# Deformation conditions and 40Ar/39Ar dating of thrusting recorded by clay minerals: An example of the Lakora thrust (west-central Pyrenees)

Ahmed Abd Elmola, Patrick Monié, Pierre Labaume, Antonio Teixell, Delphine Charpentier, Martine Buatier

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Ahmed Abd Elmola, Patrick Monié, Pierre Labaume, Antonio Teixell, Delphine Charpentier, et al. Deformation conditions and 40Ar/39Ar dating of thrusting recorded by clay minerals: An example of the Lakora thrust (west-central Pyrenees). Journal of Structural Geology, 2022, 165, 10.1016/j.jsg.2022.104744. insu-03863348

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1 Deformation conditions and <sup>40</sup>Ar/<sup>39</sup>Ar dating of thrusting
2 recorded by clay minerals: an example of the Lakora thrust
3 (west-central Pyrenees)

4 Ahmed Abd Elmola<sup>a,b</sup>, Patrick Monie<sup>c</sup>, Pierre Labaume<sup>c</sup>, Antonio Teixell<sup>d</sup>,
5 Delphine Charpentier<sup>a</sup>, Martine Buatier<sup>a</sup>.

6 <sup>a</sup> Chrono-Environnement, Université Bourgogne Franche-Comté, CNRS, Besançon, France.

7 <sup>b</sup> Laboratoire IC2MP, Université de Poitiers, CNRS, Poitiers, France

8 <sup>c</sup> Géosciences Montpellier, Université de Montpellier, CNRS, Montpellier, France.

9 <sup>d</sup> Departament de Geologia, Universitat Autònoma de Barcelona, Bellaterra, Spain

10
11 * Corresponding author: Ahmed (First name) Abd Elmola (Family name)
12 Primary e-mail: abd_elmola.ahmed@yahoo.com
13 Secondary e-mail: ahmed.abd.elmola@univ-poitiers.fr
14 Previous address (where the work has been done): Chrono-Environnement, 16 Route de Gray,
15 25030 Besançon, France.

16 Present address : Laboratoire IC2MP, 4 Rue Michel Brunet, 86073 Poitiers, France
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26 **Abstract**

27 Newly-formed clay minerals in fault zones can provide crucial information about conditions,
28 mechanisms, and timing of deformation at low-grade metamorphism. The present study focuses
29 on the major Lakora thrust in the west-central Pyrenees, that exposes an assemblage of Paleozoic-
30 Upper Cretaceous rocks of the North Pyrenean Zone in the hanging wall and Campanian-
31 Maastrichtian turbidites of the Axial Zone cover in the footwall. Detailed structural, petrological,
32 mineralogical, and chemical investigations were performed on fault rock samples from the
33 Lakora thrust footwall to unravel the deformation conditions related to thrusting. In addition, an
34 attempt to date the thrust activity was made through direct $^{40}$Ar/$^{39}$Ar step-heating on different
35 fractions of muscovite/illite. The structural and petrographic observations suggest a brittle-ductile
36 deformation in two stages along the fault zone: a first stage of bedding-parallel extension and
37 bedding-perpendicular shortening that is marked by boudinage associated with a bedding-parallel
cleavage (S0-S1) and a second stage of oblique shearing as marked by reverse faults, folds and
38 oblique cleavage (S2). Pressure-solution and dissolution-recrystallization in the presence of fluids
39 are the main deformation mechanisms that enhanced the neoformation of chlorite and illite
40 associated with cleavage development during the two stages of deformation. The crystallinity,
41 polytypism, and chemistry of illite show that the synkinematic clay minerals with the S2-stage
42 formed under upper anchizone conditions (~250-300°C). Such thermal conditions are higher than
43 those expected from the inferred maximal burial (~6 km), thus arguing for a possible circulation
44 of hot fluids along the thrust footwall. $^{40}$Ar/$^{39}$Ar step-heating results confirm that the newly-
45 formed illite is more abundant in the finer fractions of the most deformed sample, with a total gas
46 age of c. 61.7 ± 0.5 Ma in the finest <0.2µm fraction. This age is older than the early-mid Eocene
47 age expected from structural constraints for the emplacement of the Lakora thrust sheet due to the
48
mixing of newly-formed illites with small amounts of inherited micas and/or contamination by
excess argon carried by fluids.

**Keywords:** $ ^{40}\text{Ar}/^{39}\text{Ar} $ geochronology; Clay minerals; Sedimentary basins; Thrust faults;
Pyrenees; Lakora thrust.

## 1. Introduction

Clay minerals, which are major constituents of sedimentary rocks, are an important source of information and are commonly used to unveil the geological processes responsible for their formation (Frey, 1987; Meunier, 2005; Abd Elmola, 2017). The texture arrangement, chemical composition, and crystal structure of synkinematic clay minerals in fault zones can provide important information about the mechanisms (ductile vs. brittle) and conditions (P-T-t and grade of diagenesis-metamorphism) of deformation (Buatier et al., 2012, 2015; Lacroix et al., 2012; Torgersen et al., 2015; Aldega et al., 2019; Abd Elmola et al., 2020; Curzi et al., 2020a, 2020b; Tartaglia et al., 2020). For example, the illite/muscovite ‘crystallinity’ index is a commonly used method for evaluating the diagenetic and low-temperature metamorphic history of the host rocks (Kubler, 1968; Frey, 1987; Kisch et al., 2004; Ferreiro Mählmann et al., 2012; Warr and Cox, 2016). In addition, illite/mica shows irreversible polytype transformation from $1M_d$ to $1M$ to $2M_1$ with the increase of diagenesis and metamorphism grade (Weaver, 1989; Lonker and Fitz Gerald, 1990; Pevear, 1999; Meunier, 2005). Therefore, characterizing the crystallinity and polytypes of newly-formed mica/illite in fault zones can allow assessing the thermal conditions of fault activity (Velde, 1965; Moore and Reynolds, 1997). Furthermore, K-bearing clay minerals can be used to date fault activity when they form simultaneously with deformation (McDougall and Harrison, 1999; Viola et al., 2018; Aldega et al., 2019; Curzi et al., 2020a, 2020b; Nordbäck et al., 2022). However, dating thrust activity at diagenetic to low-grade metamorphism conditions
by the <sup>40</sup>Ar/<sup>39</sup>Ar method is not straightforward. This is because the dated sample usually contains a mixture of mica/illite with different ages, polytypes, and origins (McDougall and Harrison, 1999). Moreover, fluid circulation in fault zones may also result in <sup>40</sup>Ar loss or gain, thus producing misleading ages (Spikings and Popov, 2021). Therefore, a detailed multiscale structural, petrographical, mineralogical, and chemical characterization of the samples selected for dating is very critical to (i) identify the different mica/illite generations, (ii) determine a strategy (e.g., grain-size fractionation) to separate them, (iii) study the possible impact of fluid circulations, and (iv) evaluate the significance of the obtained <sup>40</sup>Ar/<sup>39</sup>Ar ages (e.g., Clauer, 2013; Abd Elmola et al., 2018; Tartaglia et al., 2020).

The Pyrenean orogen is a collisional belt formed by the convergence of the Iberian and Eurasian plates during the Late Cretaceous to Early Miocene (e.g., Roure et al., 1989; Muñoz, 1992; Mouthereau et al., 2014; Teixell et al., 2018) (Fig. 1a). Previous studies conducted in fault zones within the range (e.g., McCaig et al., 2000; Trincal et al., 2014, 2017; Abd Elmola et al., 2017, 2018; Bellahsen et al., 2019; Caldera et al., 2021) highlighted the brittle-ductile mechanisms of deformation and the prominent role of phyllosilicate neoformation associated to fluid circulation. Dating tectonic activity in the Pyrenees still largely relies on studies of growth strata in the foreland basins (e.g., Teixell, 1996; Vergés et al., 2002; Muñoz et al., 2013; Labaume et al., 2016), and on thermochronological analyses (e.g., Sinclair et al., 2005; Jolivet et al., 2007; Metcalf et al., 2009; Bosch et al., 2016; Labaume et al., 2016). However, these approaches may be challenging to apply to thrusts in the inner belt where growth strata are missing due to erosion or non-deposition, and thermochronological results are debatable because of the stacking of various thrusts. Direct dating of newly-formed phyllosilicates by the <sup>40</sup>Ar/<sup>39</sup>Ar method may be a useful tool to overcome these difficulties. However, the method has been used so far in a limited
number of thrusts in the Pyrenean orogen (e.g., Rahl et al., 2011; Abd Elmola et al., 2018). In this
paper, we focus on the Lakora thrust, which is a major south-vergent fault in the west-central
Pyrenees (Teixell, 1996; Labaume and Teixell, 2020) (Fig. 1a). We combined microstructural,
mineralogical, and chemical observations to unravel the conditions and mechanisms of
deformation which prevailed during the activity of the Lakora thrust. We also performed
$^{40}$Ar/$^{39}$Ar radiometric clay dating to give constraints on the age of thrust activity, which remained
loosely defined due to the lack of preserved growth strata (Teixell, 1993; Bosch et al., 2016;
Teixell et al., 2016).

Our results allowed us to define the mechanisms of brittle-ductile deformation during the Lakora
thrusting under thermal conditions of the upper anchizone that are coherent with previous results
in the studied area. On the other hand, the Palaeocene $^{40}$Ar/$^{39}$Ar radiometric age obtained from
the present study, although falling in the age interval inferred regionally for the Lakora thrust,
appears too old to be consistent with the structural constraints on the studied outcrop.

## 2. Geological setting

The Pyrenees belt comprises three main structural domains, which from north to south are
represented by: 1) the North Pyrenean Zone, which corresponds to a Mesozoic extensional basin
inverted during the Pyrenean compression, 2) the Axial Zone, which corresponds to imbricated
south-vergent basement thrust units that expose Paleozoic rocks deformed during the Hercynian
orogeny, and 3) the South Pyrenean Zone, characterized by a south-vergent detached fold-and-
thrust system in the Meso-Cenozoic cover (e.g., Teixell et al. 2018, and references therein) (Fig.
1a).
The Lakora thrust corresponds to the southern leading edge of the North-Pyrenean Zone pop-up in the west-central Pyrenees (Fig. 1a). The studied Lakora thrust outcrop is located along the cross-section shown in Figure 1b. In the St-Engrâce area, where the thrust is classically defined, the thrust crops out as a gently-dipping fault system which hanging wall comprises a basal slice of Triassic rocks (labeled the Larrau Triassic in Fig. 1b) and a main sheet of Albian to Santonian conglomerates and breccia resting on Paleozoic metasediments (the Iguntze massif in Fig. 1b) (Teixell, 1993, 1996; Teixell et al., 2016). To the south, the basal thrusts of these two units merge into a single one forming the sole thrust of the Lakora klippe, which corresponds to an extension of the Igunze massif and exposes the hanging wall cut-off of the Santonian breccia. The Iguntze massif represents the upper part of the Iberian margin of the Mesozoic rift (Durand-Wackenheim et al., 1981; Boirie and Souquet, 1982; Souquet et al., 1985). To the north, the central part of the rift is represented by the Mauléon Basin, detached on the Upper Triassic (Keuper) layer and comprising a folded Jurassic and Lower Cretaceous carbonate succession overlain by thick (up to around 5000 m) Albian to Santonian syn- to post-rift marls and flysch (Saspiturry et al., 2019). The Lakora thrust system roots in the Triassic décollement level and probable basement thrusts underlying the Mauléon Basin (see discussion in Labaume and Teixell, 2020).

The Upper Cretaceous of the Lakora thrust footwall corresponds to the cover of the Axial Zone and comprises Cenomanian to Santonian limestones followed by Campanian-Maastrichtian marls and turbidites (Fig. 1c). Below the Lakora klippe, where the studied outcrop is located, the thrust is parallel to bedding in the Maastrichtian turbidites, defining a large footwall flat. South of the klippe, the footwall Maastrichtian flysch and marls are overlain by the Paleocene and Lower Eocene flysch of the northern edge of the Jaca South-Pyrenean foreland basin (Labaume et al., 1985, 2016; Teixell, 1996) (Fig. 1b). The relationships between the Lakora thrust and the folds
and thrusts in the Upper Cretaceous-Tertiary to the south are elusive on the BRGM geological map (Casteras and Souquet, 1970), but in the Pic d'Orhy area, 11 km west of the cross-section shown in Figure 1b, mapping constraints strongly suggest a connection of the Lakora thrust with a thrust unit comprising the Campanian-Maastrichtian to Lower Eocene (Ducasse and Velasque, 1988). This thrust unit (referred to as the Orhy thrust unit hereafter) is shown above the topography in Figure 1b. In the absence of other constraints, it is shown with a 4.5 km displacement, similar to the one observed in the Pic d'Orhy area (Casteras and Souquet, 1970).
Structural and stratigraphic constraints used for Figure 1b cross-section construction imply a minimum displacement of around 16 km on the Lakora thrust at the level of the Santonian breccia hanging wall cut-off (see Teixell, 1993, for the detailed restoration of an analogous section located a few km to the east). Consequently, we assume that the main part of the Lakora displacement (11.5 km) may have occurred on a branch of the thrust passing above the Orhy thrust unit and on a hypothetical back-thrust at the top of the Santonian breccia. More to the south (Isaba area in Fig. 1b), we assume that tectonic repetitions of a breccia key-bed in the lower Eocene flysch, with a total displacement of about 5 km (Labaume et al., 1985), are connected to the Orhy thrust unit.

The Lakora thrust is underlain by the Larra thrust, which affects the footwall Upper Cretaceous limestones to the north and rises in the series to the south (Fig. 1b, c). Both thrusts are deformed/folded by the basement culmination of the underlying Gavarnie thrust.

The Lakora thrust results from the inversion of the North Pyrenean extensional basin margin and corresponds to the northernmost and older thrust of the south-vergent Pyrenean thrust system. Based on the kinematics of the Pyrenean convergence and indirect relationships with foreland basin deposits, the thrust activity is inferred to have begun during the latest Cretaceous in the
162 North-Pyrenean Zone, then to have propagated southward in the South-Pyrenean foreland basin
163 until the mid-Eocene (Teixell, 1996; Teixell et al., 2016). However, the precise thrusting age
164 remains poorly defined in the study area due to the lack of syn-tectonic growth strata. The Larra
165 thrust formed as a footwall splay of the Lakora thrust, probably during the late Lutetian-
166 Bartonian, as it is suggested by the age of the youngest sediments affected by the Larra thrust in
167 the Jaca basin (Teixell, 1996). The Axial Zone antiform resulting from the Gavarnie thrust
168 emplacement deformed both the Lakora and Larra thrusts. The Gavarnie thrust was active during
169 the mid to late Eocene-early Oligocene, as it is shown by tectono-sedimentary relationships in the
170 Jaca basin (see discussion in Labaume et al., 2016). This is also supported by the $^{40}$Ar/$^{39}$Ar dating
171 on a lower branch of the Gavarnie thrust in the Axial Zone (Abd Elmola et al., 2018) and zircon
172 (U-Th)/He thermochronology data in the Lakora thrust hanging wall and footwall (Bosch et al.,
173 2016).

174 3. Sampling and analytical techniques

175 Samples have been collected from the footwall Maastrichtian turbidites in deformed intervals
176 close to the main thrust contact below the Lakora klippe (Figs. 1c and 2; for detail of the sample
177 structural location and GPS coordinates, see Supplementary Figures S1 and S2). However, the
178 thrust surface, located at the top of the outcrop, could not be observed due to bad exposure, and
179 the samples were collected a few meters to 15 meters from its inferred location. The samples
180 were investigated by X-ray diffraction (XRD), optical microscope, and scanning electron
181 microscope (SEM) coupled with energy-dispersive X-ray spectroscopy. Then, the $^{40}$Ar/$^{39}$Ar step-
182 heating on muscovite/illite was applied to different size fractions of two representative samples
183 showing different strain intensities.
### 3.1. X-ray diffraction

The bulk and clay mineralogy of the samples were characterized at the ITINAM laboratory (Université de Bourgogne Franche-Comté) using a Bruker D8 Advance diffractometer system, Cu-Kα radiation equipped with a fast LynxEye linear detector. For the bulk composition characterizations, the X-ray diffractograms of the random powder mounts were acquired between 5 and 70°2θ at a step size of 0.02, a number of steps 3450, a time/step 1s at 40 kV and 40 mA.

Samples were hand-ground to ~ 50 µm, and the < 16 µm fraction obtained by sedimentation was submitted to ultra-centrifugation to extract the 2-0.5 µm, 0.5-0.2 µm, and <0.2µm fractions. For the clay fractions, the oriented specimens on glass slides were analyzed in air-dried conditions between 3 and 30°2θ at a step size of 0.01, a number of steps 2228, and time/step 0.5s. No expandable or interstratified minerals were found, as confirmed by ethylene-glycol saturation.

For mica/illite polytypes determination, the X-ray diffractograms of random powder mounts were acquired between 20 and 32°2θ at a step size of 0.02, a number of steps 3450, and a time/step 1s.

### 3.2. Scanning electron microscopy

Petrographic characterizations were performed on polished thin sections using optical microscopy and a JEOL JSM 5600 scanning electron microscope at the FEMTO-ST Institute (Université de Bourgogne Franche-Comté). The chemical composition of clay minerals was analyzed on carbon-coated polished thin sections, with SEM equipped with an Energy Dispersive X-ray Spectroscopy (EDX) detector and XFlash 4030 Silicon drift detector. The analytical conditions for the quantitative EDX analysis of clay minerals were as follows: acceleration voltage 15 kV, current beam 1 nA, counting time 50 s, working distance 11 mm, and analytical area of ~ 2µm.
206 **3.3. Kübler Index**

207 The Kübler Index (KI) was measured on mica/illite $d_{001}$ reflection at around 10 Å in oriented
208 preparations of the <2 µm fractions to determine the diagenetic-metamorphic grade conditions.

209 Interlaboratory variations have been reduced by following the recommendations of Kisch (1991)
210 and using the widely applied standardized Crystallinity Index Scale (CIS) procedure
211 recommendations for measurements. The calibration of the crystallinity results to the CIS scale
212 was achieved by measuring the crystallinity of the four standards proposed by Warr and Rice
213 (1994). Our measurements (y) were transformed into CIS values (x) according to the equation: y
214 = 0.6574 x + 0.0082, (R² = 0.9993). The evaluation of metamorphic conditions boundaries was
215 performed by using the boundaries of the anchizone originally defined by Kübler (1967) at
216 0.25°2θ and 0.42°2θ. To be compatible with the original anchizone boundaries of Kübler (1967),
217 the CIS values were calibrated to the traditional Kübler-Kisch scale by using the equation: Kübler
218 Index (CIS) = 1.1523 * Kübler Index (Basel lab) + 0.036, which is presented in Figure 1 of Warr
219 and Ferreiro Mählmann (2015).

220 **3.4. <sup>40</sup>Ar/<sup>39</sup>Ar step-heating**

221 Two samples (i.e., 15 LAK 1 and 15 LAK 3) with pelitic lithology have been selected for the
222 <sup>40</sup>Ar/<sup>39</sup>Ar dating, representative of different intensities of deformation observed in the thrust
223 footwall damage zone (see Supplementary Figures S1 and S2 for sample location). Sample
224 preparation was performed without any chemical treatment to avoid any impact on the clay
225 mineral crystal structure. The bulk, 50-16 µm, 16-2 µm, 2-0.5 µm, 0.5-0.2 µm, and <0.2 µm
226 fractions of the two samples were compacted mechanically, and about 1 to 2 mg of these
227 fractions were encapsulated in quartz tubes at 10<sup>-8</sup> bar to evaluate <sup>39</sup>Ar recoil effects (e.g., Dong
228 et al., 1995; van der Pluijm et al., 2001). The samples were then irradiated in the core of the Triga
Mark II nuclear reactor of Pavia (Italia) with several aliquots of the Taylor Creek sanidine standard as flux monitors. Argon analyses were performed at the Géosciences Montpellier laboratory (Montpellier University, France) with a multi-collector mass spectrometer (Argus VI from Thermo-Fisher). For more details about the <sup>40</sup>Ar/<sup>39</sup>Ar step-heating analytical procedure, see the methodology part in Abd Elmola et al. (2018).

## 4. Results

### 4.1. Field observations

On the studied outcrop, the thrust surface and the footwall turbidites dip about 20° to the NNE, thus defining a footwall flat, while the breccia in the hanging wall is sub-vertical (Figs. 1c, and 2). The turbidites comprise alternations of plane-parallel, up to decimeter-thick, fine-grained sandstone beds and pelite layers.

Deformation in the footwall turbidites is marked by various mesoscale features (Fig. 3a to e and Supplementary Figure S2). Conspicuous is mostly parallel to the bedding cleavage in the pelites and finest sandstones, hereafter named S0-S1. Slight obliquity to the bedding of the S0-S1 cleavage is nevertheless locally observed, with a northward dip with respect to bedding (Fig. 3b). The sandstone beds are often affected by boudinage, with boudins moulded by the S0-S1 cleavage in the pelites (Fig. 3b). They also exhibit stylolites parallel to sandstone lamination and with bedding-normal peaks (Fig. 3f). Extension calcite veins normal to bedding are frequent in sandstone layers, particularly at boudin necks (Fig. 3b, e). The vein calcite is generally fibrous with fibers sub-parallel to bedding (Fig. 3e). Deformation is also marked by south-vergent reverse faults and folds and a north-dipping S2 cleavage oblique to bedding and axial planar of folds (Fig. 3a, c, d). The S2 cleavage is generally discontinuous and restricted to fold hinges and
their close vicinity. Reverse faults and folds are often located at boudin necks, attesting they are late features with respect to boudinage (Fig. 3d).

## 4.2. Petrography and microstructures

The petrography and microstructures are mainly described here from the optical microscope and SEM images shown in Figure 3f for sandstones and Figures 4, 5, and 6 for pelite.

The sandstone sample presented here was collected about 10 m below the main thrust surface and is characterized by bedding-parallel stylolites, which mark the S0-S1 cleavage at the macroscopic scale (Fig. 3e) and appear as dark seams of insoluble material under the optical microscope (Fig. 3f). Quartz grains have variable shapes, but the elongated ones are sub-parallel to the cleavage, and some of them feature cleavage-normal extension fractures sealed by calcite (cf. domain labeled 1 in Fig. 3f). Quartz grains in sandstone often feature strain fringes of fibrous calcite cement with corrosion of quartz at the contact with calcite. Calcite fibers are generally sub-parallel to cleavage (domain 2 in Fig. 3f), but some calcite fringes show curved fibers indicative of quartz grain rotation (domain 3 in Fig. 3f).

Figures 4, 5, and 6 show two representative pelite samples (also used for the <sup>40</sup>Ar/<sup>39</sup>Ar dating) collected about 5 m below the main thrust surface. The less-deformed 15 LAK 1 sample features the S0-S1 cleavage (Fig. 4a, c, e), while the more-deformed 15 LAK 3 sample features a cm-scale fold hinge with an axial planar S2 cleavage overprint the S0-S1 cleavage (Fig. 4b, d, f). The mineralogical composition of both pelite samples mainly comprises phyllosilicates and quartz and subsidiary rutile, pyrite, and albite (Figs. 4, 5, and 6). The phyllosilicate minerals are micas and chlorite in the form of discrete grains or mica/chlorite stacks. Two size classes of micas and chlorite can be identified, with large (~50 µm) and small (<5 µm) grains.
The S0-S1 cleavage in pelite is a pervasive grain-shaped foliation marked by the preferred orientation of phyllosilicate plates and the long axis of often elliptic quartz grains, which feature corroded surfaces filled with small grains of mica and chlorite (Figs. 4e, 5a, c, and 6a, c). In general, small micas and chlorite crystallize parallel to the S0-S1 cleavage (Figs. 6a, c), but they can also display random orientation in the pressure shadow of quartz grains or aggregates associated with small quartz or rare feldspar grains (Figs. 5c, and 6e, f). The S2 cleavage is a discrete crenulation cleavage (Fig. 4f) that separates microlithons in which the S0-S1 is folded (Fig. 5b, d). Microfolds are well marked by phyllosilicates asymptotically bent along the S2 cleavage (Fig. 5d). Both large and small mica and chlorite grains are concentrated along the S2 shearing plane and are associated with Fe-oxides (Fig. 5b, f). Chlorite/mica stacks that follow the S0-S1 cleavage in the 15 LAK 1 sample are folded, kinked, and display smaller grain sizes in the more-deformed 15 LAK 3 sample (Fig. 6b, d, f). Small mica and chlorite crystals (< 1µm) are more abundant in the more-deformed sample. Most of those crystals that are associated with the S2 cleavage display orientation parallel to the latter (Figs. 5f and 6d). However, small micas are also observed in open spaces along the cleavage planes (S2), where they grew perpendicular to the cleavage (Fig. 6d).

### 4.3. Bulk composition and clay mineralogy, polytypism, and crystallinity

The XRD results of bulk powder of five representative pelite samples collected < 10 m below the main thrust surface (including the 15 LAK 1 and 15 LAK 3 samples described above) are presented in Figure 7. The XRD results of all the studied samples are presented in Supplementary Figure S3. The results show that all the samples have similar mineralogical assemblages composed of quartz, chlorite, mica/illite, albite, and rutile, confirming the microscope observation presented above. The only difference is the absence of calcite in the 15 LAK 1 and
15 LAK 3 samples. The $d_{060}$ reflections at 1.50 Å and 1.55 Å correspond to a dioctahedral clay phase (muscovite/illite) and a trioctahedral phase (chlorite), respectively. The oriented preparations of the fractions 16-2 µm, 2-0.5 µm, 0.5-0.2 µm, and < 0.2 µm of the 15 LAK 1 and 15 LAK 3 samples confirm that muscovite/illite and chlorite are present in all the fractions (Fig. 8). The muscovite/illite is characterized by the following reflections: 10.02 Å, 5.00 Å, and 3.33 Å corresponding to $d_{001}$, $d_{002}$, and $d_{003}$ basal reflections, respectively. Reflections at 14.16 Å, 7.08 Å, 4.73 Å, and 3.54 Å correspond to $d_{001}$, $d_{002}$, $d_{003}$, and $d_{004}$ reflections of chlorite. The $d_{002}$ and $d_{004}$ reflections of chlorite are more intense than the $d_{001}$ and $d_{003}$ reflections, suggesting Fe-rich chlorite (Brown and Brindley, 1980; Moore and Reynolds, 1997). In both 15 LAK 1 and 15 LAK 3 samples, the $d_{001}$ reflection peak of muscovite/illite (~10 Å) shows an asymmetrical shape with a shoulder on the low angle side of the diffractograms suggesting a peak overlapping between more than one mica component. Besides chlorite and muscovite/illite, the XRD patterns show quartz, a small amount of rutile, and albite that decrease in the finest fractions. The small pyrite crystals observed on SEM images were not detected by XRD.

The XRD diffractograms (2θ=32°) of the randomly oriented preparations of bulk powder and finer size fractions (2-0.5 µm, 0.5-0.2 µm, and < 0.2 µm) of the 15 LAK 1 and 15 LAK 3 samples are shown in Figure 9. In both samples, the results indicate that the $2M_1$ muscovite/illite is the most abundant polytype with a relatively very low amount of $1M$ polytype, as confirmed by their peak intensity ratio (Moore and Reynolds, 1997). The $1M_d$ polytype is not present in the samples as no hump shape was observed at the $d_{003}$ muscovite/illite reflection position, which is a characteristic of the $1M_d$ mica polytype. With decreasing grain size, the intensity of the characteristic peaks of the $1M$ mica polytype decreases, as demonstrated in Figure 9. In the finest
318 < 0.2 µm fraction, the 2M₁ polytype reflections are the most abundant phase with almost no or
319 very low proportion of the 1M polytype.

320 The Kübler index (KI) calculations of the d<sub>001</sub> muscovite/illite reflections for five representative
321 samples are presented in Table 3. The samples have Kübler index values ranging from 0.31°2θ to
322 0.49°2θ (Kübler-Kisch scale) that correspond to diagenetic-anchizone grade conditions. The less-
323 deformed 15 LAK 1 sample yields the highest KI value of 0.49°2θ (Kübler-Kisch scale),
324 corresponding to diagenetic conditions. In comparison, the more-deformed 15 LAK 3 sample has
325 the lowest value of 0.31°2θ, which indicates upper-anchizone conditions. This classification of
326 diagenetic-metamorphism grade is based on the anchizone grade boundaries of 0.25°2θ to 0.42
327 °2θ, and 0.335°2θ as a limit between upper and lower anchizone on the KF scale (Kübler, 1967;
328 Warr and Ferreiro Mählmann, 2015).

329
4.4. *Chemical composition of clay minerals*

330 Quantitative SEM-EDX point analyses were performed on the two groups of chlorites and
331 mica/illite identified in both the 15 LAK 1 and 15 LAK 3 samples. The corresponding structural
332 formulas were calculated from the normalized oxides weight percent (wt.%) (based on 11 and 14
333 oxygens for mica and chlorite, respectively) (Tables 1 and 2) (Meunier, 2005; Bergaya, 2008;
334 Wilson, 2013). Total Fe has been considered as Fe<sup>3+</sup> for mica and as Fe<sup>2+</sup> for chlorite. In addition,
335 Mg was assumed to be only present in the octahedral layer as it was not possible to calculate the
336 part that may be present in the interlayer sheets of both mica and chlorite (Bailey, 1988; Meunier
337 and Velde, 2004; Abd Elmola et al., 2020).

338 The chemical analyses of mica/illite particles are plotted in 4Si-M⁺-R²⁺ ternary diagram where
339 4Si = (Si/4), M⁺ =Na⁺ +K⁺ + 2Ca²⁺; R²⁺ = (Mg + Fe⁺² + Mn) (Meunier and Velde, 1989) (Fig.
340 10a). In both samples, the large mica grains have high Al₂O₃ and low MgO contents ranging from
341 33 to 37% and 0.5 to 0.9%, respectively, and 7-10 % of K₂O (Table 1). Their chemical
342 compositions plotted in the 4Si-M⁺-R²⁺ ternary diagram are located close to the muscovite end
343 member (Fig. 10a). The small micas have Al₂O₃ and MgO contents of 33 to 38% and 0.5 to 1%,
344 respectively, and K₂O content of 5-7%. They have a chemical composition located close to the
345 compositional field of illite (Fig. 10a) with relatively low K₂O content compared to the large
346 muscovite grains (Table 1).

347 The chemical analyses of chlorite particles are plotted in Mg-Al(VI)-Fe ternary diagram where
348 Al(VI) = octahedral aluminum (Meunier and Velde, 1989) (Fig. 10b). The large chlorite grain
349 compositions in both samples are located between the sudoite, amesite, and chamosite end
350 members (Fig. 10b), with relatively high Fe content: FeO ranges from 26 to 33%. In both
351 samples, SiO₂, Al₂O₃, MgO, and TiO₂ contents in the large chlorite grains range from 32 to 37%,
352 22 to 24%, 7 to 9%, and 1 to 2%, respectively (Table 2). In the less-deformed 15 LAK 1 sample,
353 the large chlorite grains have slightly higher Al and less Fe contents compared to chlorite from
354 the more-deformed 15 LAK 3 sample. On the other hand, the small-grained chlorite in both
355 samples is more homogenous. Compared to the large chlorite, the small grains have almost no
356 TiO₂, relatively more iron content (~ 34-37%), and chemical composition more shifted to the
357 field of the chamosite end member (Table 2 and Fig. 10b).

358 4.5. $^{40}$Ar/$^{39}$Ar step-heating data

359 Step-heating experiments were performed on the bulk and the fractions 50-16 µm, 16-2 µm, 2-0.5
360 µm, 0.5-0.2 µm, and < 0.2 µm of both the 15 LAK 1 and 15 LAK 3 samples. A summary of the
361 calculated ages is presented in Table 4, and complete isotopic data are presented in
362 Supplementary Table S1. The age spectra are corrected from $^{39}$Ar recoil effects; the recoil
363 fractions are plotted at the first step on the spectra with a zero age (Fig. 11).
364 **4.5.1. Sample 15 LAK 1 (less-deformed)**

365 In this sample, the coarse fractions (bulk, 50-16 µm, and the 16-2 µm) produced spectra with 366 three distinct segments: a hump-shaped profile for the low-temperature steps followed by a 367 staircase profile with steadily climbing apparent ages at intermediate temperatures and a peak- 368 like at higher temperatures (Fig. 11). By contrast, the spectra of the fine fractions (2-0.5 µm, 0.5- 369 0.2 µm, and < 0.2) have a relatively flat shape for the last part of the degassing experiment. The 370 maximum age which is reached during the last step-heating experiments decreases with 371 decreasing grain-size and ranges between c. 303.2 ± 1.5 Ma (50-16 µm) to c. 110.4 ± 0.9 Ma (< 372 0.2 µm fraction) (Table 4). The total gas age is, in general, directly proportional to the grain-size 373 and decreases from c. 140.4 ± 1.0 Ma (50-16 µm) to c. 71.3 ± 0.5 Ma (< 0.2 µm fraction). In 374 addition, the amount of <sup>39</sup>Ar recoil increases as well when decreasing the grain size; the finest 375 fraction (< 0.2 µm) has the highest value of 25.4% of the total <sup>39</sup>Ar (Table 4).

376 **4.5.2 Sample 15 LAK 3 (more-deformed)**

377 Similar to the less-deformed 15 LAK 1 sample, the bulk and the coarse fractions of the more- 378 deformed sample produced staircase-shaped spectra (Fig. 11). The fine fractions (2-0.5 µm, 0.5- 379 0.2 µm, and < 0.2 µm) also display relatively flat spectra for the high-temperature steps. The 380 maximum ages range between c. 274.9 ± 1.4 Ma (16-2 µm fraction) to c. 103.2 ± 1.0 Ma (< 0.2 381 µm fraction) (Table 4). The total gas ages range between c. 129.9 ± 1.1 Ma (16-2 µm fraction) to 382 c. 61.7 ± 0.5 Ma (< 0.2 µm fraction). The < 0.2 µm fraction experienced the highest amount of 383 <sup>39</sup>Ar recoil compared to the coarser fraction with a value of 27.8%.

384 These results show that the fine fractions from both samples have younger total gas ages than the 385 coarse fractions. Moreover, all the fractions from the more-deformed 15 LAK 3 sample have 386 younger ages compared to the same fractions from the less-deformed 15 LAK 1 sample. Finally,
387 the finest fraction (< 0.2 µm) in 15 LAK 3 has the youngest age compared to all the fractions
388 from both samples.

389 **5. Discussion**

390 **5.1. Stages and mechanisms of deformation associated with thrusting**

391 Structural observations in the Lakora thrust footwall fault zone show that deformation occurred in
392 two stages, as shown in Fig. 12. The first stage corresponds to bedding-parallel extension and
393 bedding-perpendicular shortening marked by the boudinage of the competent sandstone beds,
394 while the continuous S0-S1 cleavage developed mostly parallel to bedding in the more ductile
395 pelite beds, moulding the boudins. Bedding-normal extension veins at boudin necks and bedding-
396 parallel stylolites in sandstone beds, as well as calcite fibers parallel to the S0-S1 cleavage in
397 strain fringes around quartz grains, are also related to this deformation stage. However, slight
398 local obliquity of cleavage with respect to bedding and curved calcite fibers in some strain fringes
399 suggest that deformation was not purely coaxial but may have involved a component of top-to-
400 south shearing associated with the development of the S0-S1 cleavage. The second deformation
401 stage corresponds to top-to-south bedding-oblique shearing marked by reverse faults, folds, and
402 axial planar S2 crenulation cleavage, the latter mostly present in the folded domains.

403 Both stages of deformation are characterized by brittle-ductile mechanisms associating fractures
404 and faults with cleavage formation, the latter involving deformation of inherited phyllosilicates,
405 pressure-solution of calcite, quartz, and feldspars, and neoformation of phyllosilicates (see next
406 section for discussion of phyllosilicate origin). Mass transfer triggered by pressure-solution also
407 resulted in calcite precipitation along fractures and faults. However, the XRD and SEM results
408 show similar bulk mineralogy of all samples, composed of abundant quartz, chlorite,
409 muscovite/illite, and subsidiary rutile, pyrite, and albite. The only difference is the presence of
calcite in some samples with different proportions and the presence of Fe-oxides (hematite) in the more-deformed 15 LAK 3 sample. The scarce presence of Na-feldspars in the matrix (associated with clay minerals) and the absence of K-feldspars (despite K availability) suggest that these unstable minerals were rare in the sediments or they have been dissolved/transformed into clay minerals during diagenesis and deformation. Gratier et al. (2013) and Janssen et al. (2016) suggested similar mechanisms in active faults of the Western Alps and San Andreas, where the dissolution of soluble minerals (feldspars, quartz, and calcite) and neoformation of phyllosilicates in the presence of fluids have been described.

The prominent fluid-assisted ductile mechanism involved in the deformation suggests an important component of slow rate and aseismic activity of the Lakora thrust, as proposed for other fault zones featuring similar deformation mechanisms (e.g., Gratier and Gamond, 1990; Renard et al., 2000; Gratier et al., 2011; Lacroix et al., 2011).

## 5.2. Origin of clay minerals along the Lakora thrust

The clay minerals observed along the Lakora thrust are Fe-rich chlorite and dioctahedral mica/illite based on XRD results. The Fe-rich nature of chlorite is indicated by the $d_{002}$ and $d_{004}$ reflections that are more intense than $d_{001}$ and $d_{003}$, while the dioctahedral nature of mica/illite is indicated by $d_{060} \sim 1.50$ Å, and both are confirmed by the SEM-EDX analyses (Moore and Reynolds, 1997). Based on their textural arrangement and chemistry, these chlorite and mica/illite can be classified into three types: detrital, diagenetic, and newly-formed. Detrital chlorite and mica are characterized by very large grains and strong deformation features marked by folding and kinked morphology. XRD and SEM-EDX analyses confirm that the detrital chlorites are Fe-rich, and detrital micas have a muscovite-like composition. The $TiO_2$ content detected in the chemical composition of the detrital chlorite suggests that at least part of these detrital clay
minerals was initially of hydrothermal or metamorphic origin (Abd Elmola, 2017 and references therein). This hypothesis is in agreement with other studies that demonstrated the formation of TiO₂ rich-chlorite (~1-2% TiO₂ in the present study) in hydrothermal and/or metamorphic environments through the replacement of biotite precursor (e.g., Parry and Downey, 1982; Eggleton and Banfield, 1985; Jiang and Peacor, 1994; Yuguchi et al., 2015). The presence of diagenetic clay minerals is marked by the presence of large chlorite/mica stacks that are parallel to bedding in the less-deformed 15 LAK 1 sample. Several studies have described these stacks in sediments that have experienced burial diagenesis (e.g., Milodowski and Zalasiewicz, 1991; Jiang and Peacor, 1994; Bartier et al., 1998; Randive et al., 2015; Abd Elmola et al., 2017). According to these authors, chlorite/mica stacks are formed by the alteration of detrital biotite into chlorite through a layer-by-layer transformation, and then muscovite progressively fills the chlorite cleavages. These chlorite/mica stacks are affected by deformation induced by the Lakora thrust activity as marked by the folding, kinking, and smaller grain size of these stacks in the more-deformed 15 LAK 3 sample, as shown in Figure 6b and described in section 4.2.

The petrographic observations show that the smallest mica and chlorite grains occur along the S0-S1 and S2 cleavages in close association with deformation features. These particular occurrences and their morphology and chemistry that differ from those of the large detrital and diagenetic phyllosilicates argue for synkinematic growth during both deformation events. In the less-deformed 15 LAK 1 sample, the presence of small phyllosilicates in pressure shadows of quartz grains, where they fill irregularities of dissolved quartz grain surfaces, suggests that their formation is related to the pressure solution process and hence is syn-kinematic to the S0-S1 cleavage (Fig. 6c). Their occurrence in aggregates associated with small quartz and (rare) feldspar grains also suggests that they formed following a dissolution-recrystallization
mechanism (Figs. 5c, and 6e). In the more-deformed 15 LAK 3 sample, the abundant small mica and chlorite grains aligned along the S2 cleavage are interpreted as syn-kinematic to this second stage of deformation (Figs. 5f and 6d), while the micas, which grew perpendicular to the S2 cleavage in opened spaces along the cleavage attest to a post-kinematic, fluid-assisted crystallization/growth (Fig. 6d). These fluids may have contained dissolved ions from the host rocks that were probably accumulated by diffusion in open spaces, fractures, and veins. The presence of Fe-oxides (hematite) and newly-formed Fe-rich chlorite aligned along the S2 cleavage (Figs. 5b, 5f, and 6d) and within the spaces opened along the cleavage (Fig. 6d) reflects that these fluids were probably rich in iron and originated from the dissolution of detrital Fe-rich minerals (biotite and/or chlorite) (e.g., Abd Elmola et al., 2020).

Another scenario could be a flow of hot fluids expelled from the deep parts of the belt and drained into the thrust footwall along the fault planes, cleavages, and fractures (e.g., Abd Elmola et al., 2017). This is consistent with the high Cl/K ratio (0.070) of the fine fraction <0.2 µm compared to all the larger grain-size fractions (~ 0.001 to 0.004) (Cl/K=0.559x³⁸ArCl/³⁹ArK; Supplementary Table S1), which indicates a more prominent interaction of the finest clay minerals with metamorphic fluids. A similar mechanism has been demonstrated by Abd Elmola et al. (2018) in the Axial Zone of the Pyrenees and by Villa et al. (2014) and Villa (2016) in the French-Italian Alps.

To conclude, our results indicate that newly-formed clay minerals recorded the progressive southward emplacement of the Lakora nappe during the two stages of deformation and that brittle-ductile deformation/shearing has enhanced their formation/growth in the studied fault zone. Furthermore, pressure-solution and dissolution-recrystallization in the presence of fluids are the main mechanisms that controlled this deformation and the enrichment of newly-formed
479 clay minerals in the fault zone. Thus, crystallinity, polytypism, and isotopic compositions of
480 newly-formed micas can be used to estimate the deformation conditions and provide time
481 constraints for the fault activity. However, as discussed below, the crystallization of newly-
482 formed illite during the two stages of deformation and their mixing with inherited micas make the
483 interpretation of the $^{40}$Ar/$^{39}$Ar results very challenging.

484 5.3. *Thermal conditions of thrusting*

485 The Kübler index values of representative samples from the Lakora footwall turbidites indicate
486 diagenetic to anchizone grade conditions. This difference in the crystallinity values is attributed
487 to the presence of different proportions of detrital, diagenetic, and newly-formed mica/illite in the
488 measured samples/fractions.

489 The less-deformed 15 LAK 1 sample registered diagenetic conditions as indicated by the KI
490 value of 0.49°2θ. These thermal conditions are consistent with Bosch et al. (2016)
491 thermochronological data. Reset U-Th/He zircon ages in both the hanging wall and footwall of
492 the thrust indicate that temperatures above 200°C were maintained until the early-mid Oligocene.

493 The more-deformed 15 Lak 3 sample recorded the lowest KI value (0.31°2θ), indicating upper-
494 anchizone conditions ~250-300°C (Abad, 2007, and references therein). These thermal conditions
495 are consistent with the peak temperatures of 220 to 250°C deduced from Raman Spectrometry of
496 Carbonaceous Material (RSCM) by Saspiturry et al. (2020, 2022) in the Maastrichtian turbidites
497 of the Lakora thrust footwall south of the Iguntze massif. The higher strain degree or possible
498 focused hot fluid flow affecting this sample could explain the lower (better) crystallinity index
499 compared to the less-deformed 15 LAK 1 sample (Kubler, 1968; Eberl and Velde, 1989; Ferreiro
500 Mählmann et al., 2012).
The finer clay fractions of the two samples are composed of newly-formed illite with an abundant 2M₁ polytype (Fig. 9). This indicates that the thermal conditions during thrusting were ~250-300°C (i.e., formation temperature of the 2M₁ polytype; Velde (1965), Srodon and Eberl (1984), and Meunier and Velde (2004)). Such thrusting thermal conditions are less favorable to the formation of the 1M polytype as it usually grows at <200°C (Velde, 1965; Srodon and Eberl, 1984). This suggests that the origin of the 1M mica polytype observed in the coarse and intermediate fractions of the samples is probably diagenetic.

### 5.4. $^{40}$Ar/$^{39}$Ar step-heating data

As mentioned above, the majority of the newly-formed illite-like grains in the studied samples are of the 2M₁ polytype, as confirmed by the XRD results (Fig. 9). This precludes using the mica polytype-age correlation method in the present study (e.g., van der Pluijm et al., 2001; Šegvić et al., 2021) since this method uses a binary mixing model between an inherited 2M₁ polytype and an authigenic 1M polytype to calculate the different age components. However, the petrographic and textural observations suggest that discrimination between the different muscovite/illite generations is possible through their grain size (van der Pluijm et al., 2001, 2006; Abd Elmola et al., 2018). As in the two studied samples, the detrital muscovite grains have a larger grain size than the newly-formed ones, the coarse fractions should have older ages compared to the fine fractions. The question arises whether the finest grain size fractions can provide an unequivocal answer to the age of the Lakora thrusting. For this, we have to consider the argon budget in the different grain-size fractions resulting from mixing different argon reservoirs due to the coexistence of inherited, diagenetic and authigenic micas. Moreover, due to significant recoil effects in the finest fractions, only total gas ages (TGA), including $^{39}$Ar recoil, rather than retention age (RA), excluding $^{39}$Ar recoil, will be considered in the discussion below.
524 **5.4.1. Significance of the <sup>40</sup>Ar/<sup>39</sup>Ar step-heating data**

525 In both 15 LAK 1 and 15 LAK 3 samples, the spectra of the bulk and the coarse fractions have a
526 staircase-shaped profile, while the fine fractions (i.e., < 2 µm) have less discordant age spectra
527 (Fig. 11). These staircase-shaped spectra and the absence of a plateau, or plateau-like segment
528 can be explained by different effects: mixing of several mica generations, partial argon resetting
529 related to dissolution-recrystallization effects, grain comminution, and increase of lattice defect
530 density favoring short-circuit argon diffusion, and argon recoil (Dong et al., 2000; Verdel et al.,
531 2011). The discordant age profile recorded by the first heating increments in all fractions from
532 both samples is due to <sup>39</sup>Ar recoil (e.g., Dong et al., 1995) that reaches ~30%. The coexistence of
533 chlorite grains with muscovite/illite in all fractions could be a second reason for the discordant
534 age patterns recorded by the low-temperature steps. Part of recoiled <sup>39</sup>Ar from the rim of
535 muscovite could be trapped by adjacent chlorite grains during irradiation (Kunk et al., 2005;
536 Ghanem et al., 2016; Abd Elmola et al., 2018).

537 The last heating steps of the coarse fractions of the two samples can place minimum constraints
538 on the apparent age of the inherited muscovite present in the Maastrichtian turbidites (Verdel et
539 al., 2012). The maximum age reached during step-heating experiments in the less-deformed 15
540 LAK 1 sample is c. 303.2 ± 1.2 Ma. These detrital grains are probably derived from the low-
541 grade metamorphic basement rocks (Hercynian basement) outcropping in the eastern Pyrenees
542 and the southern foreland at Maastrichtian times (Vacherat et al., 2017). For all the dated grain-
543 size fractions, the maximum ages and total gas ages of the more-deformed 15 LAK 3 sample are
544 younger than those of the less-deformed 15 LAK 1 sample (Table 4, Fig. 11), indicating more
545 pronounced recrystallization of inherited (detrital or diagenetic) micas during shearing in the fault
546 zone (Kirschner et al., 1996). In both samples, the fine fractions generally have younger total gas
ages than the coarse fractions suggesting a preferential abundance of newly-formed illite in the
fine fractions. The finest fraction (< 0.2 µm) of the more-deformed sample displays the youngest
total gas age of c. 61.7 ± 0.5 Ma compared to c. 71.3 ± 0.5 Ma for the same fraction of the less-
deformed sample. This is consistent with the petrographic observations that show that authigenic
illite is more abundant in the more-deformed sample. However, the age of 71 Ma obtained in the
less-deformed sample cannot represent the age of illite growth during the first deformation stage
and related S0-S1 cleavage as it is older than the Maastrichtian host rocks. This indicates that
small detrital-diagenetic micas that did not experience full argon resetting during the thrusting are
still present in the dated < 0.2 µm fraction.

The question here is whether the lower Paleocene age (c.62 Ma), which falls in the age interval
inferred regionally for the Lakora thrust, (i) corresponds to the Lakora activity at this locality or
(ii) it is overestimated due to the coexistence of two newly-formed illite populations that formed
during the two stages of deformation and/or the preservation of some inherited argon in the tiny
detrital-diagenetic grains and/or excess argon carried by fluids.

Compared to the less-deformed sample, the more-deformed sample experienced deformation
conditions that were more favorable in intensity, temperature, and fluid interaction for a higher
resetting of the argon chronometer in the pre-kinematic micas (Villa et al., 2014; Bosse and Villa,
2019). However, there is no direct and indisputable isotopic argument to exclude that this age
corresponds to the emplacement of the Lakora thrust sheet during the second phase of
deformation. This assumption might be true if all the newly-formed illite formed prior to the S2
cleavage were totally reset. Alternatively, it may be a mixed age between the different neo-
formed illite populations and, therefore, too old for the S2 illite growth. This age may also be too
old if inherited micas were not fully reset and persisted in the finest fraction (< 0.2 µm), as it is
the case in the less-deformed sample which is only affected by the first deformation stage.

Finally, fluid flow may be a source of excess argon that was probably trapped in late-authigenic
illite (crystallized in the open spaces of the S2 cleavage) and thus resulting in an overestimated
age.

In the next section, the 61.7 ± 0.5 Ma age obtained from the finest fraction of the more-deformed
sample is discussed in the light of the tectono-stratigraphic constrains. Based on this discussion,
the final activity of the Lakora fault would be early-mid Eocene in age, i.e., at least 6 Ma younger
than the $^{40}$Ar/$^{39}$Ar age of the 15 LAK 3 sample. It is considered that such age difference cannot
be explained by the mixing of the different newly-formed illite populations but implies the
contribution of small amounts of inherited micas and/or excess argon carried by fluids. The
dating of more finer grain size fractions (e.g., <0.1 µm or 0.05 µm) could eventually provide a
clearer answer to this problem of partial vs. complete argon resetting of detrital micas and/or
mixing between authigenic and inherited phases (e.g., Mancktelow et al., 2015).

### 5.4.2. $^{40}$Ar/$^{39}$Ar age vs. structural and thermochronological constraints

A precise determination of the age of Lakora thrust activity from tectono-sedimentary constraints
is difficult due to the absence of preserved growth strata in the study area. However, some limits
can be deduced from mapping constraints.

The section in Figure 1b shows that the Lakora thrust is the emergence to the south of the thrust
system underlying the Mauléon Basin, which corresponds to the main part of the North Pyrenean
Zone on this transverse. The Mauléon Basin is separated from the Iguntze massif by the Licq
fault, interpreted as an ancient major diapiric structure welded by the southward displacement of
the Mauléon Basin above the décollement in the Triassic evaporitic layer during the early stage of
the Pyrenean compression (Labaume and Teixell, 2020). In this view, once the Licq fault was
welded around the end-Cretaceous, the displacement could be transferred during the earliest
Tertiary to the Iguntze massif, the latter interpreted as the short-cut of a former extensional block
of the mid-Cretaceous margin, and to a slice of Triassic Keuper facies (the Larrau Triassic in Fig.
1b) located more to the south (Teixell, 1993). In the Lakora klippe area studied here, the structure
implies that displacement on the Lakora thrust cannot be older than the Maastrichtian, owing to
the age of the footwall turbidites and that it was active at least into the Early Eocene if we
consider its connection with the Orhy thrust unit to the south. The inferred connection of the
Orhy thrust unit with a first generation of thrust units in the Eocene flysch (Labaume et al., 1985)
suggests that the thrust may have remained active during the Lutetian. In summary, the Lakora
thrust corresponds to a thrust structure possibly initiated from the latest Santonian in the North-
Pyrenean Zone and which propagated southwards during the latest Cretaceous-early Tertiary with
final displacement during the mid-Eocene.

Although the Paleocene age of 62 Ma obtained in the present study is consistent with this period
of activity, it appears incompatible with the stratigraphic continuity proposed in the Figure 1b
cross-section between the Maastrichtian turbidites dated here and the Paleocene layer exposed
about 2 km southward. Indeed, this continuity, as well as the inferred presence of the Paleocene
in the eroded Orhy thrust unit, imply that the dated turbidites were located close (a few hundred
meters) below the sea-bottom at Paleocene times, suggesting that the high-temperature thrust
deformation observed in these turbidites should have been younger. This contradiction may be
resolved if we consider that the Paleocene age records an early deformation when the thrust front
propagated in the Maastrichtian turbidites, i.e., when the thrust sheet comprising the thick Iguntze
Paleozoic, and its Cretaceous conglomerate/breccia cover were still located more to the north.
Accordingly, the high temperature recorded being probably due to the flow of hot fluids expelled
from the deep parts of the belt and drained into the thrust footwall. However, it is improbable that
the deformations observed in the fault zone at the studied outcrop may have occurred prior to the
Lakora thrust sheet emplacement. Assuming that the first deformation stage resulted from a
vertical loading, thus it can hardly be attributed to another origin than the overlying Lakora thrust
sheet (Fig. 12-1). Moreover, the ~250°-300° C temperature related to the thrust deformation is
not specific to the Lakora thrust fault zone, but it corresponds to the regional temperature
recorded by RSCM in the Lakora thrust footwall, where it decreases along strike from ~300°C in
the Ossau valley, 30 km east of the Lakora klippe, to ~250°C on the transverse studied here, to
160-180°C 30 km more to the west (Saspiturry et al., 2022). According to the latter authors, these
peak temperatures represent post-collisional heating of the thrust footwall in response to the
Tertiary thickening of the continental crust, and their evolution along strike corresponds to a
segmentation of the structure of the orogenic prism by transfer zones. Based on the estimation of
sediment volumes transferred by erosion to the foreland basin, Teixell (1998) estimated the
paleoburial above the Lakora klippe to be around 6 km, mainly made of the Upper Cretaceous to
Eocene flyschs involved in the Lakora thrust sheet. This burial would be insufficient to explain
the 250°C peak temperature recorded at the Lakora thrust footwall, giving an argument for hot
fluid flow originating from the belt interior. Peak temperatures higher than those expected from
estimated paleoburial have also been recorded in other areas in, or close to, the Axial Zone
(Labaume et al., 2016; Bellahsen et al., 2019; Caldera et al., 2021). This led these authors to a
similar interpretation of hot fluid flow derived from the belt interior, the high temperature
possibly being a relic of the thermal anomaly associated with the Mesozoic thinning of the
Pyrenean rift. However, even if hot fluid flow may have contributed to the regional peak
temperature in the Upper Cretaceous flysch, it is unlikely that this peak was reached when the
flysch was still in the foreland domain before its burial below the thrust units.
Hence, considerations of the structural and thermal evolution of the Lakora thrust footwall make it unlikely that the Paleocene age determined by the Ar dating is consistent with the age of the main shear deformation observed in the thrust footwall, but they rather argue for a lower-mid Eocene age. As it is already explained above, the underlying Larra and Gavarnie thrusts are younger than the Lakora thrust, interpreted to have been active during the Lutetian-Bartonian and the Priabonian-Rupelian, respectively (Teixell, 1996; Labaume et al., 2016). The Lakora-Larra transition may have occurred in the late Ypresian/early Lutetian, but both thrusts having low dips across the stratigraphy, although some overlap cannot be excluded.

As explained in the previous section, dating finer grain size fractions may help in understanding the origin of Ar age overestimation in the present study. Based on Saspiturry et al. (2022), studying two new localities where the Lakora thrust outcrops may also provide better conditions for dating the thrust activity: (i) more to the west of our study area, where thermal conditions are lower (160-180°C) and hence favorable to the crystallization of 1M illite polytype (LAK3 zone, Fig. 3 in Saspiturry et al., 2022), and ii) more to the east, where thermal conditions seem higher and favorable to a more complete reset of the K-Ar chronometer (LAK1 zone, Fig. 3 in Saspiturry et al., 2022).

## Conclusions

The structural, petrographical, and mineralogical results obtained in this study suggest that the Lakora thrust fault zone experienced brittle-ductile deformation with two stages of deformation in the presence of fluids. The newly-formed clay minerals (i.e., Fe-rich chlorite and illite, the latter with abundant $2M_1$ and a small proportion of 1M) in the Lakora thrust are the results of pressure-solution and dissolution-recrystallization mechanisms. Based on the polytypes and crystallinity of these synkinematic minerals, the thermal condition of thrusting corresponds to the
663 upper-anchizone (~250-300°C). These thermal conditions are consistent with previously
664 estimated regional peak paleo-temperatures in the Lakora thrust footwall. The brittle-ductile
665 deformation regime under low-grade conditions is consistent with the structural level inferred in
666 previous works for the studied area during the Eocene-Oligocene Pyrenean collision, i.e., at least
667 6 km of burial with hot fluid flow explaining the excess of temperature.

668 The $^{40}$Ar/$^{39}$Ar step-heating results yield a Paleocene age (c. 61.7 ± 0.5 Ma) for the finest illite
669 grain-size fraction (<0.2 µm) of the more-deformed sample in the thrust fault zone. Although this
670 age is consistent with the overall timing frame of the Lakora thrusting in the west-central
671 Pyrenees, it is not in agreement with the structural constraints in the study area, which suggest
672 that the burial and temperature conditions requested to develop the observed deformations were
673 not realized before at least the Early Eocene. This discrepancy may be due to the preservation of
674 a small amount of partially reset inherited micas and/or the presence of illite particles that are
675 contaminated by excess argon in the dated fraction, thus leading to an overestimated age.

676 This result shows that $^{40}$Ar/$^{39}$Ar dating of deformation in low-grade metamorphic conditions is a
677 real challenge due to the complexity of such systems (i.e., polyphase tectonic activity, fluid
678 circulation, inherited and excess argon, and possible incomplete mineralogical reactions). Based
679 on the present study results, combining the $^{40}$Ar/$^{39}$Ar step-heating method with detailed
680 multiscale structural work and textural-mineralogical-chemical characterizations is highly
681 recommended and the only way to validate or not the $^{40}$Ar/$^{39}$Ar radiometric ages.

682 Acknowledgments

683 This work was supported by the INSU program (TelluS-SYSTER). Access to SEM facilities of
684 the MIMENTO center was made through cooperative projects in the French Renatech network.
685 Special thanks to Alain Meunier and Daniel Beaufort for the discussion about clay minerals.
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1008
1009
1010 **Figure captions**

1011 **Figure 1:** (a) Structural map of the Pyrenees with the location of the study area (blue line: cross-section in Fig.1b; thick black lines: thrust and oblique faults; red lines: the three major thrusts discussed in this paper (i.e., the Lakora, Larra, and Gavarnie thrusts); star symbol: studied outcrop); Map modified from Teixell et al. (2000). (b) Cross-section of the study area illustrating the Lakora thrust structural setting (the trace of the section is shown in (a)). (c) Detail of cross-section in (b) showing the stratigraphic and structural location of the studied outcrop marked by a star symbol. In (b) and (c), the Lakora thrust system is shown with thicker lines; in (b), note that it comprises two branches to the north and south, but a single one in the central part, i.e., below the Lakora klippe where the present study is located. The section in (b) is modified from Labaume et al. (1985) for the southern part and from Teixell and Garcia Sansegundo (1994) for the middle part, the latter being detailed in (c). For the northern part, the section integrates structural interpretations of sections located a few km to the east by Teixell (1993) and Labaume and Teixell (2020).

1024
1025 **Figure 2:** The Lakora thrust outcrop, (a) Photograph showing the hanging wall and footwall of the thrust fault zone. The white dashed line indicates the bedding direction. (b) A simplified stratigraphic column showing the Lakora and Larra thrusts position with the sample locations marked by a red star (simplified after Teixell et al. 2000). The samples are labeled as 15 LAK 1 to 12.

1030
1031 **Figure 3:** Deformation structures observed in the Maastrichtian turbidites of the Lakora thrust footwall damage zone. (a) Top-to-south reverse faults and folds deforming bedding and bedding-
1033 parallel cleavage. (b) Boudinage and bedding-parallel cleavage moulding the boudins; note that cleavage is nevertheless locally at low-angle to bedding. (c) Folds and axial planar cleavage. (d) Reverse faults and folds post-dating boudinage (stars label boudin necks deformed by fold hinges). (e) Thin-section scanner image of a bedding-normal calcite vein and bedding-parallel stylolites in a sandstone layer. (f) microscope image (cross-polarized light) of sandstone with dissolution seams (stylolites) marking the bedding-parallel cleavage; (1): quartz grain with extensional crack sealed by calcite; (2): calcite fringes with fibers parallel to cleavage; (3): calcite fringes with curved fibers indicative of quartz grain rotation (arrows). C: calcite; Qz: quartz; S0: bedding; S0-S1: bedding-parallel cleavage in pelite layers, S2: oblique cleavage in pelite layers; St (S1): dissolution seam (stylolite) marking the bedding-parallel cleavage; V: vein. See Supplementary Figure S1 for the location of pictures (c: shows the sampling site of 15 LAK 3; e: sample 15 LAK 5; f: sample 15 LAK 10).

1045 **Figure 4:** (a, b) Photographs showing the outcrops from where the two samples, 15 LAK 1 and 15 LAK 3, have been collected. White dashed lines indicate the bedding-parallel cleavage (S0-S1), yellow dashed thick line indicates the fold axial plane, and yellow dashed thin lines indicate axial-planar crenulation cleavage (S2). (c, d) Thin-section scanner images of the two samples showing cleavage attitude. (e, f) Optical microscope images (cross-polarized light) of the matrix of the two samples with zooming to show mica grains. Red dashed arrows indicate the zooming locations. Mca = mica, Qz = quartz.

1052 **Figure 5:** Backscattered electron (BSE) images of the two samples (a, c, e) 15 LAK 1 and (b, d, f) 15 LAK 3. In (a), large quartz, mica, and chlorite grains are parallel to bedding, while in (b), chlorite and mica grains are concentrated along the oblique S2 crenulation cleavage and associated with Fe-oxides. (c, e) High magnification BSE images showing large, elongated grains
and the chlorite/mica stacks in the matrix of the 15 LAK 1 sample. Some small grains of mica fill the dissolved surfaces of quartz grains (c). In the 15 LAK 3 sample, the chlorite and mica are located and folded along the two cleavages, S0-S1 and S2, as shown in (d). Zooming on the fine-grained chlorite and mica that follow the S2 cleavage and are associated with Fe-oxides (F). Chl = chlorite, Mca = mica, Qz = quartz, Rt = rutile, Pyr = pyrite, and Alb = albite, Fe-ox = Fe-oxides.

**Figure 6:** Backscattered electron (BSE) images of samples 15 LAK 1 and 15 LAK 3 at high magnification. (a, b, c, d) illustrate two groups/types of chlorite and mica/illite: (i) large chlorite and mica grains (~20-50 µm) parallel to the horizontal cleavage in 15 LAK 1 and following the S0-S1 and S2 cleavage in 15 LAK 3. (ii) Small chlorite and mica/illite are present in the pressure-shadows of quartz grains (a, c) and observed in the matrix displaying random orientations (e, f). Some authigenic small mica crystals are growing in open spaces with an orientation perpendicular to the shearing (i.e., S2) (d). Chl = chlorite, Mca = mica, Qz = quartz, Rt = rutile, Pyr = pyrite.

**Figure 7:** X-ray diffraction patterns of randomly oriented preparations of the bulk powder of the Lakora samples (15 LAK 1, 3, 4, 6, and 9). Chl = chlorite, Mca = mica, Qz = quartz, Rt = rutile, Alb = albite, and Ca = calcite.

**Figure 8:** X-ray diffraction patterns of oriented preparations (air-dried conditions) of the fractions 16-2 µm, 2-0.5 µm, 0.5-0.2 µm, and <0.2 µm. The fractions are from samples 15 LAK 1 and 15 LAK 3. Chl = chlorite, Mca = mica, Qz = quartz, Rt = rutile and Alb = albite.

**Figure 9:** X-ray diffraction patterns of randomly oriented preparations of the bulk and the fractions 2-0.5 µm, 0.5-0.2 µm, and <0.2 µm of the samples 15 LAK 1 and 15 LAK 3. The
1078 mica/illite polytypes (2*M*₁ and 1*M*) peak reflections are highlighted by red and green vertical dashed lines, respectively. Chl = chlorite, Qz = quartz, Rt = rutile, and Alb = albite.

1080 **Figure 10:** Ternary diagrams, (a) (4Si - M⁺ - R²⁺) and (b) (Mg- Al⁽⁶⁾⁻ - Fe) of the chemical analyses of large and small chlorite and mica/illite grains (Meunier and Velde, 1989) where 4Si = (Si/4), M⁺ = Na⁺ + K⁺ + 2Ca²⁺; R²⁺ = (Mg + Fe)²⁺ + Mn; and Al⁽⁶⁾⁻ = octahedral aluminum. Black circles represent standard pure clay phases, and the analyzed samples are highlighted by colored circles and triangles shaped. Montmorillonite Si₄ O10 (Mg₀.₃ Al₁.₇) (OH)₂ Na₀.₃; Muscovite (Si₃ Al) O10 (Al₂) (OH)₂ K; Illite (Si₃.₃ Al₀.₇) O10 (Al₁.₉ Mg₀.₁) (OH)₂ K₀.₈; Biotite (Si₃ Al) O10 (Mg₃ (OH)₂ K; Sudoite (Si₃Al) O₁₀ (Mg₂ Al₃) (OH)₈; Clinochlore (Si₃Al) O₁₀ (Mg₅ Al) (OH)₈; Amesite (Si₂ Al₂) O₁₀ (Mg₂ Al₂ Fe₂) (OH)₈; Chamosite (Si₃ Al) O₁₀ (Fe₅ Al) (OH)₈.

1088 **Figure 11:** ⁴⁰Ar/³⁹Ar step-heating results of the Lakora samples: 15 LAK 1(blue color) and 15 LAK 3 (brown color). The results show the age spectra for the bulk powder, 50-16 µm, 16-2 µm, 2-0.5 µm, 0.5-0.2 µm, and <0.2 µm fractions, respectively. TGA refers to the total gas age.

1091
1092 **Fig. 12:** Summary of the mesostructures resulting from the two stages of deformation observed in the Maastrichtian turbidites of the Lakora thrust footwall damage zone (cf. Figs. 3 to 6).

1094 **Table captions**

1095
1096 **Table 1:** Representative normalized chemical analyses of mica/illite and the corresponding calculated structural formulae.

1098 **Table 2:** Representative normalized chemical analyses of chlorite and the corresponding calculated structural formulae.
1099
1100 **Table 3:** The FWHM (raw data), the calibrated KI (CIS), KI (Basel lab) values in °2θ, and the
corresponding crystallinity zone of representative samples.

1102 **Table 4:** Summary of the <sup>40</sup>Ar/<sup>39</sup>Ar data.

1103
<figure><img src="image_2.png" /></figure>

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

<figure><img src="image_4.png" /></figure>
<figure><img src="image_5.png" /></figure>

<figure><img src="image_6.png" /></figure>
<figure><img src="image_7.png" /></figure>

a

SW-NE

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

b

SW-NE

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

c

WSW-ENE

S0-S1

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

d

SW-NE

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

e

1 cm

St (S1)

S0

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

f

50 µm

O2

O2

O2

C

2

O2

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

3

1
15 LAK 1

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

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

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

15 LAK 3

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

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

<figure><img src="image_19.png" /></figure>
15 LAK 1

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

15 LAK 3

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

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

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

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

<figure><img src="image_25.png" /></figure>
15 LAK 1

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

15 LAK 3

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

c

<figure><img src="image_28.png" /></figure>

d

<figure><img src="image_29.png" /></figure>

e

<figure><img src="image_30.png" /></figure>

f

<figure><img src="image_31.png" /></figure>
<figure><img src="image_32.png" /></figure>

1115

1116
<figure><img src="image_33.png" /></figure>
<figure><img src="image_34.png" /></figure>
<figure><img src="image_35.png" /></figure>

<figure><img src="image_36.png" /></figure>
<figure><img src="image_37.png" /></figure>

Cumulative ³⁹Ar Released [ % ]

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

Cumulative ³⁹Ar Released [ % ]

<figure><img src="image_39.png" /></figure>

Cumulative ³⁹Ar Released [ % ]

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

Cumulative ³⁹Ar Released [ % ]

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

Cumulative ³⁹Ar Released [ % ]

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

Cumulative ³⁹Ar Released [ % ]
SSW

NNE

<figure><img src="image_43.png" /><figcaption>1. Bedding-parallel stretching and bedding-normal shortening with a component of top-to-south shearing:
- Boudinage of sandstone layers associated with bedding-normal calcite veins and bedding-parallel stylolites
- Penetrative cleavage (sub) parallel to bedding (S1-S0) in pelite layers</figcaption></figure>

<figure><img src="image_44.png" /><figcaption>2. Top-to-south shearing: bedding slip, reverse faulting, folding and axial planar crenulation cleavage (S2)</figcaption></figure>

1125

1126
<table><caption>**Table 3** The FWHM (raw data), the calibrated KI (CIS), KI (Basel lab) values in °2θ and the corresponding crystallinity zones of some representative samples</caption><thead><tr><th>Sample</th><th>FWHM °2θ</th><th>KI (CIS) °2θ</th><th>KI (Basel lab) °2θ</th><th>Crystallinity zone</th></tr></thead><tbody><tr><td>15 LAK 1</td><td>0.40</td><td>0.60</td><td>0.49</td><td>Diagenetic zone</td></tr><tr><td>15 LAK 3</td><td>0.27</td><td>0.40</td><td>0.31</td><td>Upper anchizone</td></tr><tr><td>15 LAK 4</td><td>0.36</td><td>0.54</td><td>0.43</td><td>Diagenetic zone</td></tr><tr><td>15 LAK 6</td><td>0.39</td><td>0.58</td><td>0.47</td><td>Diagenetic zone</td></tr><tr><td>15 LAK 9</td><td>0.32</td><td>0.47</td><td>0.38</td><td>Lower anchizone</td></tr></tbody></table>
<table><thead><tr><th>Sample</th><th>Grain size (µm)</th><th>Maximum age (Ma)</th><th>Retention age (Ma)</th><th>Total gas age (Ma)</th><th><sup>39</sup>Ar recoil (%)</th></tr></thead><tbody><tr><td rowspan="6">15 LAK1</td><td>Bulk</td><td>298.6 ± 1.2</td><td>151.9 ± 1.1</td><td>129.9 ± 1.0</td><td>15.6</td></tr><tr><td>50-16</td><td>303.2 ± 1.2</td><td>162.9 ± 1.2</td><td>140.4 ± 1.0</td><td>15.0</td></tr><tr><td>16-2</td><td>277.6 ± 0.7</td><td>162.1 ± 1.2</td><td>138.2 ± 1.0</td><td>15.4</td></tr><tr><td>2-0.5</td><td>133.5 ± 0.6</td><td>110.7 ± 0.9</td><td>91.8 ± 0.7</td><td>17.6</td></tr><tr><td>0.5-0.2</td><td>110.5 ± 0.8</td><td>94.5 ± 0.7</td><td>76.6 ± 0.6</td><td>19.4</td></tr><tr><td>&lt;0.2</td><td>110.4 ± 0.9</td><td>94.9 ± 0.7</td><td>71.3 ± 0.5</td><td>25.4</td></tr><tr><td rowspan="6">15 LAK3</td><td>Bulk</td><td>265.1 ± 2.4</td><td>119.6 ± 0.9</td><td>105.8 ± 0.8</td><td>12.4</td></tr><tr><td>50-16</td><td>233.9 ± 1.3</td><td>133.0 ± 1.0</td><td>119.3 ± 0.9</td><td>10.9</td></tr><tr><td>16-2</td><td>274.9 ± 1.4</td><td>144.9 ± 1.1</td><td>129.9 ± 1.0</td><td>18.1</td></tr><tr><td>2-0.5</td><td>141.4 ± 2.6</td><td>106.0 ± 0.8</td><td>87.3 ± 0.7</td><td>19.0</td></tr><tr><td>0.5-0.2</td><td>105.9 ± 0.4</td><td>86.9 ± 0.6</td><td>68.9 ± 0.5</td><td>21.4</td></tr><tr><td>&lt;0.2</td><td>103.2 ± 1.0</td><td>84.8 ± 0.7</td><td>61.7 ± 0.5</td><td>27.8</td></tr></tbody></table>