base NN3–top NN3; 19–18 Ma), and Sphenolithus heteromorphus (base NN4–top NN5; 18–13.5 Ma).

A total of 21 out of 34 analyzed samples allowed biostratigraphical constraints from dinoflagellate cyst assemblages (Table S4). Species that allowed precise stratigraphic calibrations are Cousteaudinium sp. (upper D16–top D17; 20.5–14.8 Ma), Ectosphaeropsis burdigalensis (lower D16–lower D18; 23.2–13.8 Ma), Hystrichosphaeropsis obscura (lower D17–top D19; 18.65–7.5 Ma), and Systematophora placacantha (Eocene–D18; Eocene–13 Ma).

Only 2 out of 40 analyzed samples provided a biostratigraphical dating of sedimentary sections based on foraminiferal assemblages (Table S5). In addition, 13 foraminiferal assemblages were already available from the literature (Lamiraux, 1977; Latreille, 1969; Mein, 1985; Aguilar et al., 2004) and were implemented in this study using recent biostratigraphical charts (Table S5). The species allowing firm stratigraphical constraints are Globigerinoides sicanus (MMi4a–MMi4c; 16.1–14.87 Ma, sensu Lirer et al., 2019), Paragloborotalia bella, and Praeorbulina circularis (N5b–N9, respectively, 20.4–14.0 Ma; N8b–N9a, 15.9–14.5 Ma; sensu BouDagher-Fadel, 2015).

Also, 21 mammal assemblages available from the literature (Fig. S4; see also Kalifi, 2020, and references therein for more details) were revealed to be useful to constrain the age of the sea retreat (Figs. S5 to S17) between zones MN6 and MN9 (15.2–9.5 Ma).

Figure S4 compiles a synthesis of the Miocene biostratigraphy chart of the study area.

Within the Chartreuse massif, a major thrust bringing Lower Cretaceous units on top of younger Cretaceous units has long been distinguished in the literature. It runs from Chambéry to the west of Grenoble and was successively referred to as the “chevauchement de la Chartreuse orientale” (Gidon, 1964, 1995) (φOr, Fig. 1B), the central Chartreuse thrust zone (Butler and Bowler, 1995), the “chevauchement subalpin principal” (Gidon and Arnaud, 1978; Gidon et al., 1978), the subalpine front (Deville and Chauvière, 2000), and the Chartreuse oriental thrust (Philippe et al., 1998). Here we refer to this thrust as the Chartreuse oriental thrust (φOr, Fig. 1B). It shows a large offset of about 10 km (Deville and Chauvière, 2000) at the latitude of Chambéry. Along its southern portion this thrust brings the Néron syncline affecting “Urgonian” (Barremian–Aptian) limestones on top of the Proveyzieux syncline filled with Miocene deposits. In its central part, it brings Mesozoic sediments on top of the Charmant Som anticline (CS, Fig. 1B), while further north Cretaceous strata are found over Oligocene deposits (Doudoux et al., 1992a, b; Gidon and Barféty, 1969). The Chartreuse oriental thrust is affected by numerous right-lateral NE–SW faults (Fig. 1B). North of Chambéry, the thrust connects with the western front of the Bornes and Bauges massifs characterized by Mesozoic units thrust over Oligocene and Miocene deposits (φB, Fig. 1B).

The prolongation of this thrust towards the south across the Isère valley is more contentious. The valley has sometimes been interpreted as the location of either a fold saddle (e.g., Gignoux and Moret, 1952) or possibly a fault (Gidon and Arnaud, 1978) that would offset the structures outcropping on both sides. The most recent detailed study concluded that there is no offset in the downstream part of the valley, but possibly a tear fault in the upstream (Gidon, 1995). South of the valley, the Moucherotte thrust that brings Cretaceous sediments atop Miocene of the Villard-de-Lans syncline has a complex 3D geometry, connecting with the poorly exposed Perrières Fault through the Bruzier right-lateral fault (Debelmas, 1965; Gidon, 2020b) (Fig. 7A). We combine new field observations with the published geological maps and other publications to produce a new structural map, with corresponding cross sections (Fig. 7) and interpretation. North of the Isère valley, the Proveyzieux syncline is a complex structure since it corresponds in fact to two synclines separated by a pinched anticline (the Proveyzieux anticline in Figs. 6A and 7A). South of the Isère valley, the Sassenage fold is also an anticline conformably capped by Miocene deposits (Fig. 6B). On each side of the Isère valley, the two anticlines have close axis direction: N198∘, 03∘ for Proveyzieux and N181∘, 06∘ for Sassenage (Fig. 7A). In both cases, the eastern limb of the anticline is overthrusted by Cretaceous sediments including Urgonian limestones (Moucherotte thrust in the south, Néron thrust in the north), strongly suggesting that the thrust system is continuous across the Isère valley (Gidon, 1995) (Figs. 7A and S18B). Therefore, we propose that the Proveyzieux and Sassenage anticlines correspond to the same structure. The precise trace of the thrust is unclear and hidden by Quaternary sediments, but the Moucherotte–Bruzier–Perrières thrust appears to be offset from the Neron thrust by ∼ 2km along a left-lateral NW–SE strike-slip fault along the Isère valley (Figs. 7A and S18B). Such a fault is compatible with the general direction of shortening (∼ E–W). A major thrust thus appears to be continuous from the western front of the Bauges massif to the western front of the Moucherotte range (from north to south φB, φOr, φMo) and is referred to here as fault zone 1 (FZ1) (Fig. 1B).

Although the westward Moucherotte thrust has long been described, its detailed geometry is still discussed. We conducted a detailed mapping based on field observations at 210 stations (Table S6) combined with a detailed analysis of previous maps and publications as well as one aircraft flight in order to constrain the 3D geometry of the Moucherotte range geology.

On the western flank of the Moucherotte range, the Cretaceous beds of the Moucherotte anticline are thrusted on top of the Miocene of the Villard de Lans syncline, above the Sassenage anticline (Figs. 6B and 7D) (Gidon, 1981). In the Comboire range, at the foot of the eastern flank of the Moucherotte massif (locality 1, Fig. 7A), gently dipping to the west Upper Jurassic strata rest above marls. These marls, previously mapped as Jurassic (Vif geological map; Barféty et al., 1967), are rather Early Cretaceous in age based on the occurrence of Berriasella (Gidon, 2020a). This implies a duplication of the Mesozoic series and thus the emergence of a westward-dipping thrust. This thrust most probably connects in depth with the Moucherotte thrust, making the Moucherotte range a klippe (Fig. 7D).

A few kilometers further south, near the Col de l'Arc (locality 2, Fig. 7A), the Cretaceous series are also duplicated. To the west, Urgonian beds of the Moucherotte anticline rest on top of the Upper Cretaceous (Figs. 6C and 7E). The hanging wall structure is complex with a syncline of Upper Cretaceous sediments in the footwall and a boxed anticline overlain by a monocline in the hanging wall (Fig. 6D). To the east of the Col de l'Arc, below the Crête des Crocs (locality 3, Fig. 7A), the Lower Cretaceous dips steeply to the east (∼ 70∘, Fig. 6C) and is allochthonous above more gently dipping Urgonian limestones. Further to the north the Urgonian limestones and Upper Cretaceous are seen below the Lower Cretaceous through the St. Ange window (locality 4, Fig. 7A) (Fig. 6E) (Barféty et al., 1967). The top of the Moucherotte range is thus a klippe that can be traced ∼ 1 km farther to the south to the Pierre Vivari (locality 5, Fig. 7A).

Further south, the Moucherotte klippe has been completely eroded, and the Cornafion summit (locality 6, Fig. 7A) corresponds to the inverted eastern limb of a N10∘-trending syncline in the footwall of the thrust (Figs. 6E and 7F), while further east the Upper Jurassic and Lower Cretaceous series are duplicated, constituting the Pieu klippe (Figs. 6G and 7F) (Barféty et al., 1967). At the western edge of this klippe, Upper Jurassic limestone allochthonous strata form a bent-fault anticline above the autochthonous Jurassic and Cretaceous strata (Fig. 6H) suggesting thrusting toward the west. We interpret this klippe as the southern and lower extension of the Moucherotte one (Fig. 7F).

The Moucherotte klippe probably did not extend south of the Double Brêche fault (locality 7, Fig. 7A), south of which the footwall syncline is no longer observed (Fig. 7A). Further south, the Mesozoic formations are weakly deformed, locally exhibiting westward-verging anticlines such as the Grande Moucherolle (Figs. 6I and 7G).

No evidence of the westward-verging Moucherotte thrust is found east of the Pieu klippe, implying that it roots east of the Conest range, which is the southwestern prolongation of the external Belledonne cover (Figs. 1B; 7A, D, E, F; 11D). Such geometry, resulting in apparent normal motion along the westward-dipping portions of the thrust (i.e., east of the Pieu klippe and Comboire range), strongly suggests that the thrust was folded after its activity.

Apparent offsets measured along NW–SE cross sections increase from south to north: 1.1 km along the Pieu section (Fig. 7F) and 3.3 km along the Pic St. Michel and Moucherotte sections (Fig. 7E, D).

As discussed above, we interpret the Moucherotte thrust as the prolongation of the Chartreuse oriental thrust, or FZ1, locally call the Néron thrust, across the Isère valley. In the same way, the Corenc–Jalla folded-thrust system (locality 8, Fig. 7A) (Blanchet and Chagny, 1923; Gignoux and Moret, 1952), also termed φc in the geological map of Grenoble (Gidon and Arnaud, 1978), probably corresponds to the northern prolongation of the eastern flank of the Moucherotte thrust outcropping in Comboire south of the Isère valley (Gidon, 1981) (Fig. 7A, B). This thrust system has been affected by the Ecoutoux anticline, a late eastward-verging fold that also affects the autochthonous units (Blanchet and Chagny, 1923; Gidon, 1981) (Fig. 7A, B). Apparent offset of the base of the Lower Cretaceous across the Neron–Corenc thrust (FZ1) measured along a NW–SE cross section amounts to 2.8 km (Fig. 7B).

Barféty and Gidon (1996) further consider FZ1, which they termed φ1, to be a major thrust which re-emerges to the east of the Isère valley on the western edge of the Belledonne massif (Fig. S19A, D). There, along the Grésivaudan valley on top of the Belledonne basement, φ1 dips 20–40∘ to the west, being sub-parallel to the Mesozoic series it affects, and duplicates the Dogger (Barféty and Gidon, 1996) (Fig. S19A, D). φ1 thus appears as a top to the west fault with an apparent normal motion, but that is an additive contact and thus more probably a tilted thrust (Deville et al., 1994; Barféty and Gidon, 1996). Further to the north, at the eastern edge of the Bornes massif, the Dogger and Lias are tightly folded in a top-to-the-west ∼ 4 km thick shear zone (Barféty and Barbier, 1983; Barféty and Gidon, 1996; Doudoux et al., 1982, 1999) that also possibly corresponds to a thrust tilted by the later uplift of the Belledonne range (Doudoux et al., 1999) (Fig. S19A, B, C).

Whilst further field study would be necessary to support this interpretation, we propose that the Chartreuse oriental thrust (FZ1) runs continuously from the western front of the Bornes to the western Moucherotte thrust in the south and from the east of Moucherotte to east of the Bornes, defining a ∼ 120 km long thrust nappe (Fig. 1B). This thrust sheet was termed the Aravis–Mt. Granier unit (Bellahsen et al., 2014) or the subalpine nappe (Pfiffner, 2014). These authors further suggested that this thrust roots in the basement east to the Belledonne massif, which would have been uplifted after motion along the thrust ended (also see Doudoux et al., 1982; Lacassin et al., 1990; Menard and Thouvenot, 1987).

In the Vercors massif, the main thrust west of FZ1 is the Rencurel thrust (φRe, Figs. 1B and 7A) which outcrops near La Balme de Rencurel (Fig. 8A). There, the thrust exhibits Barremian–Aptian limestones thrusted above overturned Miocene molasse deposits (Fig. 8B). The apparent offset for the top of the Urgonian is of 1.5 km along this section (Fig. 8C). Miocene strata are limestones bearing Chlamys praescabriusculus, which are upper Burdigalian in age (Fig. S9). At this location, the fault strikes N3∘, 40∘ E with slickensides having a pitch of 80∘ S (Fig. 8B), indicating an almost pure reverse faulting. Analysis of the dips in the Lower Cretaceous marls indicates that the hanging wall of the fault is characterized by two anticlinal structures separated by a syncline (Fig. 8C). Further east, two other thrusts with smaller offsets (0.8 and 0.9 km) are observed parallel to the Rencurel thrust (Fig. 8C). According to the published geological map (Barféty et al., 1967), the thrust extends southwards for 6 km until it branches at the tip of a NW–SE left-lateral fault (Fig. 1B). Towards the north, the thrust is continuous until the Col de Montaud (MC, Fig. 7A), with Miocene deposits observed in the footwall. At this location, it was proposed that the thrust connects with the thrust at the western front of the Vercors (Watkins et al., 2017). However, the fault does not cut across the footwall Miocene deposits (Gidon and Arnaud, 1978) but more likely turns from N–S to NNE–SSW to reach Veurey–Voroize where Upper Jurassic sediments are found thrusting atop the Miocene (Fig. 1B; Fig. 7A). We suggest that the Rencurel thrust prolongates toward the NE to the Voreppe thrust across the Isère valley (Fig. 7A). Indeed, the seismic profile 91VER1 that follows the Isère valley (Fig. 9A) reveals a major thrust at this location (Fig. 9C) with a 1.5 km apparent offset of the basal Lower Cretaceous reflector. Hence, we interpret the Rencurel and Voreppe thrusts as the same tectonic structure as previously proposed by Gidon et al. (1978) and Dumont and SPIA (2020).

Further north, the Voreppe thrust is mapped as a continuous structure bringing Jurassic sediments on top of Miocene deposits of the Voreppe syncline (Fig. 1B) (Gidon and Arnaud, 1978). North of Saint-Laurent-du-Pont, this thrust terminates, but other thrusts continue further to the north: to the west two faults merge together before bounding the western limb of the l'Epine anticline (EP, Fig. 1B), while to the east Miocene and Cretaceous deposits are vertical in the western flank of the Outherans (OU) anticline (Fig. 1B). The seismic profile 91CHA1 stands across the four faults (Fig. 9A, B). Its interpretation suggests that the two western faults (2b and 2c) merge at depth just above the basement and that they could also merge with the Voreppe thrust (2a) just east of the eastern limit of the seismic profile (Fig. 8B). We thus interpret the western faults (2b and 2c) as the northern prolongation of the Voreppe thrust (2a), with an 8 km long left-lateral step-over (Fig. 1B). Within the step-over, short NW–SE left-lateral faults are reported on the 1 : 50 000 map (M. Gidon, 1970) (Fig. 1B) that may concur with the westward shift of the shortening from the Voreppe thrust to the l'Epine thrust. Based on the basal Lower Cretaceous reflector, the apparent offsets are 2.2 km for the Voreppe thrust (2a), 2.3 km for the fault (2b), and 1.1 km for the fault (2c) (Fig. 9B). Following this interpretation, the FZ2 thrust system would be shifted westwards and continue northward as the (2b) thrust bordering the western flank of the Epine (EP), Grand Colombier (GC), and Crêt du Nu anticlines (Fig. 1B).

The seismic profile 91CHA1 shows that the western flank of the Outherans anticline is bounded by the Outherans blind thrust (OU), which is almost parallel to the three other ones and has a 1.3 km apparent offset (Fig. 9B). To the south, the Outherans (OU) anticline corresponds to the Chartreuse median anticline of Gidon (1964, 1990) (CS, Figs. 1B and 10A), while to the north it most likely connects with the Gros Foug (GF) anticline across the Chambéry valley (Fig. 1B). The geometry of the Gros Foug anticline can be observed on seismic profiles 88SV01 and 88SV03 (Fig. 10C, D) and supports an eastward-dipping Gros Foug thrust with an apparent offset between 2 and 2.5 km of the basal Upper Jurassic limestone reflector. This thrust appears to root in the Triassic detachment level above the basement (Fig. 10C, D) and is probably the northern extent of the Outherans thrust between the Chartreuse oriental (FZ1) and the Voreppe thrust (FZ2) (Fig. 10A).

The seismic profiles also reveal another thrust fault between FZ1 and FZ2: the Salève (SAL) thrust bounding the Salève (SA) anticline (Figs. 1B and 10A, B) that also roots above the basement in the Triassic. The apparent offset of the SAL fault is evaluated to be 2.5 km based on the basal Barremian–Aptian limestone reflector in seismic profile 88SV01 (Fig. 10D) and on the basal Upper Jurassic limestone reflector in seismic profile SV03 (Fig. 10C). The seismic profiles do not allow for the visualization of the connection between FZ1 and the SAL faults, and two hypotheses have been proposed: (i) FZ1 and the SAL faults are connected at depth (Deville and Sassi, 2006; Beck et al., 1998); (ii) FZ1 and the SAL faults are distinct (Guellec et al., 1990; Doudoux et al., 1982).

The western edge of the Vercors and Chartreuse massifs corresponds to open anticlines characterized by steep westward-dipping forelimbs and gently dipping back-limbs. Such anticlines are continuous from Pont-en-Royans in the Vercors massif (PR) to the Chailles anticline in the Chartreuse massif (CA), with the Ratz (RA) anticline in an intermediate position (Fig. 1B). Seismic profiles crossing these structures (Fig. 9A) reveal that these anticlines formed in the hanging wall of a major thrust zone: the Royans–Ratz thrust (FZ3). This thrust zone corresponds to a single fault bringing Barremian–Aptian to Upper Jurassic limestones above Miocene strata. According to the basal Upper Jurassic limestone reflector, the apparent offset is estimated to be 1.5 km along the 91CHA1-2 and 91VER1 seismic profiles (Fig. 9B, C). Alternatively, the seismic profile 82SE01 (Fig. 9D) shows that the main fault (FZ3) has an apparent offset of 4.5 km but that other minor faults also occur. FZ3 roots in the Triassic detachment level above the basement (Fig. 9B, C). According to the published geological maps the Chailles anticline appears to connect in the north with the Tournier anticline (TO) (Fig. 1B), suggesting that FZ3 could connect with the Jura internal thrust (φIJu) (Fig. 1B).

To the west of FZ3, seismic profiles reveal a blind thrust system with two thrusts likely connected at depth, referred to as FZ4 here, with a total apparent offset of the basal Upper Jurassic limestone reflector of ∼ 100–300 m in the north (91CHA1-2; 91VER1, Fig. 9B, C) and ∼ 1.24 km in the south (82SE01, Fig. 9D). To the south, this thrust system bounds the Monts-du-Matin anticline (MM, Fig. 9A). The thrusts crosscut inherited normal faults (i.e., the Saint-Nazaire fault – SN, Fig. 9D) as already reported by Deville et al. (1992) and root in the Triassic detachment level (Fig. 9).

West of FZ4, the seismic profile 82SE01 (Fig. 9D) highlights the presence of an additional blind thrust system that corresponds to the St. Lattier anticline (St-La, Fig. 1B) (Deville et al., 1992), referred to as FZ5 here. The extension of FZ5 to the north is not well constrained but may coincide with thrusts exhibiting small offsets (<100 m) that are rooted on an inherited normal fault seen along profile 91VER1 ∼ 40 km further north (Fig. 9C). FZ5 is not observed on the seismic profile 91CHA1-2 in the north (Fig. 9B).

Five main westward-verging thrust zones have been identified that all mostly trend NNE–SSW on an updated structural map of the subalpine ranges (Fig. 1B) and are presented on four regional cross sections (Fig. 11). Folds between the thrusts have the same general trend, and the cross sections have been drawn in a WNW–ESE direction.

More specifically, subsurface data and surface data (field observations and geological maps at 1 : 50 000) were used to build cross section A west of FZ1, while east of FZ1, the construction of the cross section relies on the analysis of 1 : 50 000 scale geological maps and the ECORS-CROP profile (Guellec et al., 1990; Roure et al., 1990; Mugnier and Marthelot, 1991) recently revised by Pfiffner (2014). The Cenozoic filling consists of Oligocene and Eocene deposits and is involved in superposed nappes (Doudoux et al., 1982). In the northern part of the zone, west of Cluses, the Upper Jurassic limestones are found three times vertically along the BZN1 borehole (Fig. 1B for location) (Charollais and Jamet, 1990). Seismic imaging suggests the presence of a thrust uplifting the Brizon basement high (Guellec et al., 1990). To the east, we have also integrated the Montjoie valley synthetic section of Gidon (2019) (Fig. S19A, B, C) for the Aravis and Belledonne chain relationships. In order to build the westernmost part of section B up to FZ1, subsurface data and surface data (field observations and geological maps at 1 : 50 000) were used. Between FZ1 and the summit of the Grand Manti (Fig. 11B), the construction of the cross section relies on detailed analyses of 1 : 50 000 geological maps. To the east of the Grand Manti, the data originate from the northernmost section available in Barféty and Gidon (1996). The results presented in Figs. 6 and 7 and in Barféty and Gidon (1996) (Fig. S19A, D) were used to build sections C and D (Fig. 10C, D) east of FZ2, while subsurface data (Fig. 9C, D) and results presented in Fig. 7 (for section D) were used to the west of FZ2. Sections B and C have been balanced and restored using the MOVE software (Fig. 18).

In this interpretation, FZ1 has been refolded and roots east of the external Belledonne with large parts of the subalpine ranges (Moucherotte, eastern Chartreuse, Beauges, Bornes) being a klippe. West of FZ1, FZ2 (2a, 2b, 2c) runs from south to north from the Vercors to the l'Epine and the Grand Colombier–Crêt du Nu. Seismic profiles (Fig. 9B) and structural mapping (Fig. 1B) highlight the fact that FZ2 has not been refolded as FZ1, but most probably branches at the top of the basement on a flat décollement in the Triassic evaporites that root further east below the external Belledonne massif in the Belledonne basal thrust (Doudoux et al., 1982; Bellahsen et al., 2014, 2012; Butler, 2017; Deville and Chauvière, 2000; Deville et al., 1992, 1994). FZ1 and FZ2 are only 4 to 9 km apart in the Vercors–Chartreuse but 39 km apart in the Jura–Bauges (Fig. 1B). Such widening is the combination of a fan-like map shape of the folds and offsets by left-lateral NW–SE faults. The greater distance between the thrust and the inner part of the belt in the Jura is likely to represent a more efficient décollement in the north, where Triassic evaporitic series are thicker (Philippe et al., 1996; Lickorish et al., 2002; Deville, 2021). Anticlines are found in the FZ2 hanging wall and thus in the FZ1 footwall (Figs. 1B and 11), but these are related to the activity of FZ2 according to Guellec et al. (1990) and Doudoux et al. (1982). FZ3 is continuous from the Vercors to the western Chartreuse. Further north, according to published geological maps, it appears to turn to the west and merge with the “chevauchement majeur du Jura interne” (φIJu, Fig. 1B) (Philippe, 1995). The map pattern is similar to that of FZ2 and interpreted in the same way. As there is no significant basement culmination between FZ2 and FZ3, FZ3 is interpreted to follow the same décollement level and to root underneath the Belledonne external massif as FZ2. FZ4 is only documented from seismic profiles, and its apparent offset decreases toward the north. The general fault pattern might suggest that FZ4 could connect with the “chevauchement majeur du Jura externe” (φEJu, Fig. 1B) through the NW–SE Jura transfer zone followed by the Rhône river (J, Fig. 1B) (Philippe, 1995). FZ5 is only documented west of the Vercors. As there is no significant basement culmination between FZ5 and FZ2, FZ5 and FZ4 are interpreted to follow the same décollement level and to root underneath the Belledonne external massif. Eastward-verging thrusts are locally described. East of the Royans (RO, Fig. 1B) in the Vercors and in the crêt du Nu in the Jura (CN, Fig. 1B), they appear to be linked with the underlying westward-verging décollement (Philippe, 1995). In Grenoble (Ecoutoux anticline, Figs. 7A, B and 11C), they appear to be related to a late deformation phase (Barféty and Gidon, 1996).

As described above, the NW–SE (N145∘) left-lateral southern Jura transfer zone appears to offset FZ3 and possibly FZ4 (Fig. 1B). The step-over between the Voreppe (2a) and the Epine (2b) thrusts stands in the prolongation of this zone and shows small NW–SE left-lateral faults (Fig. 1B), suggesting that it is linked to the same process. Other similar faults are only found at the southern tip of the Rencurel thrust (2a, N140∘) and possibly along part of the Isère valley near Grenoble offsetting FZ1 (Fig. 1B). Numerous NE–SW (∼ N65∘) right-lateral faults affect the southern Bauges, the Chartreuse, and the northern Vercors (Fig. 1B). These faults offset FZ1 and possibly FZ2 but not FZ3 (Fig. 1B). Both the left-lateral and the right-lateral faults are compatible with an ∼ WNW–ESE compression as the main trusts and folds.

In the footwall of FZ1, in the Rumilly syncline located to the north of the oriental domain (zone A, Fig. 5), the Oligo-Miocene succession starts with sub-concordant onlaps on the Mesozoic substratum, according to seismic profiles (Fig. 10). The thickness of the Oligo-Aquitanian continental deposits decreases westward (Fig. 10C, D). Indeed, these deposits are ∼ 200 m thick west of the Rumilly syncline (Fig. 9C, D) (Enay et al., 1970; P. Gidon, 1970), while to the east, they reach 1838 m between the footwall of FZ1 and the hanging wall of the SAL fault (SLV2 well data, Fig. 10A, B), and 1716 m at the footwall of the SAL fault (SV-101 well data, Fig. 10A, D). Thus, on both sides of the SAL fault, which appears west of FZ1 (Fig. 10B), the thickness of Oligocene deposits is relatively similar (Fig. 10B, D), suggesting that the Oligo-Aquitanian depocenter was located in the footwall of FZ1. This, along with sub-concordant Oligocene strata on the Mesozoic substratum, suggests that the SAL fault likely activated after the deposition of both Oligocene and early Miocene sediments (21–18 Ma). This is further consistent with field analyses revealing that the basal Miocene molasse deposits (upper Aquitanian–lower Burdigalian) are conformably lying on continental Aquitanian deposits dipping vertically in the eastern flank of the Rumilly syncline (Alby-sur-Chéran, 3, Fig. 3). Altogether, this implies that both Oligocene and Miocene sediments were deformed by the SAL fault after deposition, then after the early Burdigalian.

To the south of the Rumilly syncline in the Grésy-sur-Aix sedimentological section (4, Figs. 3 and 12), the sequence S2a basal boundary (dated to 18.05 ± 0.3 Ma based on correlation with the Forezan section 5; see Figs. 2 and 3 for location) was extrapolated along-strike to the north. In the 88SV01 and 88SV03 seismic profiles (Fig. 9D, C), this major stratigraphic surface coincides with a high-amplitude reflector. This reflector is characterized by low-angle toplaps and onlaps that either suggest (i) an angular unconformity, which would materialize the onset of a tectonic phase, or (ii) an erosive surface above the underlying deposits, which corresponds to a classic feature for sequence boundaries.

The Grésy-sur-Aix section outlines three examples of syntectonic sedimentation within sequence S2a (Fig. 12). First, sequence S2a is abnormally thick and has a very high sedimentation rate of 0.72 ± 0.32 mm yr-1 (∼ 750 m of sediments deposited over ∼ 0.85 Myr between 18.05 ± 0.25 at the base to 17.2 ± 0.15 Ma at the top. For the top of the S2a sequence, which was recorded during the Chron C5Cr, the calculation excludes an age older than the Chron C5Cr starting at 17.23 Ma) compared to the underlying sequences S1a–S1b that are only 200 m thick have an average sedimentation rate of 0.06 ± 0.01 mm yr-1 (deposited over ∼ 3.5 Myr between 21.45 ± 0.3 and 18.05 ± 0.25 Ma) (Fig. 12A). This suggests a significant increase in accommodation space at the initiation of sequence S2a, likely related to a regional tectonic event. Second, a 15 m thick interval (390–405 m, Fig. 12A) with a regional-scale continuity (also described 18 km to the south in section 5 at ∼ 380 m, Fig. 2) containing disorganized monogenic clasts of various sizes (centimeters to pluri-meters; Fig. 12C) and “ball and pillow” structures (Fig. 12D) suggests an earthquake-disturbed layer (i.e., seismites, F25; sensu Kalifi et al., 2020). Third, the occurrence of an angular unconformity observed in tidal flat deposits (Fig. 12A, B) is characterized by a decrease in bedding dip from 18 to 7∘ up-section and an erosive surface exhibiting westward-directed onlaps above it. This angular unconformity argues for a compressive tectonic event.

The timing of the angular unconformity is further well constrained by magnetostratigraphy and has been associated with Chron C5Dn at 17.35 ± 0.15 Ma (Fig. 2). Thus, sequence S2a most probably records the onset of a compressive phase. In this interpretation, in the 88SV01 and 88SV03 seismic profiles (Fig. 9D, C), the S2a sequence boundary characterized by low-angle toplaps and onlaps is a tectonically driven angular unconformity and thus dates the initiation of compression at 18.05 ± 0.25 Ma.

Strontium, biostratigraphical, and magnetostratigraphical dating were applied to sedimentological sections 3, 4, 5 (oriental domain, zone A), and 13 (median domain, zone C) from the southern Jura synclines immediately north of Chambéry (Fig. 3). Sequence stratigraphy interpretations enable E–W correlations in lower Miocene deposits between sections 3 and 13 (Fig. 13) by using sequence boundaries and maximum flooding surfaces. The thicknesses of sequences S1a–S1b do not exceed 200 m in sections 3, 4, and 5 (Fig. 13) and are even absent to the west in section 13 (Fig. 13). During sequence S2a, the maximum thickness is ∼ 750 m and is recorded in section 4 (Fig. 13), while to the west, in section 13 (Fig. 13), sequence S2a is 145 m thick. Furthermore, in section 4 (Fig. 13), the average sedimentation rate increases sharply from 0.06 ± 0.01 mm yr-1 during sequences S1a–S1b to 0.72 ± 0.32 mm yr-1 during sequence S2a (Fig. 12A). This firmly demonstrates that a depocenter localized close to section 4 (Fig. 13) appeared during sequence S2a. Subsequently, the thickest accumulation of the following sequence (S2b) lies further west at the Forezan locality (275 m, section 5, Fig. 13). This lateral variation of the thickness is associated with significant lateral facies variation characterized by a dominance of proximal marine deposits to the east (950–1015 m, section 4, Fig. 13), while to the west, sequence S2b is mainly represented by distal marine deposits (700–920 m, section 5, Fig. 13), thereby suggesting a westward migration of the depocenter between sequences S2a and S2b. This is also consistent with a progressive decrease in accommodation rates from ca. 1.17 to 0.17 mm yr-1 between sequences S2a and S2b in section 4 to the east (Fig. 13). Finally, the maximum thickness of sequence S3 is recorded west of section 5 (Fig. 13) at section 13 (Fig. 13). There, the sequence S3 is 390 m thick, while in the Rumilly–Chambéry syncline area (Fig. 13), the sequence S3 was probably much thinner. Furthermore, during deposition of sequence S3, lateral variation of the thickness is associated with significant lateral facies variation. At Loisieux (section 13, Figs. 13 and S7), sequence S3 is mainly characterized by distal marine deposits, while it exhibits coarse-grained deltaic deposits to the east (sections 4, 5, Figs. 13 and S5). These proximal deposits suggest that the Rumilly–Chambéry syncline area was progressively exhumed, which is also suggested by the continuation of the progressive decrease in accumulation rates recorded between sequences S2a and S2b. Thus, the S3 depocenter is most probably located in the Loisieux locality (section 13, Fig. 13), further highlighting the westward migration of the depocenter (Fig. 13).

In the 91CHA1-2 seismic profile (Fig. 14A), the Miocene sequences were calibrated using the PA-1 well log (Fig. 14B, C). The thicknesses of sequences S4 and S5 increase from the hanging wall to the footwall of FZ4, which is likely associated with a decrease in bedding dip up-section (Fig. 14A), suggesting growth strata in response to FZ4 activity. On the other hand, the underlying S3 deposits are isopachous both in the footwall and the hanging wall of FZ4, giving an initiation of FZ4 after the deposition of sequence S3. The end of tectonic activity on FZ4 is not so well constrained since seismic data do not allow a detailed observation of S6 deposits, although sequence S6 apparently seals (conformably) the underlying deposits, which might suggest the end of tectonic activity. This could be in accordance with the fact that FZ4 offsets solely S4 deposits, while S5 deposits are simply folded, which probably indicates a progressive decrease in the tectonic activity from sequences S4 to S5. All together, these observations suggest that FZ4 activity was recorded by the sequences S4 and S5 (∼ 15.0 to ∼ 12.0 Ma).

Further east, beyond the artifact located at the emergence of FZ3 (Massieu, Fig. 14A), the identification of sequence S4 is proposed based on similar seismic facies (Fig. 14D). In the hanging wall of FZ3, the basal Miocene deposits are conformably lying on the Mesozoic bedrock (on top of the Chailles anticline). This is consistent with fieldwork observations made 5 km to the north in the Chailles gorges (Fig. 14A, C), where the basal Miocene molasse deposits have been attributed to sequence S2a (Fig. 14E), suggesting that FZ3 was inactive during deposition of the sequence S2a. On the FZ3 hanging wall, the overlying sequence (S2b?–S3) geometries are observed on the 91CHA1-2 seismic profile (Fig. 14A), and their thicknesses decrease gradually when approaching the Chailles anticline together with a decrease in bedding dip. In the up-section, S2b(?), S3, and S4 sequences are also characterized by eastward-directed onlaps on the folded S2a deposits (Fig. 14A), thus suggesting growth strata related to FZ3 activity between these sequences. Downstream to the Chailles valley (Fig. 14C), above the S2a sequence, high in the cliff, growth strata (Fig. 14F) are again suggested based on apparent eastward-directed onlaps onto S2a deposits, in agreement with the interpretation of the seismic profile. However, these outcropping onlaps does not allow for the identification of the timing of the growth strata more precisely.

In the Bas-Dauphiné basin (Fig. 1B), the Miocene molasse deposits are poorly deformed and only the upper part of the Miocene succession crops out. So, the three SE–NW transects (Fig. 3) presented here are proposed using the interpretation of borehole data using sequence stratigraphy (Fig. 15) and our updated chronostratigraphy. Correlations are proposed by using the elevation of the main sequence boundaries (in m a.s.l. – meters above sea level).

The Bas-Dauphiné basin is separated into two tectonic zones. West of the Montmiral high (Fig. 15B, C), no compressive structures were found (Kalifi, 2020; Couëffé and Tourlière, 2008). On the other hand, FZ4 and FZ5 blind thrusts are found to the east of the Montmiral and L'île Cremieu highs (Fig. 15A, B, C). In this area, the S2a, S2b, and S3 sequences exhibit very few thickness variations (∼ 100 m for S2a–b sequences between VAF-1 and VAF-2 wells, Fig. 15B, and ∼ 50–70 m for S3 sequence between SLF-1, MO-1, MO-2, and MO-3 wells, Fig. 15C), despite the presence of the FZ5 thrusts within transect C and possibly B. On the western edge of the Bas-Dauphiné basin, the absence of S2a–S2b deposits (to the northwest of PA-1, VAF-2, and MO-3 wells, Fig. 15A, B, C) and the thickness variations of the S3 sequence to the west of the Montmiral high (Fig. 15B, C) are attributed to a complex inherited paleo-topography (Kalifi, 2020) along the Oligocene western European Rift (Debelmas, 1974; Curial, 1986; Bergerat, 1987; Ziegler, 1988, 1990, 1994; Bergerat et al., 1990; Sissingh, 2003).

East of the Montmiral high, significant variations in depositional thickness are observed for the S4 sequence with a maximum located in the footwall of FZ4. In transect A (Fig. 15A), S4 sequence maximum thickness of 215 m is recorded in the PA-1 well, while to the west, the sequence S4 thickness decreases progressively, reaching 53 m in the CH-1 well. In transect B, sequence S4 maximum thickness is recorded in the VAF-1 well, with a thickness of at least 161 m (since its top is eroded), while to the west, sequence S4 is 118 m thick in the BRI-1 well above the Montmiral high (Fig. 15B). In transect C, sequence S4 maximum thickness is recorded in the SLF-1 well, with a minimum of 190 m (eroded top), while to the west, sequence S4 is 129 and 116 m, respectively, in the MO-2 and MO-3 wells (Fig. 15C).

The sequences S5a–S5b also exhibit variation in thickness. In transect A, S5a–S5b maximum thickness of 210 m is recorded in the PA-1 well, while it decreases progressively to the west, reaching 82 m in the CH-1 well (Fig. 15A). In transect B, S5a–S5b maximum thickness is difficult to estimate because the top of the VAF-1 and VAF-2 successions is eroded to the east of the Montmiral high (Fig. 15B). However, a probable depocenter is located to the west of FZ5 in the FA-1 well, with a thickness of 153 m (Fig. 15B). In transect C, the S5a–S5b depocenter can be located between the Montmiral high and FZ5 in the MO-2 well (Fig. 15C). There, the sequences S5a–S5b are 204 m thick, while to the west, they are ∼ 165 m thick above the Montmiral high (MO-2 and MO-3 wells, Fig. 15C). To the west, the interpretation of the DP-108 well is difficult as only lithological data are available but might suggest another depocenter with a thickness of 205 m (Fig. 15C).

No significant thickness variations are observed for the S6 sequence, which is ∼ 105 m thick along transect A (Fig. 15A), ∼ 145 m thick along transect B (Fig. 15B), and ∼ 70 m thick along transect C (Fig. 15C). The variation in thickness for S7 and S8 sequences cannot be estimated as they are rarely present in well logs (Fig. 15).

To the east, the La Bresse basin is bordered by the l'île Cremieu high in the south and by the N–S-striking Jura thrusts in the north (Fig. 1B). In the La Bresse basin, the lower Miocene is absent, and the area is finally flooded at ∼ 12.0 Ma only during the S6 sequence (Fig. S13). The S6 sequence (∼ 12.0 to ∼ 10.8 Ma, Fig. 4B) and the transgressive deposits of the S7 sequence (∼ 10.8 to ∼ 10 Ma, Fig. 4B) are characterized by marine deposits, while the regressive system tract of sequences S7 and S8 exhibits purely continental deposits, implying a major regression starting during sequence S7.

At the Jura front, the BY-101 well log (Fig. 3) shows Mesozoic units thrusted atop Miocene deposits (Dumont, 1983). The same Miocene succession outcrops at the eastern border of the La Bresse basin, at Jujurieux (Dumont, 1983; I2, Fig. 3), and corresponds to marine deposits of sequences S6 and S7 based on Sr ages (Table S2), which are consistent with the magneto-biostratigraphical results of Aguilar et al. (2004). Therefore, sequences S6 and S7 are thrusted by the Jura thrust at the BY-101 locality. Thus, the onset of the Jura thrust is either synchronous or younger than the deposition of S6–S7 marine deposits (∼ 12 to ∼ 10 Ma).

At the Jura front, the continental deposits belonging to the S8 sequence according to mammal-based datings (I4, I5, I6 localities; Fig. 3) could also be partly deposited during the S7 regressive tract and are characterized by a 200 m thick continental clay and marl succession (Demarcq, 1970). This important thickness may suggest the existence of a tectonically controlled depocenter in the footwall of the Jura thrust.

The westward thinning of the Oligocene and Miocene S1a–S1b succession (∼ 21.0 to ∼ 18 Ma) observed in the Rumilly syncline implies a regular subsiding area between ∼ 29 ± 2 Ma (Rigassi, 1957; base of the Oligocene deposits dated by palynostratigraphy westwards of Annecy, locality 3; Fig. 10A, C) and 18.05 ± 0.15 Ma (S2a sequence boundary, Fig. 10A, C). Moreover, the westward-directed onlaps on the Mesozoic substratum, as well as the younger age of the base of the Oligocene sediments at the west of the Rumilly syncline (∼ 23 Ma based on mammals, Gaudant et al., 2002, locality 7; Fig. 10A, C), suggest a regular and slow westward migration of the depocenter. In a foreland basin, this geometry is consistent with a foredeep depozone located between the poorly subsiding proximal flank of the forebulge and the footwall of the active (tectonic) orogenic front, where the maximum of subsidence is recorded in response to the interplay between topographic loads and long-wavelength lithospheric deflection in relation with subduction process (DeCelles and Giles, 1996; Schlunegger and Kissling, 2015). The progressive decrease in depositional thicknesses towards the west might imply proximity to the forebulge, whereas the maximum depositional thickness recorded to the east of the Rumilly syncline suggests proximity to the orogenic front. This interpretation is consistent with that of Deville et al. (1994), who previously interpreted the Oligocene sedimentary succession in this area as deposited in a passive foreland flexural basin. The upper Oligocene–lower Miocene tectonic front is therefore located east of the Rumilly syncline. Since marine Miocene deposits have never been described east of FZ1, we believe that it already formed a morphostructural barrier during the first marine sequences (S1a–S1b; 21.0 to ∼ 18 Ma, Fig. 4B, C) and was thus active prior to ∼ 21.0, possibly as early as 29 ± 2 Ma. However, in the early stage, the active front could also have been located east of FZ1, for example along the Entrevernes thrust (E) (Fig. 1B) or even further east, and other arguments are still needed to decipher the ante-21.0 Ma structural history.

Southwest of the Rumilly syncline, the E–W correlation between sedimentological sections 3, 4, 5, and 13 (Fig. 3) highlights important thickness variations (Fig. 13), leading us to propose the following chronology. At 18.05 ± 0.25 Ma, sequence S2a exhibits a prominent acceleration in accumulation rates, especially at the Grésy sur Aix locality (section 4, Fig. 13), suggesting the rapid production of accommodation space probably linked to a regional tectonic event. Interestingly, the occurrence of seismites involving a major earthquake is observed within the S2a transgressive tract (Fig. 12). Seismic profiles from the Rumilly syncline (Fig. 10) outline the S2a basal sequence boundary as corresponding to high-amplitude reflectors characterized by low-angle S1 toplaps and S2 onlaps, thereby advocating for an angular unconformity. Together, this suggests a tectonic event associated with the activation of the SAL thrust at 18.05 ± 0.25 Ma, since the sequence S2a depocenter is located at the footwall of this fault (Fig. 17). Subsequently, the angular unconformity highlighted in tidal flat deposits and associated with Chron C5Dn (17.53–17.23 Ma, Fig. 12A, B) recovered at the top of the S2a regressive tract, most probably reflects another regional tectonic event with the depocenter abruptly migrating westwards at that time at the footwall of the GF fault (section 5, Fig. 13). These two lines of evidence therefore support the hypothesis of the onset of the GF fault during Chron C5Dn (Fig. 17). At ∼ 16.2 Ma (S3 transgression), the depocenter migrated westwards again and lies at the footwall of FZ2 (section 13, Fig. 13). There, during deposition of sequence S3, thick distal marine deposits were sedimented, whereas coarse-grained deltaic deposits prevailed to the east of FZ2 (sections 4, 5, Fig. 13), suggesting a more proximal and uplifted area. Thus, FZ2 was activated during deposition of sequence S3 (Fig. 17).

At the hanging wall of FZ3, the base of sequence S2a is concordant on the folded Urgonian, suggesting that folding did not start before the deposition of the S2a sequence. Sequences S2b(?)–S3 exhibit growth strata geometries suggesting a westward thickening of sedimentary deposits (Fig. 14C, D), which can be attributed to an anticline formation in the hanging wall of FZ3. Without further stratigraphic constraints, however, we can only speculate that FZ3 was active at least since deposition of sequence S3 (∼ 16.2 Ma, Fig. 17). On the eastern edge of the Bas-Dauphiné basin at the footwall of FZ4, the presence of growth strata geometries between sequences S4 and S5 (Fig. 14A), as well as the occurrence of a sequence S4 depocenter between FZ4 and FZ5 (Fig. 15), implies a flexural subsidence at the footwall of the active FZ4, therefore constraining the activity of FZ4 between the onset of sequence S4 and the end of sequence S5, so between ∼ 15.0 and ∼ 12.0 Ma (Figs. 4A, B and 17). This is also in agreement with the underlying S2 and S3 deposits, as well as the overlying S6 deposits, which all appear isopachous (Figs. 14 and 15), and thus suggests the absence of tectonic activity during deposition of these sequences. In the FZ5 footwall the presence of a sequence S5a–S5b depocenter (∼ 14.0 to ∼ 12.0 Ma) (Fig. 15C) suggests a continuous westward migration of the depocenters (Fig. 17) that could have been controlled by FZ5 activity at that time. This hypothesis is consistent with the general trend observed since the base of S2a at zone A 18.05 ± 0.3 Myr ago, which shows the successive activation of the fault zones from east to west (SAL fault, GF fault, FZ2, FZ3, FZ4, and finally FZ5). Thus, the thrusts that are interpreted to root below the Belledonne massif in the Belledonne basal thrust were activated in-sequence (Fig. 17).

In the Rumilly–Chambéry synclines the thick S2a and S2b sequence (18.05 ± 0.15 to 16.3 ± 0.3) marine deposits of the depocenter that developed in the footwall of the SAL and GF (sections 4, 5, Figs. 13 and S5) are capped by a rapid shallowing-upward succession, followed by S3 gravel-rich braided-river deposits (Fig. S8). This evolution suggests a final immersion linked with the fold-and-thrust propagation (Heller et al., 1988; DeCelles and Giles, 1996; Sinclair, 1997) and thus that the SAL and GF faults were active during the S2 and S3 sequences (until ∼ 15 Ma). However, the lack of timing control within the S3 and overlying sequences at this location precludes any precise constraint on the timing of the end of motion on SAL and GF (Figs. 13 and S8).

In the footwall of FZ2 near Grenoble (zone D, 20, 22, Fig. 3), the sedimentary succession of the S4–S5 sequences of the Voreppe syncline is characterized by three stacked Gilbert deltas (GDs). GD1a and GD1b are 25 and 40 m thick, respectively, belong to the S4 (∼ 15.0 to ∼ 14.0 Ma), and GD2 belongs to the S5 (∼ 14.0 to ∼ 12.0 Ma) (Fig. S8). As these stacked Gilbert deltas attest a continuous creation of slope and accommodation space (Ricketts and Evenchick, 2007) we interpret them as reflecting continuous activity of FZ2, thus until at least ∼ 12 Ma (Kalifi et al., 2021). GD2 is much thinner than GD1b (20 m against 40 m, respectively) and shows an abrupt transition between foresets and ∼ 10 m thick topsets themselves overlain by ∼ 5 m high foresets interpreted as a mouth bar in a distal deltaic position, suggesting a final transgression corresponding to S5b (Kalifi et al., 2021) (Fig. S8). This unit is finally overlain by 50 m of planar-stratified conglomerates interpreted as braided fluvial deposits likely evolving from deltaic to fluvial hydrodynamics. This transition could correspond to the final shallowing-up succession of the Gilbert deltas' stacking pattern, thus suggesting a progressive decrease in FZ2 activity until the end of the S5 sequence (∼ 12.0 Ma) (Kalifi et al., 2021).

To the west, in the footwall of FZ3 and FZ4, the sequence S6 deposits have a constant thickness in the Bas-Dauphiné basin (Fig. 15) and the sequence S6 seals the underlying S4 and S5 growth strata (Fig. 14A). This suggests an end of FZ3 and FZ4 activity in the Vercors and Chartreuse massifs at ∼ 12.0 Ma (top S5 sequence).