# Onset of Iberian-European plate convergence: Late cretaceous flexural response of a hot lithosphere
(Aquitaine Basin, France)

Benoit Issautier, Eric Lasseur, Nicolas Saspiturry, Paul Angrand, Simon Andrieu, Olivier Serrano

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## To cite this version:

Benoit Issautier, Eric Lasseur, Nicolas Saspiturry, Paul Angrand, Simon Andrieu, et al., Onset of Iberian-European plate convergence: Late cretaceous flexural response of a hot lithosphere (Aquitaine Basin, France). Tectonophysics, 2022, pp.229504. 10.1016/j.tecto.2022.229504. hal-03782544

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1 Onset of Iberian-European plate convergence: Late Cretaceous

2 flexural response of a hot lithosphere (Aquitaine Basin, France)

3 **Authors:** Issautier Benoit¹ ; Lasseur Eric¹ ; Saspiturry Nicolas², Angrand Paul³ ; Andrieu

4 Simon¹ ; Serrano Olivier

5

6 ¹BRGM-French Geological Survey, 3 Avenue Claude Guillemin, 45100 Orléans, France

7 ²Géosciences Montpellier, CNRS, University of Montpellier, Montpellier, France

8 ³Geosciences Environnement Toulouse (GET), Université de Toulouse, UPS, Univ. Paul Sabatier,

9 11 CNRS, IRD, 14 av. Edouard Belin, F-31400 Toulouse, France

10 **Corresponding author:** Issautier Benoit (b.issautier@brgm.fr)

11 **Abstract:**

12 The Aquitaine Basin is closely linked to the evolution of the Pyrenees, providing precious
13 evidences of the early stages of the Pyrenean orogeny in the Late Cretaceous. Although the timing
14 and geometry of the Early Cretaceous rifting stage is well constrained in the Pyrenees and
15 surrounding north Pyrenean basins, the subsequent early inversion stage has resisted
16 characterization. In this study, we used well log correlations and interpretation of seismic reflection
17 profiles to determine the Late Cretaceous evolution of sedimentary record in the Aquitaine Basin,
18 spanning the transition from the postrift stage to the beginning of convergence.

19 We identified two main deformation stages during the Late Cretaceous with radically different
20 basin's geometry between the western and eastern portion of the Aquitaine Basin. The Western
21 Aquitaine Domain (Pau area) seems to remain under this thermal subsidence regime, inherited from
22 the postrift stage initiated in the Turonian, with little to absent deformation stages. In contrast, the
23 Eastern Aquitaine Domain records early inversion, as indicated by the occurrence of a well-defined
24 Campanian flexural basin. In the central Aquitaine domain (Tarbes to Saint Gaudens area),
25 Coniacian-Santonian period is marked by the sharply increasing distinction between a very low
26 subsiding North Aquitaine Domain (condensed sedimentation), and a strongly subsiding South
27 Aquitaine Domain. We interpret it as a consequence of lithospheric buckling resulting from far-
28 field compressive strain between the Eurasia and Africa plates.

29 During the Campanian, northward convergence of the Iberia plate led to the formation of a flexural
30 basin with a narrower, steeper geometry than the subsequent Eocene basin that formed in nearly
31 the same region during the Pyrenean orogeny. We attribute this difference to the development of
32 the foreland over a not thermally reequilibrated lithosphere in Campanian time, roughly 10 Myr
33 after rifting in a context of active thermal subsidence. This inheritance of a rifted state conferred a
34 yet non-equilibrated lithosphere being thus of very low rigidity. and a susceptibility to .
35 Sedimentary basins formed during both the Coniacian-Santonian and Campanian to early
36 Maastrichtian were thus accommodated by short-wavelength flexural deformation of the
37 lithosphere, and they were subject to a distinctive deformation pattern and an immediate response
38 to the onset of compressive strain.

39 **Keywords:** Pyrenees, Iberia-Eurasia convergence, Late Cretaceous, hot lithosphere, flexural basin,
40 inheritance, far-field deformation
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# 1. Introduction

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The Pyrenean mountain belt is the result of Late Cretaceous to Miocene north-south plate
convergence and collision between Iberia and Europe (Puigdefàbregas and Souquet, 1986; Vergès
et al., 1995; Olivet, 1996; Rosenbaum et al., 2002; Sibuet et al., 2004; Teixell et al., 2018). To the
north, the Aquitaine basin which is the retro-wedge records the inversion of the Early Cretaceous
North Pyrenean rift basins due to the northward migration of the Iberia plate (Souquet et al., 1977;
Ducasse and Vélasque, 1988,) as early as Late Santonian? time (Desegaulx and Brunet, 1990;
Mouthereau et al., 2014; Ford et al., 2016; Rougier et al., 2016; Vacherat et al., 2017; Angrand et
al., 2018). The North Pyrenean Basins result from extreme thinning of continental crust and mantle
exhumation (Jammes et al., 2009, Lagabrielle et al., 2010, Masini et al., 2014; Teixell et al., 2016).
This history implies a significant change in the thermal regime of these basins, marked by high
temperatures (>600°C) at their bases, as suggested by HT-LP metamorphism in the Internal
Metamorphic Zone (Ravier, 1957; Golberg and Leyreloup, 1990; Boulvais et al., 2006; Clerc et
al., 2015, 2016; Ducoux et al., 2019; Angrand et al., 2021). In the southern part of the Aquitaine
Basin, the legacy of the extensional stage was a hot and stretched lithosphere before and at the
onset of Iberia-Europe convergence. After the rifting stage ended in the Cenomanian, the Aquitaine
domain entered in a postrift stage, typified by cooling and thermal subsidence, which continued
during the Late Cretaceous at least (Brunet, 1991; Desegaulx and Brunet, 1991; Angrand et al.,
2018). In the North Pyrenean Zone, the thermal regime inherited from the rifting stage may have
remained elevated until Santonian time (Saspiturry et al., 2020a) or even the Eocene (Vacherat et
al., 2014; Angrand et al., 2018). The character of such a postrift stage and the response of this
probably still hot lithosphere to compression 10 Ma only after stretching requests further
investigation.
Field evidence places the onset of convergence between Iberia and Europe during latest Santonian to Campanian time, as documented in the Southern Pyrenees (Puigdefàbregas and Souquet, 1986; Berástegui et al., 1990; Puigdefàbregas et al., 1992; Bond and McClay, 1995; Arbués and Berasteguy, 1996; Ardèvol et al., 2000; Beaumont et al., 2000; Vergès et al., 2001; Filleaudeau et al., 2011), the Corbières region (Billotte et al., 2005), and Provence (Leleu et al., 2009). Cretaceous plate kinematic models based on reconstruction of magnetic anomalies in the Atlantic have been strongly debated for many years. Several authors have placed the initiation of convergence in the Santonian (Roest and Srivastava, 1991; Olivet, 1996; Rosenbaum et al., 2002; Macchiavelli et al., 2017).

Current understanding is that compression affected the Pyrenean realm less than 10 Myr after the end of rifting (Mouthereau et al., 2014; Vacherat et al., 2014; Ford et al. 2016; Angrand et al., 2018; Grool et al., 2018; Dielforder et al. 2019). The Aquitaine foreland Basin developed from the late Santonian/early Campanian to early Miocene times in two major compressive phases, i.e. Late Cretaceous and Eocene-Oligocene (Ford et al., 2016; Rougier et al., 2016; Grool et al., 2018; Vacherat et al., 2017; Ternois et al., 2019; Ortiz et al., 2020). These two phases of active subsidence are separated by a quiet phase of little subsidence from late Maastrichtian to Thanetian times (Ford et al. 2016; Dielforder et al. 2019). While the orogenic phase has been widely investigated, the Cenomanian to late Santonian period recording the transition from rifting to compression needs to be further investigated. Indeed, the rifting event created a complex structural and thermal inheritance that should leave evidence of a particular response during the early orogeny.

The purpose of this work is to analyze the Late Cretaceous paleogeography and sedimentary thickness variations and architecture to highlight the timing and style of deformation of this hot stretched lithosphere and tentatively use this record to constrain the geodynamic evolution during
the transition from post rift to compression. Through this effort we can better characterize the
transition between the end of rifting and the onset of deformation that formed the structure of the
Aquitaine foreland basin.

## 2. Geological setting

### 2.1. Structural framework of the Aquitaine foreland basin

The asymmetric Pyrenean realm is characterized by the Ebro pro-foreland basin to the south and the Aquitaine retro-foreland basin to the north (Boillot and Capdevilla, 1977; Williams and Fischer, 1984; Séguret and Daignières, 1986). The Aquitaine Basin can be divided on the basis of major tectonic structures into three structural and paleogeographic domains referred here as the North Aquitaine Domain (NAD), the South Aquitaine Domain (SAD), and the East Aquitaine Domain (EAD) (Figure 2A). The Aquitaine Basin is overthrusted by the North Pyrenean Zone along the North Pyrenean Frontal Thrust (NPFT; Choukroune and ECORS Team, 1989; Daignières et al., 1994). In the Central and Eastern Pyrenees, the southernmost Aquitaine basin comprises the Sub-Pyrenean Zone, which overthrust the autochtonous basin along the Sub-Pyrenean Thrust (SPT; Ford et al., 2016; Rougier et al., 2016). The North Pyrenean Zone (Figure 2B) records the inversion of the Early Cretaceous rift system (Ducasse et Vélasque, 1988; Teixell et al., 2016; Espurt et al., 2019; Saspiturry, 2019a; Angrand et al., 2021).

The Aquitaine Basin has undergone a complex polyphase history that generated a substantial structural inheritance, which includes the major Hercynian basement structures (Figure 2A): the "Flexure Celtaquitaine" (BRGM et al., 1974) and the "Sillon Houiller" and its southern branch the Toulouse-Fault (De Saint Blanquat et al., 1990; Bouhallier et al., 1991). Three subsequent stages of rifting have affected the combined Pyrenean-Aquitaine domain: (1) late Paleozoic to Early Jurassic
110 (Hettangian) rifting (BRGM et al., 1974; Curnelle et al., 1983) recording the collapse of the
111 Variscan belt (Saspiturry, 2019; Saspiturry et al., 2019b), (2) Jurassic tectonic quiescence (Delfaud
112 and Henry, 1967; BRGM et al., 1974), and (3) late Aptian to Albian rifting (Delfaud and Villanova,
113 1967; Arnaud-Vanneau et al., 1979; Bouroullec et al., 1979; Peybernès, 1979, 1982; Peybernès and
114 Souquet, 1986; Serrano et al., 2006; Masini et al., 2014; Issautier et al., 2020; Ducoux et al.,
115 2021). At the same time, hyperextension affected the surrounding basins: (1) the V-shaped Bay of
116 Biscay margins (Thinon, 1999; Ferrer et al., 2008; Tugend et al., 2014), (2) the Parentis Basin (e.g.
117 Bois and ECORS Scientific team, 1990; Ferrer et al., 2008), (3) the Mauléon Basin (Jammes et al.,
118 2009; Lagabrielle et al., 2010; Masini et al., 2014; Saspiturry et al., 2019a), (4) the Ballongue Basin
119 (Debroas, 1987; Angrand et al., 2021), (5) the Baronnies Basin (Debroas et al., 1990), and (6) the
120 Boucheville and St. Paul de Fenouillet Basins (Clerc and Lagabrielle, 2014; Clerc et al., 2016).
121 These hyper-extended rift basins were characterized by a syn-rift high-temperature and low-
122 pressure metamorphism (Montigny et al., 1986; Golberg and Maluski, 1988) and a relatively steep
123 geothermal gradient (Vacherat et al., 2014; Hart et al., 2017; Saspiturry et al., 2020a). A magmatic
124 depocenter of thick syenite volcanism (Serrano et al., 2006) existed north of the Bigorre Basin (the
125 eastern termination of the Baronnies Basin), possibly supplied by the Sub-Pyrenean basement fault,
126 which was later reactivated as a Pyrenean thrust (Rougier et al., 2016). During the Late Cretaceous,
127 the NAD spanned the deposition of a shallow carbonate platform represented by the Jouansalles
128 and Gensac Formations (Serrano, 2001). This carbonate platform graded to a deep basin in the
129 SAD containing turbidites fed mainly from the North Aquitaine Platform (BRGM et al., 1974;
130 Razin, 1989). The SAD contains several extensional structures that were subsequently inverted
131 forming, among other features, the Audignon, Antin-Maubourguet, Pecorade, Pau, and Gensac
132 anticlines, which influenced the Cretaceous and Cenozoic evolution of the SAD (Serrano et al.,
133 2006). They correspond for most of them to clearly identified salt-related structures, some of them
being inverted during the Late Cretaceous, more particularly in the EAD likewise the Gensac anticline. Borehole data and seismic reflection profiles show a Late Cretaceous thining trend across this fold limb, therefore salt tectonics certainly influenced sedimentation on a local scale (Serrano et al. 2006; Espurt et al. 2019; Fillon et al., 2021).

## 2.2. Role of Late Triassic evaporites during the Pyrenean orogeny

The Pyrenean-Aquitaine domain is a salt-rich domain where evaporites were deposited in a wide Late Triassic-Earliest Jurassic rift system all over North Africa to South France (Leleu et al., 2016; Angrand et al., 2020) and their thickness might exceed 1000m in well preserved basins (Curnelle, 1983; Ortí et al., 2017; Scotese and Schettino., 2017). During the Early Cretaceous rifting, these evaporites are believed to have played a significant role in controlling the basin's thermicity and geometry (Mauriaud, 1987; Mediavilla, 1987; Masini et al., 2014; Clerc et al., 2016; Corre et al., 2016; Ducoux et al., 2019; Duretz et al., 2019; Issautier et al., 2020; Labaume and Teixell, 2020; Lagabrielle et al., 2020; Izquierdeo-Lavall et al., 2020; Saspiturry et al., 2021b; Ducoux et al., 2021; Ford and Verges, 2020). Evolution of knowledges and modelling of salt-tectonics pointed the importance of evaluating the impact of evaporitesoccurences/distribution and inherited salt-controlled structure to understand how rift basins are subsequently inverted (e.g. Jackson and Hudec, 2017). Recently, the Pyrenees has been strongly investigated in this scope. Main results demonstrate that the belt-geometry is strongly conditioned by pre-orogenic salt distribution as well as crustal to upper lithospheric structures inherited from the previous rift events. Key examples are found both in the North Pyrenean Zone, where syn-rift basin(s) (Mauléon and salt mini-basins of the Chainon Béarnais) are uplifted as pop-up gliding along evaporites preserving relatively well their initial geometry (Labaume et al., 2020; Saspiturry et al., 2020b). Furthermore, unravelling the salt-controlled pre-compression folded basins such as the Chainon béarnais and easternmost NPZ
through halokinesis represents a major step in assessing the Pyrenean shortening rates as demonstrated by Izquierdo-Lavall et al. (2020) and Ford and Verges (2020). A similar study within the South-Pyrenean Zone where pre-compression Ager salt mini-basin is evidenced provides new shortening values steeming for the folding inheritance and the authors to conclude that estimations are similar (Burell and Teixell, 2021). Furthermore, in the SPZ, the existence of early Mesozoic horst and grabben significance controls both the extensional and compressional deformation distribution between a southern and northern domains (Hudec et al., 2021).

## 2.3. Timing of convergence initiation

The earliest accepted evidence of active deformation between Iberia and Europe is recorded in the South Pyrenean Zone as the initiation of the Turbón ramp fold-Boixols thrusts during latest Santonian time (Puigdefàbregas and Souquet, 1986; Puigdefàbregas et al., 1986, 1992; Berastegui et al., 1990; Bond and McClay, 1995). This event corresponds to the inversion of the Early Cretaceous Organyà extensional basin, which imposed a flexure that defined an east-west elongated Pyrenean foredeep basin to the south (Deramond et al., 1993; Whitechurch et al., 2011). This flexural basin was initially filled with 1 km of Santonian-Campanian turbidites of the Valcarga Formation and Campo breccias (Simó, 1985, 1986, 1989). These rocks clearly indicate the initial growth of the Turbon ramp, which provoked the collapse of the Calcaires des Cañons carbonate platform into the flexural basin. Accelerated thrust activity is recorded by the 1-km-thick Arén Formation (Simo, 1985, 1986, 1989), in which paleocurrent directions indicate westward sediment transport (Puigdefàbregas and Souquet, 1986; Simó, 1986; Ardèvol et al., 2000; Rosell et al., 2001). The filling of the South-Pyrenean Boixols basin was influenced, at least in late Campanian to Maastrichtian time, by the growth of thrusts and folds (Bond and McClay, 1995). Progradation of the Arén Sandstone ended with the deposition of fluvial facies of the grey and red Garumnian
180 facies (e.g. Rosell et al., 2001; Gómez-Gras et al., 2016), marking the overfilling of the Cretaceous
181 flexural basin.

182 The eastern branch of the North Pyrenean Zone records the abrupt formation of an east-west
183 elongated deep basin in the early Campanian, where a former carbonate platform is overlain by
184 turbidites (Déramond et al., 1993; Billotte et al., 2005). As with the flexural basin in the South
185 Pyrenean Zone, turbidites were followed by southwestward delta progradation ending within
186 Campanian- Maastrichtian fluvial deposits that represent and general infilling of the basin (Plaziat,
187 1981; Bilotte et al., 2005). This deltaic setting supplied the western Aquitaine domain foredeep
188 basin with thick turbidites deposits (Ricateau and Villemin, 1973; Bilotte, 1985; Bilotte et al.,
189 2005).

190 Recent work on the Axial Zone 'Calcaires des Canons' Formation demonstrates a significant
191 deformation stage as soon as the Coniacian. It is materialized by the uplift and exposure of part of
192 the future Axial Zone while the former Iberian necking zone records important accommodation
193 space creation (Andrieu et al., 2021). The authors evokes two possible mechanisms: (1) onset of
194 the Iberia-Europe convergence or (2) a far-field origin. However, the debate remain open since
195 most of the large-scale kinematic studies and crustal-balanced cross-section restorations published
196 along the Pyrenees argue for an initiation of the syn-rift exhumed mantle domain closure starting
197 during the late Santonian resulting in the gentle inversion of the inherited syn-rift normal faults
198 (Mouthereau et al., 2014; Rougier et al., 2016;

199 **3. Data and Methodology**

200 In this section, we present the data used in the study (3.1), the facies and biostratigraphy analysis
201 realized on the dataset (respectively 3.2 and 3.3) and to finish, the methodology used (3.4 and 3.5).
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**3.1. Material**

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This study relies on the interpretation of five seismic reflection profiles to outline the large-scale basinal geometry of the NAD, the SAD and EAD, and correlation along a transect of wells to refine the sedimentary geometry and better assess the timing of deformation. For sake of clarity, seismic reflection profiles are labeled as: (i) Mas D'Azil - Lezat (REG_BRGM2011_FV07_PSTM), (ii) West-East Pailhes (REG_BRGM2011_80MIR01), (iii) Rousse - Theze (composite section with 89MLN04 – 82MLN10_101 – 82SVG09), (iv) Saoulous - Recurt (TOTAL_LaNoue67-87-71-CG-17S-MIGRATION) and to finish (v) Tournay - Trouley-Labarthe (composite section with TOTAL_RGF_PYRAMID_P1-91-Z4-RT-MIGRATION – TOTAL_RGF_PYRAMID PUT-87-PUT-10-RT – 85CAS-06-PSTM – 79CBE-04-PSTM).

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The seismic interpretation was carried out using the GVERSE® GeoGraphix®Suite that allows for each CDP of the seismic profile to export its X, Y, time, amplitude and depth values. Seismic/ well tie is obtained by using the time/depth curves acquired in the boreholes and derived from the recording of checkshots. The time/ depth conversion was performed using a mean velocity field calculated on the basis of all the time/depth data of the calibration wells, integrating their deviations

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Data used to calibrate the subsurface data (seismic reflection profile and well-logs) are of two kind: (i) literature (Bilotte, 1985 ; Serrano, 2001 ; Grool et al. 2018) and drilling report (Landes-de-siougos, Bazordan-1 and Dreuilh-4 and Dreuilh-5 wells), and (ii) new data (Rennes-les-Bains outcrop section, Calavante and Jacque-1 wells).

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To determine the basin scale sedimentary record of the Late Cretaceous sequence, each of these three domains (NAD, SAD, EAD) is represented by a type-section which we correlated in
terms of sequence stratigraphy to construct the framework of the study. This first task provides a
raw evolution of the Aquitaine domain throughout the Late Cretaceous, and will serve as guidelines
for the seismic reflection interpretation and boreholes correlations.

These four type-sections corresponds to Landes-de-Siougos well (NAD), Bazordan well (SAD) and Rennes-les-Bains (EAD) outcrop section plus Dreuilh-4 and Dreuilh-5 wells (EAD) (Figure 2). Landes-de-Siougos information (facies, biostratigraphy and well-logs) were published by Serrano (2001); Bazordan log was realized using drilling report information (lithologies, biostratigraphy, well-logs). Rennes-les-Bains was logged for the Cenomanian-Santonian interval, and completed with the work of Bilotte (1985) on biostratigraphy and facies analysis for the Campanian-Maastrichtian interval. A fourth section was added in the EAD, closer to our study area: Dreuilh-4 well (from Grool et al., 2018) completed with Dreuilh-5 Cenomanian to Coniacian cores description by Bilotte (1985).

Because resolution is different in seismic reflection profiles interpretation and well
correlations, two set of data calibration were used. A database of roughly 20 wells was established
for the seismic reflection profile interpretation, with, in each borehole, lithostratigraphy attribution
coherent with the up-to-date “Référentiel Lithostratigraphique Pyrénéens” produced in the RGF
research project leaded by BRGM. Orthogonal short-distance projection (less than 500m) was used
to plot these wells on the studied seismic reflection profiles.

242 **3.2. Sedimentary facies**

Description of sedimentary facies, in facies association (FA) and depositional environments
(DE) is a key for subsurface analysis. Espacially, it allows the calibration of well-logs, a
prerequisite for their correlation. In this scope, we (i) logged the Rennes-les-Bains outcrop section
(Figure. 2B) in detail for the Cenomanian to early Coniacian interval and used the outcrop descriptions of Bilotte (1985) for the late Coniacian to Maastrichtian interval; (ii) performed microfacies analysis on cuttings in Calavante-1 and Jacque-1 wells which we compared to well-logs signature, and (iii) beneficiated from the core description and associated well logs from Serrano (2001). Finally, we used sedimentary descriptions within the Bazordan-1 drilling report to complete the dataset. All these data are summarized in Figure 3.

Facies were determined from visual characterizations of components (skeletal and non-skeletal grains), textures, sedimentary structures, sorting, and grain size. We employed the classic wave zonation, using the fair-weather wave base to separate inner ramp and mid-ramp environments and the storm-wave base to separate mid-ramp and outer ramp environments (Plint, 2010). Depositional environment subdivisions and interpretations were based on the standard facies zones of Burchette and Wright (1992) for carbonate ramps, modified by Flügel (2010, p. 666).

Facies were grouped in seven facies associations (FA) (Figure 3). These encompass settings from distal to proximal: basin (FA1), outer ramp (FA2), mid-ramp (FA3), inner ramp including shoreface (FA4), lagoon (FA5), bay to estuary (FA6), and continental (FA7).

## 3.3. Biostratigraphy

Biostratigraphic markers in the Rennes-les-Bains sedimentary section from Bilotte (1985), in the NAD (Landes de Siougos Serrano 2001), in the well drilling report (Bazordan-1) and newly acquired (Calavante-1 and Jacque-1 wells; Figure 4) were used to date the depositional cycles and sequence stratigraphy surfaces. We used the biostratigraphic zones defined in the Late Cretaceous of the EAD by Bilotte (1985) for benthic and planktonic foraminifer, rudist, and ammonite faunas, and the zones of Massieux et al. (1985) for charophytes. These are presented in Supplementary Material 1.
269 **3.4. Sequence stratigraphy**

270 The depositional sequences considered in this study correspond to second-order stratigraphic
cycles, representing approximately 10 Myr (Guillocheau, 1991), and were identified and correlated
271 based on the evolution of the sedimentary environment. Each cycle is bounded by recognizable
272 and correlatable isochron surfaces. Depending on the depositional environment, two sorts of
273 surfaces might be chosen: maximum regressive surfaces or maximum flooding surfaces. The
274 maximum flooding surface is defined by the most landward position of the shoreline (Galloway,
275 1989; Catuneanu et al., 2009), and the maximum regressive surface corresponds to the most
276 basinward position of the shoreline (Catuneanu et al., 2009). We chose here to bound our cycles
277 by maximum regressive surfaces because they are more apparent in the utilized data. Besides these
278 two surfaces, we used the sequence boundaries as markers of accelerating base level fall (Vail,
279 1987).
280

281 **3.5. Reconstruction of sedimentary geometries**

282 We reconstructed sedimentary geometries at basin scale based on a multi-step approach.
283 First, sedimentary facies analysis leads to the recognition of several depositional profiles that are
284 bounded by major paleogeographic reorganizations. Second, well logs are calibrated through core
285 data, biostratigraphy, and well log signatures? electrofacies classification. Third, reference type
286 sections with a sound biostratigraphic framework are identified to recognize time surfaces to be
287 correlated between wells and in seismic profiles. Furthermore, these data provide refined ages for
288 the major sedimentary cycles described in the results. In the following sections, we first identify
289 and correlate the second-order cycles, based on the sequence stratigraphy framework of Serrano
290 (2001) and Bilotte (1985).
Five seismic reflection profiles and a well-log correlation transect are presented to constrain the Late Cretaceous geometries of the Aquitaine domain (Figure 2B). The Mas d'Azil-Lezat and the perpendicular West-East Pailhes seismic reflection profiles are in the EAD. The Clarens-St. Médard, Tournay-Trouley Labarthe, and Rousse-Theze seismic reflection profiles constrain geometries within and between the NAD and the SAD. The Tournay-Sadeillan well-log transect highlights in detail the Coniacian-Santonian sequence in the SAD.

## 4. Establishment and correlation of main depositional sequences

Following previous sequence stratigraphy studies (Serrano, 2001; Bilotte 1985, 2005), we defined three main transgressive and regressive cycles in the Upper Cretaceous depositional sequence. The first includes the Cenomanian and Turonian stages, the second the Coniacian and Santonian stages, and the third the Campanian and Maastrichtian stages (Figure 5). These three cycles are consistent with Serrano. (2001), and their correlation across the sedimentary type sections is shown in Figure 5. The different facies associations (FA) in these sedimentary sections, and the lines of evidence defining them, are assembled in tabular form in Figure 3.

### 4.1. Cenomanian-Turonian cycle

#### 4.1.1 Cenomanian transgression

In the NAD, the Cenomanian Pilo Formation corresponds to inner-ramp facies (F5a in Figure 3) while in the SAD, the Recurt Formation consists of interlayered volcanic and outer ramp carbonate deposits (Figure 5) implying deeper environments. Maximum flooding surface is well recorded in the NAD with outer-ramp facies deposition (F2a in Figure 3 and Figure 5), while? relative aggradation of these same facies in Bazordan-1 well smooths the flooding peak.
In the EAD, the early Cenomanian is also characterized by the deposition of inner-ramp facies as outlined by the Dreuilh-5 cores and the Rennes-les-Bains section that respectively show detrital muddy limestones and grainstone with lagoon fauna (F5a and F4a facies in Figure 3 and Figure 5). These facies correspond to the Calcaires à Préalvéolines et Caprines Formation (Figure 5; Dreuilh-4 well) and the Calcaires d'Escudiés Formation (Figure 5; Rennes-les-Bains). There, retrogradation culminates with the deposition of rudists biconstructions (F4b in Dreuilh-5 core Figure 3 and Figure 5) and condensed surface caping outer-ramp marls of the Fontaine Salée Formation (F2a in Rennes les bains in Figure 3 and Figure 5).

### 4.1.2. Turonian regression

The Turonian corresponds to the gradual return of shallow carbonate platform deposits, as part of the cycle's regression. In the NAD and SAD, it is respectively materialized by the Jouansalles Formation (NAD) and the uppermost part of the Recurt Formation (SAD), which regression culminates with development of inner ramp deposits (shoreface F4a in Landes de Siougos; lagoon F5a in Bazordan-1 well and Figure 5). As suggested by the Dreuilh-4 well, in the EAD the regression culminates within the Calcaires à Rudistes du Bois de la Forêt Formation (inner-ramp facies in Bilotte, 1985, p160). In contrast, the Grès de la Sals Formation in Rennes-les-Bains attests for inner delta environment with the deposition of conglomeratic level (F6-7a in Figure 3 and Figure 5).

## 4.2 Coniacian-Santonian cycle

### 4.2.1 Coniacian transgression

The Coniacian marks a general retrogradation attested by the transition to mid-to-outer-ramp deposits (F3c facies in Figure 3) over the Turonian deposits in the NAD (Landes-de-Siougos and
334 Bazordan wells figure 5). In the EAD, transgression is sharp passing from inner-ramp deposits
335 (Dreuilh-5 and 4 well, Figure. 5) and shoreface deposits (Rennes les basins section in Figure 5) to
336 outer ramp to basin deposits deposits.

337 4.2.2 Santonian regression

338 The Santonian regression initiates in facies similar to the underlying Coniacian open-marine
deposits (outer to mid-ramp). They grade and culminate in the SAD and NAD to shoreface deposits
339 (F4a in Figure 3, Figure 5 in Landes-de-Siougos wellFigure 5), as well as in the Rennes-les-Bains
340 section where thick sandstone beds are interpreted as the Santonian MRS (F6d facies in Figure 3,
341 Figure 5 in Uppermost Montagne des Cornes Formation). Thick aggradation of outer-ramp deposits
342 within Dreuilh-4 well (Flysch Senonien Formation) enables a clear indentification of the MRS
343 which was picked at the transition between basinal marl deposits and detrital turbidites of the
344 Plagne Formation (Figure 5).
345

346 4.3 Campanian-Maastrichtian cycle

347 The Campanian-Maastrichtian cycle is asymmetric since only the regressive interval is
348 preserved (Figure 5). It implies that the base of the cycle corresponds to a stacked MRS/MFS as
349 suggested by Bilotte (1985). Early Campanian deposits correspond in both the NAD to i) Pé-Marie
350 Formation's outer ramp facies (F2a/F3c Figure 3; Figure 5 in, Landes-de-Siougos well), and SAD
351 to: and (ii) basin deposits of the Flysch Campanien Formation in Bazordan well (Figure 5). In the
352 EAD, The Grès de Labarre Formation records a progressive transition from estuarine to fluvial
353 depositional environments (Figure 5). A similar evolution is observed in the Rennes-les-Bains
354 section with the deposition of the fluvial Grès d'Alet Formation (Campanian) over the marine
355 deposits, passing to a succession of red marls (floodplain deposits) (Figure 5).
The existence of this Campanian-Maastrichtian cycle slightly differs from Bilotte (1985, 356 357 2005) sequential framework. Indeed, in the Eastern Pyrenees, this author consider two 358 megasequences: Cenomanian to Turonian and Coniacian to Maastrichtian. Yet, as mentioned 359 previously (section 2.4), a significant event is recorded in the Eastern Pyrenees during the early 360 Campanian. It is marked by a major reorganization of the Aquitaine domain paleogeography which, 361 in our opinion requires to consider the Campanian-Maastrichtian interval as a distinct 362 megasequence. 363

# 5. Analysis of Late Cretaceous sedimentary geometries

## 5.1. East Aquitaine Domain sedimentary geometries

Two seismic reflection profiles were interpreted within this domain: Mas d'Azil-Lezat and 365 West-East Pailhes. These lines are roughly orthogonal and intersect at the Pailhes-1 well (Figure 366 2B). 367

### 5.1.1. Se Well calibrations

The Mas d'Azil-Lezat seismic reflection profile (FV07 line) is 42 km long and oriented 369 north-south. It is calibrated at its southern end by the Mas d'Azil-1 well (Figure 6 370 Figure 7), which 371 penetrates thick Campanian calcareous turbidites (Plagne Fm. in Figure 6) passing upward to 372 Maastrichtian deltaic to fluvial deposits (Grès de Labarre and Marnes d'Auzas formations in Figure 373 6). To the north, the lowermost Cretaceous deposits in the Pailhes-1 well (Figure 6) are detrital 374 carbonate facies of Coniacian-Santonian age that directly overlie Permian strata (Figure 6). These 375 cretaceous deposits consist of alternating carbonate facies (packstone to wackestone) and sandstone 376 to siltstone rich in lignite in a sequence typical of deltaic progradation. This sequence is capped by 377
377 Campanian turbidites that grade upward to Maastrichtian marginal marine to continental facies
378 (Grès de Labarre and Marnes d'Auzas Formations; Bilotte, 1990). Further north, the Lezat-1 well
379 (Figure 6) intersects the Marnes d'Auzas Formation, resting unconformably on Paleozoic dolomite.

380 **5.1.2. Review of existing interpretations in the area**

381 Closed to the Mas d'Azil – Lezat seismic reflection profile), recent seismic interpretation
382 from Ford et al. (2016) evidences the onset of the Sub-Pyrenean Thruts formation during the
383 earliest Eocene, scelled wright before the Oligocene to the south of Pailhes-1 well. Furthermore,
384 they document the wedge geometry of the Coniacian to Maastrichtian sequences characterized by
385 southward thickening and northward thinning of both sequences. There, the authors demonstrate
386 the development of the Campanian-Maastrichtian flexural basin through acceleration of tectonic
387 subsidence in the Sub-Pyrenean Zone, high sedimentation rates with deep-water facies displaying
388 northward onlap onto the Paleozoic basement.

389 **5.1.3. EAD sedimentary geometries**

390 The main structure visible on the Mas d'Azil-Lezat profile is the Sub-Pyrenean Frontal
391 Thrust, passing through Mas d'Azil well which separates two radically different units (Figure 6
392 and Figure 7). In the SPT footwall, the profile reveals a southward-opening wedge-shaped basin in
393 which Cenomanian and Turonian deposits are absent (Figure 7). The basal Cretaceous deposits are
394 made up of a Coniacian-Santonian sequence that onlaps an erosive surface over Permian deposits
395 (north of the Pailhes-1 well in Figure 7). The Coniacian-Santonian sequence is characterized by a
396 significant southward thickness increase as documented in the Pailhes-1 well (Figure 7). The
397 multiple onlap markers attest of the gradual northward extent of the area of sedimentation (Figure
398 7). The Campanian sequence shows a similar trend with significant thickness reduction northward
399 and pinch a few kilometers to the north of Pailhes-1 well, whereas the Maastrichtian deposits
400 extend northward nearly the full length of the profile (Figure 7). The Maastrichtian sequence
401 follows the northward thickness reduction trend with thickness twice thinner in Lezat-1 than in
402 Pailhes-1 well (Figure 7) and onlap onto Paleozoic basement rocks).

403 The Cenozoic sequence rests unconformably above the Cretaceous sequence, as shown by
toplaps in the Cretaceous series near the Paihles-1 well (Figure 7). Whereas the Maastrichtian
404 deposits are continental facies that mark the final infilling of Cretaceous basin, the Cenozoic facies
405 record a resuming of shallow marine deposits (carbonate platform) followed by an upward
406 transition to continental deposits (Molasses) These Cenozoic deposits consist of syn- to post-
407 orogenic sediments, and the sequence displays a massive increase in thickness with clear onlaps in
408 the vicinity of the Lezat-1 well and northward (Figure 7). The area of sedimentation thus enlarges
409 northward, associated with an increase of preservation of continental deposits during syn orogenic
410 phase ( Al Reda et al., 2021) (Figure 6 and Figure 7).

412 The West-East Pailhes seismic reflection profile (Figure 8) is 40 km long and is oriented
413 roughly N110-120° (Figure 2B). north of the Sub Pyrenean thrust. As with the Mas d'Azil-Lezat
414 profile, the Cenomanian-Turonian sequence is absent, and Coniacian deposits rest directly upon a
415 Paleozoic substratum. The Coniacian-Santonian deltaic interval is pinching out sharply to the west
416 and is characterized by successive westward progradations highlighted by downlap terminations
417 (Figure 8). The transition between the Rennes-les-Bains outcrop section that displays a thick
418 proximal deltaic sequence in the Late Santonian (Formation de la Montagne des Cornes) passing
419 in Dreuilh-4 to distal open ramp to basinal deposits (Senonian Flysch) is a proximal equivalent of
420 the overall progradation trend observed in this seismic reflection profile (Figure 5; Figure 8).

421 The westward regressive trend culminates in the Campanian sequence. The Campanian
422 deposits are more constant in thickness than the Coniacian-Santonian sequence (Figure 8) and show
evidence of successive progradation prisms expressed by large scale clinoforms filling the available space, expressed by turbidites at the bottomsets and front to fluvial deposits in topset as evidenced in the Pailhes-1 well The truncations below the Campanian surface east of Pailhes-1 well suggest an erosion prior to the onlap by turbidites facies, which we interpret as a submarine erosion related to the installation of a turbiditic system. The Maastrichtian sequence likewise has a relatively constant thickness except at the eastern end of the profile where the thickness decreases, which is partly explained by the erosional truncatures visible below the Tertiary section there (Figure 8). This more isopachous trend is consistent with the deposition of continental (coastal to alluvial plain) deposits over a large extension, witnessing a significant continental-deltaic depositional profile elongation in Maastrichtian times. Very little thickness change compared to the Campanian sequence suggests poor subsidence variations during the Campanian-Maastrichtian.

To sum up, these two profiles lie in a domain where the initiation of a flexural basin is well documented (see previous sections), and they shed light on the major structural and basin evolution trends of the EAD. First, the abrupt thickening of the Upper Cretaceous sequence south of the Sub Pyrenean Pyrenean Frontal Thrust suggests that this steeply dipping structure controls the domains on either side of the thrust (Figure 7). The thickening of the Coniacian-Santonian and Campanian sequences toward the North Pyrenean Frontal Thrust indicates the development of a more rapidly subsiding southern domain. The Maastrichtian sequence appears to be more isopachous except where it is significantly eroded by the Cenozoic sequence (Figure 7). The available space created from Coniacian-Santonian (onset of sedimentation) is rapidly filled up during Campanian interval by deltaic progradations (Figure 8). This pattern is very likely related to a sediment supply exceeding the creation of available space. The nearly flat geometries within the Maastrichtian sequence are consistent with the development of a widespread coastal plain landward of the delta
446 front observed in the Campanian sequence, signifying the final infilling of the basin. Moreover,
447 poor thickness variations between the Campanian and Maastrichtian deposits implicates very little
448 deformations in Maastrichtian times.

449 These two profiles document the formation of a Late Cretaceous flexural basin in the EAD.
450 The process consists of drowning in the Campanian and emplacement of turbidites in the distal
451 area (the Pailhes-1 well, Figure 7), along with progressive significant northward onlap
452 (sedimentary area enlargement, Figure 7) in a basin undergoing significant southward thickening.
453 This available space is rapidly filled up by westward deltaic progradations in an E-W elongated
454 trough. This interpretation is consistent with current knowledge of the eastern Pyrenees, which
455 entails a Campanian age for the formation of the flexural basin (e.g., Bilotte 1985 Christophoul et
456 al. 2003; Ford et al., 2016; Rougier et al., 2016). The global thickness reduction and erosive trend
457 at the northernmost end of the profile indicate the outer-rim of the flexural basin characterized by
458 erosion of the underlying facies and very weak subsidence. The Campanian pinchout was located
459 a few kilometers north of the Pailhes-1 well where the facies sequence thins (Figure 7). The
460 Cenozoic flexural basin outer rim, in contrast, was roughly 30 km to the north (Figure 7). The
461 Cenozoic flexural basin thus has a different geometry as indicated by the major enlargement of
462 sedimentation area compared to the Campanian-Maastrichtian flexural basin (Figure 7). The
463 implication is a relative narrow Late Cretaceous flexural basin and a much wider and thicker
464 Cenozoic flexural basin. The reason for this discrepancy is addressed in section 7.

465 The three reflection profiles in the following section shed light on the sedimentary geometry
466 to the west of this flexural basin as well as the evolution of the narrow Coniacian-Santonian
467 sequence and
468

**5.2. South and North Aquitaine Domain sedimentary geometries**

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Three N-S trending seismic reflection profiles were interpreted within these two domains: Rousse-Theze, Clarens-St. Médard, and Tournay-Trouley Labarthe (Figure 2B). We also constructed a well-correlation profile near the Tournay-Trouley Labarthe profile.

472

**5.2.1. Seismic line and well calibration**

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The Rousse-Theze seismic reflection profile (Figure 2B) is 40 km long and is calibrated from south to north by the Rousse-2, Lons-1, Theze-1, and Theze-301 wells (Figure 9, Supplementary Material 2 for details). The southern domain, i.e Lons-1 to Rousse-2 area is represented by deep basin deposits which consists from base to top to the base of slope Brèches de Soumoulou Formation (Cenomanian-Turonian), passing to turbidites of the Hopital Saint Blaise Formation (Coniacian-Santonian) with clear northward thinning. Turbidites deposits pursue in Campanian (Campanian Flysch Fm.) times and Maastrichtian times (Marnes d'Auzas Fm.) and they migrate northward onto the platform. Lons-1 well itself shows no sign of sedimentary activity from Cenomanian to Santonian times as the first deposits correspond to Campanian turbidites. On the seismic reflection profile, south of Lons-1 well (included), the turbidites deposits were grouped as undifferentiate Late Cretaceous because of the lack of lithology contrasts and subsequent difficulty to correlate a reflector. The platform itself is occupied from the Cenomanian to Santonian by shallow carbontaes deposits (Pilo Fm. And Jouansalles Fm.) passing to Campanian outer-ramp facies and Maastrichtian outer-ramp to turbidites deposits (Marnes d'Auzas Fm. To Couches e Dumes. Fm).

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The Clarens-St. Médard seismic reflection profile is 40 km long and located in the eastern part of the SAD (Figure 2B). It is calibrated by the Recurt-1 and Saoulous-1 wells (Figure 11,
490 Supplementary Material 2 for details). The largest structure is the Gensac salt-anticline, bounded
491 in its northen edge by a south-dipping thrust. The southern domain (Recurt-1 well) spans deposition
492 of Cenomanian subaerial to shallow volcano-sedimentary deposits of the Recurt Fm passing
493 vertically to shallow mixt detrital-carbonates of the Gensac Fm (Coniacian-Santonian), and basin
494 turbidites (Campanian Flysch Marneux Fm.) to prodelta marls (Maastrichtian Marnes d'Auzas
495 Fm.). More proximal facies are encountered in the northern domain (Saloulous-1 well) where
496 succession is typical of the NAD (Jouansalles Fm – Pé-Marie Fm. – Couches de Dumes Fm.) except
497 that Cenomanian deposits are absent.
The Tournay–Trouley Labarthe seismic reflection profile (Figure 2B) is 25 km long and is calibrated from south to north by the Tournay-1, Calavante-1, Osmets-1, and Troulay-Labarthe-1 wells (Figure 13, Supplementary Material 2 for details). It is bordered to the north by the southern flank of the Antin-Maubourguet anticline (Figure 2B). This structure corresponds to the northern edge of the Early Cretaceous Arzacq rift basin, separating it from the Tartas basin (Issautier et al., 2020), and it corresponded during the Albian rifting stage to a salt diapir that raised a ridge between the two basins (Mauriaud, 1987). Similarly to Clarens-St Médard profile, a southern and northern domain might be distinguished. The first one spans deposition of Turonian Jouansalles Fm. over the Jurassic (gap of the Cenomanian), passing to turbidites and pro-delta deposits of respectively the Flysch marneux Fm. (Campanian) and Marnes d'Auzas Fm. (Maastrichtian). The northern domain shows typical shallow carbonate NAD succession except that JOuansalles Fm. is replaced by mict detrital-carbonates of the Gensac Fm. (Turonian to Santonian). In order to refine the stratigraphic resolution given by seismic profiles, we constructed a well-correlation profile parallel to the Tournay–Trouley Labarthe seismic reflection profile (Figure 2B and Figure 15). It integrates five supplementary wells (from south to north, Goudon-1, Jacque-1, Lameac-1, Mazerolles-1, and Sadeillan-1) and it must be noticed that the Lameac-1 well is placed on top of an Albian salt diapir: the Laméac diapir, part of the Antin-Maubourguet anticline. 5.2.2. Review of existing interpretations in the area

Rougier et al. (2016) provide in the surrounding of our Récurt-Saoulous profile a study highlighting the geometry of the northern Bigorre Basin (NPZ), Comminges Basin (Sub-Pyrenean Zone) and Aquitaine basin. Their seismic profile interpretation and crustal-balanced cross-section argue for an Eocene activity of the SPT, while the NPFT was active during the development of the
Late Cretaceous flexural basin as deformation is sealed by late Maastrichtian marls. Their study demonstrates the northward and diachronous migration of dépocentre from the NPZ to the Aquitaine Platform during the two orogenic stages (Late Cretaceous and Eocene) separated by a tectonic-quiescent stage (Ford et al., 2016; Rougier et al., 2016; Grool et al., 2018). In the same area, Fillon et al. (2021) propose, based on unpublished data (Y.Philippe, 1998) a large-scale interpretation of the Recurt-Saoulous profile. It underlines the existence of a salt-anticline whose top is calibrated by the Recurt-1 well. Thickness variations from either side of the anticline suggest that this salt-structure developed mainly during the Early Cretaceous rifting and was subsequently slightly reactivated in latest Cretaceous times before secondary growth in Eocene times (compression paroxysm). They also evidence that the Late Cretaceous sequence thins northward of the Saoulous-1 well but due to the large stratigraphic interval represented, age of the sequence preserved northward is not precisely provided. In the same area, seismic (LR06 seismic profile) and well interpretation of Espurt et al. (2019) evidences the growth of the Gensac salt anticline during the Early Cretaceous and Campanian to Eocene as witnessed by the associated depocenter migration from either side of the structure.

Finally, the Tournay - Trouley-Labarthe profile was recently presented in Fillon et al. (2020). The large-scale interpretation indicates the existence of a salt-anticline in the vicinity of Osmets well (Osmets- Laméac structure, Serrano et al., 2006), which growth occurs in Early Cretaceous times. This interpretation is witnessed by the developement of mini-salt basins from either side of the salt-anticline. The Late Cretaceous shows a slight thickness increase from North to South, suggesting the end of halokinesis at that time. Late Cretaceous deposits show thickness increase in the vicinity of the NPFT, interpreted as a low angle thrust.
543

5.2.2. Sedimentary geometry and main Late Cretaceous cycles

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In the western part of the SAD (Rousse-Theze profile, Figure 2B), two domains can be distinguished: Arzacq basin to the north and Grand Rieu High to the south (Figure 10). In the Arzacq basin, the Cenomanian to Santonian deposits form an undifferentiated thin sequence, not exceeding 100 ms two-way traveltime thickness (Figure 10). It thus appears that this site was a relatively homogeneous carbonate ramp until the Campanian. The Campanian-Maastrichtian sequence is more argillaceous and open-marine (Figure 9), signaling a deepening of the carbonate platform, and clinoforms are visible in the vicinity of the Theze-301 well attesting for a more abrupt slope at the transition to basinal facies that develop to the south. To the north of the Lons-1 well, several possible erosional truncations below the Cenozoic deposits imply that the thickness of the Campanian-Maastrichtian sequence is reduced by Cenozoic erosion (Figure 10c).

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The Grand Rieu domain, in contrast, contains an undifferentiated homogeneous Upper Cretaceous sequence made up of turbidites (Figure 10). The only exception is the presence of breccias at the base of the sequence, interpreted in the literature as base-of-slope breccias (Serrano et al., 2006). This interpretation supports the existence of a steep slope between the carbonate platform in the Arzacq Basin and the Grand Rieu domain breccias, which is now slightly inverted as a thrust since the Grand-Rieu turbidite basin gently overthrusts the Arzacq Basin. Thus, the Grand Rieu high represents an unrifted crustal domain (Angrand et al., 2018) that acted as a significant paleogeographic limit separating the Aquitaine Basin and the North Pyrenean Zone.

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In the central and eastern part of the SAD (Figure 2), the sedimentary record is more easily differentiated. For simplicity, we will describe the sequences in the Clarens-St. Médard (Figure 12) and Tournay-Trouley Labarthe (Figure 14) profiles together.
The Cenomanian-Turonian cycle in the Clarens-St. Médard profile (Figure 12) is represented by relative shallow succession of volcano-sedimentary facies (see Recurt-1 well in Figure 11) yet in a relative thick sequence which pinchout northward as witnessed by the Coniacian-Santonian sequence of Saoulous-1 well resting directly on Lower Cretaceous deposits. In the Tournay-Trouley Labarthe profile (Figure 14), the Cenomanian-Turonian cycle is restricted to the north of the Calavante well, where it corresponds to very shallow to marginal deposits (sandy marls rich in lignite and organic debris), thinning southward.

An unconformity separates the Cenomanian-Turonian and Coniacian-Santonian sequences, as illustrated by the toplap terminations in the Clarens-St. Médard profile (Figure 12C). There, Coniacian deposits are transgressive to the north of the Saoulous-1 well witnessed by onlap termination in the vicinity of Saoulous-1 well, (Figure 12C). Deposits correspond to shallow carbonate of the Jouansalles Formation. To the south of the Recurt-1 well (Figure 12), a thickening of the Coniacian-Santonian sequence is visible (Gensac Formation), which together with the unconformity on top of the Turonian sequence suggest the onset of a salt anticline at the Recurt locality, laterally interpreted as the Gensac anticline. Deposits correspond in the Recurt-area to mixt detrital-carbonate proximal facies of the Gensac Formation.

In contrast, the Tournay-Trouley Labarthe seismic reflection profile (Figure 14), has a relatively homogeneous trend in the central SAD despite a slight southward increase in thickness of the Late Cretaceous deposits. Between the Calavante-1 and Osmets-1 wells, numerous downlap terminations are visible in the Coniacian-Santonian sequence (Figure 14C), corresponding to mainly southward-prograding sedimentary prisms. In the Sub Pyrenean Zone (south of the Sub-Pyrenean Thrust), the deposits are composed of undifferentiated turbidites (Figure 14).
As with the Coniacian-Santonian sequence, the Campanian cycle forms a wedge-shaped basin in the central and eastern SAD, opening toward the south in both the Clarens-St. Médard and Tournay-Trouley Labarthe profiles (Figure 12 and Figure 14). Onlap terminations in the Saoulous-1 area onto a slightly unconformable base of Campanian surface illustrate this notable early Campanian transgression (Figure 12 C-D). The development of this Campanian basin is associated with a relatively sharp drowning of the underlying Coniacian-Santonian platform overlain by flysch facies over the South Aquitaine Domain (Figure 11 and Figure 13).

In the distal part of Tournay-Trouley Labarthe seismic reflection profilethick series of deep water facies are onlapping to the north (north of Calavante), they represent the infilling of the deeper more subsident part of the basin and onlap of the slope during infilling. (Figure 13, Figure 14).

The Maastrichtian sequence is relatively homogeneous throughout the two seismic reflection profiles (Figure 12 and Figure 14). A slight reduction in thickness is locally apparent and results from Cenozoic erosion, as indicated by erosional truncations in Tournay-Trouley Labarthe profile (Figure 14C). In the latter, downlaps and toplaps are visible between the Calavante-1 and Osmets-1 wells, suggesting gently inclined progradations (Figure 14C). Depositional environment associated with these low-angle clinoformes correspond to mid-ramp carbonate facies (FA3) in the Osmets-1 and Trouley-Labarthe-1 wells (Figure 13), compatible with a gently-inclined ramp profile.

In the Tournay-Trouley Labarthe seismic reflection profile, the Cenozoic sequences in the Tournay-1 and Calavante-1 wells are dramatically thicker than the Campanian-Maastrichtian sequence (800 m vs. 300 m in the Calavante-1 well, Figure 13). The presence of outerramp, turbidite facies and distal prodelta facies marks the development of the Eocene foredeep basin.
To better assess the age and nature of the Upper cretaceous evolution in the Tournay-Trouley
Labarthe seismic reflection profile (Figure 14) we constructed a well-correlation profile shown in
Figure 15.. Although very thin (see thicknesses within Figure 15), the Cenomanian-Turonian
transgression-regression cycle is evident as Cenomanian transgression between the Osmets-1 and
Sadeillan-1 wells, and the initial depositional environment corresponds to the fluvial and tide-
dominated deltaic facies (FA6 in Figure 3) and probable bay setting (lithofacies F6b in Figure 3).
Lagoon deposits (F5a in Figure 3) between the Osmets-1 and Mazerolles-1 wells are consistent
with this transgression. The maximum flooding surface is within the mid-ramp facies (FA3 in
Figure 3) and indicates the presence of a trough between the Osmets-1 and Sadeillan-1 wells
(Figure 15). The Turonian regression is preserved as a widespread inner ramp facies (FA5 and FA6
in Figure 3), except in the distal profile (Tournay-1 and Calavane-1 wells) and in Sadeillan-1 well
where uppermost Santonian-lowermost Campanian deposits rest directly on the Cenomanian
maximum flooding surface (Figure 15).

The Coniacian-Santonian cycle, is characterized in the Osmets-1 to Mazerolles-1 wells by
aggradation of the underlying inner ramp facies (F5a in Figure 3). Signs of the onset of this
transgressive cycle are visible only in the most distal wells (Goudon-1, Calavante-1, and Tournay-
1). In the Tournay-1 well, these shoals are overlainby mid-ramp facies (F4a and F3a in Figure 3)
that mark the top Coniacian maximum flooding surface (Figure 15). The Coniacian transgressive
sequence is several times as thick at the south end of the profile (Osmets-1 to Tournay-1 wells) as
at the north end (Trouley Labarthe-1 to Lameac-1 wells), 190 m vs. 45 m (Figure 15). Note that in
the two northernmost wells (Mazerolles-1 and Sadeillan-1), the Coniacian sequence is removed by
erosion at a Santonian sequence boundary (Figure 15). Indeed, the Santonian regressive cycle
begins above the Coniacian maximum flooding surface within the inner ramp facies (F5a in Figure
3) in the Calavante-1 to Lameac-1 wells and in the mid-ramp facies (F3a) in the Tournay-1 well.

The regression abruptly accelerates as indicated by the sharp progradation of shallow proximal facies (F6a in Figure 3) throughout most of the profile (Figure 15). The absence of this downward shift in the most proximal area (Sadeillan-1 well) argues for subaerial exposure following an erosional stage that removed the Turonian and Coniacian deposits. In the Calavante-1 well, the absence of the radioactive sandstone lithofacies (Figure 15) suggests a lateral shift toward the inner ramp facies (F5a in Figure 3). The top these sandstone marks the Santonian maximum regressive surface and the subsequent transition into the Campanian (Figure 15). The Santonian regressive cycle is represented by a thick sequence in the distal area (nearly 500 m thick in the Goudon-1 well), but is poorly preserved (84 m thick) in the Lameac-1 well and absent in the most proximal wells (Figure 15).
The Campanian transgression begins within inner ramp facies (F5a in Figure 3) as recorded in the proximal area (Figure 15; Goudon-1 to Sadeillan-1 wells); these deposits rest unconformably above Cenomanian deposits in the Sadeillan-1 well. The transgression accelerates with the development of mid-ramp (F3b and F3c in Figure 3) to outer ramp facies (F2a in Figure 3) between the Osmets-1 and Sadeillan-1 wells, passing in the distal part of the profile to basinal turbidites (Fa1 in Figure 3; Goudon-1 to Tournay-1 wells). The exact nature of Campanian-Maastrichtian facies in the Goudon-1 well is uncertain, but they probably represent the transition between the outer ramp (FA2 in Figure 3 and basin (FA1 in Figure 3 facies. It must be noticed that transition from platform to basin depositional environment lies in relative short-distances, i.e 6km (Figure 15). The reduction in thickness between the Tournay-1 and Calavante-1 wells reflects the northward onlap of the turbidites as seen on the adjacent seismic reflection profile (Figure 14). The Maastrichtian sequence in the well-correlation profile confirms the progradations observed in the Tournay-Trouley Labarthe seismic reflection profile (Figure 14)
as indicated by the southward progression of the mid-ramp facies
toward the Goudon-1 well (Figure 15). This progradation appears to
be responsible for the triggering of a mega-turbidite observed in the
Calavante-1 well (Figure 4 and Figure 15), possibly explained by fault
activity to the south of the profile (i.e SPFT).
6. Tectono-sedimentary
evolution of the Late Cretaceous Aquitaine Basin

Here we use the aforementioned stratigraphic descriptions to present an integrated history of
the Aquitaine Basin during the Late Cretaceous. Figure 16 corresponds to schematic diagram of
the stratigraphic record and the inferred tectonic events that explain the sedimentary record of the
Upper Cretaceous (Figure 16).

672
673
**6.1. Cenomanian to Turonian: Tectonic stability following the Early Cretaceous**
**rifting stage**

The overall Cenomanian-Turonian sequence is marked by the absence of tectonic evidences
nor significant halokinesis. This time-period is thus interpreted as a relatively stable tectonic period
following the antecedent Early Cretaceous rifting stage. Nevertheless, the occurrence and
preservation of this sequence is uniuall. In the western portion of the SAD, clear platform-basin
geometry exists in the Rousse-Theze seismic reflection profile (Figure 10). Here, the shallow
carbonate facies of the Pilo Formation develop in a relative isopachous sequence up to the north of
the Lons-1 well which lacks any Cenomanian to Santonian deposits and marks a by-pass area up
to Campanian times. In Rousse-2 well, base of slope breccias develop (Serrano et al., 2006). Their
clasts composition suggest that the sedimentary pile has been denudated until reaching the Triassic
sediments (clasts composed of Early Cretaceous, Jurassic and Triassic rocks, with a majority of late Jurassic clasts). This pattern implicates the existence of a significant slope in between the Rousse-2 well and the domain localized to the north of the Lons-1 well. This talus is passively filled with Late Cretaceous turbidites. This description supports the presence of a wide portion of slope, acting as a by pass zone, that spreads from Pau to Tarbes (AGSO and BRGM, 2018).

In the central and eastern SAD (Figure 2), the localization and preservation of Cenomanian-Turonian deposits is more complex. The Clarens-Saint Médard seismic reflection profile (eastern SAD, Figure 12) shows the development of relative shallow carbonate during the late Cenomanian to Turonian. This sequence thins towards the north of the SAD, where it does not exist anymore in Saoulous-1 well. In the same area, Rougier et al. (2016) suggest an exposure domain north of the SPT that might explain the absence of Cenomanian-Turonian deposits described in our work. It is not clear whether the sequence was never deposited nor eroded prior to Coniacian-Santonian truncation. The first scenario implies the development of an isostatic rebound of the northern SAD during the post-rift stage while subsidence develops in the hyperextended domain of the former Early Cretaceous rift system. The second hypothesis implies a deformation stage or eustatic fall prior to the deposition of the Coniacian-Santonian sequence. The Tournay - Troulay Labarthe seismic reflection (central SAD, Figure 14a,d) and well-correlation (Figure 15) profiles indicate the absence of Cenomanian-Turonian deposits in their southern edge (Calavante area). The very proximal facies (base of Jacque-1 well) disappear rapidly toward the south, conjugated to the significant thickness reduction in the vicinity of the Calavante-1 well draw a southward onlap toward an exposed domain localized to the north of the SPZ (Figure 15 and Figure 16). This paleogeography supports the statement of Rougier et al. (2016) and implies the presence of a
complex topography inherited from the former rifting stage, i.e. salt structures and syn-rift
structural network.

## 6.2. Coniacian to Santonian: Tilting of the eastern South Aquitaine Domain

The Coniacian-Santonian sequence is time of a major paleogeographic reorganization clearly evidenced in the EAD and central/eastern SAD where this sequence is regionally characterized by a significant wedge shape with southward increase in thickness (Figure 16).

Similarly to the Cenomanian-Turonian sequence, the organization of the Coniacian-Santonian deposits is different within the western and central/eastern SAD (Figure 16).

In the first one, the Rousse-Theze seismic reflection profile depicts a thin and shallow carbonate ramp in a relative isopachous sequence onto the Arzacq domain. On the contrary, the northern NPZ spans the deposition of turbidites deposits. This pattern signifies that the late syn-rift platform-basin transition remains steep and short, as for the underlying Cenomanian-Turonian deposits. Yet, the deposition of turbidites over the base-of slope breccias witness a northward migration of the basinal series and the gradual infill of the basin. To sum-up, the western SAD displays a gently southward-thickening geometry, reaching its maximum thickness in the inverted Mauléon basin (Figure 16, left). This north to south gradient of subsidence seems to correlate with the structural inheritance of the Early Cretaceous rifting stage as the maximum of subsidence are encountered above the remaining thinnest continental-crust (i.e Mauléon hyperextended Basin). Our model highlights the key role of long-term sustained post-rift thermal subsidence proposed by Angrand et al. (2018).

In contrast, In both the EAD and the central/eastern SAD the Coniacian-Santonian sequence corresponds to a wedge-shaped basin (Figure 16). The latter is slightly tilted by late Santonian to
early Campanian deformation as indicated by erosional truncations below the Campanian sequence in the Clarens-St. Médard seismic reflection profile (Figure 12). This wedge-shaped geometry is particularly marked in the Clarens-Saint Médard (Figure 8) and Tournay-Trouley Labarthe (Figure 12 and Figure 14) seismic reflection profiles and the well-log correlation transect (Figure 15). Here, the Coniacian deposits are five times thicker in the distal portion than in the proximal domain. In contrast, we demonstrated that the Cenomanian-Turonian cycle is preserved in the northern part of the Tournay-Trouley profile (Figure 15) and in the Recurt area (Figure 12), conformly to Rougier et al. (2016). Both depocenters are separated by exposure surfaces (Figure 16), evidenced by southward onlaps in Tournay-Troulay Labarthe seismic reflection profile. There, this pattern is crucial because it imposes a relative flat Cenomanian-Turonian depositional profile, followed by non-symmetric available space creation. Such a mechanism can only be explained by basinward tilting of the substratum during the Coniacian. This tilting increases the available space in the southeastern SAD, while the northeastern SAD and the NAD are either slightly subsiding or subaerially exposed (Figure 16)1; The following Santonian regressive cycle shows a thickness trend similar to the one identified within the Coniacian. Nevertheless, a lowstand systems tract develops in the upper part of the sequence. It is highlighted by an abrupt influx of detrital material within the carbonate platform (Figure 15). This lowstand system increases in thickness southward while the proximal part (in the Sadeillan-1 well) is exposed, eroding the underlying Coniacian deposits. These observations indicate a suppression of available space in the proximal zone while the most distal zone record the deposition of a large thicknesses of sediments (Figure 16).
Previous studies (Mauriaud, 1987; Serrano et al., 2006) have emphasized a very slight reactivation of the Laméac salt diapir (Figure 2) during the Coniacian-Santonian. This is consistent with our observations of a thickness reduction on top of this same diapir. One could evoke the role of salt-anticline growth in Coniacian-Santonian time as a parameter driving the basinward tilt of the platform, although Mazerolles-1 and Sadeillan-1 wells are quite far from this diapir (10 km for Sadeillan-1) and they indicate the most extreme thickness reduction (Figure 15). Similarly, the zone where accommodation space is being created, i.e southern part of the profile (Tournay-1 and Calavante-1 wells) lacks any kind of halokinesis (Figure 14). Similarly, the Clarens-St. Médard profile (Figure 12) does not show any diapiric structures in the northern part of the SAD and yet the Coniacian-Santonian sequence still pinches out north of the Saoulous-1 well. Consequently, although diapir reactivation probably contributes to local thickness variations, it cannot be the only mechanism responsible for the large-scale trend: suppression of available space in the northernmost part of the SAD and creation of available space in the southern SAD. Such discrepancies between the western SAD and both the EAD – and the central/eastern SAD evoke a strong segmentation of the basin during the Coniacian-Santonian sequence.6.3. Campanian: Initiation of flexural subsidence

The Mas d'Azil-Lezat seismic reflection profile in the EAD (Figure 7) illustrates the development of a Campanian flexural basin as indicated by three features: (1) the Campanian turbidites resting on Santonian open marine facies, (2) the northward onlap of the Campanian sequence and (3) the unconformable relationship between the Santonian and the Campanian
769 sequences on the border of the peripheral relief that implies a tilting of the previous deposits prior
770 to the sedimentation of the Campanian sequence. The eastern SAD turbidites are clearly derived
771 from the EAD (Figure 5). This transport indicates the formation of a trough trending E-W, parallel
772 to the orogenic structure (Déramond et al., 1993; Billotte et al., 2005), with sediment shed from the
773 initial relief in the East Pyrenean segment of the orogenic belt (Vacherat et al., 2017 and references
774 therein).

775 In the entire SAD, the Campanian sequence differs from the Coniacian-Santonian one. It is
776 characterized in its northern part by thin and stable outer ramp deposits grading toward the North
777 Pyrenean Zone to thick deep-basin turbidites (Figure 16). The Tournay-Trouley Labarthe well-log
778 profile in the central SAD shows, as in the EAD, the platform-basin transition migrates
779 significantly northward onto the European margin during the Campanian, i.e., the Goudon-1 well
780 (Figure 15). The exactly same pattern is emphasized in the eastern SAD (Clarens ST Médard)
781 where Flysch deposits migrate northward onto the previous shallow carbonate formations (Figure
782 12). This newly formed flexural basin has relatively narrow dimensions, being limited in its
783 northen portion to a few kilometers north of SPFT

784 Rougier et al. (2016) show in the Petites Pyrénées (easternmost SAD) that the North
785 Pyrenean Frontal Thrust was active during the Campanian and ceased activity during the late
786 Maastrichtian. In the EAD (Mouthoumet area), Bilotte (1985) demonstrated that the earliest
787 Campanian is characterized by the generation of a steep slope north of the NPFT and its northward
788 migration. Indeed, the platform-basin limit was controlled by the existence of the future North
789 Pyrenean Frontal Thrust during Cenomanian-Santonian time, while the steep slope migrates onto
790 the European margin during the Campanian. This migration is attested by the occurrence of
791 turbidites onlapping unconformably overlying the Santonian rudist shoals, signifying the formation
792 of the Campanian flexural basin in response to Europe-Iberia convergence (Rosenbaum et al., 2002; Macchiavelli et al., 2017). This pattern is fully consistent with our interpretation of the 793 Clarens-St. Médard seismic reflection profile (Figure 8). Both the growth of the North Pyrenean 794 Frontal Thrust and the steep slope migration onto the European margin support the hypothesis of 795 lithospheric loading that controls the flexural deformation of the SAD. 796

797 The flexural Campanian-Maastrichtian basin is only clearly apparent in the EAD and 798 central/eastern SADIn the western SAD, the Campanian-Maastrichtian sequence in the Rousse- 799 Theze profile (Figure 10) only shows a deepening trend and very little signs of a flexural 800 deformation (Figure 16), except the Campanian progradations identified in the Theze wells 801 cinvinity (Figure 10). The platform- basin transition is at the same position as during Cenomanian- 802 Turonian and Coniacian-Santoninan sequences.. Again, the late syn-rift structural framework 803 controls the geometry of the NAD-SAD. Post-rift thermal subsidence seems to remain the main 804 driver of subsidence (Angrand et al., 2018). 805

## 6.4. Maastrichtian: Flexural quiescence and passive infill of the basin

806 The Maastrichtian sedimentary sequence records a general infilling of the EAD et eastern 807 SAD. It is nearly isopachous, indicating a widespread regional subsidence accompanied by a 808 significant regressive trend. In the EAD, depositional environments pass from delta to continental 809 (coastal plain to alluvial plain), demonstrating that the flexural basin is being overfilled, i.e., the 810 sediment supply significantly exceeds the creation of accommodation space (Sinclair and Allen 811 1992). This pattern might be explained by a global subsidence decrease due to (i) slowdown of 812 Iberia-Europe convergence and subsequent drop in the flexural subsidence and/or (ii) pursuit of 813 thermal subsidence (Angrand et al., 2018) yet in a decreasing trend as expected in post-rift stages. 814
The absence of wedge shape geometry during Maastrichtian together with overfilling conditions suggest a lower flexural induced subsidence. This is congruent with studies showing that Maastrichtian stage corresponds to a period of tectonic quiescence, except in the eastern SAD where the North Pyrenean Frontal Thrust is active until the late Maastrichtian (Rougier et al., 2016).

## 7. Discussions

The study highlights two different deformation stages recorded in the Late Cretaceous: a first of Coniacian-Santonian age followed by a Campanian-Maastrichtian one. Because they have different sedimentary record expressions we chose here to present this long-lasting deformation (Coniacian to Campanian-Maastrichtian) in two stages althoguth it might be considered as a deformation continuum.

The Campanian-Maastrichtian deformation stage corresponds to the clear development of a flexural basin related to the Iberia-Europe convergence and subsequent syn-rift basins' inversion. Indeed, plate kinematic reconstructions tend as well as geological evidences to agree on a late Santonian to Campanian age for the onset of Iberia-Europe convergence (Choukroune and Mattauer, 1978; Roest and Srivastava, 1991; Olivet, 1996; Rosenbaum et al., 2002; Stampfli and Kozur, 2006; Macchiavelli et al., 2017). Field evidence in the Pyrenees (Figure 17) implies a late Santonian age for the onset of convergence (Turbón-Boixols thrusts; Puigdefàbregas and Souquet, 1986; Puigdefàbregas et al., 1992). In the Northeast Pyrenees (Figure 17, Corbières), a flexural basin develops in Campanian time (Bilotte et al., 2005) synchronous with the first folds in Provence (Figure 17, Leleu et al., 2009). Recent thermochronology data from the Agly-Salvezeines area (Figure 17, Eastern Pyrenees) suggest an early cooling age (75 Ma) that is interpreted as early
orogenic crustal thickening (Ternois et al., 2019). These Maastrichtian ages are likely to reflect the
growth of relief features in the easternmost part of the belt (Filleaudeau et al., 2012).

Whereas the Campanian deformation stage is clearly due to Iberia-Eurasia convergence and
subsequent flexural basin formation, the Coniacian-Santonian is believed to be a postrift period
before the onset of any Pyrenean compressive deformation. Therefore, the drivers of this
deformation must be questioned.

## 7.1 Coniacian-Santonian deformation stage: Origin of the deformation and geodynamic implications

The Coniacian-Santonian deformation stage, particularly in the central and eastern SAD, is
marked by the southward tilting of the European margin and the development of wedge shape
geometry (Figure 16). The geometry is related to a reduction of available space on the proximal
margin and a significant increase of the available space in the distal area, toward the former
necking-zone. Although salt diapirism can be cited as a local contributor to these available space
variations (slight reactivation of the Antin diapir in the Laméac area), we show that this trend
develops over much larger extent and also occurs at the transition between the NAD and the SAD,
where no salt diapirs exist. Consequently, the simultaneous uplift of the northernmost SAD (and
NAD) and the subsidence of the southern SAD are not controlled by local brittle faults or
reactivation of inherited structures. This deformation must be widespread given that the exposed
lands and the thickest sequence are at least 30 km apart (Figure 15, ) ie a large part of the south
European plate margin. We investigate now the synchrrounous deformation in the Pyrenean-
Aquitaine domain to better characterize the wave length of this deformation stage.
857 **7.1.1 Onset of Iberia-Europe convergence?**

858 In the near field, several authors describe a major deformation stage in Coniacian time. On
859 the NAD, drowning of the North-Arcachon block (northern Parentis Basin, Figure 2) in the Early
860 Coniacian has been linked to the northward shift of the steep slope separating the platform and
861 basin, interpreted as a late-rifting extensional event (Figure 17; Platel, 1987, 1996). In the
862 meantime, in the Pays Basques area, the Mid-Late Coniacian boundary is characterized by abrupt
863 drowning of the platform (Saint Jean de Luz Basin). This deepening coincides with a general onlap
864 of hemipelagic facies onto the slightly tilted reef deposits, followed by a sequence of westward-
865 directed turbidites (Figure 17; Razin, 1989). This event is interpreted as a late extensional stage
866 (Razin, 1989), similar to the events recorded in the NAD by Platel. (1987, 1996). Onto the Iberian
867 margin's necking zone, extensional salt tectonics initiated during the mid-Coniacian in the Cotiella
868 Basin, where southward-tilted blocks associated with strata growth (Figure 17; López-Mir et al.,
869 2014a, 2014b) suggest a northward gliding of the sedimentary cover. Recent work demonstrates a
870 sharp increase in the Coniacian accommodation creation rate on the Iberian margin necking zone
871 coinciding with an uplift of the proximal margin (Andrieu et al.2021) thus a northward tilting of
872 the Iberian margin. Santonian accommodation rates increase threefold on the necking zone while
873 the proximal margin significantly subsides (after being uplifted in Coniacian time) (Andrieu et
874 al.2021). These vertical movements in the necking zone of the Iberian margin, combined with the
875 general absence of normal faults in the entire Pyrenean-Aquitaine domain lead Andrieu et al.
876 (2021) to exclude extensional tectonics, and to propose two scenarios: (i) onset of Iberia-Europe
877 convergence as soon as Coniacian times, and (ii) a far-field origin.

878 In the hypothesis of a Coniacian early convergence onset, expected plate's deformation is (i)
879 the uplift of the margin(s) resulting from the mechanical load of the plate(s) boundary; and (ii) the
subsidence of the former necking zone and hyperextended domains. The deformation recomposed by Andrieu et al. 2021 fits this idealized response, i.e uplift and northward tilting of the Iberian margin, accompanied by halokinesis and important subsidence of the necking zone. The results presented in this paper show strong similarities, as we evidence, for the central and eastern SAD, a southward tilting of the European margin and exposure of its northernmost edge (NAD), attested by a wedge shape with southward available space increase (toward the SPZ) development onto the margin (Figure 16).

Yet, at that time, no significant shortening resulting from pure north-south compression might be evoked as hyperextended domains closure only initiate during latest Santonian times with orthogonal inversion of the former syn-rift normal faults (Mouthereau et al., 2014; Rougier et al., 2016; Teixell et al., 2016; Grool et al., 2018).

One could argue that far-field deformation resulting from larger Africa-Eurasian plates motion might control this deformation stage, and we propose to investigate this possibility.

### 7.1.2 Africa-Europe oblique convergence : an Alpine far-field deformation ?

The “mid” Cretaceous period is marked at large scale by the Eo-alpine deformation and the first stage of closure of the Alpine Tethys (e.g., Handy et al., 2010, 2015; Angrand et al., 2020; Angrand and Mouthereau, 2021), as shown in Figure 17 and Figure 18. A 84 Ma event (Late Santonian) is currently considered as a time of relative plate motion change between Africa and Europa from divergent to convergent (e.g. Jolivet et al., 2021). In detail, this motion evolves from divergent to convergent through a period including a strike slip component particularly in the Western Tethys (Jolivet et al., 2021), not completely resolved due to long time magnetic quiescence during the Upper Cretaceous normal superchron (Aptian to Santonian).
During the Early Cretaceous, the southern part of the Alpine Tethys closes along the western and northern margin of Adria, due to northward kinematics of Adria and Africa relative to stable Europe (Figure 18; e.g., Angrand and Mouthereau, 2021, and references therein). Northwestern boundary of Adria evolved into a transpressive plate boundary during the Late Cretaceous, although precise structures remain poorly constrained. The subsequent far-field deformation of the transpressive Iberia-Adria and Europe-Adria plate boundaries may be responsible for compressive deformation, far into the European plate.

Late Cretaceous compression is recorded in Africa and Europe up to the north Sea. Multiple field-studies evidence a Coniacian deformation stage, thus before the inferred deformations in the Pyrenees: in North Aquitaine platform, Provence, External and Inner Alps as well as in a more far field settings although their kinematics are debated (Central to Northern Europe, Paris Basin).

In the most distal domain, far-field deformations spread from Turonian to Santonian times, with for instance the inversion of the Lower Saxony, grabbens (in Germany; Baldshuhn et al., 1991) the Roer Valley Graben shoulder (Southern Netherlands, Gras and Geluck, 1999), the polyphase Sole Pit Basin (North Sea, Betz et al., 1987; Van Hoorn, 1987). Similarly, in Western Europe (Paris Basin), Lasseur (2007) demonstrates a major flexure deformation stage within the same period (Late Turonian to the Early Campanian), interpreted as the onset of the Alpine cycle's compressive stress. Contemporaneous far field deformations are also recorded in Africa, culminated with the so called Santonian phase (Guiraud et al., 2005, Jolivet et al., 2016 2021). These far-field deformations, while distant from thousand of kilometers of the main plate-boundaries stress the large-scale consequences of major deformation onto strain propagation.

Closed to the Pyrenean domain (External Alps: Devoluy; Provence: South Provence Basin; see Figure 1), there are evidences of a tectonic deformation stage as soon as Coniacian time (Figure
17, Floquet and Hennuy, 2003; Hennuy, 2003; Grosheny et al., 2015) although . In the Devoluy area, deformation stage is characterized by a large-scale tilting of the latest Turonian platform, followed by the deposition of early Coniacian unconformably slope-apron deposits (Austrian Ante-Gossau episode in Figure 17; Grosheny et al., 2015). These authors interpret it as the consequence of early Alpine compression (Figure 17; Winkler and Bernoulli, 1986; Faupl and Wagreich, 1992; Wagreich, 1995). Southward the enlargement of the South Provence Basin accompanied by the deposition of terrigenous turbidites also witness a significant deformation stage (Figure 17; Floquet and Hennuy, 2003; Hennuy, 2003). Although evidence of normal faults is scant, the presence of wedge-shaped turbidite bodies leaning back on scarps suggests the creation of steep submarine slopes that were preserved beneath the Coniacian deposits. While no consensus yet exists on the mechanisms leading to this deformation stage in the Devoluy and in Provence, geologic evidence suggests a transpressive regime (Marroni and Treves, 1998; Schreiber et al., 2011; Grosheny et al., 2015).

The Iberian and European margins, which uplift in Coniacian-Santonian times (Andrieu et al., 2021 and this study) are separated by a more than 100 km highly subsiding domain (former necking zones and hyperextended domains), a distance that we consider compatible with lithosphere buckling processes driven by the oblique convergence of Africa with respect to Europe (Figure 18).

To be correct, this far-field origin hypothesis requires a strong mechanical coupling between (Africa-)Adria and Iberia-European plates. This implies that the oceanic domains (i.e., the southern Alpine Tethys) must have been of lesser significance or already closed (Figure 18; Angrand and Mouthereau, 2021) than in some reconstructions (e.g., Jolivet and Faccenna, 2000; van Hinsbergen et al., 2020).
Current knowledges and observations cannot determine whether the Coniacian-Santonian deformation stage is controlled by Iberian-Europe plate convergence or lithosphere buckling related to “Alpine” far-field deformation. Both hypothesis converge toward the same conclusion: the inversion of the former Pyrenean-Aquitaine Albian-Cenomanian rift occurs as soon as Coniacian times, implying a very brief 10 My post-rift stage.

## 7.2. Segmentation between the eastern and western South Aquitaine domains

On the basis of seismic reflection profiles, we demonstrate that the Campanian-Maastrichtian flexural regime as well as the Coniacian-Santonian far-field deformation stages are poorly recorded in the vicinity of Pau area (Figure 16), where Late Cretaceous basins and sedimentary geometries do not show any increase in thickness that we can relate to flexural subsidence. This pattern argues for a long-lasting post-rift subsidences. On the contrary, early inversion is recorded in the eastern domain since the Coniacian (Figure 16). As a reminder, in the central and eastern domain, a major Coniacian-Santonian lithosphere buckling marks the critical inversion from the Cenomanian-Turonian post rift stage to the Campanian-Maastrichtian syn-orogenic flexural stage. This discrepancy between an eastern and western domain argues for strong lateral change within the Aquitaine foreland basin, as it was previously proposed by Angrand et al. (2018). These authors used 2D flexural modelling to suggest that the subsidence in the Western Aquitaine Basin is mainly controlled by post-rift cooling regime and associated thermal subsidence. Our results based on sequential basin analysis fully agree with this hypothesis. We might therefore interrogate on such a short-distance subsidence along-strike variation. Recent results on both the North Pyrenean Zone and the Aquitaine Basin (Canérot, 2017; Saspiturry et al., 2019; Saspiturry et al., 2019b; Issautier et al., 2020; Lehujeur et al., 2021) detail the existence of NE-SW transfer zones in the Arzacq and Mauléon basins. In the central Aquitaine Basin Angrand et al. (2018) suggest that both the crust
and the foreland is segmented by NE-SW structures. Such crustal to lithospheric structures might play a significant role on the westward strain migration induced by the syn-orogenic loading, and they could yield to a severe along-strike mechanical compartmentalization of the Aquitaine foreland basin.

## 7.3. Impact of a non-reequilibrated hot lithosphere on early compression in the eastern Aquitaine domain

The Campanian period corresponds to the formation of a widespread flexural basin as a response to compression induced by Iberia-Europe convergence. In the EAD and central/eastern SAD, the Mas d'Azil-Lezat and Tournay-Trouley Labarthe seismic reflection profiles document a major difference in geometry between the Eocene and Campanian flexural basins. These two flexural basins are separated by a Late Maastrichtian Paleocene quiet tectonic subsidence phase evidenced by relative flat sedimentary geometries and lack of fault-activity (Ford et al., 2016; Rougier et al., 2016; Grool et al., 2016).

At the scale of the Aquitaine Basin, the Cenozoic foredeep is characterized by the northward migration of the foredeep and the outer rim of the flexural basin, because of the growth of the Pyrenean orogen (Figure 7, Figure 12 and Figure 14). The striking difference from the Campanian foreland is that the foredeep is much larger, as shown by the 30 km increase in the N-S wavelength during the Eocene visible in the Mas d'Azil-Lezat profile (Figure 7). Similarly, in the SAD, the Tournay-Trouley Labarthe profile shows an abrupt thickness increase associated with the platform-basin transition in the Campanian flexural basin and a more gradual increase in the Cenozoic basin (Figure 14).
The flexural response of a hot, thinned lithosphere is believed to produce relatively steep and narrow foreland basins (Stockmal et al., 1986; Sinclair et al., 1991; Watts, 1992). The inversion of rifted basins into flexural basins is believed to be strongly controlled by the state of the lithosphere and thus its elastic thickness (Watts, 1992; Haddad and Watts, 1999; Lin and Watts, 2002; Horton et al., 2016). When the time spans between the end of rifting and inversion is short (poor lithosphere reequilibration), the extensional template strongly controls how the compressive deformation localizes (Ziegler et al., 1995). In the Pyrenean domain, only 10 Myr separate the Albian-Cenomanian extension from the Campanian inversion. We suggest that the short wavelength and high amplitude of the Campanian foreland basin is related to the emplacement of the flexural basin on a rifted crust in the southern Aquitaine basin (Angrand et al., 2018) and on in the North Pyrenean Zone, the former proximal and hyper-extended domains, respectively. In this scheme, inherited rift structures possibly governed the localization of the Campanian foredeep (Angrand et al., 2018). Similarly, a tectonic control of the foredeep due to localization of the forebulge by inherited structures has also been proposed for the southern Andes foreland basin (Fosdick et al., 2014). When the Pyrenean orogeny enters in its main stage in the Eocene, the orogenic load increases and the flexural basin migrates northward onto unextended lithosphere. Together with the increase of topographic load, because of the growth of the Pyrenean orogen, this leads to a larger and more gently inclined foredeep basin. The control of the Aquitaine foreland basin by inheritance, addressed first by Desegaulx et al. (1991) and more recently by Angrand et al. (2018), is strengthened by our results.

Because temperature is thought to be an important influence on the elasticity thickness (rigidity) of a plate (Burov and Diament, 1995; Hyndman et al., 2009; Mouthreau et al., 2013; Tesauro et al., 2018), we should consider the impact of the high geothermal gradient and thick
sedimentary succession inherited from the rifting stage on the ensuing compressive deformation. In the Pays Basques, the Turonian was a period of extension (Razin, 1989), and recent studies indicate that the high geothermal gradient inherited from the Albian only started to decrease at the onset of rift inversion (Eocene in Vacherat et al., 2014; Latest Santonian in Saspiturry et al., 2020a). Thus the Coniacian-Santonian deformation stage is characterized by buckling of a lithosphere in a thermally non-equilibrated state, strongly prone to flexure, inherited from the Albian-Cenomanian hyperextended rifting stage. The Campanian growth of a submarine orogenic wedge (Angrand et al., 2021) loads this weak lithosphere while it is still in thermal reequilibration, forming a steep and narrow foreland basin. From the Late Cretaceous to the Eocene, the European lithosphere had time to reequilibrate, hence producing a larger flexural basin during the Eocene.

## 8. Conclusion

The Pyrenean-Aquitaine realm is a remarkable place to unravel the process of hyperextended lithosphere inversion because the elapsed time between breakup and rift closure is very short. Because the postrift stage is brief in the Pyrenean realm, the lithosphere there is still relatively hot at the beginning of rift inversion.

Our study reveals a complex polyphase inversion in the Aquitaine and North Pyrenean basins. Following a brief (10 Myr) Cenomanian-Turonian postrift stage, an initial deformation stage is recorded within the Coniacian-Santonian interval characterized by uplift of the European margin and responsible for the development of a detrital lowstand systems tract in the EAD. At the same time, the southern and southeastern parts of the SAD basins undergo marked subsidence. We demonstrate that this deformation stage is similar to coeval dynamics on the Iberian margin documented in recent work. We interpret this regional deformation, with its wavelength of roughly
1037 100 km, as a stage of lithosphere buckling that results from far-field compressive deformation in
1038 the Alpine Tethys realm.

1039 The Campanian-Maastrichtian cycle records the development of a flexural basin, classically
1040 interpreted as the first Pyrenean foreland basin. Compared to the Eocene flexural basin formed
1041 during the climax of the Pyrenean orogeny, this Late Cretaceous basin displays a steeper slope, a
1042 sharper platform-basin transition, and a much smaller dimension. These differences argue for
1043 strong control by lithospheric inheritance in which the Campanian basin develops above a hot,
1044 thermally unbalanced lithosphere with a low rigidity. Because major Pyrenean structuration occurs
1045 in the early Eocene after the Maastrichtian-Paleocene quiescence stage (14My time laps), we may
1046 assume that when main compression initiates in Ypresian time, the lithosphere had gained strength
1047 and thermal equilibrium, allowing the development of a wider and smoother flexural basin.

1048 Acknowledgments

1049 This study was conducted within the framework of the GAIA project (Téréga, Agence de l'Eau
1050 Adour-Garonne and BRGM), and the Orogen integrated geologic research project funded by
1051 TOTAL, BRGM, and INSU. Seismic reflection profile and wells were provided by MINERGIES
1052 (http://www.minergies.fr/fr) and as so we greatly thank them. We would like to thank the Orogen
1053 project managers Isabelle Thinon (BRGM) Sylvain Calassou (Total), Emmanuel Masini (former
1054 Total; M&U), Olivier Vidal (CNRS). We also warmly thank Andrew Alden for the English review,
1055 Mary Ford for her precious help in reviewing and helping and structuring this paper, and the
1056 reviewers for their valuable work.
1057

<figure><img src="image_3.png" /></figure>

1058

1059

1060

1061

1062

1063

Figure 1: Location of the study (black frame). The Aquitaine Basin hosts a precious sedimentary record of surrounding complex domains which suffered major rifting stages (Bay of Biscay, Pyrenees and Mediterranean Sea) and orogeny processes (Pyrenees, Alps). The map was compiled with the million-scale map (Chantraine et al., 2003) ; Le Maire et al., (2021) for the Bay of Biscay Margin domains : Serrane, 1999 and Canva et al., 2020 for the Mediterranean Margin domains.
<figure><img src="image_4.png" /></figure>

STRATIGRAPHY

<figure><img src="image_5.png" /></figure>

Quaternary

<figure><img src="image_6.png" /></figure>

Cenozoic

<figure><img src="image_7.png" /></figure>

Upper Cretaceous

<figure><img src="image_8.png" /></figure>

Lower Cretaceous

<figure><img src="image_9.png" /></figure>

Jurassic

<figure><img src="image_10.png" /></figure>

Triassic

<figure><img src="image_11.png" /></figure>

Paleozoic

TECTONIC STRUCTURES

<figure><img src="image_12.png" /></figure>

Southward thrusts

<figure><img src="image_13.png" /></figure>

Northward thrusts

<figure><img src="image_14.png" /></figure>

Anticline

<figure><img src="image_15.png" /></figure>

Geophysical lineation

GEOGRAPHY

<figure><img src="image_16.png" /></figure>

Boreholes

<figure><img src="image_17.png" /></figure>

Study sections

<figure><img src="image_18.png" /></figure>

Main cities

<figure><img src="image_19.png" /></figure>

BOREHOLES

Ca1 - Calavante-1

Go1 - Goudon-1

Ja1 - Jacque-1

La1 - Lameac-1

Le1 - Lezat-1

Ma1 - Mas d'Azil-1

Mz1 - Mazerolles-1

Os1 - Osmets-1

Os2 - Osmets-2

Pa1 - Pailhes-1

Lo1 - Lons-1

Pa2 - Pau-2

Pa4 - Pau-4

Re1 - Recurt-1

Ro1 - Rousse-1

Ro2 - Rousse-2

Sa1 - Saoulous-1

Sd1 - Sadeillan-1

Th1 - These-1

Th301 - Theze-301

Tl1 - Trouley-Labarthe-1

To1 - Tournay-1

<figure><img src="image_20.png" /></figure>

CROSS-SECTIONS

<figure><img src="image_21.png" /></figure>

WESTERN AQUITAINE (SERRANO ET AL., 2006)

<figure><img src="image_22.png" /></figure>

CENTRAL AQUITAINE (FORD ET AL., 2006)
Figure 2: (A) Structural framework of the Aquitaine Basin. The main structures within the South Aquitaine Domain are modified after Serrano et al. (2001). (B) Locations of interpreted seismic reflection profiles, wells, and field-type sections. NPFT = North Pyrenean Frontal Thrust; TF = Toulouse Fault; FCA = Celtaquitaine Flexure; NPF = North Pyrenean Fault; LT = Lakhoura Thrust; NPZ = North Pyrenean Zone; SPZ = South Pyrenean Zone.

<table><thead><tr><th>FACIES ASSOCIATIONS (POSITION ON THE DOWNDIP PROFILE)</th><th>LITHOFACIES</th><th>NON-BIOCLASTIC COMPONENTS</th><th>BIOCLASTIC COMPONENTS</th><th>SEDIMENTARY AND BIOGENIC STRUCTURES</th><th>SORTING, GRAIN SIZE</th><th>FACIES AND DEPOSITIONAL ENVIRONMENT (dep.env.) INTERPRETATIONS</th><th colspan="3">ELECTROFACIES</th></tr><tr><th></th><th></th><th></th><th></th><th></th><th></th><th></th><th>GAMMA RAY</th><th>NEUTRON</th><th>RESISTIVITY</th></tr><tr><th></th><th></th><th></th><th></th><th></th><th></th><th></th><th>0</th><th>0</th><th>1500</th></tr><tr><th></th><th></th><th></th><th></th><th></th><th></th><th></th><th>100</th><th>2</th><th>20000</th></tr></thead><tbody><tr><td rowspan="3">FA7: CONTINENTAL</td><td>F7a: sandy marls</td><td>Quartz</td><td>Pollen, dynophyta, dinosaur eggs</td><td>Mottling, pedogenetic limestone nodules, sandy to conglomeratic channels</td><td>Poorly sorted</td><td>Bioclastic components, and sedimentary features suggest a flood plain dep. env.</td><td></td><td></td><td></td></tr><tr><td>F7b: channelized sandstone to conglomerate</td><td>Quartz</td><td>Dinosaurs eggs fragments, oncoids</td><td>Channels</td><td>None</td><td>Bioclastic components and sedimentary features suggest a fluvial dep. env.</td><td></td><td></td><td></td></tr><tr><td>F7c: lake limestone to clay</td><td>None</td><td>Characca gyroponites, dinosaurs eggs fragments</td><td>None</td><td>None</td><td>Bioclastic component and fine granulometry suggest a very calm continental dep. (lake)</td><td></td><td></td><td></td></tr><tr><td>FA6-7: ESTUARY TO FLUVIAL</td><td>F6a: polygenic brecchia</td><td>Quartz, paleozoic basement clasts</td><td>None</td><td>Basal erosive channels</td><td>Poorly sorted and very coarse (centimetric to decimetric)</td><td>Very coarse granulometry, poor sorting and erosive channels indicate a very high energy. Often associated to F6, it suggests a continental position</td><td></td><td></td><td></td></tr><tr><td rowspan="4">FA6: MIXED FLUVIAL-TIDE DOMINATED DELTA</td><td>F6b: Coarse radioactive sandstone</td><td>Dolomite, Pyrite, Quartz</td><td>Bryozoans, miliolids, echinids, ostracods, lamellibranchs</td><td>None</td><td>Poorly sorted coarse</td><td>Fauna suggests a relative proximal dep. env. probably dominated by fluvial influx (grain-size/sorting) and hyperpycnal flow (heavy mineral placers) interpreted as littoral sandstone by Serrano et al. (2006)</td><td></td><td></td><td></td></tr><tr><td>F6c: Argileaceous siltstone and wackstone / packstone intercalations</td><td>Quartz, Lignite, Pyrite</td><td>Prealvolinas, Miliolids, Orbitolinids</td><td>None</td><td>None</td><td>The fine grained texture, fauna and lignite clasts suggest a very proximal dep. env. The colour implies a reduction environment such as a bay</td><td></td><td></td><td></td></tr><tr><td>F6d: quartz sandstone to conglomerate</td><td>Quartz</td><td>None</td><td>Planar bedding, erosive channels, current ripples, clay interbeds, bioturbations</td><td>Moderately sorted</td><td>Sedimentary structures indicate a high hydrodynamic dep. env. dominated by tide current as suggested by clay drapes and planar bedding (Tidal flat)</td><td></td><td></td><td></td></tr><tr><td>F6e: oncolite wackestone to packstone</td><td>Quartz, glauconite, micas</td><td>Miliolids, bivalves, echinids and ostracods</td><td>Planar and tangential cross-bedding in megaripiles, sigmoides, rare bioturbations</td><td>Moderately to poorly sorted, fine to very coarse</td><td>Coarse granulometry, fauna, sedimentary structures suggest high energy conditions above the fair weather wave base in a delta, locally subject to tides influences (sigmoides): probable Mouth bar</td><td></td><td></td><td></td></tr><tr><td>FA5: LAGOON (INNER RAMP)</td><td>F5a: oncolite wackestone to packstone</td><td>Dolomite</td><td>Oncoids, gastropods, miliolids, ostracods</td><td>Microbialites</td><td>Poorly sorted</td><td>Well-preserved gastropods, microbialites, milioids and absence of sedimentary structures and muddy textures suggest a calm and restricted dep. env.</td><td></td><td></td><td></td></tr><tr><td rowspan="2">FA4: SHOREFACE (INNER RAMP)</td><td>F4a: bioclastic grainstone</td><td>Peloids, rare quartz</td><td>Bivalves, corals fragments, bryozoans, alveolids, and miliolids</td><td>Planar and tangential cross-bedding in megaripiles, rare bioturbations</td><td>Moderately sorted, coarse</td><td>Grainstone texture, fauna and sedimentary structures suggest high energy conditions above the fair weather wave base</td><td></td><td></td><td></td></tr><tr><td>F4b: rudist rudstone to boundstone</td><td>None</td><td>Rudists, algae, miliolids and other foraminifera</td><td>None</td><td>Poorly sorted</td><td>Isolated rudist bioconstructions developing in high energy environments in the shoreface</td><td></td><td></td><td></td></tr><tr><td rowspan="3">FA3: MID RAMP</td><td>F3a: bioclastic clayey wackestone</td><td>Glauconite, peloids</td><td>Crinoids, bivalves, pithonellae, coral fragments, foraminifera</td><td>Abundant bioturbations and thalassinoides planolites</td><td>Well sorted</td><td>Texture, abundant vertical burrows, and fauna suggest moderate energy conditions along a mid-ramp, below the fair weather wave base</td><td></td><td></td><td></td></tr><tr><td>F3b: marl/bioclastic packstone to grainstone alternations</td><td>Lignite fragments</td><td>None</td><td>Bioturbations, thalassinoides and planolites, erosive gullies</td><td>Moderately sorted</td><td>Marl/limestone alternations, and erosive gullies suggest high energy conditions (storms) alternating with calm periods</td><td></td><td></td><td></td></tr><tr><td>F3c: silty marl/sandstone alternations</td><td>Quartz, micas, peloids</td><td>Bivalves, echinoderms</td><td>Hummocky and swaley cross stratifications, wave ripples</td><td>Moderately sorted</td><td>Hummocky and swaley cross-stratifications and varying granulometry suggest high energy conditions (storms) alternating with calm periods</td><td></td><td></td><td></td></tr><tr><td>FA2: OUTER RAMP</td><td>F2a: marl/micritic limestone alternations</td><td>Glauconite</td><td>Crinoids, bivalves, bryozoans</td><td>Abundant bioturbations zoophyos, planolites, chondrites, teichichnus and rare thalassinoides</td><td>Very well sorted, very fine</td><td>Texture, bioturbations, very fine granulometry and absence of sedimentary structures suggest calm and distal depositional conditions along an outer ramp</td><td></td><td></td><td></td></tr><tr><td>FA1: BASIN</td><td>F1a: calcareous turbidites, marl/ granular limestone alternations</td><td>Micas</td><td>Spicules, miliolids, sidorelites</td><td>None</td><td>None</td><td>F/lysch facies deposited in a very low energy, basin environment. The fauna are exported from shallow platform.</td><td></td><td></td><td></td></tr></tbody></table>

Figure 3: Lithofacies and electrofacies classifications used in this study. Facies descriptions were based on the Rennes-les-Bains outcrop section and the Landes-de-Siougos well, which furnished the data for the electrofacies calibration (Serrano, 2001). Seven facies associations (FA) were created, ranging from basinal (FA1) to continental (FA7).
CALAVANTE 1: 12-1943

<figure><img src="image_23.png" /></figure>

DEPOSITIONAL
ENVIRONMENT

<figure><img src="image_24.png" /></figure>

LITHOLOGY

<figure><img src="image_25.png" /></figure>

SEQUENCE STRATIGRAPHY

<figure><img src="image_26.png" /></figure>

FAUNA

BX* Boix Martinez

JACQUE 1: 12-2078

<figure><img src="image_27.png" /></figure>

Pseudomedia drorimensis

Praestorssella rostrae
Pseudocyclammina
Abrardia mosae
Cuneolina cf. pavonia
Nummoloculina
Pyrcnorotalia longifolia BX*
Rotorbinella cf. campaniola BX*
Late Santonian

Pyrenerotalia longifolia
Pararotalia tuberculifera
Iberotalia reicheli A type
Early Santonian

Whiteinella
Heterohelix reussi
Prealveolina gr. cretacea
Ovalveolina ovum
Mid to Late
Cenomanian

1076 Figure 4: Biostratigraphy and depositional environments for Calavante-1 and Jacque-1 wells.

1077 Biostratigraphy allows to refine stratigraphic attributions and highlights relative thick and shallow

1078 Coniacian-Santonian sequences. This sequence is marked by the development of proximal deltaic

1079 facies influenced by fluvial and tidal processes. In both wells, first turbidites deposits are identified

1080 within the Campanian.
<figure><img src="image_28.png" /></figure>

1081

1082 Figure 5: Correlation of the type sections for the three subdivisions of the Aquitaine domain. The three major transgression-regression cycles, of Cenomanian-Turonian, Coniacian-Santonian, and Campanian-Maastrichtian age, are bounded by maximum regressive surfaces.
<figure><img src="image_29.png" /></figure>

Figure 6: Well calibration of the Mas-d'Azil – Lezat seismic reflection profile (Fig.7). It illustrates the emplacement of the SPT cross-cutting the thick Campanian Plagne Fm, as well as a significant southward thickening of the Late Cretaceous deposits. The location of the NPFT is not to scale.
<figure><img src="image_30.png" /><figcaption>Figure 7: Mas d'Azil-Lezat seismic reflection seismic reflection profile; see location in Fig.2B.

The Cretaceous and Cenozoic basins correspond to wedge-shaped basins opening southward toward the North Pyrenean Frontal Thrust. The onlap of the proximal system progresses gradually</figcaption></figure>

1089

1090

1091

1092
1093 northward from the Coniacian to Maastrichtian, then abruptly migrates 10 km northward during
1094 the Cenozoic (Eocene).

<figure><img src="image_31.png" /></figure>

1096 Figure 8: West-East Pailhes seismic reflection profile; location in Fig. 2B. The first Upper
1097 Cretaceous sequence consists of Coniacian-Santonian deposits displaying westward progradation
1098 that significantly increase within Campanian.
<figure><img src="image_32.png" /><figcaption>Figure 9: Well calibration of the Rousse-Theze seismic reflection profile (Fig. 10). It illustrates a southern domain (Lons to Rousse wells) dominated by turbidites (Campanian Flysch, Hopital Saint Blaise Fm) to base of slope facies (Brèches de Soumoulou Fm.) and outer ramp deposits (Marnes d'Auzas Fm.). On the contrary, the stratigraphic succession between Lons and Theze wells is characterized by relative shallow-marine carbonate developping following a transgressive trend.</figcaption></figure>

1099

1100

1101

1102

1103

1104
<figure><img src="image_33.png" /><figcaption>Figure 10: Rousse-Theze seismic reflection seismic reflection profile; see location in Fig. 2B. The Arzacq Basin shows a relative thin Upper Cretaceous sequence while it significantly increases toward the North-Pyrenean Zone. 10C is a detailed illustration of the clinoformes identified in</figcaption></figure>
1109 between Theze-1 and Theze-301 wells. They attest of a progradational trend within the Campanian-

1110 Maastrichtian on the outer-rim of the Pecorade anticline.

<figure><img src="image_34.png" /></figure>

1111

1112 Figure 11: Well calibration of the Clarens-Saint Medard seismic reflection profile. Cenomanian-

1113 Turonian deposits are restricted to the south of Saoulous-1 well. They correspond to volcano-

1114 sedimentary facies. They grade vertically to shallow carbonate of the Gensac and Jouansalles

1115 formations which thickness increase southward. The Campanian is marked by a deepening of the

1116 depositional environments and the northward migration of the turbiditic deposits (Flysch Marneux

1117 Fm.) over the underlying shallow-marine carbonate deposits. The same pattern is identified in the
1118 Maastrichtian sequence with turbidites in the southern part of the profile: Recurt-1 well (Marnes de Nay Fm.) and shallow-marine carbonate in the northern part of the profile: Saoumis-1well (Couches de Dumes Fm).

<figure><img src="image_35.png" /></figure>

1121

1122 Figure 12: Clarens-St. Médard seismic reflection seismic reflection profile; see location in Fig. 2B.
1123 It displays a wedge-shaped basin accommodating Cenomanian to Campanian deposits with northward onlaps, suggesting an increase of the subsidence throughout the Late Cretaceous.
1124
1125 and 12D illustrate the erosional truncations and onlap markers onto the South Aquitaine

1126 Domain in respectively its southern and northern portions.

<figure><img src="image_36.png" /></figure>

1127

1128 Figure 13: Well calibration of the Tournay – Trouley-Labarthe seismic reflection profile (Fig. 14).

1129 The Cenomanian (Pilo Fm.) and the Turonian part of the Jouansalles Fm. thin toward the south

1130 (see figure. 4: the Turonian facies no longer exist in the Calavante-1 well). The Coniacian-

1131 Santonian sequence is marked by the development of the Jouansalles Fm. to the south of the

1132 Calavante-1 well, and the Gensac Fm. to the north of it. It attests of a clastic influx in the northern

1133 part of the profile. Finally, the Campanian-Maastrichtian sequence is marked by a clear northward
1134 migration of the turbiditic facies (Flysch Marneux) and outer-ramp deposits (Pé-Marie Fm., Marnes de Nay Fm., Couches de Dumes Fm.) over the underlying shallow carbonates.

1135

<figure><img src="image_37.png" /></figure>

1136

1137 Figure 14: Tournay–Trouley Labarthe seismic reflection seismic reflection profile; see location in

1138 Fig. 2B. The interpretation shows a southward onlap of the Cenomanian-Turonian sequence

1139 (Calavante-1 well) and a northward onlap of the Coniacian to Campanian sequence. The southern
1140 part of the profile is characterized by thickness increase coinciding with the facies transition from
1141 platform to basin. Prominent toplaps are visible below the base of the Campanian sequence. 14 C
1142 and D illustrate respectively the clinoformes identified in the vicinity of southern Osmets, and
1143 the onlap markers in the southern part of the profile (Tournay-Calavante). This dynamic reflects a
1144 difference in the sedimentary dynamic between the southern SAD where facies are deposited in a
1145 retrograding trend (onlap markers) and the northern SAD recording suppression of available space
1146 and subsequent south-directed regression.

<figure><img src="image_38.png" /></figure>

<figure><img src="image_39.png" /></figure>
Figure 15: (Top) Sedimentary geometry; (Bottom) Distribution of facies associations within the Tournay–Trouley Labarthe well-log profile; see location in Fig. 2B. The top of the profile is horizontalized on the Cretaceous–Paleogene sequence boundary. The Cenomanian–Turonian deposits onlap southward while the Coniacian–Santonian sequence displays two southward-opening prisms that pinch out northward. The upper prism is characterized by a sharp southward progradation of radioactive sandstone (littoral lithofacies). The Campanian sequence unconformably overlies the Coniacian–Santonian deposits.

1155

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

1156

Figure 16: Schematic diagram showing the Late Cretaceous evolution of the SAD, modified after Saspiturry et al. (2019) (left) and Rougier et al. (2016) (right). The domain differs in its tectono-sedimentary evolution between a western domain in the Pau area and an eastern domain in the Tarbes to Corbières area (see locations in Fig. 1B). The eastern domain and western domains show
1161 radical different basin's dynamic. The first one records an inversion stage in Conaician-Santonian

1162 followed by installation of a Campanian-Maastrichtian flexural basin while western SAD remains

1163 under post-rift regime throughout the Late Cretaceous.

<figure><img src="image_41.png" /></figure>
Figure 17: Chart of Late Cretaceous tectono-sedimentary events in the Pyrenean and Alpine domains. Several tectonic pulses are mentioned in this chart, the strongest ones and most widely recorded are the Coniacian-Santonian and Campanian events.

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

1168

Figure 18: Sketches showing the geodynamic state of the Alpine-Pyrenean realm (A) At the end of Albian-Cenomanian rifting (94 Ma) and (B) At the onset of the latest Santonian-early Campanian
1171 inversion (84 Ma) in the Aquitaine Basin. During the earlier period, the northwest Adria region is
1172 in a convergent regime while extension is occurring in the Bay of Biscay domain. In the latest
1173 Santonian, Africa-Europe convergence is more oblique, controlling a compressive stage in the
1174 Pyrenees. Modified from Angrand et al. (2021).

1175 SUPPLEMENTARY 1

1176 SUPPLEMENTARY MATERIAL 1

<table><thead><tr><th colspan="2">EPOCH</th><th colspan="10">LATE CRETACEOUS</th></tr><tr><th colspan="2">STAGES</th><th colspan="3">CENOMANIAN</th><th colspan="3">TURONIAN</th><th colspan="3">CONIACIAN</th><th colspan="2">SANTO NIAN</th><th colspan="2">CAMPANIAN</th><th colspan="2">MAASTRICHTIAN</th></tr><tr><th colspan="2">SUB-STAGES</th><th>EARLY</th><th>MID.</th><th>LATE</th><th>EARLY</th><th>MIDDLE</th><th>LATE</th><th>EARLY</th><th>MIDDLE</th><th>LATE</th><th>SANTO NIAN</th><th>LATE</th><th>EARLY</th><th>MIDDLE</th><th>LATE</th><th>EARLY</th><th>LATE</th></tr></thead><tbody><tr><td rowspan="2">Benthic foraminifera</td><td>Pseudocyclammina rugosa</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Praealveolina gr. cretacea</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td></td><td>Nummofallotia cretacea</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td rowspan="2">Planktonic foraminifera</td><td>Dicarinella concavata</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Globotruncana arca</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Rudists</td><td>Hippurites (H.) requieni</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td rowspan="8">Ammonites</td><td>Coilopoceras requienianum</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Tissotia ewaldi</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Peroniceras subtricarinatum</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Gauthiericeras margae</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Parabevahites emscheris</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Texanites gallicus</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Pseudoschloenbachia aff. inconstans</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Placenticeras syrtale</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td></td><td>Paralenticeras caneroti</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td></td><td>Glyptoxoceras souqueti</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Charophytes</td><td>Septorella ultima</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr></tbody></table>

1177

1178

1179 **Supplementary Material 1.** Time distribution of benthic and planktonic foraminifer, rudist, ammonite and charophyte fauna from Cenomanian to Maastrichtian time in the eastern Pyrenees (Bilotte, 1985; Ogg et al., 2016). Stage boundary ages from Ogg *et al.* (2016).

1182
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1737 Winkler, W., Bernoulli, D., 1986. Detrital high-pressure/low-temperature minerals in a late Turonian flysch sequence of the eastern Alps (western Austria): implications for early Alpine tectonics. Geology 14, 598-601.

1740 Y.Philippe, 1998. Bilan du permis La Noue après traitement 2D (bassins de Tarbes et Commonges)

1741