# New evidences of subaerial volcanism as a trigger for the Kilian event (Aptian-Albian transition) and major climatic changes from offshore Morocco (DSDP Site 545)

Elodie Bracquart, Guillaume Charbonnier, Sylvain Garel, Thomas Munier, Thierry Adatte, Julien Danzelle

► To cite this version:

Elodie Bracquart, Guillaume Charbonnier, Sylvain Garel, Thomas Munier, Thierry Adatte, et al.. New evidences of subaerial volcanism as a trigger for the Kilian event (Aptian-Albian transition) and major climatic changes from offshore Morocco (DSDP Site 545). Global and Planetary Change, 2022, 218, pp.103959. 10.1016/j.gloplacha.2022.103959. hal-03829141

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New evidences of subaerial volcanism as a trigger for the Late Aptian Kilian event and
major climatic changes from offshore Morocco (DSDP Site 545)

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Elodie Bracquart<sup>1,2*</sup>, Guillaume Charbonnier<sup>1,3</sup>, Sylvain Garel<sup>1,4</sup>, Thomas Munier<sup>1</sup>, Thierry Adatte<sup>3</sup>, Julien Danzelle<sup>1,5</sup>.

1. UMR CNRS 7193 Institut des Sciences de la Terre de Paris (ISTeP), Sorbonne Université, Paris, France.

2. Institut des Sciences de la Mer de Rimouski (ISMER), Université du Québec à Rimouski, Québec, Canada.

3. Institute of Earth Sciences, Géopolis, University of Lausanne, CH-1015 Lausanne, Switzerland.

4. Energy & Geoscience Institute, The University of Utah, Salt Lake City, UT 84108, USA

5. Department of Earth Sciences, University of Cambridge, Downing Street, Cambridge CB2 3EQ, UK

* Corresponding author.

E-mail address: elodie.bracquart@uqar.ca

Abstract

The Aptian-Albian transition is characterized by the presence of four organic-rich
intervals all linked with major disturbances of the carbon cycle. Among these, the second
subevent of the oceanic anoxic event (OAE) 1b, also called Kilian event, has a latest Aptian
age and is characterized by a negative carbon isotope excursion. On the contrary of the
Paquier OAE1b subevent, the Kilian event was not extensively studied and its triggering
mechanisms and its consequences on regional climates remain uncertain. Here we present a
geochemical (Rock-Eval, Mercury) and mineralogical (clay minerals) study on late Aptian
sediment cores from the Deep-Sea Drilling Program (DSDP) Site 545 (offshore Morocco,
Eastern North Atlantic) where a Kilian level equivalent is present. Mercury (Hg) content
results show two main peaks in the late Aptian, immediately below and above the Kilian level
 equivalent. The absence of correlations between Hg and total organic carbon and Hg and
 CaCO₃ indicates that the mercury content is mainly related to subaerial volcanic activity,
 probably issued from the emplacement of the Kerguelen large igneous province in the Indian
 Ocean. Mineralogical analyses show that the clay content, in our studied interval, mainly
 reflect the weathering and climatic conditions in Northwest Africa. The evolution in clay
 mineral assemblages depicts three distinctive climatic intervals: (i) a seasonally contrasted
 semi-arid climate below the first Hg anomaly; (ii) a moister and warmer climate during the
 Kilian level equivalent deposition; and (iii) a rapid transition toward more arid climate post
 Kilian level equivalent deposition that is linked with the second Hg anomaly (large increases
 in illite and palygorskite). Therefore, our study confirms that the Kerguelen volcanic activity
 was probably the trigger of the Kilian event and shows that the resulting changes in the
carbon cycle caused two major climate changes before and after the Kilian event and a long
term aridification of Northwest Africa.

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Keywords: Kilian oceanic anoxic event (OAE1b), clay minerals, mercury, volcanism, climate
change

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1. Introduction

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The Cretaceous is marked by the recurrence of high amplitude carbon-isotope
excursions and multiple deposition of laminated organic-rich layers in marine sediments,
interpreted as the consequence of considerable palaeoclimatic and palaeoenvironmental
changes (e.g., Jenkyns, 2010; Föllmi, 2012). The Aptian/Albian boundary interval is
characterized by the succession of four distinct laminated organic-rich sediments initially
described in the Vocontian Basin ("Jacob", "Kilian", "Paquier", and "Leenhardt", Bréhéret,
1995), and their equivalents in the western Tethys, Atlantic, and Pacific Ocean basins (e.g.,
Herrle et al., 2003a, b, 2015; Li et al., 2016; Benamara et al., 2020). These levels are
synchronous with important carbon-isotope shifts, suggesting major disturbances in the global
carbon cycle (Herrle et al., 2004). Palaeotemperature and surface water productivity proxies
suggest that these specific levels are associated with intense warming and an increase in
primary productivity (hyperthermal events) (Herrle et al., 2003a, b; Bottini and Erba, 2018).
This time interval coincides with one of the largest and more significant species turnovers of
planktonic foraminifera (Premoli Silva and Sliter, 1999; Leckie et al., 2002; Huber and
Leckie, 2011). These levels and their environmental and climatic perturbations correspond to
the four events of the Aptian/Albian Oceanic Anoxic Event 1b (OAE1b, Trabucho-Alexandre
et al., 2011; Coccioni et al., 2014).

The second event of the OAE1b (i.e. Kilian event, latest Aptian) is defined by a
negative shift in marine carbonate and organic carbon isotope records (e.g., Herrle et al.,
2004). It has been identified in the Northwestern Tethys (Herrle et al., 2004; Sabatino et al.,
2015; Benamara et al., 2020), in the Eastern Tethys (Li et al., 2016), in the Atlantic Ocean
(Herrle et al., 2004; Trabucho-Alexandre et al., 2011), in the Eastern Pacific (Navarro-
Ramirez et al., 2015), in the shallow-marine platform in Mexico (Madhavaraju et al., 2017),
and in the Canadian Arctic Sverdrup Basin (Herrle et al., 2015). Several causes have been
proposed for this event, such as the volcanic activity of the southern Kerguelen Plateau
(Coffin et al., 2002; Sabatino et al., 2018), enhanced emission from marine methane clathrates
(Wagner et al., 2007; Zhang et al., 2016), and cyclic evolution of a dynamic monsoonal
system related to orbital forcing (Herrle et al., 2003a, b). However, the climatic consequences
of the Kilian event are still debated. Particularly, a few studies investigate the change in the
continental hydrological cycle and their links with nutrient and detrital input into the ocean
during this key period of environmental changes.

The objectives of this study are to (i) document changes in continental weathering
conditions during the latest Aptian, and (ii) determine if enhanced volcanic activity triggered
the Kilian event. For this, we performed high-resolution mineralogical (clay minerals) and
geochemical (Rock Eval organic carbon/pyrolysis measurements, mercury) studies from the
latest Aptian record of the DSDP Site 545 from the Mazagan Plateau (Eastern North
Atlantic).

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## 2. Geological setting

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The Mazagan Plateau is located off the coast of Morocco (Fig. 1), about 200 km
southwest of Casablanca (33°39.86′N; 09°21.88′W) (Hinz and Winterer, 1984). It ends in a
crescent-shaped escarpment: the Mazagan Escarpment and is bounded to the west by the
Seine Abyssal Plain (Ruellan et al., 1985). This Plateau constitutes the outer edge of the
continental passive margin of Northwest Morocco formed by extension during the Triassic
due to the Central Atlantic opening. Site 545 was drilled in 1981 during Leg 79 at the foot of
the steep Mazagan Escarpment. During the Late Aptian, Site 545 was located at a
paleolatitude of about 20°N (Hay et al., 1999) and a paleo-water depth of about 2000 m
(Hofmann et al., 2008).
Aptian sediments are mainly green to grey nannofossil clays ~110 m thick (Hinz and Winterer, 1984). The Kilian level interval includes lightly laminated black shales ~80 cm thick, ranging from 389.85 to 390.65 mbsf (Leckie et al., 2002). The biostratigraphy of this interval has been well defined by the calcareous nannofossils (i.e. nannofossil biozones NC8A and NC8B) and foraminifera assemblages (i.e. *Paraticinella eubojaouensis* Zone) that give a Late Aptian age for the Cores 41 to 45 (Herrle et al., 2002; Leckie et al., 2002; McAnena et al., 2013). This core interval benefits from high-resolution carbon isotope stratigraphy, which shows a 1.5‰ negative carbon excursion in the Core 42 that corresponds to the Kilian event (Herrle et al., 2004; McAnena et al., 2013).

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## 3. Material and methods

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### 3.1. Material

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For this study, a total of 216 samples were collected from DSDP Site 545. Samples were all cleaned, and their outer rims (0.5 mm) were removed to minimize pollution.

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### 3.2. Methods

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#### 3.2.1. Rock-Eval parameters and carbonate content

About 60 mg (±1 mg) of 65 dried and powdered samples were analysed with a Rock-Eval 6 Turbo (Vinci-Technologies) at the ISTeP laboratory (Sorbonne University). The method used was the one described in Behar et al. (2001) to obtain Rock-Eval parameters (total organic carbon: TOC, hydrogen index: HI, oxygen index: OI, and temperature of the pyrolysis oven recorded at the maximum of hydrocarbons production: Tmax). The amount of
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$CaCO_3$ was also obtained with the same device by multiplying the value of mineral inorganic carbon (MinC) with the ratio between the molar mass of $CaCO_3$ (100.1 g/mol) and the molar mass of C (12 g/mol), i.e. 8.34. Analytical uncertainties are estimated to be < 0.1%, based on 8 replicates and intern standard analysis (IFP160000 shale).

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### 3.2.2. Mercury

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Mercury concentrations of 216 samples were determined using a Zeeman R-915F (Lumex, St. Petersburg, Russia) high-frequency atomic absorption spectrometer at the University of Lausanne. Analyses are based on the direct thermal evaporation of Hg from solid samples. Measurements were systematically conducted on two aliquots. The accuracy was confirmed by the analysis of certified reference materials (GSD-11 standard, Chinese alluvium: 72.0 ppb; Zintwana et al., 2012) with a correlation coefficient of 0.99 and a standard residual deviation of 5% (±3.6 ppb for the standard).

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### 3.2.3. Clay mineralogy

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The clay mineral assemblages of 41 samples were obtained by the X-ray diffraction (XRD) method. The samples were carried out at the ISTEP laboratory (Sorbonne University). After grinding with a mortar and pestle, about 3g of each sample was treated with 0.2N hydrochloric acid (HCl) to remove carbonates. Successive washing in osmosed water allowed clay defloculation. The <2 µm clay fraction was extracted with a syringe after decantation for 95 min, according to the Stoke's Law (Holtzapffel, 1985). After centrifugation, the clay fraction was deposited on an oriented glass slide and was spread homogeneously using a thin slide in the same orientation so that the clay minerals are uniformly oriented to be analysed by X-ray. Three X-ray diagrams were measured using a Bruker D2 Phaser diffractometer (30 kV, 10 mA) with CuKα radiation, a LynxEye detector and a Ni filter: one after air drying, one
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after ethylene glycol solvation for 12h minimum and one after heating at 490°C for 2h
(Holtzapffel, 1985; Moore and Reynolds, 1997) (Fig. 2).

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Clay minerals were identified based on the diffractograms using the main X-ray
diffraction peaks (001) of each mineral: R0 type illite-smectite (I-S) mixed-layers: 17 Å on
glycolated runs, illite: 10 Å, 5 Å and 3.33 Å peaks, palygorskite: 10.4 Å and 6.44 Å; and
kaolinite: 7.18 Å and 3.58 Å (disappears when heated) (Fig. 2). The diffractograms are
analysed using the MacDiff® software (Petschik, 2000). Due to a possible deviation in the
position of the peaks, the diffractograms obtained are all corrected using the quartz peak
located at 4.26 Å. Clay minerals were quantified using the area of the main peaks of each
mineral on glycolated diffractograms. As the main peak of palygorskite and illite overlap, a
deconvolution procedure was applied. This method has an error margin of approximatively
±5% for the relative proportions of clay minerals in the clay fraction.

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## 4. Results

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### 4.1. Total organic carbon contents and carbonate contents

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The TOC contents range between 0.8 and 4.0 wt.%. Among these values, three
significant peaks stand out compared to others. The first peak (3 wt.%) is located between
403.31 and 401.08 meters below seafloor (mbsf). The second, between 390.73 to 390.26 mbsf
corresponds to one of the samples of the Kilian level equivalent (KLE) and exhibits a
maximum TOC of 4.0 wt.%. Finally, the third one (3.2 wt.%) is observed between 381.47 and
379.61 mbsf (Fig. 3A). Between these peaks, the TOC contents are relatively stable, around
1.5 wt.%.

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The carbonate contents ($CaCO_3$) generally vary inversely with TOC, between 20 and 50 wt.% (Fig. 2). Samples from the KLE show relatively low $CaCO_3$ values with a minimum of 20 wt.% at 390.51 mbsf (Fig. 3A).

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## 4.2. Mercury contents

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Mercury contents range from 25 to 67 ppb throughout the section. Four intervals of enrichment in Hg concentrations relative to the background values (~35 ppb) are observed (Fig. 3A): (i) a first one between 407.6 and 406.8 mbsf (from 30 to 58 ppb), (ii) a second between 401.5 and 396.0 mbsf (maximal value of 67 ppb); (iii) a third one in the KLE at the onset of the negative $\delta^{13}$C excursion, between 394.0 to 390.5 mbsf (maximal value of 57 ppb); and (iv) a last one at the top of the series between 383.0 to 382.0 mbsf, (maximum value of 63 ppb).

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## 4.3. Clay mineralogy

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Clay mineral assemblages (fraction <2 µm) are composed of R0 I-S mixed layers (45 to 70%, average to 57.5%), illite (14 to 33%, average to 24.5%), palygorskite (7 to 26%, average 15.1%), and kaolinite (1 to 5%, average to 3.0%) (Fig. 3B). The distribution of these minerals allows the core to be divided into three intervals (i1 to i3). At the base, the first interval (424.19 to 406.19 mbsf) is defined by a high amount of R0 I-S mixed layers (~60%), a low amount of palygorskite (~15%) and illite (~22%), and a very low amount of kaolinite (~2%). The second interval (406.19 to 390.51 mbsf) is characterized by an increase in R0 I-S mixed layers (up to ~70%), a significant decrease in palygorskite (from 15 to 7%) and a slight increase in illite and kaolinite. The transition between interval i2 and i3 corresponds to the

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Kilian event. At the top of the series, the third interval (390.51 to 379.15 mbsf) is marked by a decrease in R0 I-S mixed layers (from 70 to 40%) and a significant increase in palygorskite (up to 25%), and illite (from 23 to 35%) (Fig. 3B).

# 5. Discussion

## 5.1. Hg sequestration processes in marine sediments

Mercury anomalies in sediments have been highlighted in several studies as a proxy for studying past volcanic events (e.g., Sanei et al., 2012; Sial et al., 2013; Grasby et al., 2015, 2019; Percival et al., 2015, 2016; Font et al., 2016; Charbonnier and Föllmi, 2017; Charbonnier et al., 2017). Volcanic activity represents the main source of mercury to the atmosphere. During an eruption, mercury is emitted in its elementary gaseous form (Hg⁰) and transported to distal sites due to its relative long residence time (0.5-1 year; Selin et al., 2008). After oxidation (under its Hg²⁺ form), it becomes soluble and can be deposited in the oceanic domain through wet or dry deposition where it will be finally absorbed into sediments.

However, the enrichment of Hg in marine sediments is more complex because it is also controlled by other mechanisms. Mercury has a strong affinity for different specific host phase such as OM, clay minerals, or sulfides (Outridge et al., 2007; Sanei et al., 2012; Sial et al., 2013; Percival et al., 2015; Shen et al., 2019). In addition, degradation of the permafrost, forest fires, coal combustion, or increased soil erosion may represent others significant sources of Hg scavenged in sediments (e.g., Berquist, 2018; Grasby et al., 2019). Finally, the interpretation of the original Hg signal is also complicated by moderate to intense weathering processes and postdepositional degradation of OM (Charbonnier et al., 2020a; Park et al.,
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2022). A careful evaluation of the origin of Hg sequestration processes in marine sediment is crucial when exploring deep-time Hg signatures as a proxy of distal volcanism.

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### 5.1.1. Impact of postdepositional organic degradation

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Intense weathering and oxidation of the originally organic-rich laminated sediments may have strong consequences on the original signature of Hg/TOC ratio classically used as a proxy for volcanic activity. These processes affect the structure of the organic-rich sediments and induced a severe loss of the original TOC content (nearly 100% of TOC), which results in anomalously high Hg/TOC ratios (>1000 ppb/wt.%; Charbonnier et al., 2020a).

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At Site 545, the $T_{max}$ values range from 411 to 426 °C, which indicate that the OM is immature, i.e. in a zone where the diagenesis temperatures range from 10 to 80 °C (Fig. 4). This implies a low to moderate degree of thermal diagenesis, suggesting both an original composition and structure of the organic-rich layers. This interpretation is well supported by the significant presence of irregular R0 I-S minerals, which are particularly sensitive to high temperature changes during burial diagenesis (Dellisanti et al., 2010; see also Section 5.2.1.).

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### 5.1.2. Impact of organic matter and detrital input

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In marine environments, Hg has a strong affinity for OM, which is indicated by the numerous observations of strong covariations between Hg concentrations and TOC contents (e.g., Sanei et al., 2012; Grasby et al., 2013; Percival et al., 2015). At Site 545, the correlation coefficient between Hg and TOC ($R^2$=0.049) is however very low (Fig. 5A), suggesting that Hg anomalies are not controlled by OM deposition alone. Besides, we also observe a good temporal correspondence between the overall Hg content and the Hg/TOC anomalies (Fig. 5A), confirming this inference. For example, explicit Hg/TOC positive anomalies, reaching 55.8 and 32.4 (ppb/wt.%) at 396.0 and 381.5 mbsf respectively, correspond to samples with

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distinct maxima in Hg contents (Fig. 3A). The only exception to this good correlation between the Hg and Hg/TOC signals is located within the KLE, where higher Hg contents (up to 57 ppb) correspond to lower Hg/TOC values (Fig. 3A). In this particular case, the large increase in OM content in the sediments due to the high productivity conditions (Wagner et al., 2007) and the increase in preservation conditions (Hofmann et al., 2008) strongly affect the Hg/TOC ratio.

Moreover, one must check whether changes in detrital input and runoff influence overall Hg contents (Sial et al., 2013). The Hg contents is thus here compared with the carbonate phase ($CaCO_3$ contents) (Fig. 5B), because the nannofossil claystone nature of the sediments indicates an inverse relationship between carbonate and the total clay content, which also derives from variations in the local detrital inputs. The variations in Hg contents do not correspond to those in $CaCO_3$ ($R^2=0.042$), suggesting that detrital Hg inputs are not significant (Fig. 5B).

Consequently, the Hg anomalies are not related to changes in detrital input, primary productivity, redox conditions, and OM preservation alone.

### 5.1.3. Volcanic pulses: regional versus global influence

A series of Hg/TOC anomalies occurs below and above the KLE in the DSDP Site 545 (Fig. 3A). In analogy to previous studies in which anomalous Hg/TOC peaks were interpreted as indicative of an increase in volcanic activity we suggest that these enrichments are related to rapid increase in atmospheric Hg concentrations associated with a massive release of volcanic Hg⁰ into the atmosphere. A similar positive Hg/TOC anomaly is recorded just below the KLE in the Umbria Marche Basin (Poggio le Guaine section; Sabatino et al., 2018). Recently, positive excursions of Hg concentrations were also observed at the Aptian-Albian transition in the Araripe Basin (South Atlantic) (Benigno et al., 2021) and in the lacustrine
sediments from the Jiuquan Basin (NW China) (Zhao et al., 2022). However, uncertainties in the age control hampers to establish direct correlation with Hg anomalies recorded from the Northwestern Tethyan realm. Interestingly, pronounced negative spikes in Os isotopic records are observed below the KLE in the Tethyan (Umbrian Marche Basin) and Pacific (DSDP Site 463) pelagic site, suggesting multiple volcanic events (Matsumoto et al., 2020). These data appear consistent with the new Hg results at Site 545 and reinforce the hypothesis that the latest Aptian was characterized by intense subaerial volcanic pulses.

Interestingly, at the opposite to the distal sites recording Hg enrichments at the onset of the Kilian event, the base of the Selli oceanic anoxic event 1 a (OAE1a, early Aptian) is characterized by limited increases in Hg/TOC enrichments in the Tethyan Realm (Charbonnier and Föllmi, 2017) and the absence of Hg spikes in the Boreal Realm (Percival et al., 2021). It has often been proposed that the volcanic outgassing related to the submarine Greater Ontong Java LIP triggered the early Aptian environmental perturbations (Méhay et al., 2009; Charbonnier and Föllmi, 2017). By contrast to the submarine LIPs eruptions, which limited the Hg cycle perturbation to the proximal sites due to the low residence time of Hg (~20-30 years) regarding the ocean mixing time (Scaife et al., 2017; Charbonnier et al., 2018a; Percival et al., 2018), the subaerial eruptions contribute significantly to increase aerial flux of Hg into the stratosphere allowing its global distribution.

The Kerguelen Plateau and Broken Ridge form one of the most voluminous large igneous provinces (LIPs) of the Cretaceous, associated with the continental breakup of the Indian-Australian and the Antarctic plates (Coffin and Edholm, 1994; Coffin et al., 2002). It was divided into six domains: the Skiff Bank, the Ellan Bank, the Broken Ridge, the Northern Kerguelen Plateau, the central Kerguelen Plateau, and the Southern Kerguelen Plateau (Coffin et al., 2002). Geochronological data reported for the Southern Kerguelen Plateau indicate a major peak in magmatic output between 119 and 110 Ma (Whitechurch et al., 1992; Duncan,
298 2002; Coffin et al., 2002), which corresponds to the late Aptian-early Albian interval
299 according to GTS2020 (Aptian-Albian transition: 113.2 Ma; Gradstein et al., 2020). The
300 Southern Kerguelen Plateau was erupted in a subaerial environment (Coffin et al., 2002;
301 Mohr et al., 2002). The evidence of explosive subaerial phases has critical implications in
302 terms of global climate and environment (Frey et al., 2000, 2002). By contrast to the OAE1a,
303 the subaerial activity of the Southern Kerguelen Plateau could represent a potential candidate
304 to explain source of Hg input into the atmosphere up to the distal site such as the DSDP Site
305 545 (this study), the Umbria Marche section (Sabatino et al., 2018), the Araripe Basin
306 (Benigno et al., 2021), or the Jiuquan Basin (Zhao et al., 2022).

307 Alternatively, the late Aptian is also a crucial interval marked by significant
308 paleogeographic changes due to geodynamic evolution of the Atlantic Ocean and the
309 reorganization of the peri-Tethyan plates (e.g., Stampfli and Hochard, 2009). This active
310 tectonic reorganization is associated to intense magmatic activities: (i) the Peri-Atlantic
311 alkaline province (central High Atlas, Morocco) linked to the opening of the Central Atlantic
312 (Guezal et al., 2011, 2015), (ii) the Vardar intra-oceanic back arc system (Balkans) linked to
313 the closure of the Paleo-Tethys (Graziano et al., 2016), and (iii) the Waddenzee volcanic
314 centre (northwest Netherlands) (Zimmerle, 1989). Regional explosive volcanic activity is also
315 documented by the presence of volcanogenic materials in late Aptian sediments such as
316 bentonites layers from the Vocontian Basin ("Niveau Van Gogh, Cézanne, and Matisse")
317 (Beaudoin et al., 2012), pyroclastic material from the Apennine carbonate platform (Graziano
318 et al., 2016), or tuff horizon (Wöhrum succession) from the Lower Saxony Basin (Mutterlose
319 et al., 2003). However, the source, the dispersion mode, or the chronostratigraphic correlation
320 between the volcanic ashes remain largely discussed and are not yet well established (e.g.,
321 Graziano et al., 2016). Particularly, the timing of these pulses is poorly constrained due to the
322 lack of radiometric dates and the uncertainty of the geological time scale. Regional explosive
volcanism may have a significant Hg flux that could produce regional Hg anomalies, which
can also explain the positive Hg/TOC anomalies recorded in the latest Aptian sediment from
Site 545. Consequently, superimposed to the Southern Kerguelen activity a regional source of
Hg⁰ to explain the Hg cycle perturbation recorded in the Mazagan Plateau cannot be totally
excluded.

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329 5.2. Origin of clay minerals

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The use of clay minerals as an indicator of climate conditions requires confirmation of
their detrital origin. The mineralogical composition of detrital clays in marine sediments
result from several processes such as (i) hydrolysing condition depending on climate, (ii)
lithology of source rocks, (iii) differential settling on the ocean and (iv) sea-level fluctuation
(Gibbs, 1977; Chamley, 1989; Thiry, 2000; Adatte et al., 2002). However, diagenetic and
hydrothermal processes can also affect clay minerals assemblages. Indeed, burial diagenesis
can transform detrital clays while hydrothermal activity, porosity and low sedimentation rates
can form authigenic clays (Šucha et al., 1993; Charbonnier et al., 2020b).

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340 5.2.1 Diagenetic influence on clay minerals

The temperature increase during burial can transform any clay minerals, affecting the
detrital clay assemblages. The illitisation of smectite is a common transformation during
burial diagenesis. Transformation of smectites into R1-type illite/smectite mixed-layers (50-
90% of illite layers) occurs at 60-70°C (Šucha et al., 1993; Chamley, 1994; Merriman and
Frey, 1999; Lanson et al., 2009). Dellisanti et al. (2010) show that transformation of smectites
into R1-type illite/smectite mixed-layers coincides with the maturity of the OM. The
significant occurrence of R0-type illite/smectite mixed-layers in the Aptian deposits of Site
348 545 suggests a weak impact of burial diagenesis on clay assemblages. This hypothesis is supported by the depth of Aptian deposits that does not exceed 440 mbsf, which prevents, in a normal geothermal gradient, an impact of the burial diagenesis. Moreover, the presence of smectites is associated with low T<sub>max</sub> values (411°C to 426°C), indicating immature OM (Espitalié et al., 1985). Yet, if the smectite had suffered from thermal diagenesis, the associated OM would be at least early mature. Therefore, it is likely that no strong burial diagenesis affected clay minerals (Dellisanti et al., 2010).

355 Early diagenetic and hydrothermal processes can also disturb detrital clay assemblages. The formation of authigenic kaolinite can occur in sandstone and carbonates (Worden and Morad, 1999). Thin deposits of claystone in Site 545 exclude authigenesis of kaolinite. The high sedimentation rate and the absence of correlation between lithology and clay mineralogy reject an authigenic origin of smectites (Chamley and Debrabant, 1982).
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359 Marine origin of palygorskite could result in the transformation of Fe-smectite close to volcanic rocks due to hydrothermal activity in a deep-marine environment (Chamley, 1989).
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362 Palygorskite can also be formed as an alternative of authigenic smectite in a favourable chemical environment depleted in cations. This origin for palygorskite in Central Atlantic cretaceous sediments is supported by Brosse (1982), but Chamley and Debrabant (1982) do not show chemical changes that would allow the formation of authigenic palygorskite in the Aptian sediments of Site 545. SEM observations performed on samples from Site 545 do not reveal palygorskites resulting from the transformation of smectites (Fig. 2B). These observations allow us to suggest, as Chamley and Debrabant (1982), for that the clay minerals in the late Aptian sediments of Site 545 have a dominant detrital origin.
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371 5.2.2. Differential settling processes
372

The distribution of clay minerals in deep marine deposits may be controlled by the differential settling processes, which is related to physical segregation of minerals during transport of particles (e.g., Thiry, 2000; Adatte et al., 2002). Kaolinite and illite are dense and coarse minerals that settle preferentially closer to the coastline in shallow-water setting. By contrast, smectite is a fine-grained mineral that settles much more slowly in deep marine deposits. The absence of high R0 I-S contents (>80 %) and the presence of illite in moderate proportions (up to 33%) support that the differential settling processes did not play a major role in clay mineral distribution. Furthermore, the low proportion of kaolinite (<5 %) may be well explained by climatic conditions (paragraph 5.2.3.) and not necessarily only by differential settling processes as recently observed in the DSDP Proto North Atlantic sites during the Valanginian (Charbonnier et al., 2020c).

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### 5.2.3. Detrital sources of clay minerals

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Clay minerals could be transported over long distances although Northwest Africa seems to have been the main detrital source of the Mazagan Plateau. The clay mineral assemblages of Aptian deposits of DSDP Site 545 are similar to nearby DSDP Sites 369, 370, 398, 416, 547 and to the Chichaoua section (Chamley, 1979; Brosse, 1982; Chamley and Debrabant, 1982; Daoudi et al., 1989; Pletsch et al., 1996). The abundances of chlorite and illite usually reflect the erosion of continental masses. Therefore, the absence of chlorite and moderate proportions of illite (<35%) at Site 545 indicate a weak erosion of the basement. However, the high proportion of smectite (40 to 70%) reflects moderate hydrolysis conditions in a warm and seasonally humid climate, also confirmed by the low proportion of kaolinite (<5 %). The dominance of smectite could also indicate a poor drainage condition on flattened continental masses. This geomorphological conditions in the landmasses could be favourable to the formation of palygorskite in peri-marine environment under warm and evaporative
conditions (Pletsch et al., 1996). Then, palygorskite can be transported into the oceans by rivers and wind. Changes in clay assemblages of Aptian deposits in Site 545 likely mostly reflect continental climate of adjacent landmasses influenced by the geomorphological configuration of North Africa (Mélières, 1978; Chamley, 1979; Chamley and Debrabant, 1984).

402

403 5.2.4. Late Aptian climate changes

The three intervals (i1 to i3) depicted thanks to the clay-mineral assemblages results can be interpreted as distinct climatic periods. Due to a very low concentration of kaolinite throughout the series (~ 4%), it is not possible to interpret these fluctuations in terms of climatic variations, except that there was no major hydrolysing climate or significant enhanced runoff during the Late Aptian (Chamley, 1989) (Fig. 6).

The first climatic period (i1; Fig. 6) is characterized by strong proportions of R0 I-S mixed layers associated with low concentrations of palygorskite and illite. In marine sediments, R0 I-S mixed layers can derive from several sources: (i) reworked from soils formed under warm and seasonally contrasted semi-arid climate and (ii) submarine weathering of volcanic glass (Chamley, 1989; Deconinck and Chamley, 1995). While illite is commonly interpreted as a physical weathering product of crystalline rocks under rather arid conditions and then reflects the intensity of continental erosion and run-off, palygorskite is often used as a proxy for semi-arid or arid conditions (Chamley, 1989). Here, Site 545 being far from known late Aptian volcanic provinces, high concentrations of R0 I-S mixed layers were interpreted as terrigenous and not volcanogenic. Therefore, the mineralogical assemblage of interval i1 can reflect a seasonally contrasted semi-arid climate (Fig. 6). At the end of i1, a peak of Hg/TOC precedes a warming highlighted by a decrease in δ¹⁸Ocarb and an increase in TEX₈₆ and SST (Herrle et al., 2004; Wagner et al., 2007; McAnena et al., 2013)
422

(Fig. 6). This suggests that a first volcanic pulse initiated a change in climatic conditions that could explain the transition from i1 to i2.

423

424

During the second climatic period (i2; Fig. 6), the massive release of C-depleted carbon into the ocean-atmosphere reservoir due to the first volcanic pulse result in a negative shift of δ¹³Cₚcarb (Fig. 6). This shift is correlated to a drop in δ¹⁸O, which indicate the establishment of a warmer climate and, therefore, an enhanced continental run-off. At DSDP Site 545, the resulting intensification of the hydrological cycle is well observed in the mineralogical assemblage. Indeed, strong concentrations in R0 I-S and medium proportions of illite associated with the lowest concentrations of palygorskite of the series indicate an enhanced hydrological cycle (Chamley, 1989). The interval i2 was therefore linked with a semi-arid climate slightly more humid than during i1 deposition, and favorable to a massive input of continental nutrients into the ocean, particularly phosphorus (Wagner et al., 2007). The addition of CO₂ from volcanism and nutrients from both hydrothermal activity and continental weathering in the oceans would have increased the primary productivity, particularly in the upwelling areas as the Northwestern African coast. This increase in primary productivity, promoting dysoxia (Wagner et al., 2007), combined with a higher sedimentation would have favoured the preservation of OM, and thus allowed the deposition of the KLE black shales.

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The third climatic period begin at the base of the KLE (i3; Fig. 6). It is characterized by a significant and drastic change in the mineralogical assemblage with a drop in the abundances of R0 I-S mixed layers and a strong increase in palygorskite and illite concentrations. This reversal of mineralogical trends compared to interval i2 reflects a rapid transition toward a more arid climate post Kilian event. Besides clay mineralogy, exoscopy of quartz grains within the silt and sand fractions through the Kilian event at Site 545 showed a dominant aeolian origin (Trabucho Alexandre et al., 2011). Therefore, clay mineral

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assemblages at Site 545 confirm that Northwestern Africa, was a hot desert or semi-desert environment during the latest Aptian.

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450

### 5.3. The Kilian event: a climate turnover

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452

At a longer-term (>2 Myr), the late Aptian climate was characterised by a global cooling phase particularly evidenced by calcareous nannofossil assemblages typical of cold surface waters (Mutterlose et al., 2009), by physical evidence of glaciation such as glendonite or dropstones (De Lurio and Frakes, 1999), and by long-term increase in whole rock and belemnite $\delta^{18}$O values (Clarke and Jenkyns, 1999; Bodin et al., 2015). This interpretation is confirmed and detailed by sea surface temperatures derived from TEX$_{86}$ data from Site 545 (McAnena et al., 2013), which suggest a cooling of 5°C during this long-time interval (~2 Myr). This cooling phase had a strong impact on the marine biology as documented by a synchronous change in carbon cycle depicted by higher $\delta^{13}$C values (McAnena et al., 2013), higher nutrient index (Herrle et al., 2003a, b), a drop in planktonic species in the foraminifera assemblages and a rise in radiolarians and benthic organisms (Leckie et al., 2002), depicting a higher export production probably due to enhanced oceanic convection.

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The termination of this cooling phase and the return to pre-cooling temperature values broadly corresponds to the end of the i1 phase (Fig. 6). This temporality suggests that temperature and weathering conditions are well correlated in the region and that the warming phase is accompanied with more humid conditions. At the 43-44 cores transition (≈407 mbsf), both TEX$_{86}$ and TOC values increase (McAnena et al., 2013; this study), whereas primary productivity indicators stay high (Leckie et al., 2002). This transition could thus be a consequence of an enhanced continental weathering as marked by the gradual transition to more humid conditions depicted by clay mineralogy (Fig. 6).

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466

467

468

469

470

471
A possible cause for this long-term temperature increase can be attributed to the emplacement of the Kerguelen LIP in the Indian Ocean at that time emitting large amount of CO₂ in the atmosphere (Sabatino et al., 2018; This study). Therefore, high temperatures and CO₂ concentrations would have been the reasons for the increase in weathering conditions in land. Elevated CO₂ concentrations, and the consequent lowering in pH (i.e., ocean acidification), is also a possible cause for the dominance of radiolarians over planktonic foraminifera derived from a reduced carbonate availability (Coccioni et al., 2014; Matsumoto et al., 2020). Besides, high weathering conditions would have boosted the flux of nutrients from lands to the surface waters, increasing the net primary productivity and favouring radiolarian productivity (Herrle et al., 2003a, b). The relationship between a subaerial LIP activity and weathering conditions is here even more precisely established as the evolution in Hg/TOC is temporally coherent with changes in the clay minerals assemblage at Site 545. Therefore, pulses of the Kerguelen LIP volcanic activity seem to have had a direct impact on the late Aptian climate and on primary productivity. The hyperthermal Kilian event, could have been an ultimate consequence of environmental changes caused by a volcanic pulse in a context of already elevated CO₂ concentrations in the oceans and atmosphere, particularly enhanced since the i1-i2 transition. The associated negative δ¹³C excursion may therefore be caused by the ¹³C-depleted CO₂ and CH₄ emissions partly linked with the dissociation of methane clathrate caused by the rise in the oceanic temperatures (Wagner et al., 2007; Zhang et al., 2006).

Increasing amounts of organic carbon storage in sediments is thought to be the main cause of the post-Kilian short-term climate cooling (Wagner et al., 2007; McAnena et al., 2013). At DSDP Site 545, this phase is also associated with an aridification on lands, particularly marked by an increase in palygorskite in the clay mineral assemblages, indicating a detrital flux highly derived from evaporitic-sabkha environment (Fig. 6). The Moroccan
497

margin being located near the transition between the evaporitic climate belt and the mid-
latitude warm humid belt during the Aptian (Chumakov et al., 1995), it is probable that the
variations in weathering conditions were associated with latitudinal migration of this
transition.

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500

501

Further evidences of a regional aridification at the Aptian-Albian transition were
observed in the Iberian Peninsula with the development of desert sedimentary facies
(Rodriguez-Lopez et al., 2006) and the dominance of physical weathering derived in the clay
mineral assemblages and palynological data (Heimhofer et al., 2008). Moreover, large amount
of palygorskite distributed mainly on North-Africa and Iberia confirm the presence of a
regional desert environment in the area (Pletsch et al., 1996). Furthermore, in the Vocontian
and Paris Basins (France), the Aptian-Albian transition is also characterized by cooler and
dryer climatic conditions followed by more intense hydrolysing conditions during the Albian
(Corentin et al, 2020; Deconinck et al., 2021). Those observations confirm a northward
migration of the arid-evaporitic belt at the Aptian-Albian transition. This migration is also
marked by an intensification of the arid conditions at lower latitudes, and of humid conditions
at mid-latitudes (Corentin et al., 2020).

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507

508

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510

511

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513

Such a scenario of subtropical aridification conjugated to a more humid climate in
equatorial and mid-latitudes (Chumakov et al., 1995; Hay and Floegel, 2012) is consistent
with an increase of atmospheric CO₂ concentration as predicted by numerical paleoclimatic
models (i.e. Bice and Norris, 2002) and observed through the Aptian/Albian boundary (Barral
et al., 2017).

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515

516

517

518

Recently, an estimation of the volume and degassing rate of CO₂ emission from the
Southern Kerguelen Plateau proved it is one order magnitude lower than other LIPs such as
Siberian or Central Atlantic Magmatic Province (Jiang et al., 2022), both associated to major
biology crisis. Besides, compared to the major OAEs of the Mesozoic (>500 kyrs-long

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521
Toarcian Event, OAE1a and OAE2; Huang et al., 2010; Charbonnier et al., 2018b), the Kilian event is shorter (90-120 kyrs; Huang et al., 2010; Leandro et al., 2022) and is not associated with an extensive development of anoxic conditions in the oceans. Its main characteristic is a change in the global climate regimes and the development of regional extreme humid conditions favoring the deposition of sapropel-like sediments, caused by a mainly subaerial volcanic activity.

The difference between the major OAEs and the Kilian event within OAE1b seems to result in the intensity of the triggering volcanic activity and/or the type of volcanic activity, as the CO₂ influx seems to be the major factor in the development of an OAE (Jenkyns et al., 2010). OAE1a and OAE2 are associated with the emplacement of large submarine LIPs, the Ontong-Java for the OAE1a and the Caribbean and/or the High-Arctic LIPs for the OAE2. For those two events, the liberated nutrients (mainly Fe) from the hydrothermal activity associated is seen as a factor boosting the primary productivity and developing marine anoxic conditions (Snow et al., 2005). Similarly to the Kilian event, the Toarcian Event is associated with the subaerial activity of a LIP, the Karoo-Ferrar. A distinction can be brought by the high Hg anomalies measured through the Toarcian Event (Percival et al., 2021), probably indicating a more intense emission of volcanic gases.

## 6. Conclusion

We investigated high-resolution geochemical (Rock Eval, Hg) and mineralogical (clay mineral assemblages) data in the late Aptian sediments of the DSDP Site 545 located in the Mazagan Plateau, which includes the Kilian level equivalent.

Two main Hg anomalies occurs immediately below and above the Kilian level in the DSDP Site 545. The absence of significant covariations between Hg and TOC and between
547 Hg and CaCO₃, and the high Hg/TOC trends suggest that these enrichments are mainly linked
548 with volcanic outgassing related to the subaerial phase of the Southern Kerguelen Plateau.

549 Clay mineralogical assemblages are dominated by R0 type illite-smectite mixed-layers
550 (R0 I-S), illite, and palygorskite. The mineralogical distribution has a dominant detrital origin
551 mainly related to local climatic conditions. Three distinct climatic periods can be recognized
552 along the late Aptian:

553 • (i1) a seasonally contrasted semi-arid climate below the first Hg anomaly (R0 I-S and palygorskite);

555 • (i2) a moister climate during the Kilian level equivalent deposition (high proportions of R0 I-S);

557 • and (i3) a rapid transition toward more arid climate post Kilian level equivalent deposition that is linked with the second Hg anomaly (large increases in illite and palygorskite).

560 Our study confirms that the intensification of the southern Kerguelen volcanic activity
561 and the massive input of volcanic CO₂ into the atmosphere triggered the hyperthermal Kilian
562 event. Elevated pCO₂ levels and associated warming may have accelerated the hydrological
563 cycle that boosted the flux of nutrients from continents to ocean, eventually increasing
564 primary productivity and inducing a major marine carbonate production crisis. Finally, this
565 enhanced primary productivity boosted the burial of organic carbon in marine sediments that
566 consumed a significant amount of atmospheric CO₂, which likely contributed to the post
567 Kilian short-term drier and cooler climate. Therefore, the late Aptian represents a key interval
568 in the Cretaceous climate changes between cold climatic conditions that have caused
569 intermittent polar glaciation and the extreme hyperthermal climatic events of the
570 Aptian/Albian boundary (OAE1b).

571
572 Acknowledgments

573
574 We would like to thank François Baudin for Rock Eval analysis, Omar Boudouma and Laurent Riquier for SEM analysis and Karl Föllmi for IODP samples request. We are grateful to Prof. Stephen E. Grasby and an anonymous reviewer for their constructive comments and useful suggestions. We thank also the Swiss National Science Foundation (project 200021_168994) and the University of Lausanne for financial and logistic support.

579
580 References

581
582 Adatte, T., Keller, G., Stinnesbeck, W., 2002. Late Cretaceous to Early Palaeocene climate and sea-level fluctuations: the Tunisian record. Palaeogeography, Palaeoclimatology, Palaeoecology 178, 165-196.

585 Barral, A., Gomez, B., Fourel, F., Daviero-Gomez, V., Lécuyer, C., 2017. CO₂ and temperature decoupling at the million-year scale during the Cretaceous Greenhouse. Scientific Reports 7, 8310 DOI: 10.1038/s41598-017-08234.

588 Beaudoin, B., Fries, G., Bulot, L., 2012. Van Gogh, Cézanne et autres impressionnistes : ben-
tonites aptiennes/albiennes du SE de la France ... et plus loin. 2012. Réunions
589 thématiques du Groupe français du Crétacé, Les évènements de l'Aptien-Albien. Paris,
590 France, pp. 9-15.

592 Begnino, A.P.A., Saraiva, A.A.F., Sial, A.N., Lacerda, L.D., 2021. Mercury chemostratigraphy as a proxy of volcanic-driven environmental changes in the Aptian-Albian transition, Araripe Basin, northeastern Brazil. Journal of South American Earth Sciences 107, 103020.
596 Behar, F., Beaumont, V., De, B., and Penteado, H.L., 2001. Rock-Eval 6 technology: perfomances and develoments: Oil and Gas Science and Technology revue IFP, v. 56, p. 111-134.

599 Benamara, A., Charbonnier, G., Adatte, T., Spangenberg, J.E., Föllmi, K.B., 2020. Precession-driven monsoonal activity controlled the development of the early Albian Paquier oceanic anoxic event (OAE1b): Evidence from the Vocontian Basin, SE France. Palaeogeography, Palaeoclimatology, Palaeoecology 537, 109406.

603 Bergquist, B.A., 2018. Mercury, volcanism, and mass extinctions. Proceedings of the National Academy of Sciences of the United States of America 114, 8675-8677.

605 Bice, K. L., Norris, R.D., 2002. Possible atmospheric CO2 extremes of the warm mid-Cretaceous (late Albian-Turonian). Paleoceanography 17, 1070.

607 Bodin, S., Meissner, P., Janssen, N.M.M., Steuber, T., Mutterlose, J., 2015. Large igneous provinces and organic carbon burial: Controls on global temperature and continental weathering during the Early Cretaceous. Global and Planetary Change 133, 238-253.

610 Bottini, C., Erba, E., 2018. Mid-Cretaceous palaeoenvironmental changes in the western Tethys. Climate of the Past Discussions doi.org/10.5194/cp-2018-34.

612 Bréhéret, J.G, 1995. L'Aptien et l'Albien de la Fosse vocontienne (bordures au bassin) : Evolution de la sedimentation et enseignements sur les evenements anoxiques. François Rabelais, Tours.

615 Brosse, E., 1982. Contribution à la minéralogie et à la géochimie des sédiments pélagiques profonds. Comparaison des "Black-shales" du Crétacé dans l'Atlantique Central Nord et des dépôts du Malm et du Crétacé en Briançonnais. Thèse Doctorat Ingénieur Ecole des Mines de Paris, France 474 pp.

619 Chamley, H., 1979. North Atlantic clay sedimentation and paleoenvironment since the late Jurassic: In: Talwani, M., Hay, W., Ryan, W.B.F., Eds., Deep drilling results in the
621 Atlantic Ocean: Continental margins and paleoenvironment, Maurice Ewing cer.

622 American Geophysical Union 3, 342-361.

623 Chamley, H., 1989. Clay Sedimentology. Springer-Verlag, Berlin. 623 pp.

624 Chamley H., 1994. Clay Mineral Diagenesis. In: Parker A., Sellwood B.W. (eds) Quantitative

625 Diagenesis: Recent Developments and Applications to Reservoir Geology. NATO ASI

626 Series (Series C: Mathematical and Physical Sciences), vol 453. Springer, Dordrecht.

627 https://doi.org/10.1007/978-94-011-0189-9_5

628 Chamley, H., Debrabant, P., 1982. L'Atlantique Nord à l'Albien : influences américaine et

629 africaine sur la sédimentation. Compte Rendu Académie Science de Paris 294, 525-528.

630 Chamley, H., Debrabant, P., 1984. Mineralogical and geochemical investigations of

631 sediments on the Mazagan Plateau, northwestern African margin (Leg 79, DSDP). Pp.

632 497-508 in: Init. Rep. Deep Sea Drill. Proj. 89, (Hinz, K., Winterer, E.L. et al., editors).

633 U.S. Gov. Print. Off. Washington.

634 Charbonnier, G., Föllmi, K.B., 2017. Mercury enrichments in lower Aptian sediments support

635 the link between Ontong Java large igneous province activity and oceanic anoxic

636 episode 1a. Geology 1, 63-66.

637 Charbonnier, G., Morales, C., Duchamp-Alphonse, S., Westermann, S., Adatte, T., and

638 Föllmi, K.B., 2017. Mercury enrichments indicate volcanic triggering of Valanginian

639 environmental change: Scientific Reports DOI:10.1038/srep40808.

640 Charbonnier, G., Godet, A., Bodin, S., Adatte, T., Föllmi, K.B., 2018a. Mercury anomalies,

641 volcanic pulses, and drowning episodes along the northern Tethyan margin during the

642 latest Hauterivian-earliest Aptian. Palaeogeography, Palaeoclimatology, Palaeoecology

643 505, 337-350.
644 Charbonnier, G., Boulila, S., Spangenberg, J.E., Adatte, T., Föllmi, KB., Laskar, J. 2018b. Obliquity pacing of the hydrological cycle during the Oceanic Anoxic Event 2. Earth and Planetary Science Letters 499, 266-277.

647 Charbonnier, G., Adatte, T., Föllmi, K.B., Suan G., 2020a. Effect of intense weathering and post-depositional degradation of organic matter on Hg/TOC proxy in organic-rich sediments and its implications for deep-time investigations. Geochemistry, Geophysics, Geosystems. 21, e2019GC008707

651 Charbonnier, G., Desmares, D., Bardin, J., 2020b. Authigenic, volcanogenic, and detrital influences on the Cenomanian-Turonian clay sedimentation in the Western Interior Basin: Implications for palaeoclimatic reconstruction. Cretaceous Research 106, 104228.

655 Charbonnier, G., Duchamp-Alphonse, S., Deconinck, J.F., Adatte, T., Spangenberg, J.E., Colin, C., Föllmi, K.B., 2020c. A global palaeoclimatic reconstruction for the Valanginian based on clay mineralogical and geochemical data. Earth Sciences Reviews 202, 103092.

659 Chumakov, N.M., Zharkov, M.A., Herman, A.B., Doludenko, M.P., Kalandadze, N.N., Lebedev, E.A., Ponomarenko, A.G., Rautian, A.S., 1995. Climate belts of the Mid-Cretaceous time. Stratigraphy Geology Correlation 3, 241-260.

662 Clarke, L.J., Jenkyns, H.C., 1999. New oxygen isotope evidence for long-term Cretaceous climatic change in the Southern Hemisphere. Geology 27, 699-702.

664 Coccioni, R., Sabatino, N., Frontalini, F., Gardin, S., Sideri, M., Sprovieri, M., 2014. The neglected history of Oceanic Anoxic Event 1b: insights and new data from the Poggio le Guaine section (Umbria-Marche Basin). Stratigraphy 11, 245-282.

667 Coffin, M.F., Eldholm, O., 1994. Large igneous provinces: crustal straucture, dimensions, and external consequences. Reviews of Geophysics 32, 1-36.
669 Coffin, M.F., Pringle, M.S., Duncan, R.A., Gladczenko, T.P., Storey, M., Müller, R.D.,
670 Gahagan, L.A., 2002. Kerguelen Hotspot Magma output since 130 Ma. Journal of
671 Petrology 43, 1121–1139.

672 Corentin, P., Deconinck, J.F., Pellenard, P., Amédro, F., Bruneau, L., Chenot, E., Matrion, B.,
673 Huret, E., Landrein, P., 2020. Environmental and climatic controls of the clay mineralogy
674 of Albian deposits in the Paris and Vocontian Basins (France). Cretaceous Research 108,
675 104342.

676 Daoudi, L., Deconinck, J.F., Beauchamp, J., Debrabant, P., 1989. Minéraux argileux du
677 bassin d'Agadir (Maroc) au Jurassique supérieur-Crétacé. Comparaison avec le domaine
678 Est-Atlantique voisin. Ann. Soc. géol. Nord 108, 15–24.

679 Deconinck, J.-F., Chamley, H., 1995. Diversity of smectite origins in Late Cretaceous
680 sediments: example of chalks from Northern France. Clay Minerals, 30, 365-379.

681 Deconinck, J.F., Boué, D., Amédro, F., Baudin, F., Bruneau, L., Huret, E., Landrein, F.,
682 Moreau, J.D., Santoni, A.L., 2021. First record of early Aptian Oceanic Anoxic Event 1a
683 from the Paris Basin (France)-Climate signals on a terrigenous shelf. Cretaceous
684 Research 125, 104846.

685 De Lurio, J.L., Frakes, L.A., 1999. Glendonites as a paleoenvironmental tool: Implications for
686 early Cretaceous high latitude climates in Australia. Geochimica et Cosmochimica Acta
687 63, 1039-1048.

688 Dellisanti F., Pini G.A., Baudin F. 2010. Use of Tmax as a thermal maturity indicator in
689 orogenic successions and comparison with clay minerals evolution. Clay Minerals, 45/1,
690 115-130.

691 Duncan, R. A. 2002. A timeframe for construction of the Kerguelen Plateau and Broken
692 Ridge. Journal of Petrology 43, 1109-1119.

693 Föllmi, K.B., 2012. Early Cretaceous life, climate and anoxia. Cretaceous Research 35, 230-
694 257.

695 Font, E., Adatte, T., Sial, A.N., de Lacerda, L.D., Keller, G., Punekar, J., 2016. Mercury anomaly, Deccan volcanism, and the end-Cretaceous mass extinction: Geology 44, 171–174.

698 Frey, F.A., Coffin, M.F., Wallace, P.J., Weis, D., Zhao, X., Wise, S.W., Wähnert, V., Teagle, D.A.H., Saccocia, P.J., Reusch, D.N., Pringle, M.S., Nicolaysen, K.E., Neal, C.R., Müller, R.D., Moore, C.L., Mahoney, J.J., Keszthelyi, L., Inokuchi, H., Duncan, R.A., Delius, H., Damuth, J.E., Damasceno, D., Coxal, H.K., Borre, M.K., Boehm, F., Barling, J., Arndt, N.T., Antretter, M., 2000. Origin and evolution of a submarine large igneous province: the Kerguelen Plateau and Broken Ridge, southern Indian Ocean. Earth and Planetary Science Letters 176, 73–89.

705 Frey, F.A., Weis, D., Borisova, A.Y., Xu, G., 2002. Involvement of continental crust in the formation of the Cretaceous Kergeulen Plateau: new perspectives from ODP Leg 120 Sites. Journal of Petrology 43, 1207–1239.

708 Frizon de Lamotte, D., Fourdan B., Leleu S., Leparmentier F., de Clarens P., 2015. Style of rifting and the stages of Pangea breakup. Tectonics, 34, 1009-1029.

710 Gibbs, R. J., 1977. Clay mineral segregation in the marine environment, J. Sediment. Petrol., 47, 237–243,

712 Gradstein, F.M., Ogg, J.G., Schmitz, M.D., Ogg, G.M., 2020. The Geologic Time Scale 2020, 2 volumes, 1390 pp. doi.org/10.1016/C2020-1-02369-3.

714 Grasby, S.E., Sanei, H., Beauchamp, B., and Chen, Z.H., 2013, Mercury deposition through the Permo-Triassic biotic crisis: Chemical Geology, 351, 209–216.

716 Grasby, S.E., Beauchamp, B., Bond, D.P.G., Wignall, P., and Sanei, H., 2015, Mercury anomalies associated with three extinction events (Capitanian crisis, latest Permian
718 extinction and the Smithian/Spathian extinction) in NW Pangea: Geological Magazine,
719 13 p.

720 Grasby, S.E., Them II, T.R., Chen, Z., Yin, R., Ardakani, O.H., 2019. Mercury as a proxy for
721 volcanic emissions in the geologic record. Earth Science Reviews 196, 102880.

722 Graziano, R., Raspini, A., Bartiromo, A., 2016. Late Aptian palaeoclimatic turnovers and
723 volcanism: Insights from a shallow-marine and continental succession of the Apennine
724 carbonate platform, southern Italy, Sedimentary Geology, 339, 188-217.

725 Guezal, J., El Baghdadi, M., Barakat, A., Raïs, J., 2011. Le magmatisme jurassique-crétacé de
726 Béni-Mellal (Haut-Atlas central, Maroc) : géochimie et signification géodynamique.
727 Bulletin de l'Institut Scientifique, Sciences de la Terre, 33, 17-23.

728 Guezal, J., El Baghdadi, M., Barakat, A., 2015. The Jurassic–Cretaceous volcanism of the
729 Atlas of Beni-Mellal (Central High Atlas, Morocco): evidence from clinopyroxene
730 composition. Arabian Journal of Geosciences, 8, 977-986.

731 Hay, W.W., Floegel, S., 2012. New thoughts about the Cretaceous climate and oceans. Earth
732 Science Reviews 115, 262–272.

733 Hay, W.W., DeConto, R.M., Wold, C.N., Wilson, K.M., Voigt, S., 1999. Alternative global
734 Cretaceous paleogeography. Geological Society of America Special Papers, 332, 1-47.

735 Heimhofer, U., Adatte, T., Hochuli, P.A., Burla, S., Weissert, H., 2008. Coastal sediments
736 from the Algarve: low-latitude climate archive for the Aptian-Albian. International
737 Journal Earth Science 97, 785-797.

738 Herrle, J.O., Pross, J., Friedrich, O., Hemleben, C., 2003a. Short-term environmental changes
739 in the Cretaceous Tethyan Ocean: micropalaeontological evidence from the Early
740 Albian Oceanic Anoxic Event 1b. Terra Nova 15, 14–19. https://doi.org/10.1046/j.1365-
741 3121.2003.00448.x

742 Herrle, J.O., Pross, J., Friedrich, O., Kößler, P., Hemleben, C., 2003b. Forcing mechanisms
743 for mid-Cretaceous black shale formation: evidence from the Upper Aptian and Lower
744 Albian of the Vocontian Basin (SE France). Palaeogeography, Palaeoclimatology,
745 Palaeoecology 190, 399-426. https://doi.org/10.1016/S0031-0182(02)00616-8

746 Herrle, J.O., Kößler, P., Friedrich, O., Erlenkeuser, H., Hemleben, C., 2004. High-resolution
747 carbon isotope records of the Aptian to Lower Albian from SE France and the Mazagan
748 Plateau (DSDP Site 545): a stratigraphic tool for paleoceanographic and paleobiologic
749 reconstruction. Earth and Planetary Science Letters, 218, 149-161.

750 Herrle, J.O., Schröder-Adams, C.J., Davis, W., Pugh, A.T., Galloway, J.M., Fath, J., 2015.
751 Mid-cretaceous High Arctic stratigraphy, climate, and Oceanic Anoxic Events. Geology
752 43, 403-406.

753 Hinz, K., and Winterer, E.L., et al. (Shipboard Scientific Party), 1984. Initial reports of the
754 Deep Sea Drilling Project 79, US Government Printing Office, Washington DC, 81-177.

755 Hofmann, P., Stüsser, I., Wagner, T., Schouten, S., Sinnighe Damsté, J.S., 2008. Climate-
756 Ocean coupling off North-West Africa during the Lower Albian: the Oceanic Anoxic
757 Event 1b. Palaeogeography, Palaeoclimatology, Palaeoecology, 262, 157-165.

758 Holtzapffel, T., 1985. Les minéraux argileux : préparation, analyse diffractométrique et
détermination. Société géologique du Nord, 12.

760 Huber, B.T., Leckie, R.M., 2011. Planktic foraminiferal species turnover across deep-sea
761 Aptian/Albian boundary sections. Journal of Foraminiferal Research, 41, 53-95.

762 Huang,, C., Hinnov, L., Fischer, A.G., Grippo, A., Herbert, T. 2010. Astronomical tuning of
763 the Aptian stage from Italian reference sections. Geology 38, 899-902.

764 Jenkyns, H.C., 2010. Geochemistry of oceanic anoxic events: REVIEW. Geochemistry,
765 Geophysics, Geosystems 11, n/a-n/a. https://doi.org/10.1029/2009GC002788.
766 Jiang, Q., Jourdan, F., Olierook, H.K., Merle, R.E., Bourdet, J., Fougerouse, D., Godel, B.,
767 Walker, A.T., 2022. Volume and rate of volcanic CO2 emissions governed the severity
768 of past environmental crisis. PNAS 119, 1-10.

769 Lanson, B., Sakharov, B.A., Claret, F., Drits, V.A., 2009. Diagenetic smectite to illite
770 transition in clay rich sediments: A reappraisal of X-ray diffraction results using the
771 multi specimen method. American Journal of Science 309, 476-516.

772 Leandro, C.G., Savian, J.F., Kochhann, M.V.L., Franco, D.R., Coccioni, R., Frontalini, F.,
773 Gardin, S., Jovane, F., Figueiredo, M., Tedeschi, L.R., Janikian, L., Almeida, R.P.,
774 Trindade, R.I.F. 2022. Astronomical tuning of the Aptian stage and its implications for
775 age recalibrations and paleoclimatic events. Nature Communication 13, :2941|
776 https://doi.org/10.1038/s41467-022-30075-3.

777 Leckie, R.M., 1984. Mid-Cretaceous planktonic foraminiferal biostratigraphy off central
778 Morocco, Deep Sea Drilling Project Leg 79, Sites 545 and 547. In: Hinz, K, Winterer,
779 E.L., et al. (eds.), Initial Reports of the Deep Sea Drilling Project, Washington (U.S.
780 Govt. Printing Office), 79, 579-620.

781 Leckie, R.M., Bralower, T.J., Cashman, R., 2002. Oceanic anoxic events and plankton
782 evolution: Biotic response to tectonic forcing during the mid-Cretaceous: Oceanic
783 anoxic events and plankton evolution. Paleoceanography 17, 13-1-13-29.

784 Li, X., Wei, Y., Li, Y., Zhang, C., 2016. Carbon isotope records of the early Albian oceanic
785 anoxic event (OAE) 1b from eastern Tethys (southern Tibet, China). Cretaceous
786 Research, 62, 109-121.

787 Madhavaraju, J., Lee, Y.I., Scott, R.W., González-León, C.M., Jenkyns, H.C., Saucedo-
788 Samaniego, J.C., Ramasamy, S., 2018. High-resolution carbonate isotopic study of the
789 Mural Formation (Cerro Pimas section), Sonora, México: Implications for early Albian
790 oceanic anoxic events. Journal of South American Earth Sciences 82, 329-345.
791 https://doi.org/10.1016/j.jsames.2017.10.019.

792 Matsumoto, H., Kuroda, J., Coccioni, R., Frontalini, F., Sakai, S., Ogawa, N.O., Ohkouchi,

793 N., 2020. Marine Os isotopic evidence for multiple volcanic episodes during Cretaceous

794 Oceanic Anoxic Event 1b. Scientific Reports DOI: 10.1038/s41598-020-69505-x.

795 McAnena, A., Flögel, S., Hofmann, P., Herrle, J., Griesand, A., Pross, J., Talbot, H.,

796 Rethemeyer, J., Wallmann, K., Wagner, T., 2013. Atlantic cooling associated with a

797 marine biotic crisis during the mid-Cretaceous period. Nat. Geosci. 6, 558–561.

798 Mélières, R, 1978. X-ray mineralogy studies, DSDP Leg 41. In Lance-lot, Y., Seibold, E., et

799 al., Init. Repts. DSDP, 41: Washington (U.S.Govt. Printing Office), 1065-1086.

800 Méhay, S., Keller, C.E., Bernasconi, S.M., Weissert, H., Erba, E., Bottini, C., Hochulli, P.A.

801 2009. A volcanic CO2 pulse triggered the Cretaceous Oceanic Anoxic Event 1a and a

802 biocalcification crisis. Geology 37, 819-822.

803 Merriman, R. J., Frey, M. 1999. Pattern of very low- grade metamorphism in metapelitic

804 rocks. In Frey, M., and Robinson, D., eds., Low-grade metamorphism. Oxford,

805 Blackwell Science, p. 61–107.

806 Moore, D.M., Reynolds, R.C.J., 1997. X-Ray diffraction and the Identification and Analysis

807 of Clay Minerals. Oxford University Press, New York, 378 pp.

808 Mohr, B.A.R., Wähnert, V., and Lazarus, D., 2002. Mid-Cretaceous paleobotany and

809 palynology of the central Kerguelen Plateau, southern Indian Ocean (ODP Leg 183, Site

810 1138). In Frey, F.A., Coffin, M.F., Wallace, P.J., and Quilty, P.G. (Eds.), Proc. ODP,

811 Sci. Results, 183, 1–39.

812 Mutterlose, J., Bornemann, A., Lupplod, F.W., Owen, H.G., Ruffell, A., Weiss, W., Wray, D.,

813 2003. The Vöhrum section (northwest Germany) and the Aptian/Albian boundary.

814 Cretaceous Research 24, 203–252.

815 Mutterlose, J., Bornemann, A., Herrle, J., 2009. The Aptian–Albian cold snap: evidence for
816

“mid” Cretaceous icehouse interludes. N. Jb. Geol. Paläont. (Abh.) 252, 217–225.

817

Navarro-Ramirez, J. P., Bodin, S., Heimhofer, U., Immenhauser, A., 2015. Record of Albian to early Cenomanian environmental perturbation in the eastern sub- equatorial Pacific.

818

819

Palaeogeography, Palaeoclimatology, Palaeoecology, 423, 122-137.

820

Outridge, P.M., Sanei, H., Stern, G.A., Hamilton, P.B., Goodarzi, F., 2007. Evidence for control of mercury accumulation rates in Canadian High Arctic lake sediments by variations of aquatic primary productivity. Environmental Science & Technology, 41, 5259–5265.

821

822

823

824

Park, J., Stein, H.J., Georgiev, S.V., Hannah, J.L., 2022. Degradation of Hg signals on incipient weathering: Core versus outcrop geochemistry of Upper Permian shales, East Greenland and Mid-Norwegian Shelf. Chemical Geology, 608, 121030.

825

826

827

Percival, L.M.E., Witt, M.L.I., Mather, T.A., Hermoso, M., Jenkyns, H.C., Hesselbo, S.P., Al-Suwaidi, A.H., Storm, M.S., Xu, W., and Ruhl, M., 2015, Globally enhanced mercury deposition during the end-Pliensbachian extinction and Toarcian OAE: A link to the Karoo–Ferrar large igneous province: Earth and Planetary Science Letters, v. 428, p. 267–280.

828

829

830

831

832

Percival, L.M.E., Cohen, A.S., Davies, M.K., Dickson, A.J., Hesselbo, S.P., Jenkyns, H.C., Leng, M.J., Mather, T.A., Storm, M.S., Xu, W., 2016. Osmium isotope evidence for two pulses of increased continental weathering linked to Early Jurassic volcanism and climate change. Geology 44, 759-762.

833

834

835

836

Percival, L.M.E., Jenkyns, H., Matter, T.A., Dickson, A.J., Batenburg, S.J., Ruhl, M., Hesselbo, S.P., Barclay, R., Jarvis, I., Robinson, S.A., Woelders, L., 2018. Does large igneous province volcanism always perturb the mercury cycle? Comparing the records of oceanic anoxic event 2 and the end-Cretaceous to other Mesozoic events. American Journal of Science 318, 1-62.

837

838

839

840
841 Percival, L.M.E., Tedeschi, L.R., Creaser, R.A., Bottini, C., Erba, E., Giraud, F., Svensen, H.,
842 Savian, J., Trindade, R., Coccioni, R., Frontalini, F., Jovane, L., Mather, T.A., jenkyns,
843 H.C. 2021. Determining the style and provenance of magmatic activity during the Early
844 Aptian Oceanic Anoxic Event (OAE1a). Global and Planetary Change 200, 103461.

845 Petschick, R., 2000. MacDiff 4.2.6 manual. Available at http://www.geol-pal.uni-
846frankfurt.de/Staff/Homepages/Petschick/PDFs/MacDiff%20Manual%20E.pdf
847(consulted 03/02/20)

848 Pletsch, T., Daoudi, L., Chamley, H., Deconinck, J., Charroud, M., 1996. Palaeogeographic
849 controls on palygorskite occurrence in mid-Cretaceous sediments of Morocco and
850 adjacent basins. Clay Minerals 31, 403-416.

851 Premoli Silva, I., W. V. Sliter, 1999. Cretaceous paleoceanography: Evidence from planktonic
852 foraminiferal evolution, in Evolution of the Cretaceous Ocean-Climate System, edited
853 by E. Barrera and C. C. Johnson, Spec. Pap. Geol. Soc. Am., 332, 301-328.

854 Rodríguez-López, J.P., de Boer, P.L., Meléndez, N., Soria, A.R. Pardo, G., 2006. Windblown
855 desert sands in coeval shallow marine deposits a key for the recognition of coastal ergs;
856 mid-Cretaceous Iberian Basin, Spain. Terra Nova, 18, 314-320.

857 Ruellan, E., Auzende, J.M., Dostmann, H., 1985. Structure and evolution of the Mazagan (El
858 Jadida) plateau and escarpment off central Morocco. Oceanologica Acta, Special issue,
859 59-72.

860 Sabatino, N., Coccioni, R., Manta, D.S., Baudin, F., Vallefuoco, M., Traina, A., Sprovieri,
861 M., 2015. High-resolution chemostratigraphy of the late Aptian-early Albian oceanic
862 anoxic event (OAE1b) from the Poggio le Guaine section (Umbria-Marche Basin,
863 central Italy). Palaeogeography, Palaeoclimatology, Palaeoecology 426, 319-333.

864 Sabatino, N., Ferraro, S., Coccioni, R., Bonsignore, M., Del Core, M., Tancredi, V., Sprovieri,
865 M., 2018. Mercury anomalies in upper Aptian-lower Albian sediments from the Tethys
866 realm. Palaeogeography, Palaeoclimatology, Palaeoecology 495, 163-170.

867 Sanei, H., Grasby, S.E., and Beauchamp, B., 2012, Latest Permian mercury anomalies: 868 Geology, v. 40, p. 63-66.

869 Scaife, J.D., Ruhl, M., Dickson, A.J., Mather, T.A., Jenkyns, H.C., Percival, L.M.E., 870 Hesselbo, P., Cartwright, J., Eldrett, J.S., Bergman, S.C., Minisini, D. 2017.

871 Sedimentary mercury enrichments as a marker for submarine Large Igneous Province 872 volcanism? Evidence from the Mid-Cenomanian Event and Oceanic Anoxic Event 2 873 (Late Cretaceous). Geochemistry, Geophysics, Geosystems. 18, 4253-4275.

874 Selin, N.E., Jacob, D.J., Yantosca, R.M., Strode, S., Jaeglé, L., Sunderland, E.M. 2008. 875 Global 3-D land-ocean-atmosphere model for mercury: Present-day versus preindustrial 876 cycles and anthropogenic enrichment factors for deposition. Global Biogeochemical 877 cycles. 22, 1-13.

878 Sial, A.N., Lacerda, L.D., Ferreira, V.P., Frei, R., Marquillas, R.A., Barbosa, J.A., Gaucher, 879 C., Windmoller, C.C., and Pereira, N.S., 2013, Mercury as a proxy for volcanic activity 880 during extreme environmental turnover: The Cretaceous-Paleogene transition: 881 Palaeogeography, Palaeoclimatology, Palaeoecology, v. 387, p. 153-164.

882 Shen, J., Algeo, T.J., Chen, J., Planavsky, N.J., Feng, Q., Yu, J., & Liu, J. 2019. Mercury in 883 marine Ordovician/Silurian boundary sections of South China is sulfide-hosted and non- 884 volcanic in origin. Earth and Planetary Science Letters 511, 130-140.

885 Snow, L.J., Duncan, R.A., Bralower, T.J., 2005. Trace element abundances in the Rock 886 Canyon Anticline, Pueblo, Colorado, marine sedimentary section and their relationship 887 to Caribbean plateau construction and oxygen anoxic event 2. Paleoceanography 20.

888 Stampfli, G., C. Hochard. 2009. Plate tectonics of the Alpine realm. – In: Murphy, J. B., J. D. 889 Keppie, A. J. Hynes (Eds.). Ancient Orogens and Modern Analogues. Geol. Soc. 890 London, Special Publications, 327, 89-111.
891 Šucha, V., Kraus, I., Gerthofferová, H., Peteš, J., Sereková, M., 1993. Smectite to illite
892 conversion in bentonites and shales of the East Slovak Basin. Clay Minerals 28, 243-
893 253.

894 Thiry, M., 2000. Palaeoclimatic interpretation of clay minerals in marine deposits: an outlook
895 from the continental origin. Earth-Sciences Reviews, 49, 201-221.

896 Trabucho Alexandre, J., Van Gilst, R. I., Rodríguez-Lopez, J. P., & De Boer, P. L. 2011. The
897 sedimentary expression of oceanic anoxic event 1b in the North Atlantic.
898 Sedimentology, 58, 121-1246.

899 Wagner, T., Wallmann, K., Herrle, J.O., Hofmann, P., Stüesser, I., 2007. Consequences of
900 moderate ~25,000 yr lasting emission of light CO2 into the mid-Cretaceous ocean. Earth
901 and Planetary Science Letters, 259 (1-2), 200-211.

902 Whitechurch, H., Montigny, R., Sevigny, J., Storey, M. & Salters, V. J. M. 1992. K–Ar and
903 $^{40}$Ar/$^{39}$Ar ages of central Kerguelen Plateau basalts. In: Wise, S. W., Jr, Schlich, R., et
904 al. (eds) Proceedings of the Ocean Drilling Program, Scientific Results, 120. College
905 Station, TX: Ocean Drilling Program, pp. 71–77.

906 Worden, R. H., Morad, S., 1999. Clay minerals in sandstones: controls on formation,
907 distribution and evolution. Clay mineral cements in sandstones, 1-41.

908 Zhang, X., Chen, K., Hu, D., Sha, J., 2016. Mid-Cretaceous carbon cycle perturbations and
909 Oceanic Anoxic Events recorded in southern Tibet. Scientific Reports, 6, 39643.

910 Zhao, X., Zheng, D., Wang, H., Fang, Y., Xue, N., Zhang, H., 2022. Carbon cycle
911 perturbation and mercury anomalies in terrestrial Oceanic Anoxic Event 1b from
912 Jiuquan Basin, NW China. Geological Society London Special Publications 521, 185-
913 196.

914 Zimmerle, W. 1989. Vestiges of volcanic activity in Cretaceous sediments of Europe. In:
915 Wiedmann, J. (ed.) Cretaceous of the Western Tethys. Proceedings of the 3rd
916 International Cretaceous Symposium, Tubingen, 1987, E. Schweizerbart'sche

917 Verlagsbuchhandlung (Nägele u. Obermiller), Stuttgart, 951-987.

918 Zintwana, M.P., Cawthorn, R.G., Ashwal, L.D., Roelofse, F., and Cronwright, H., 2012.

919 Mercury in the Bushveld Complex, South Africa, and the Skaergaard Intrusion,
920 Greenland: Chemical Geology, v. 320, p. 147-155.

921

922 **Figure caption**

923

924 **Fig. 1:** Palaeogeographic map of the late Early Cretaceous (120 Ma) (modified from Frizon
925 de Lamotte et al., 2015) showing the location of the Mazagan Plateau (DSDP Site 545) and
926 key late Aptian sedimentary successions recording the Kilian event. MP: Mazagan Plateau;
927 VB: Vocontian Basin; UMB: Umbria Marche Basin; and KP: Kerguelen Plateau.

928

929 **Fig. 2: A)** Example of a diffractogram from DSDP Site 545 (sample: core 43R-1w-145-146,
930 399.45 mbsf) for untreated, glycolated and heated tests. **B)** Scanning electron microscope
931 examinations of the whole rock at Site 545 (sample: 41R-2W-5-6, 380.55 mbsf).

932

933 **Fig. 3: A)** Stratigraphic evolution of the δ¹³Ccarb (‰ VPDB; Herrle et al., 2004; McAnena et
934 al., 2013), total organic carbon (TOC, wt. %), carbonate contents (CaCO₃, %), mercury
935 contents (Hg, ppb), and Hg/TOC (ppb/ wt. %) records from the DSDP Site 545 (cores 41-45).

936 **B)** Stratigraphic evolution of the δ¹³Ccarb (‰ VPDB; Herrle et al., 2004; McAnena et al.,
937 2013), and clay mineral assemblage (R0 I/S, palygorskite, illite, and kaolinite, %) records
938 from the DSDP Site 545 (cores 41-45). The biostratigraphic framework of the core is given by
939 planktonic foraminifera and calcareous nannofossils (Leckie, 1984; Leckie et al., 2002; Herrle
940 et al., 2004; McAnena et al., 2013).
941
942 **Fig. 4:** HI/T<sub>max</sub> diagram for preserved organic matter (OM) in the late Aptian interval (cores 943 41-45).

944
945 **Fig. 5: A) TOC (wt. %) versus Hg (ppb) and B) CaCO<sub>3</sub> (%) versus Hg (ppb) diagrams for the** 946 samples from the DSDP Site 545.

947
948 **Fig. 6:** Stratigraphic evolution of the δ<sup>13</sup>C<sub>carb</sub> (‰ VPDB; Herrle et al., 2004; McAnena et al., 949 2013), the Hg/TOC ratio (ppb/wt. %, tracer for volcanism), the TEX<sub>86</sub> (°C) and the δ<sup>18</sup>O<sub>carb</sub> (‰ VPDB) compared to the distribution of R0 I/S, palygorskite and illite contents (%). The 950 biostratigraphic framework of the core is given by planktonic foraminifera and calcareous 951 nannofossils (Leckie, 1984; Leckie et al., 2002; Herrle et al., 2004; McAnena et al., 2013).

952
953
<figure><img src="image_2.png" /></figure>

Subduction Front

Thrust Front

Transform fault and Mid oceanic ridge

Deep marine

Shallow marine

Continental

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

Kerguelen Plateau Large Igneous Province (LIP)

ENVIRONMENT

Lower Cretaceous Rifts

Upper Jurassic to Lower Cretaceous Rifts

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

DSDP Site 545

Other key sections
<figure><img src="image_5.png" /></figure>

<figure><img src="image_6.png" /></figure>
A) Geochemical data

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

B) Mineralogical data

<figure><img src="image_8.png" /></figure>
<figure><img src="image_9.png" /></figure>

Immature Zone

Oil Zone

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

<figure><img src="image_11.png" /></figure>
<figure><img src="image_12.png" /></figure>

Herrle et al., 2004
McAnena et al., 2013