Diagenetic and environmental control of the clay
mineralogy, organic matter and stable isotopes (C, O) of
Jurassic (Pliensbachian-lowermost Toarcian) sediments
of the Rodiles section (Asturian Basin, Northern Spain).

Jean-François Deconinck, Juan Jose Gómez, François Baudin, Hugo Biscay, Ludovic Bruneau, Théophile Cocquerez, Olivier Mathieu, Pierre Pellenard, Anne-Lise Santoni

► To cite this version:

Jean-François Deconinck, Juan Jose Gómez, François Baudin, Hugo Biscay, Ludovic Bruneau, et al.. Diagenetic and environmental control of the clay mineralogy, organic matter and stable isotopes (C, O) of Jurassic (Pliensbachian-lowermost Toarcian) sediments of the Rodiles section (Asturian Basin, Northern Spain). Marine and Petroleum Geology, 2020, 115, pp.104286. 10.1016/j.marpetgeo.2020.104286. hal-02511232

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Diagenetic and environmental control of the clay mineralogy, organic matter and stable isotopes (C, O) of Jurassic (Pliensbachian-Lowermost Toarcian) sediments of the Rodiles section (Asturian Basin, Northern Spain)

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JEAN-FRANÇOIS DECONINCK *, JUAN JOSE GÓMEZ **, FRANÇOIS BAUDIN***

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HUGO BISCAY*, LUDOVIC BRUNEAU*, THEOPHILE COCQUEREZ*, OLIVIER

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MATHIEU*, PIERRE PELLENARD* AND ANNE-LISE SANTONI*

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* Biogéosciences, UMR 6282, UBFC/CNRS, Université Bourgogne Franche-Comté, 6 Boulevard Gabriel, F-21000 Dijon, France.

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** Departamento de Estratigrafía, Facultad de Ciencias Geológicas (UCM) and Instituto de Geociencias (CSIC-UCM), 28040 Madrid, Spain.

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*** Institut des Sciences de la Terre de Paris (ISTeP), Sorbonne Université, 4 place Jussieu, 75005 Paris, France.

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**Abstract**

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Clay mineralogical and geochemical analyses (δ¹⁸Ocarb, and paired carbonate and organic matter carbon isotope), completed by the characterization of organic matter, have been performed on the Pliensbachian sediments exposed on the Rodiles section (Asturian Basin, Northern Spain). The objectives were to precise the Pliensbachian climate evolution at a relatively low paleolatitude of the Tethyan domain by comparison with previously published high-resolution clay mineralogical and geochemical data from the NW Tethyan domain (e.g. Paris Basin and Cardigan Bay Basin, West Wales). The clay mineral assemblages are dominantly composed of illite and illite-smectite mixed-layers (I-S) associated with minor proportions of chlorite and kaolinite. Smectites are not identified what may result either to their absence at the time of sedimentation or to their progressive illitization due to the burial diagenesis. Rock-Eval data show that organic matter reaching more than 5% in black-shale
horizons is thermally mature with $T_{max}$ comprised between 444 and 454°C indicating that the section is in the oil window and that the burial temperature reached or exceeded 80 – 90°C.

The low $\delta^{18}O_{bulk}$ values comprised between -3 and -6‰ clearly results from a diagenetic influence (thermal effect due to burial) but their fluctuations compared with $\delta^{18}O$ measured on well-preserved belemnites guards show that the overall trend is however preserved.

By comparison with higher latitude sedimentary basins, the most striking feature is the very small proportions of kaolinite (generally less than 5% while the proportions of this mineral reach 40% at higher paleolatitudes), except at the Pliensbachian/Toarcian transition, where the proportion of kaolinite increases significantly. Clay minerals likely originated from the erosion of the Iberian Massif located to the South of the Asturian Basin in a semi-arid climatic belt. This may explain the scarcity of kaolinite by comparison with sedimentary basins located at higher latitude influenced by more humid climate. The increasing proportions of kaolinite by the end of the Pliensbachian suggest southward expansion of the humid belt as a consequence of cooler temperatures and the establishment of steeper latitudinal gradients of temperature likely resulting from the formation of restricted polar ice cap.

The main $\delta^{13}C$ (on both $^{13}C_{inorg}$ and $^{13}C_{org}$) Pliensbachian excursions including the negative excursion of the Sinemurian/Pliensbachian Boundary Event (SPBE) and the Late Pliensbachian positive excursion are well recognized.

Keywords: Clay minerals, Early Jurassic, Iberian Peninsula, Organic matter, Paleoclimate, Pliensbachian, Stable isotopes.

## Introduction

The Early Jurassic is characterized by significant environmental changes, the most prominent of these being the Toarcian Oceanic Anoxic Event (T-OAE) which has been extensively studied in various sedimentary basins. These studies have highlighted the important role of
global warming, attested by a dramatic increase in the temperature of ocean waters and
enhanced hydrolysis on continents (e.g. Jenkyns 1988, 2010; Hesselbo et al., 2007; Raucsik
and Varga 2008; Dera et al., 2011; Harazim et al., 2013; Hermoso and Pellenard, 2014; Korte
et al., 2015; Them et al. 2017; Müller et al., 2017; Izumi et al. 2018; Fantasia et al., 2018,
Ruebsam et al., 2019). Recent studies have shown that the Pliensbachian is also characterized
by significant fluctuations of sea water temperature (Armendáriz et al., 2012; Korte et al.,
2015; Gómez et al., 2016a; Price et al., 2016) using notably δ¹⁸O measured on several calcite
hard parts of molluscs and brachiopods. Several cold intervals occurred during the
Pliensbachian, the most prominent of these during the latest Pliensbachian, time equivalent to
the spinatum Zone (Bailey et al., 2003; Rosales et al., 2004; Dera et al., 2009; Suan et al.,
2010; Harazim et al., 2013; Gómez et al., 2016a; Bougeault et al., 2017; Arabas et al., 2017;
Rosales et al., 2018). This cold episode could have been responsible for the formation of ice at
high paleolatitudes (Suan et al., 2010), consistent with the supposed record of prominent
obliquity cycles expressed as detrital input at mid-latitudes by the end of the Pliensbachian
(Boulila and Hinnov, 2017). This is also suggested by a significant lowering of the sea-level
expressed for example by the replacement of deep facies by storm-dominated facies
indicating a shallowing trend in the Iberian basin (Val et al., 2017) or by emersion features
(Mailliot et al., 2009, Léonide et al., 2012). By contrast, the stratigraphic interval straddling
the davoei Zone is considered by most authors as a warmer period (Rosales et al., 2004; Dera
et al., 2011; Silva and Duarte, 2015; Price et al., 2016; Gómez et al., 2016a; Bougeault et al.,
2017) favored by higher pCO₂ (Mercuzot et al., 2020) but this interval is also characterized by
an acceleration of the hydrological cycle, that may be responsible for decreasing salinity
leading to low δ¹⁸O values (Deconinck et al., 2019). The Pliensbachian is therefore
characterized by alternating warm and cold episodes and by wet and drier periods respectively
highlighted by kaolinite-rich and smectite-rich intervals (Deconinck et al., 2019). Until now,
most studies dealing with high resolution clay mineralogy of Pliensbachian deposits are from
sections located at mid-paleolatitudes either in the Paris Basin (Bougeault et al., 2017) or in
the Cardigan Bay Basin (Deconinck et al., 2019). Here, we attempt, through a new high
resolution study of clay mineral assemblages of Pliensbachian sediments from the Rodiles
section (Asturias, Northern Spain), to estimate the intensity of physical and chemical
weathering as well as the intensity of hydrolysis and runoff at a slightly lower palaeolatitude
for comparison with the Paris and Cardigan Bay basins. Such climate proxies are coupled to
carbon isotope (organic and inorganic) and organic matter evolutions throughout the
Pliensbachian to evaluate both burial diagenesis and paleoenvironmental conditions.

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## I - Geological setting

The Rodiles section is cropping out along the coastal cliffs located northeast of Villaviciosa in
the eastern part of the Asturias regional autonomy (Northern Spain) (Fig. 1). The
Pliensbachian deposits were studied in the eastern part of the Rodiles Cape and the uppermost
Pliensbachian/lowermost Toarcian in the western part of the Rodiles Cape (West Rodiles
section of Gómez et al., 2008; Gómez and Goy, 2011). The well-exposed Pliensbachian
deposits (Fig. 1) consist of a succession of alternating limestones, more or less bioclastic, with
marls and black-shales, belonging to the Santa Mera Member of the Rodiles Formation
(Valenzuela, 1988; Bádenas et al., 2009). Three TOC (Total Organic Carbon) enriched
intervals corresponding to black-shales occur notably at the base of the *jamesoni* Zone, within
the *ibex* Zone and at the base of the *gibbosus* Subzone of the *margaritatus* Zone (Gómez et
al., 2016b) (Fig. 2). The section studied contains numerous fossils including bivalves,
gastropods, brachiopods, crinoids and ammonites. Ammonite fauna allows a precise
biochronostratigraphic scheme to be obtained (Comas-Rengifo, 1985; Phelps, 1985; Howarth,
2002) following the standard chronozones and subchronozones established by Dommergues

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et al. (1997) (Fig. 2). Pliensbachian sediments were deposited in an open marine external platform in lower offshore environments with sporadic intervals of oxygen deficiency.

Paleogeographical reconstruction based on comprehensive paleomagnetic data, performed by Osete et al. (2011), locates the Rodiles section at a paleolatitude of approximately 32° N for the Hettangian–Sinemurian interval, which is in good agreement with the calculations of Van Hinsbergen et al. (2015) and at a paleolatitude of almost 40° N for the Toarcian–Aalenian interval. Consequently, during the Pliensbachian, according to these authors, a paleolatitude of 34-35°N can be estimated. However, most palaeogeographic reconstructions locate the Asturian Basin at a somewhat lower latitude around 28°N (Fig. 3) and a latitudinal extension of the Iberian Massif from 28 to 22°N (e.g. Thierry et al., 2000). Whatever the paleolatitude, the Asturian Basin was located at lower latitudes than the Paris and the Cardigan Bay Basins and the continental areas likely providing detrital particles were not located in the same climatic belt (Fig. 3). In the Paris Basin, detrital particles probably originated from the London-Brabant Massif and from the Armorican/Central Massifs, and in the Cardigan Bay Basin, the Welsh High was likely the main source of terrigenous materials. These massifs were located in a humid belt (Fig. 3). By contrast, the Iberian Massif located south of the Asturian Basin in a semi-arid belt was probably the main source of clay minerals for the Asturian Basin (Fig. 3). Therefore, as climate conditions were different on landmasses supplying clay minerals, significant differences in the composition of the clay assemblages in the Asturian Basin are to be expected compared to the Paris and Cardigan Bay Basins.

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## II - Methods

### II - 1. Clay mineralogy

The entire Pliensbachian succession was sampled at an average step of 0.5 m (Fig. 2). A total of 172 samples were then analysed using X-Ray Diffraction (XRD). After a manual crush in a
mortar, powdered samples were decarbonated with a 0.2 N HCl solution. The fraction below 127 two microns (clay sized particles), was extracted with a syringe after decantation of the 128 suspension during 95 minutes following Stokes law, and then centrifuged. Clay residue was 129 then smeared on oriented glass slides and run in a Bruker D4 Endeavor diffractometer with 130 CuKα radiations, LynxEye detector and Ni filter under 40 kV voltage and 25 mA intensity 131 (Biogéosciences Laboratory, Université Bourgogne Franche-Comté). Scanning from the 132 goniometer ranges from 2.5° to 28° for each analysis. Three runs were performed for each 133 sample to discriminate clay phases: 1) Air-drying; 2) Ethylene-glycol solvation; 3) Heating at 134 490°C, as recommended by Moore and Reynolds (1997). Clay minerals were identified using 135 their main diffraction (d<sub>001</sub>) peak and by comparing the three diffractograms obtained. 136 Estimated proportions of each clay mineral were performed using the MacDiff 4.2.5 software 137 (Petschick, 2000) on glycolated diffractograms. As the main (d<sub>001</sub>) peak of kaolinite and the 138 (d<sub>002</sub>) peak of chlorite are overlapping, a deconvolution procedure was applied on the (d<sub>004</sub>) 139 peak area of chlorite (3.54 Å) and (d<sub>002</sub>) peak area of kaolinite (3.57 Å) to estimate properly 140 both mineral's proportions using the 7.1 peak (d<sub>001kaolinite</sub> + d<sub>002chlorite</sub>). Chlorite percentage was 141 calculated using the mean between (d<sub>001</sub>) and (d<sub>002</sub>) chlorite peak areas considering that the 142 chemical nature in the chlorite impacts the (d<sub>001</sub>)/(d<sub>002</sub>) ratio (Moore and Reynolds, 1997). The 143 K/I ratio based on the (d<sub>001</sub>) peak area of each mineral was systematically calculated. 144 Replicate analyses of samples and the internal procedure of quantification using MacDiff on 145 mixtures of clay minerals in known proportions allow to estimate the uncertainty of the 146 relative proportions of the clay minerals as 5%. 147

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150 II – 2. Organic matter characterization
A Rock-Eval 6 Turbo apparatus (Vinci Technologies) was used for quantitative and qualitative study of the organic matter from 20 selected samples. The method, described in details by Lafargue et al. (1998) and Behar et al. (2001), consists in a two steps analysis with programmed temperature: a pyrolysis, under inert atmosphere (N₂), followed by an oxidation. Twenty selected samples distributed over the entire section were analyzed at the Institut des Sciences de la Terre de Paris (UMR7193/Sorbonne Université). Crushed samples are first subjected to a 3 min isotherm at 300 °C at which free hydrocarbons are volatilized (peak S1). Then, a heating with a ramp of 25 °C/min leads to the vaporization of products from thermal cracking of organic matter up to 650 °C (peak S2). Pyrolysis effluents are continuously detected by a flame ionization detector (FID) and expressed in mg per g of sample. Organic CO and CO₂ are measured online as well by an infrared cell (peak S3). At the end of the pyrolysis step, samples are automatically transferred into an oxidation oven where they are subjected to a 1 min isotherm at 300 °C then a ramp of 20 °C/min up to 850 °C. Total signals of both organic and mineral CO and CO₂ are expressed in mg per g of samples. All those parameters allow the calculation of the total organic carbon (TOC). The hydrogen index (HI), corresponding to the quantity of hydrocarbon compounds released during the pyrolysis relative to the TOC (S2/TOC) in mg of HC per g of TOC as well as the oxygen index (OI) corresponding to the organic CO₂ released during the pyrolysis relative to the TOC (S3/TOC) in mg of CO₂ per g of TOC are also calculated. The HI and the OI are correlated to the H/C and O/C atomic ratios respectively, which allow determining the origin of the organic matter. Tmax is defined as the pyrolysis temperature at which the maximum amount of hydrocarbon is yielded by kerogen (Espitalié et al., 1977). Tmax increases linearly with the maturation degree of the organic matter, thus giving a rapid estimate of the thermal maturity of sedimentary basins (Espitalié, 1986).
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## II – 3. Oxygen and carbon stable isotopes and Total Organic Carbon

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The isotopic analyses of 90 bulk carbonate samples (approximately one out of two samples with an average step of 1 m) were performed at the Biogéosciences Laboratory, Université Bourgogne Franche-Comté in Dijon on a carbonate preparation device (Kiel IV) coupled with a Delta V Plus mass spectrometer (Thermofisher). Powdered samples were reacted with 20µl of orthophosphoric acid at 70 °C. The reproducibility (2σ) of the IAEA NBS19, used as an external standard, is better than 0.04‰ for the δ¹³C and 0.08‰ for the δ¹⁸O. The delta values are expressed relatively to the V-PDB (Vienna Pee Dee Belemnite).

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Organic carbon isotope compositions (δ¹³Corg) and Total Organic Carbon (TOC) contents were measured on carbonate-free residues of the same 90 samples. Sample powders reacted with HCl (2 N) at room temperature for 24 h to remove carbonate phases. Residues were rinsed with deionized distilled water until neutral, centrifuged (4500 rpm for 15 min), and dried at 50 °C overnight. Aliquots of dried decarbonated samples (~ 7–50 mg) were then weighed in tin capsules. TOC content and δ¹³Corg measurements were performed at the Biogéosciences Laboratory of the Université Bourgogne Franche-Comté (Dijon, France) on a Vario MICRO cube elemental analyzer (Elementar, Hanau, Germany) coupled in continuous flow mode to an IsoPrime stable isotope ratio mass spectrometer (Isoprime, Manchester, UK). USGS40 L-Glutamic acid (C = 40.8 wt%; δ¹³CVPDB = -26.39 ± 0.04‰) and IAEA-600 Caffeine (δ¹³CVPDB = -27.77 ± 0.04‰) certified reference materials were used for calibration. The carbon isotopic composition is expressed in delta notation and reported in per mil (‰) relative to the Vienna Pee Dee Belemnite (V-PDB) standard; external reproducibility based on duplicate analyses of samples is better than ± 0.2‰ (1σ).

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## III - Results

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### III – 1. Clay mineralogy
201 Clay mineral assemblages are dominantly composed of illite (45 to 70%), random illite-
202 smectite mixed layers (20 to 40%), minor proportions of chlorite (5 to 12%) and kaolinite
203 (traces to 12%) (see supplementary data 1). The crystallinity of illite (FWHM = Full Width at
204 Half Maximum of the 10 Å peak) ranges from 0.32 to 0.61 °2θ without any relationships with
205 depth (Fig. 4).

206 Illite-smectite mixed layers (I-S) are rich in illite sheets and correspond either to R1, R3 or a
207 mixture of both types. The mineralogical composition of the clay assemblages is relatively
208 constant along the section except in its upper part where the proportions of kaolinite increase
209 significantly around the Pliensbachian/Toarcian transition (Fig. 4). The proportions of chlorite
210 during the lower Pliensbachian (jamesoni and ibex Zones) are always lower than 10%, while
211 the proportions of this mineral reach occasionally 10 % in the davoei and margaritatus Zones
212 and a maximum around 12% in the spinatum Zone. The proportions of illite and I-S did not
213 show any clear trend along the succession. Although the proportions of kaolinite are around
214 5% throughout the section, they increase slightly in an interval from the upper part of the ibex
215 Zone to the lower part of the stokesi Subzone. The proportions of this mineral increase more
216 significantly in the upper part of the spinatum Zone and in the lowermost Toarcian
217 (tenuicostatum Zone; Fig. 4 and 5, Supplementary data 1).

218 III – 2. Rock-Eval data.

219 The total organic carbon (TOC) content measured on 20 selected samples using Rock-Eval
220 analysis ranges from 0.4 to 4.9%, half of these samples displaying TOC content higher than
221 2.5% (Table 1). These values are similar to those measured along the whole Rodiles section
222 by Gomez et al. (2016b). Remaining S2 values range from 0.3 up to 11.5 mg HC/g rock and
223 the Tmax values are well defined at the top of the S2 peak.

224 Indeed, Tmax values ranging from 444 to 454 °C (Table 1) with a mean Tmax around 450 °C,
225 indicate that the Rodiles section reaches the middle of the oil window during its burial (Fig.
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6). In spite of this level of thermal maturity, the remaining S2 values indicate a residual petroleum potential of ~4.5 kgHC/t of rock for the Rodiles section. The HI-Tmax diagram (Fig.

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6) also shows that the organic matter of the Pliensbachian-Toarcian in the Rodiles section is dominated by a marine Type II, usually deriving from a phytoplanktonic and bacterial source.

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### III – 3. Geochemical data

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CaCO₃ contents fluctuate between 20% and 80% (Fig. 7). The lowest values, between 20 and 40%, characterize the middle part of the section corresponding to the *ibex* and *davoei* Zones

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(Fig. 7; supplementary data 2). Below and above this stratigraphic interval, in the *jamesoni* Zone, and in the *margaritatus* and *spinatum* Zones, the CaCO₃ contents oscillate between 40 and 80% or even slightly more. The evolution of TOC shows values fluctuating between 0.2

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and 5% with three stratigraphic TOC enriched intervals : within the *jamesoni* Zone, the second particularly well-marked in the *valdani* Subzone of the *ibex* Zone and the last quite well marked in the *margaritatus* Zone (subnodosus – gibbosus Subzones; Fig. 7). These TOC enriched intervals correspond with the black-shales (Fig. 2) and are consistent with data previously published by Gómez et al. (2016b).

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The evolution of the δ¹⁸Ocarb is presented in Figure 7. The values fluctuate between -6 and -2.5 ‰ with an overall positive tendency especially from the *margaritatus* Zone upward. From the base of the section to the top of the *jamesoni* Zone, the values vary between -5 and -3 ‰ with a very slight tendency to increase, whereas a negative trend is observed in the *ibex* and *davoei* Zones. Then increasing values are recorded until the topmost part of the Pliensbachian.

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δ¹³Ccarb values show fluctuations from -2 to +1.5 ‰ (Fig. 7), which is in the same range than the values previously measured on belemnite guards from the same section (Armendáriz et al., 2012; Gómez et al., 2016a). Negative values are recorded in the lower Pliensbachian

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(jamesoni to davoei Zones) while positive values are observed from the *margaritatus* Zone
upward, notably in the *subnodosus/gibbosus* Subzones (Fig. 7).

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$\delta^{13}C_{\text{org}}$ values range from -29.3 to -26 ‰. These values are in good agreement with a dominantly marine origin (algal and bacterial) of the organic matter as indicated by the Rock-Eval analyses. At the base of the section, the lowest values are recorded in the *jamesoni* (*taylori-polymorphus* Subzones). In the upper part of this ammonite zone, the values increase more or less regularly until the base of the *margaritatus* Zone (*stokesi* Subzone) and then decreases sharply in the *subnodosus* Zone before a regular positive trend up-section (Fig. 7; Supplementary data 2).

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## IV - Discussion

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### IV-1. Influence of burial diagenesis on clay mineral assemblages

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Before examining the significance of clay mineral assemblages in terms of palaeoclimate and palaeoenvironment, it is essential to ensure that they are dominantly detrital in origin without any significant influence of burial diagenesis, thermal influences or authigenic growth during early or late diagenesis. In the case of the Rodiles section, the clay mineral assemblages, invariably dominated by illite, chlorite and illite-rich I-S, lead to question their detrital or diagenetic character. This type of clay mineral assemblage is often found in sedimentary series that have undergone significant thermal influences (Chamley 1989; Mangenot et al., 2018) responsible for the progressive illitization of smectite following the reaction: smectite $\Rightarrow$ I-S mixed-layers $\Rightarrow$ illite and frequently indicative of a significant thermal diagenesis linked to burial (Altaner and Ylagan, 1997; Środoń., 1999; Lanson et al., 2009; Środoń et al., 2009; Dellisanti et al., 2010). The association with chlorite is common as this mineral is stable with depth of burial. The absence of I-S R0 and smectite may be also indicative of a thermal influence related to burial, but their absence can also result from the lack of these
minerals at the time of sedimentation. This is the case for example in coeval sediments from
the Sancerre borehole, drilled in the southern Paris Basin, where the clay assemblages are
dominated by chlorite, illite and kaolinite without any smectite nor I-S R0 although the
thermal influences are negligible (Bougeault et al., 2017). In addition, except in sediments
deposited in the Cardigan Bay Basin close to the Welsh High (Deconinck et al., 2019), the
presence of smectites is relatively uncommon in European Pliensbachian sediments (Dera et
al., 2009). Consequently, the absence of smectite (or smectite-rich phases) cannot be taken as
an unequivocal evidence for a significant thermal influence. Smectite may also transform into
chlorite through chloritization process with the formation of chlorite-smectite mixed-layers
including tosudite or corrensite, but these minerals were not identified in the studied
sediments.

Rock-Eval analyses show that the organic matter from the Rodiles section is mature as the
Tmax values are comprised between 444 and 454°C (Table 1) indicating that the section
reached the oil window (Fig. 6) and that the burial temperatures have reach or exceed values
of 80 to 90 °C. As smectites turn into illite as soon as the burial temperatures reach 60 to 70
°C, it is therefore possible that smectites, if present during sedimentation, were transformed
into I-S R1, R3, or even illite during the subsequent geological history of the basin.

Kaolinite occurs in small quantities (<10%, except around the Pliensbachian/Toarcian
transition), what is surprising, kaolinite being abundant (between 20 and 40%) in
Pliensbachian sediments of several areas including the Paris and the Cardigan Bay Basins
(Dera et al., 2009; Bougeault et al., 2017; Deconinck et al., 2019). The small proportions of
kaolinite lead to wonder about the fact that kaolinite could be also partly transformed
especially in chlorite during the burial of the series or into dickite (Cuadros et al., 2014) but
this latter mineral was not identified on the Rodiles section. Kaolinite can also be converted
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into illite during burial, but this transformation occurs mainly in sandstones, whereas it is
essentially smectites that turn into illite in clays and marls (Lanson et al., 2002).

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The proportions of kaolinite increase at the Pliensbachian/Toarcian transition indicating that, despite the burial temperature exceeding 80-90°C of the series, the stability conditions of kaolinite are preserved which suggests that the maximum burial temperature did not exceed 150°C (Beaufort et al., 2015). Beyond this temperature, kaolinite is transformed into chlorite by dissolution/crystallization mechanisms. Since chlorite is scarce in the studied series, kaolinite is unlikely to have been transformed. The fluctuations of its proportions thus seem mainly of detrital origin.

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## IV-2. Diagenetic influence on oxygen and carbon isotopes

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### *Oxygen isotopes*

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On the Rodiles section, oxygen isotopes were measured on well-preserved belemnite guards by Armendáriz et al. (2012) and Gómez et al. (2016 a and b). Their results showed a range of variations of δ¹⁸Oₐₗ between -3.1 and 0 ‰ which suggest that the diagenetic imprint is negligible on the belemnite guards. Gómez et al. (2016 a and b) suggest using this biogenic material that a significant warming occurred during *ibex* and *davoei* Zones, while the *margaritatus* and *spinatum* Zones are regarded as cooler intervals. However Armendáriz et al. (2012) suggested that water mass exchange with variable salinity may be also responsible for changing δ¹⁸O values. The values obtained on the bulk sediments range between -2.5 to -6‰, much less than those from the belemnite guards (Fig. 7). These low values are generally interpreted as resulting from a diagenetic alteration (Silva et al., 2011; Silva et al., 2017; Fantasia et al., 2018). A similar difference between δ¹⁸Oₐₗ and δ¹⁸Oₕₐₗ was observed in sections from southern France and also attributed to a diagenetic alteration of the δ¹⁸Oₕₐₗ of the bulk sediment (Harazim et al., 2013). Although such a difference has also been observed

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in sedimentary series free of significant diagenetic influences (Pellenard et al., 2014), in the
Rodiles section, the low δ¹⁸Ocarb values likely reflect the thermal influence due to the burial of
the succession evidenced by Rock-Eval data. However, the lowest values are recorded in the
ibex and davoei Zones and the highest during margaritatus and spinatum Zones what is
consistent with the known relative fluctuations of temperature occurring during the
Pliensbachian (Fig. 7). By comparison to the δ¹⁸Obel, the δ¹⁸Ocarb seem generally shifted
towards more negative values, but the original relative fluctuations likely of climatic or water
mass exchange in origin seem to be partly preserved.

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Carbon isotopes

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The δ¹³Ccarb shows values in the same range (-2 to 1.5 ‰, Fig. 7) of variation than δ¹³Cbel published by Gómez et al. (2016b) and slightly lower than those published by Armendáriz et al. (2012). As δ¹³Cbel was measured on well preserved belemnite guards this suggest that the δ¹³Ccarb values are not significantly altered by diagenetic influences. In addition, the absence of a significant correlation between δ¹³C and δ¹⁸O (Fig. 8) supports the interpretation that the carbon isotopic record is not disturbed by diagenesis.

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336

337

338

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By comparison with data from other sections or boreholes from several basins, the main
carbon isotope excursions are recognized. At the base of the section, in the
taylori/polymorphus Subzone of the jamesoni Zone, the low values of δ¹³Ccarb and δ¹³Corg
likely correspond to the Sinemurian/Pliensbachian Boundary Event (SPBE) while the positive
excursion recorded in the subnodosus/gibbosus Subzones corresponds to the Late
Pliensbachian Event (LPE) as evidenced in the northern and southern margins of the Tethyan
domain (Korte and Hesselbo, 2011; Ruhl et al., 2016; Mercuzot et al., 2020). Superposed on
these main trends, minor fluctuations may result from local influences including the
preservation of organic-rich facies (black-shales) for example in the ibex Zone that may be
responsible for δ¹³C increase yet this positive excursion is also recorded in several basins

340

341

342

343

344

345

346

347

348
349

(Mercuzot et al., 2020). The identification of main previously recognized isotope events suggests that despite thermal influences due to burial of the Rodiles section, the δ¹³Ccarb signal is well-preserved.

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351

352

The comparison of the δ¹³Ccarb and δ¹³Corg curves shows that a divergence occurs in the *margaritatus* Zone. In particular, while the values of δ¹³Ccarb increase in the *subnodosus* and *gibbosus* Subzones, the values of δ¹³Corg decrease sharply. In this interval, the relatively high HI values of organic matter suggest an increase in the proportion of marine organic matter which could be responsible for the low values of δ¹³Corg (Suan et al., 2015).

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354

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356

357

## IV-3. Origin of clay minerals

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361

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Although clay minerals may be transported over long distances, according to the palaeogeographical context (Fig. 3), the Iberian Massif, part of the Variscan belt of Western Europe was probably the main detrital continental sources of clay minerals deposited in the Asturian Basin. A detrital contribution from the Armorican Massif cannot be excluded but this massif was separated from the Asturian Basin by the Bay of Biscay Trough (Thierry et al., 2000). The basement of the Iberian Massif is dominated by Proterozoic and Palaeozoic crystalline rocks including granitoids and high-grade metamorphic rocks (Arenas et al., 2016). Chlorite and illite, considered as primary minerals likely originated from the erosion of these rocks. Illite-smectite mixed-layers originated either from the moderate weathering of illite or from the illitization of smectite if we consider that this mineral was present at the time of sedimentation.

369

370

371

372

Kaolinite may derived from kaolinite bearing sedimentary rocks outcropping on the Iberian Massif or from soils developed over this massif. The very small quantities of kaolinite indicate that dry climatic conditions should prevail over the Iberian Massif, and confirm that this massif was located in a semi-arid climatic belts (Fig. 3). However, the significant increase
373

of kaolinite by the end of the *spinatum* Zone and around the Pliensbachian/Toarcian boundary
suggests that a climate change has occurred on the Iberian Massif at this time.

374

375

The latest Pliensbachian (*spinatum* Zone) is known to record a cold event and the likely
development of ice at polar latitudes (Suan et al., 2010). This cold event is followed by a
significant warming around the Pliensbachian-Toarcian transition (Suan et al., 2008; 2010).

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378

The increase in kaolinite observed at the top of the *spinatum* Zone and at the transition with the Toarcian cannot be linked solely to warming. The formation of kaolinite requests first a significant humidity, while increasing temperature accelerates hydrolyzing processes.

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380

381

Consequently there must necessarily be an increase in humidity on the Iberian massif by the end of the Pliensbachian. It is not excluded that the warming was responsible for an acceleration of the hydrological cycle. However in the Sancerre borehole in the Paris Basin kaolinite decreases at the Pliensbachian/Toarcian transition. If we assume that a "coldhouse" climate mode occurred during the time-equivalent *spinatum* Zone, with the formation of a small ice cap before the Pl-To warmer event, this coldhouse mode implies a re-organization of the climatic belts.

382

383

384

385

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388

This climate change was probably responsible for a shift of the humid belt to lower latitude, then favoring the formation of kaolinite in soils developed over the Iberian Massif. The formation of kaolinite in soils require at least 1 My before its transportation and deposition in the sedimentary basin (Thiry, 2000). The duration of the *spinatum* zone is estimated according to cyclostratigraphic studies to 1.4 My (Ruhl et al., 2016). Therefore, we suggest that the onset of kaolinite formation on the Iberian massif started during the *spinatum* Zone as soon as humid conditions appeared on the Iberian Massif without significant decrease of temperature.

389

390

391

392

393

394

395

Then kaolinite was transported and deposited by the very end of the Pliensbachian, and as a
consequence of increasing latitudinal gradients of temperature at the Pliensbachian/Toarcian
transition (Fig. 9). By contrast, the decreasing proportions of kaolinite at higher latitudes

396

397
during the *spinatum* Zone (Bougeault et al., 2017; Deconinck et al., 2019) results from colder and dryer conditions linked to a slowing down of the hydrological cycle. Consequently the cooling event occurring during the latest Pliensbachian triggered a reverse response in the proportions of kaolinite that decrease at mid-latitudes but increase at lower latitudes. These results illustrate the important role of latitudinal fluctuations of climatic belts during the Pliensbachian recently highlighted on the Southern margin of the Tethys (Baghli et al., 2020).

Interestingly, the proportions of kaolinite increase slightly in the upper part of the *ibex* Zone, and in the *davoei* Zone (Fig. 4). This trend was expected since this stratigraphic interval is characterized by increasing proportions of kaolinite in several sedimentary basins (Dera et al., 2009). It is interpreted as a result of more humid climate conditions and/or warmer temperature over the NW Tethyan area. Our results, although the increase of kaolinite is minor, show that even in the semi-arid belt, and despite the thermal diagenetic influence, this significant climate change corresponding to an acceleration of the hydrological cycle is recorded by clay minerals. This stratigraphic interval is also depleted in CaCO₃, what was already observed in the Mochras borehole (Cardigan Bay Basin, Ruhl et al., 2016). This may result from a decreasing biological productivity (demise of carbonate platforms?) or from a dilution of carbonates by higher terrigenous fluxes and organic matter content as a consequence of increasing runoff. This second hypothesis is here favored, as increasing terrigenous fluxes are generally associated with increasing nutrients enhancing productivity.

## Conclusions

Clay mineral analyses of the Rodiles section coupled with the geochemical data makes it possible to specify the intensity of the thermal diagenesis and the climatic evolution during the Pliensbachian at relatively low palaeolatitude.
Rock-Eval pyrolysis data reveal that organic matter is mature and the sedimentary succession is in the oil window. The influence of a thermal diagenesis is also marked by low values of $\delta^{18}$O$_{carb}$, but the original fluctuations of this proxy seem preserved while being shifted to low values. Therefore, the absence of smectites can be interpreted in two ways. Either the smectites deposited in the Asturias basin, but the subsequent burial of the sedimentary series led to their illitization, or during the Pliensbachian, the clay sedimentation in the Asturian basin did not include this mineral.

The low kaolinite contents reflect the erosion of a continental area, presumably the Iberian Massif, subjected to semi-arid climatic conditions throughout the Pliensbachian, except at the end of this stage where conditions become wetter on the Iberian massif in response to a significant cooling and an increase in the latitudinal temperature gradient. The onset of kaolinite formation in soils developed over the Iberian massif took place during the *spinatum* Zone before this mineral was reworked at the transition between the Pliensbachian and the Toarcian. Although very attenuated, the increase in kaolinite proportions, known in many sedimentary basins during the *ibex/davoei* Zones, is also recorded in the Asturias basin. This stratigraphic interval is known to be hot and humid, but the acceleration of the hydrological cycle may have resulted in a decrease of the $\delta^{18}$O of seawater leading to a decrease in $\delta^{18}$O in many areas. Therefore, the apparently hot interval of the *ibex/davoei* interval deserves reconsideration.

Acknowledgements: This is a contribution to the JET (Jurassic Earth Time) project. We thank Stéphane Bodin and an anonymous reviewer for their helpful and constructive remarks on the first version of the manuscript.

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657 Suan, G., Mattioli, E., Pittet, B., Mailliot S., Lécuyer C. 2008. Evidence for major
658 environmental perturbation prior to and during the Toarcian (Early Jurassic) oceanic
659 anoxic event from the Lusitanian Basin, Portugal. Paleoceanography, 23, 1.

660 Suan, G., Mattioli, E., Pittet, B., Lécuyer, C., Suchéras-Marx, B., Duarte, L.V., Philippe, M.,
661 Reggiani, L. Martineau, F. 2010. Secular environmental precursors to Early Toarcian
662 (Jurassic) extreme climate changes. Earth and Planetary Science Letters 290, 448-458.

663 Suan, G., Van de Schootbrugge, B., Adatte, T., Fiebig, J., Oschmann, W. 2015. Calibrating
664 the magnitude of the Toarcian carbon cycle perturbation. Paleoceanography, 30, 5.
665 Them, T.R. II, Gill B.C., Caruthers A.H., Gröcke D.R., Tulsky E.T., Martindale, R.C.,
666 Poulton, T.P., Smith, P.L. 2017. High-resolution carbon isotope records of the
667 Toarcian Oceanic Anoxic Event (Early Jurassic) from North America and implications
668 for the global drivers of the Toarcian carbon cycle. Earth and Planetary Science
669 Letters 459, 118–126.

670 Thierry, J. et al. (40 co-authors) 2000. Late Sinemurian. In: Dercourt, J., Gaetani, M.,
671 Vrielynck, B., Barrier, E., Biju-Duval, B., Brunet, M.-F., Cadet, J.P., Crasquin, S.,
672 Sandulescu, M. (Eds.), Atlas Peri-Tethys Paleogeographical Maps, vol. I–
673 XX.CCGM/CGMW, Paris, pp. 49-59.

674 Thiry, M. 2000. Palaeoclimatic interpretation of clay minerals in marine deposits: an outlook
675 from the continental origin. Earth Science Review, 49, 201-221.

676 Val, J., Bádenas, B., Aurell, M. Rosales, I. 2017. Cyclostratigraphy and chemostratigraphy of
677 a bioclastic storm-dominated carbonate ramp (late Pliensbachian, Iberian Basin).
678 Sedimentary Geology 355, 9–113.

679 Valenzuela, M. 1988. Estratigrafía, sedimentología y paleogeografía del Jurásico de Asturias,
680 PhD Thesis, Fac. Sci. Geol. Univ. Oviedo, Spain.

681 Van Hinsbergen, D. J. J., de Groot, L. V., van Schaik S. J., Spakman W., Bijl, P. K., Sluijs A.,
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683 Studies. PLoS ONE 10, e0126946. doi:10.1371/journal.pone.0126946.

685 Figure captions

686 Figure 1. Location maps of the Rodiles section. (a): sketched geological map of Iberia
687 showing the position of the Asturian Basin. (b): outcrops of the Jurassic deposits in the
688 Asturian and the western part of the Basque–Cantabrian basins, and the position of the
689 Rodiles section. (c): geological map of the Asturian Basin showing the distribution of
690 the different geological units and the location of the Rodiles section. (d). Field view of
691 the East Rodiles section (SPB = Sinemurian-Pliensbachian Boundary).

692 Figure 2. Simplified lithostratigraphic column (modified from Fraguas et al., 2018) and
693 location of the samples studied (ER = East Rodiles, RW = Rodiles West). ca. =
694 capricornus, dav. = davoei, fi. = figulinum, ha. = hawskerense, lur. = luridum, ma. =
695 masseanum, ml. = maculatum, rar. = raricostatum, u. Sin. = upper Sinemurian.

696 Figure 3. Palaeogeography and climatic zonation of North West Europe during the
697 Pliensbachian. AM. Armorican Massif, BM. = Bohemian Massif, CM. = Cornwall
698 Massif, IbM. = Iberian massif, IrM. = Irish Massif, LBM. = London Brabant Massif,
699 MC. = Central Massif, RM. = Rhenish Massif, SM. = Scottish Massif, WH. = Welsh
700 High.

701 Figure 4. Clay mineralogy and illite crystallinity index of the Rodiles section. I-S R1 = R1
702 type illite-smectite mixed-layers. T = Toarcian, *apyr. = apyrinum; brev. = brevispina,*
703 *ca. = capricornus, fi. = figulinum, gibb. = gibbosus, haws. = hawskerense, jam. =*
704 *jamesoni. lur. = luridum, ma. = masseanum, ml. = maculatum, S. = Sinemurian,*
705 *subnod = subnodosus*

706 Figure 5. X-ray diagram of the clay fraction (glycol-solvated) of the samples from the
707 uppermost part of the *spinatum* Zone (samples ER165 to ER 167) and the lowermost
708 Toarcian (samples ER168 and RW 4), showing the more and more prominent 002
709 peaks of kaolinite around 3.58Å. chl. chlorite, kaol. kaolinite, Q. quartz.

710 Figure 6. HI vs. Tmax diagram for selected organic-rich samples from the Rodiles section,
711 indicating that organic matter reaches the middle of the oil window.

712 Figure 7. CaCO₃ content, Total Organic carbon (TOC), δ¹³Ccarb., δ¹³Corg. δ¹⁸Ocarb., and δ¹⁸Obel. of the Pliensbachian deposits of the Rodiles section. *apyr. = apyrinum; brev. = brevispina, ca. = capricornus, fi. = figulinum, gibb. = gibbosus, haws. = hawskerense,*
715 jam. = jamesoni. lur. = luridum, ma. = masseanum, ml. = maculatum, subn. =

716 subnodosus, vald= valdani, SPBE = Sinemurian-Pliensbachian Boundary Event

717 Figure 8. Cross plot of data of the sample studied showing the absence of correlation between

718 δ¹³Ccarb and δ¹⁸Ocarb.

719 Figure 9. Sketch explaining the climate change deduced from the clay mineralogy of the

720 Rodiles section occurring from the latest Pliensbachian to the lowermost Toarcian.

721

722 Table 1. Rock-Eval data of the studied samples.

723

724 Supplementary data 1: Clay mineralogy of the studied samples

725 Supplementary data 2: Geochemistry – CaCO3, TOC, and stable isotopes.
<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>

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

Lower Pliensbachian

u. Sin.

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

Legend

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

Emerged landmasses

boreholes

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

Rodiles section
<table><thead><tr><th>Depth (m)</th><th>Stage</th><th>Biostratigraphy (ammonites)</th><th>% Chlorite</th><th>% Illite</th><th>% I-S R1</th><th>% Kaolinite</th><th>Illite °2Θ crystallinity</th></tr><tr><th>T.</th><th></th><th>Zone</th><th>Subzone</th><th></th><th></th><th></th><th></th></tr></thead><tbody><tr><td></td><td></td><td>tenuicostatum</td><td>5</td><td>10</td><td></td><td></td><td>0.35 0.45 0.55</td></tr><tr><td></td><td></td><td>spinatum</td><td>haws.<br>apyr.</td><td></td><td></td><td></td><td></td></tr><tr><td></td><td>margaritatus</td><td></td><td>gibb.<br>subnod.<br>stokesi</td><td></td><td></td><td></td><td></td></tr><tr><td></td><td>davoei</td><td>ca./fi.<br>ml.<br>lur.</td><td></td><td></td><td></td><td></td><td></td></tr><tr><td></td><td>ibex</td><td>valdani<br>ma.</td><td></td><td></td><td></td><td></td><td></td></tr><tr><td></td><td>jamesoni</td><td>taylori-<br>polymorphus</td><td>jam.<br>brev.</td><td></td><td></td><td></td><td></td></tr><tr><td></td><td>S.</td><td>raricostatum</td><td></td><td></td><td></td><td></td><td></td></tr></tbody></table>
<figure><img src="image_10.png" /></figure>
<figure><img src="image_11.png" /></figure>

Immature

Oil window

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

Ammonite
zones and
subzones

spinatum

haws.

apyr.

margaritatus

gibb.

subn.

stok.

dav.

ca.fi.

ml.

lur.

vald.

ma.

jam.

brev.

jamesoni

taylori-
polymorphus
<figure><img src="image_13.png" /></figure>
<table><thead><tr><th>Sample</th><th>S1<br>mgHC/g</th><th>S2<br>mgHC/g</th><th>Tmax<br>°C</th><th>S3<br>mgCO2/g</th><th>TOC<br>%</th><th>HI<br>mgHC/gTOC</th><th>OI<br>mgCO2/gTOC</th></tr></thead><tbody><tr><td>ER 2</td><td>0.30</td><td>0.31</td><td>444</td><td>0.52</td><td>0.43</td><td>72</td><td>121</td></tr><tr><td>ER 8</td><td>0.51</td><td>1.40</td><td>448</td><td>0.14</td><td>0.68</td><td>206</td><td>21</td></tr><tr><td>ER 9</td><td>0.96</td><td>4.34</td><td>449</td><td>0.43</td><td>2.67</td><td>163</td><td>16</td></tr><tr><td>ER 11</td><td>1.22</td><td>6.90</td><td>450</td><td>0.39</td><td>3.75</td><td>184</td><td>10</td></tr><tr><td>ER 13</td><td>1.19</td><td>6.12</td><td>448</td><td>0.35</td><td>3.19</td><td>192</td><td>11</td></tr><tr><td>ER 19</td><td>1.26</td><td>6.23</td><td>453</td><td>0.56</td><td>3.47</td><td>180</td><td>16</td></tr><tr><td>ER 76</td><td>1.41</td><td>11.48</td><td>452</td><td>0.08</td><td>4.93</td><td>233</td><td>2</td></tr><tr><td>ER 78</td><td>1.45</td><td>7.48</td><td>451</td><td>0.16</td><td>3.81</td><td>197</td><td>4</td></tr><tr><td>ER 81</td><td>0.76</td><td>3.36</td><td>450</td><td>0.09</td><td>1.72</td><td>195</td><td>5</td></tr><tr><td>ER 84</td><td>1.27</td><td>7.00</td><td>450</td><td>0.14</td><td>3.42</td><td>205</td><td>4</td></tr><tr><td>ER 95</td><td>0.60</td><td>2.47</td><td>454</td><td>0.42</td><td>1.85</td><td>134</td><td>23</td></tr><tr><td>ER 99</td><td>0.60</td><td>2.53</td><td>454</td><td>0.11</td><td>1.78</td><td>142</td><td>6</td></tr><tr><td>ER 112</td><td>0.39</td><td>1.01</td><td>450</td><td>0.22</td><td>0.92</td><td>110</td><td>24</td></tr><tr><td>ER 132</td><td>0.61</td><td>1.49</td><td>450</td><td>0.11</td><td>1.21</td><td>123</td><td>9</td></tr><tr><td>ER 135</td><td>1.16</td><td>9.53</td><td>452</td><td>0.57</td><td>4.40</td><td>217</td><td>13</td></tr><tr><td>ER 137</td><td>1.26</td><td>6.64</td><td>450</td><td>0.19</td><td>2.98</td><td>223</td><td>6</td></tr><tr><td>ER 140</td><td>0.67</td><td>3.29</td><td>448</td><td>0.51</td><td>1.70</td><td>194</td><td>30</td></tr><tr><td>ER 143</td><td>0.61</td><td>2.57</td><td>449</td><td>0.32</td><td>1.53</td><td>168</td><td>21</td></tr><tr><td>ER 145</td><td>0.26</td><td>1.07</td><td>451</td><td>0.30</td><td>0.95</td><td>113</td><td>32</td></tr><tr><td>ER 155</td><td>0.21</td><td>1.12</td><td>451</td><td>0.78</td><td>1.22</td><td>92</td><td>64</td></tr></tbody></table>