# Climatic and weathering conditions in southern high latitudes during the Turonian-Santonian interval: New insights from IODP Site U1512 (Bight Basin, Southern Australia)

Thomas Munier, Laurent Riquier, Sidonie Révillon, Hans-Jürgen Brumsack, Christian Hasler, Omar Boudouma, François Baudin

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## To cite this version:

Thomas Munier, Laurent Riquier, Sidonie Révillon, Hans-Jürgen Brumsack, Christian Hasler, et al. "Climatic and weathering conditions in southern high latitudes during the Turonian-Santonian interval: New insights from IODP Site U1512 (Bight Basin, Southern Australia). Palaeogeography, Palaeoclimatology, Palaeoecology, 2023, 613, pp.1111413. 10.1016/j.palaeo.2023.1111413. hal-03971789

HAL Id: hal-03971789

https://hal.science/hal-03971789v1

Submitted on 27 Mar 2025

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# Climatic and weathering conditions in southern high latitudes during the Turonian-Santonian interval: New insights from IODP Site U1512 (Bight Basin, Southern Australia)

Thomas Munier¹*, Laurent Riquier¹, Sidonie Révillon², Hans-Jürgen Brumsack³, Christian Hasler³, Omar Boudouma¹, François Baudin¹

¹ISTeP, UMR 7193, SU/CNRS, Sorbonne Université, 4 Place Jussieu, Paris, France.

² SEDISOR/Geo-ocean, UMR 6538, UBO/CNRS/IFREMER/UBS, Institut Universitaire Européen de la Mer, Plouzané, France.

³ICBM, University of Oldenburg, Oldenburg, Germany.

* Corresponding author: thomas.munier@sorbonne-universite.fr

## ABSTRACT

A detailed study, based on a coupled mineralogical and geochemical approach, was done on IODP Site U1512 sediments (Bight Basin, Southern Australia). The mineral assemblages of clay size fraction (<2 µm), obtained by X-ray diffraction, have been compared to inorganic (major and trace elements) geochemical signals, obtained from both whole rock and clay size fraction, and Sr/Nd isotopes, obtained from clay size fraction to determine climatic and weathering conditions during the Turonian to Santonian interval (~94-84 Ma, Late Cretaceous) in the southern high latitude band. This study reveals that the mineral assemblages of clay fraction consist of R0-type illite/smectite mixed-layers (considered as smectites), illites, and kaolinites associated with opal-CT and clinoptilolites. Electron microscopy observations have demonstrated a negligible impact of burial diagenesis on clay assemblages while authigenesis contributions of clay minerals are limited and associated with minor volcanic contributions. Clay assemblages would be thus dominantly primary in origin. Based on isotopic Sr/Nd values, clay minerals are interpreted as the results of the weathering of rocks and pedogenic blankets of Eastern Australian Highlands and from the remobilization of older sediments from nearby uplifted basins. The chemical compositions of both whole rock and clay size fraction show a relative stability during the studied interval with two exceptions: during the Middle/Late Turonian times for Si, P and Ba and at the Turonian-Coniacian boundary for Ti, K and Zr, associated respectively to an increase of biogenic fraction and to an enhanced input of terrigenous material. The clay assemblages, dominated by smectites, show weak variations along the borehole reflecting stable climatic conditions during the 10-Myr Turonian-Santonian interval. The dominance of smectites, coupled to the absence of chlorite and the low proportions of illite, which are both considered as primary clay minerals, indicates a dominant warm/temperate and sub-humid climate during the studied interval. However, changes in weathering conditions occurred during the Turonian-Santonian interval
in southern high latitude. The Early Turonian, characterized by high Chemical Index of Alteration values (>75), is followed by an interval of lower chemical weathering with enhanced biogenic silica during the Middle/Late Turonian. The slight increase of kaolinites, recorded during Coniacian and coeval with increase of Ti/Al and Zr/Al ratios, probably reflects an increase in coarser terrigenous inputs associated with a sea-level fall.

**Keywords: Continental weathering, Sub-humid Climate, Clay minerals, Sr/Nd isotopes, Major and trace elements.**

# 1. INTRODUCTION

The Cretaceous is characterized by a global increase of seafloor spreading rates, leading to the progressive opening of South-Atlantic, proto-Indian and Austral oceans in southern hemisphere (Larson, 1991a, b; Lawver et al., 1991; Eagles, 2007; Moulin et al., 2010). The increase of the mid-oceanic ridge activity and the development of large igneous provinces have resulted in the rise of atmospheric CO₂ (Arthur et al., 1985; Jarvis et al., 2011; Hong and Lee, 2012). This has led to an intensification of the greenhouse effect and a progressive increase of surface and bottom oceanic temperatures during Aptian-Cenomanian times punctuated by a thermal maximum recorded from the top of the Cenomanian to the end of the Santonian (Huber et al., 2002; Friedrich et al., 2012; O'Connor et al., 2019; Petrizzo et al., 2020). Although the Turonian-Santonian interval (~94-84 Ma) is considered as the hottest interval of the Cretaceous (Huber et al., 2018; Petrizzo et al., 2020), a short cooling episode has been identified at the Late Turonian/Early Coniacian in the Tethyan domain (Jenkyns et al., 1994; Stoll and Schrag, 2000; Wiese and Voigt, 2002) which extends to Late Coniacian in the southern hemisphere (Clarke and Jenkyns, 1999; Falzoni et al., 2016; O'Brien et al., 2017; Huber et al., 2018). The Turonian-Santonian interval is also considered as a key interval for the Gondwana breakup, especially in the southern margin of Australia. Rifting phase, which started during the Middle Jurassic along this margin, result in the creation of several offshore basins, as the Bight Basin. During Early Cretaceous (Berriasian-Albian), this basin and its sub-basins is characterized by fluvio-lacustrine sedimentary sequence (Totterdell et al., 2008). Marine incursions, which started in the Upper Albian, are marked by fine grained siliciclastic sediments as silty-clay deposits, typical of shallow shelf environment, recorded within the Blue Whale, White Pointer and Tiger supersequences in the Bight Basin (Totterdell and Bradshaw, 2004; Veevers, 2006). The transition from fluvial (Albian) to marine sedimentation (Turonian-Santonian) is associated to a period of high sea-level peaked during the early Turonian (Haq, 2014) and a phase of accelerated subsidence in the Bight Basin probably caused by crustal extension, recorded since the mid Albian (Krassay and Totterdell, 2003; Meeuws et al., 2016). The rifting phase ended by the continental breakup between Australia and
71 Antarctica recorded during the Late Santonian/Early Campanian (Totterdell et al., 2008). The 10 Myr-
72 lasting Turonian-Late Santonian/Early Campanian interval is thus marked by noticeable environmental
73 changes associated to major geodynamic and climatic upheavals and may have led to variations in
74 weathering conditions in southern margin of Australia.

75 However, unlike in the northern hemisphere, climatic conditions in the southern mid and high
76 latitudes are less constrained during the Turonian-Santonian interval. Sea surface temperatures are
77 scarce and focused only on a few oceanic boreholes in proto-Indian Ocean (Clarke and Jenkyns, 1999;
78 Huber et al., 2002; Petrizzo, 2002; Falzoni et al., 2016; Petrizzo et al., 2020). Continental climate
79 conditions, based on palynological data are limited to a few Australian basins (Eromanga and
80 Gippsland Basins) and are poorly dated (Dettmann et al., 1992; Macphail, 2007; Specht et al., 1992;
81 Fletcher et al., 2014). While the Turonian-Santonian interval is a key period for understanding
82 variations in weathering conditions induced by climate and geodynamics, the evolution of these
83 parameters in the eastern part of Gondwana, and more precisely in the Austral Ocean is still poorly
84 documented. To better constrain weathering and climate conditions associated to the breakup of the
85 eastern part of Gondwana in this time interval, a study, based on a coupled mineralogical and
86 geochemical approach, was done on a new sedimentary sequence from Site U1512 in the Bight Basin.
87 This site was drilled during International Ocean Discovery Program (IODP) Expedition 369 and
88 records a nearly complete Turonian to Santonian sedimentary succession, corresponding in age to the
89 Tiger Supersequence of Totterdell et al. (2000). This sequence, mainly consisting of silty claystone,
90 has been interpreted as deposited in the distal part (i.e., marine prodelta) of the Ceduna delta (Bight
91 Basin; Wainman et al., 2020). This latter was fed by a continental-scale fluvial system, which has
92 drained a large part of the eastern part of Australia during the Cretaceous (Raza et al., 1995; Veevers,
93 2000). Because high latitudes are more sensitive to climatic variations, they could record slight climate
94 fluctuations in a global warm context (MacLeod et al., 2020). The aim of this study is thus to identify
95 the possible expansion of the warm/temperate and humid climatic belt at high latitudes during the
96 hottest period of the Cretaceous and to reveal evolution of weathering conditions during the Turonian-
97 Santonian interval for the eastern part of Gondwana.

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## 2. GEOLOGICAL SETTING

100 Site U1512 is located at 3000 m water depth in the west of the Ceduna Sub-basin (Bight
101 Basin, Fig. 1A, Huber et al., 2019), a part of the Bight Basin. This basin corresponded to a narrow
102 epicontinental seaway bordered by Archean to Proterozoic cratons (Gawler and Yilgarn cratons),
103 Proterozoic orogenic crust (Albany-Fraser) and Proterozoic to Paleozoic sedimentary basins (Officer
104 Basin) (Fig 1B). The Bight Basin extends over a large part of the Southern Rift System in the southern
105 Australian margin. This system was developed from the Middle Jurassic (Oxfordian, ~155 Ma) with
the progressive separation of the Australian and Antarctic continents (Totterdell and Bradshaw, 2004; Veevers, 2006). The rift phase (Oxfordian to Berriasian) was followed by a first slow thermal subsidence phase (Berriasian to Albian), characterized by fluvio-lacustrine sedimentation (Totterdell et al., 2008). The Lower Cretaceous is also characterized by significant volcanic activity (e.g., Whitsunday siliceous large igneous province, Bryan et al., 2012; Barham et al., 2016) associated with the subduction on the East Australian margin. During the Cenomanian, the interruption of the subduction on this margin stopped the related volcanism (Jones and Veevers, 1983; Veevers, 2000; Gallagher et al., 1994; Russell and Gurnis, 1994; Bryan et al., 2012; Pain et al., 2012). This end of the subduction resulted in the uplift of Eastern Australian Highlands and the rearrangement of the drainage system from east to west (Bishop, 1985; Raza et al., 2009; Waschbusch et al., 1999; Veevers, 2000; Bradshaw et al., 2012). A second phase of accelerated subsidence occurred from mid-Albian to Late Santonian/Early Campanian. This phase coincided with a global sea-level rise and the first marine influence on sedimentation (Stagg et al., 1990; Totterdell and Krassay, 2003; Totterdell et al., 2008). The mid-Albian to Late Santonian/Early Campanian subsidence phase also led to an increase of accommodation space in the Bight Basin, especially in the Ceduna Sub-basin, and to high sedimentation rates of marine silt and clay deposits derived from a continental-scale fluvial system feeding the Ceduna delta. Based on previous studies (Sayers et al., 2001; Totterdell et al., 2000; Bradshaw et al., 2003; Totterdell and Krassay, 2003; Totterdell and Bradshaw, 2004; Espurt et al., 2009; MacDonald et al., 2013; MacLeod et al., 2020), the silty claystone succession, deposited in this delta, would be derived from the weathering of the uplifted eastern part of Australia including Eastern Australian Highlands and Eromanga Basin. The expansion of the rift between Australia and Antarctica led to the break-up of continental crust in the west and slow seafloor spreading during the Late Santonian/Early Campanian (Bryan et al., 2012). This progressive expansion led to a new thermal subsidence phase of continental margin, including Ceduna delta (Sayers et al., 2001; Krassay et al., 2004; Blevin and Cathro, 2008; White et al., 2013). The sediments remain deltaic until the Cenozoic, finer in the prodelta and affected by a strong marine influence in the distal zones.

Site U1512 records a well recovered (> 90%) ~660 m thick succession of Turonian to Santonian (~94 to 84 Ma) deposits overlain by 32 m of poorly-dated sediments and 10 m thick Pleistocene nannofossil ooze (Huber et al., 2019). Turonian to Santonian deposits of Site U1512 correspond to the Tiger Supersequence present throughout the Ceduna Sub-basin (Wainman et al., 2020). This supersequence, described from deposits of the Gnarlyknots 1/1A borehole (located in the delta plain, Fig. 1A), is mainly constituted of marine mudstone and silty, sandy deltaic sediments (Totterdell et al., 2000, 2008; Totterdell and Krassay, 2003; Blevin and Cathro, 2008). In Site U1512, sediments consist dominantly of grey silty claystone (prodeltaic deposits) with small amounts of biogenic components. These deposits, interpreted as corresponding to prodelta environment, appear to form a finer-grained lateral equivalent of the Tiger Supersequence, recorded in the Gnarlyknots 1/1A
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borehole (Wainman et al., 2020). This ~660 m thick succession is occasionally interrupted by coarse-
grained thin beds (2 to 20 cm) of clayey/silty, sideritic and glauconitic sandstones (Fig. 2, MacLeod et
al., 2020; Wainman et al., 2020; Watkins and Guerra, 2020). The sideritic levels are observed
throughout the borehole whereas the glauconitic sandstones are mainly observed from ~180 to 360 m
depth in a 180 m thick interval of dark gray silty claystone, depleted in carbonate (CaCO₃ < 0.5%) and
interpreted as a fining up third order sequence, associated to a phase of restriction of the prodelta basin
(Wainman et al., 2020). Then, Turonian to Santonian deposits of Site U1512 correspond to a marine
prodelta succession in which continental and marine influence are related to sea-level variations
(Wainman et al., 2020).

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The biostratigraphy of Site U1512 was established using calcareous nannofossils (Fig. 2).
Preliminary biostratigraphic data from Huber et al. (2019) were refined by Watkins and Guerra (2020)
showing nannofossils assemblages from Lower Turonian (CC10b) to Santonian (CC16) after
Gradstein et al. (2012). The Lower Turonian extends from 700.21 to 489.52 m. Nannozones CC10b
and CC11 respectively range from 700.21 to 622.22 m and from 618.75 to 489.52 m. The
Middle/Upper Turonian extends from nannozones CC12 (485.45 to 319.51 m) to CC13 (314.12 to
202.03 m, Fig. 2), also confirmed by the *Falsotruncana maslakovae* Zone from 476.79 to 428.98 m.
The nannozone CC13 is affected by a large barren interval from 292 to 220 m (Watkins and Guerra,
2020). The Lower to Middle Coniacian extends from 186.03 to 110.01 m (CC14) followed by Upper
Coniacian (103.74 to 96.28, CC15). The Santonian extends from 90.97 to 48 m (CC16).

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# 3. MATERIAL AND METHODS

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## 3.1. Clay minerals

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Clay minerals were analyzed on 120 samples using X-Ray diffraction (XRD) method. After
grinding of about 3 g in an agate mortar, samples were treated with 0.6% HCl solution to remove
carbonates. The < 2 µm fraction (clay fraction) was extracted after decantation for 95 min following
Stokes' Law with a syringe (Holtzapffel, 1985). After centrifugation, clay residue was deposited and
smeared on oriented glass slides and analyzed in a Brucker D2 Phaser X-ray diffractometer with CuKα
radiation, a LynxEye detector and a Ni filter (ISTeP laboratory, Sorbonne University, France).
Diffractometer was set at 30 kV, 10 mA and goniometer scanning from 2.5° to 28.5°. Three runs were
operated on samples to discriminate clay minerals: 1) air-drying, 2) ethylene-glycol solvation, 3)
heating at 490° (Holtzapffel, 1985; Moore and Reynolds, 1997). Clay minerals identification were
performed using main diffraction peaks and by comparison between air-drying, ethylene-glycol
solvation and heating diffractograms. Clay minerals identified were: R0 type illite-smectite mixed-
layers (15 Å on air-drying runs and 17 Å on glycolated runs), illites/micas (10 Å, 5 Å and 3.33 Å) and
kaolinites (7.18 Å and 3.58 Å, disappeared on heating runs) (Moore and Reynolds, 1997). Clay
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minerals were quantified on glycolated diffractograms using MacDiff software (Petschick, 2000). The
proportion of each clay species is estimated by semi-quantification using the area of main peaks.

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**3.2. Scanning and Transmission Electron Microscopy (SEM and TEM)**

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The morphology observations of clay minerals, identified by XRD have been realized on Scanning Electron Microscope (SEM) and Transmission Electron Microscope (TEM) on selected samples. Nine whole-rock samples were observed using a variable pressure ZEISS Supra 55 VP with a GEMINI field emission column SEM (ISTeP laboratory, Sorbonne University, France). Observations were performed on fractures made on whole-rock samples to avoid observing altered surfaces. These samples were positioned on an aluminum plate and metallized. Clay minerals were identified using their chemical composition analyzed by a Brucker QUAD energy dispersive spectrometer (EDS). Three clay fraction samples were prepared to be examined by TEM. The clay fraction (< 2 µm) was extracted following the method of clays minerals preparation previously described. A slight quantity (0.1 ml) of the preparation were mixed in 1.5 ml of 2% butylamine solution to separate minerals. A drop of the solution was disposed on a carbon formwar copper grid covered by a carbon film then observed in a Jeol JEM-2100F TEM equipped with a Brucker XFlash Detector 5030 spectrometer (EDS) (ICB laboratory, Burgundy University, France). Clay minerals were identified based on their morphology and chemical composition measured by EDS spectrometer.

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**3.3. Major and trace elements**

**3.3.1. Analysis of whole rock samples**

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Concentrations of major elements and selected trace elements were also determined by X-ray fluorescence (XRF) for 68 whole rock samples. XRF analyses were performed at the Institute for Chemistry and Biology of the Marine Environment (University of Oldenburg) with a Panalytical AXIOS Plus spectrometer. Prior to analysis, sediment sample splits were ground and homogenized in an agate ball mill. For quantitative wavelength dispersive XRF, glass beads were prepared from 700 mg of powdered sample after pre-oxidation with 1 g ammonium nitrate (NH₄NO₄) using di-lithium tetraborate (LiBO₂ or Li₂B₄O₇) as a flux (4.200 mg flux). Calibration of the XRF instrument was based on 52 international reference samples and in-house standards. Precision was determined from the pooled standard deviation of several duplicates and is better than 3% for major elements and better than 5 % for trace elements when concentration levels are above the detection limit. In order to distinguish noticeable enrichment of major and trace elements from a detrital background, the concentrations of all analyzed elements measured on whole-rock samples were normalized with respect to Al and compared to Post-Archean Australian Shale values (PAAS, Taylor and McLennan, 1985). The comparison between the samples from Site U1512 and the PAAS, which characterizes an
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average terrigenous component, reveals enrichments in biogenic or volcanogenic materials. In addition to major and trace element concentrations, the Chemical Index of Alteration (CIA) was calculated for whole rock and clay fraction samples. This index is defined by the following molar formula: CIA = $[Al_2O_3/(Al_2O_3 + CaO* + Na_2O + K_2O)] \times 100$ where CaO* refers to the carbonate- and apatite-free analyses (Nesbitt and Young, 1982). For clay fraction, as concentrations were measured on the decarbonated clay fractions, CaO* is only corrected for apatite. Lastly, the TOC values were deduced from the inorganic carbon concentrations obtained by XRF from whole rock analyses. The Total Organic Carbon was calculated by the following formula $TOC = C_{tot} - C_{min}$, with $C_{tot}$ corresponding to the total Carbon and $C_{min}$ representing the mineral carbon. $C_{min}$ is obtained from CaO concentrations, supposing all mineral carbon are incorporated in $CaCO_3$.

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### 3.3.2. Analysis of clay fraction

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In order to identify some similar trends and possible differences between the whole-rock fraction and the clay fraction, concentrations of major elements and selected trace elements (i.e., Zr, Ba) were also determined from the clay fraction for 12 selected samples following the method described in Mourlot et al. (2018). About 3 g of sample were crushed in an agate mortar. The powder was treated with 10% acetic acid and 5% hydrogen peroxide to remove carbonate and organic matter. The residual fractions were cleaned with ultrapure water (18 MΩ) and the clay fraction was extracted after decantation for 95 min following Stokes’ Law with a syringe (Holtzapffel, 1985). To remove possible persistent coarse-grained components, clay fraction was refined by centrifugation in two steps. Residual fractions were centrifuged for 2 min at 1000 rpm then clay-rich supernatants were again centrifugated for 2.5 min at 800 rpm. Clay fractions were collected after centrifugation for 40 min at 3,500 rpm. About 100 mg of treated rock powder were dissolved by a HF and $HNO_3$ solution in Teflon vials for digesting silicates then $H_3BO_3$ solution were added to neutralize HF. Spectrometer analyses were performed at the Pôle de Spectrométrie Océan (Plouzané, France). Major elements were measured using an Inductively Coupled Plasma Atomic Emission Spectrometer Horiba Jobin Yvon Ultima 2. Precision of the analyses are better than 5% for major elements except for Fe, Na (6%) and K (8%). Trace elements were analyzed using a High Resolution Coupled Plasma Mass Spectrometer Thermo Finnigan Element XR with a precision of measurements better than 5% for most elements.

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## 3.4. Sr and Nd isotopes

Sr and Nd isotope compositions were analyzed on 12 clay fractions samples. About 3 g of sample were crushed in an agate mortar. Powder was treated with 10% acetic acid during 4h and 5% hydrogen peroxide overnight to remove carbonate and organic matter respectively (Bayon et al., 2002; Corentin et al., 2022). Residual fractions were cleaned with ultrapure water (18 MΩ) then clay fraction
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was extracted following the method described for XRD. To remove possible persistent coarse elements, clay fraction was refined by centrifugation in two steps (Mourlot et al., 2018). Residual fractions were centrifuged for 2 min at 1000 rpm then clay-rich supernatants were again centrifugated for 2.5 min at 800 rpm. Clay fractions were collected after centrifugation for 40 min at 3500 rpm.

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About 150 mg of clay-rich samples were dissolved in Savillex beakers with a mixture of ultrapure distilled HF (24N) and HNO₃ (14N) for 96 hours at 120°C. Sr and Nd fractions were separated by ionic chromatography using ion exchange resins TRU-Spec, Sr-Spec and Ln-Spec (Eichrom Technologies) following column chromatography method described by Pin and Zalduegui (1997). Compositions of Sr and Nd isotopes were measured using a ThermoFisher Scientific Triton thermal ionization mass spectrometer (TIMS) at the Pôle de Spectrométrie Océan (IUEM, Plouzané, France). Sr and Nd ratios were normalized with a ⁸⁶Sr/⁸⁸Sr ratio of 0.1194 and a ¹⁴⁶Nd/¹⁴⁴Nd of 0.7219, respectively to correct signal of mass fractionation. The reliability of measurements was evaluated by the analysis of NBS 987 (⁸⁷Sr/⁸⁶Sr = 0.710245 ± 0.000011) and JNd (¹⁴³Nd/¹⁴⁴Nd = 0.512115 ± 0.000007) international standard. The analysis of international standards added to the sample series show standard deviations (2σ) under 0.003%. For this study, the average standard values measured are ⁸⁷Sr/⁸⁶Sr = 0.710263 for NBS 987 (n = 3) and ¹⁴³Nd/¹⁴⁴Nd = 0.512097 for JNd (n = 5). The variations of the ¹⁴³Nd/¹⁴⁴Nd ratios show slight amplitudes, thus the measured values are compared to the Chondritic Uniform reservoir (CHUR) value (0.512638, DePaolo, 1981). ε<sub>Nd(0)</sub> was calculated following the formulae: ε<sub>Nd(0)</sub> = [(¹⁴³Nd/¹⁴⁴Nd)<sub>sample</sub>/(¹⁴³Nd/¹⁴⁴Nd)<sub>CHUR</sub> - 1] x 10⁴ (DePaolo and Wasserburg, 1976).

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# 4. RESULTS

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## 4.1. Clay mineral assemblages

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Clay minerals assemblages of Site U1512 are composed by R0-type illite/smectite mixed-layers (I-S R0, subsequently called smectites), kaolinites and illites (Fig. 3). In detail, smectites are the predominant clay minerals (77 to 99%, Fig. 4). They are associated with kaolinites (0 to 19%) and with weak proportions of illites (1 to 10%). No significant variations in relative proportions of clay minerals are observed in the core, except some slight increases of kaolinites in the Early Turonian (568 and 576 m depth) and in the Coniacian (184 to 100 m depth) deposits. In these depths, the average proportions of kaolinite increase from 5% up to 14% in the Early Turonian and 19% in the Coniacian. Rare peaks up to 10% of kaolinites are also observed at 672 m (Early Turonian), 440.3 m and 204 m depth (Middle/Late Turonian). For illite, only one noticeable peak (9%) is observed at 376 m depth (Middle/Late Turonian). The clay-size fraction (< 2 µm) also includes quartz, opal-CT and clinoptilolites (Figs. 3 & 4). The diffraction peaks of opal-CT, observe from 52 m depth, present progressively higher intensity, which indicate an increase of proportions of this mineral along the

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borehole and become very abundant between 320 and 380 m. Then, its proportion slightly decrease, and this mineral suddenly disappears with clinoptilolites around 420 m. This demise of opal-CT is associated with higher proportion of well-crystallized quartz. The proportion of quartz increases downward until the end of the drilling and is the main minerals in 2 µm fraction between 590 and 700 m.

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## 4.2. SEM and TEM observations

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Scanning electron microscopy observations, performed on representative samples of the borehole, confirm the presence of opal-CT and clinoptilolites and allow to highlight the dominant morphology of both minerals, i.e., opal CT lepispheres (micrometric spheres of opal-CT) and well-crystallized zeolites (Fig. 5A, B, C, D, E & F). These minerals filled several dissolved shells and foraminifera tests (Fig. 5F). Most of the samples present throughout the borehole significant proportions of quartz, micas and some feldspars of large size (> 60 µm), associated with strong quantities of clays and silts (Fig. 5C). Barite and pyrite framboids, often associated with secondary gypsum, are occasionally observed in few samples between 350 to 180 m in carbonate-poor silty claystone of the Middle/Upper Turonian. These samples are also associated with highest content of zeolites. Transmission electron microscopy observations of clay minerals performed on selected samples at 72, 328 and 648 m depth confirm the presence of fleecy R0-type illite/smectite mixed-layer minerals associated with low proportions of sharp outlined illites and kaolinites (Fig. 5G & H). Some folded smectites are also observed in less abundance (Fig. 5G & I).

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## 4.3. Major and trace elements

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The sedimentary succession of the Site U1512 is dominantly composed of Al and Si-bearing minerals, as shown in the ternary SiO₂-Al₂O₃-CaO plot (Fig. 6), where all the samples (whole rock and clay fraction) plot close to the Al₂O₃-SiO₂ side. This also highlights the low Ca content (mainly <5%) for the Turonian-Santonian deposits of this site. This diagram also shows the shift of some whole rock samples towards the SiO₂ pole with respect to the PAAS (Taylor and McLennan, 1995). These samples correspond to the highest opal-CT content, especially in the interval between 420 and 250 m depth (Middle/Late Turonian). To complete these first observations and to evaluate the potential origin (detrital, biogenic, authigenic) of the major- and trace-elements measured through the U1512 site, the correlation coefficient (r-value) and associated probability (p-value) for each element against Al were determined for whole-rock samples (Table 1). Highest r-coefficients (>0.80) are obtained for Pb (0.93), Ti (0.86), Co (0.86), Zn (0.86), V (0.83); variations in these elements are thus highly influenced by the detrital content. Conversely, other elements are either poorly correlated with Al (r < 0.5), as P, Ba and Cu, or anti-correlated with Al, as Si (-0.98), suggesting major contribution of the biogenic
fraction for these elements. To study the variations of both (detrital and biogenic) contributions, the temporal evolution of selected major (Si, Ti, K and P) and trace (Zr, Ba) elements are represented in Figures 6 and 7. These elements show moderate variations with most of the values either oscillating around those the PAAS, as for Ti and Zr or are slightly higher than these, as for Si. More precisely, Ti/Al and Zr/Al ratios show values ranging respectively from 0.05 to 0.09 (average of 0.06) and 12.2 to 68.6 (average of 22.3, Fig. 7). These values are slightly above PAAS values from 697 to 254 m depth but show a gradual increase up to PAAS values (0.06 for Ti/Al and 20.9 for Zr/Al) from Late Turonian to Santonian (from 254 to 52 m depth, CC13 to CC16, Fig. 7). In the clay fraction, these ratios show same trends but are logically more depleted (average values of 0.04 and 15.8) because of the low proportion of heavy minerals in this fraction. The K/Al ratio shows values, oscillating from 0.14 to 0.32 (average 0.20) throughout the borehole. Slight increases are observed during Middle/Late Turonian (418 and 311 m depth, CC12) and during Santonian (102 to 52 m depth, CC15-CC16). The clay fraction data further shows the same trends for this ratio (average of 0.14). To complete the observations for these elements, the CIA was calculated. This index shows values ranging from 61 to 81 (average of 75) in whole rock (Fig. 6). Lower Turonian samples (439 to 683 m, CC10b-CC11) have the highest values, oscillating between 76 and 81. It is followed by a decrease during the Middle/Upper Turonian down to 61 (at 340 m, CC12). An increase of CIA occurs at the end of the Turonian (330 m, CC13) and continues in the Coniacian with values ranging from 67 to 78 in whole rock (87 to 92 in the clay fraction). A new decrease starts in the Santonian (102 m, CC15-16) with values dropping from 77 to 67 in whole rock.

For elements associated to biogenic content (e.g., Si, P and Ba), temporal variations are clearly different (Fig 8). For these elements, intervals of the highest values usually correspond to those where TOC values are higher than 1%. Si/Al ratio show variations with values ranging from 2.8 to 20.8 (average value of 5.8, Fig. 8). Compared to the PAAS value (2.9), a slight enrichment is recorded throughout the studied interval punctuated by a noticeable increase of Si/Al value up to 20.8 during a 56-meter-thick interval between 377 to 321 m depth (Middle/Late Turonian, CC12, Fig. 8). This enrichment with respect to PAAS is also observed in the clay fraction (average value of 4.2). P/Al ratio ranges from 0.003 to 0.009 with an average value of 0.005, lower than the PAAS value (0.007). This ratio shows three intervals with increasing values, from 697 to 648 m depth (Early Turonian, CC10b), from 377 to 321 m depth (Middle/Late Turonian, CC12) and from 168 to 52 m depth (Coniacian-Santonian, CC14 to CC16). In the clay fraction, P/Al show stable values oscillating from 0.001 to 0.004. Ba/Al ratio shows significant variations with values ranging from 26.5 to 295.0 (average value of 121.3), enriched compared to PAAS (65.0). This ratio shows two intervals of increasing values: one from 700 to 600 m depth (Early Turonian, CC10b-CC11) and the second from 425 to 294 m depth (Middle/Late Turonian, CC12-13). Clay fraction values show similar trend with values oscillating from 14.4 to 305.0 (average of 109.2, Fig. 8).
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### 4.4. Nd and Sr isotope composition

Values of $ε_{Nd(0)}$ and $^{87}$Sr/$^{86}$Sr vary from -15.8 to -5.9 and 0.716400 to 0.0726730 respectively (Fig. 4). These two isotopic ratios show a rather good anti-correlation (-0.64). Values of $ε_{Nd(0)}$ increase from -15.8 to -5.9 during the Early Turonian from (684 to 576 m depth, CC10b to CC11) and then remain stable around -6.5 until the beginning of the Middle/Late Turonian (~472 m depth, CC12). The $ε_{Nd(0)}$ values then decrease from -6.0 to -15.3 (472 to 348 m depth, CC12) followed by a slight increase up to -12.7 until the Santonian (Fig. 4). The $^{87}$Sr/$^{86}$Sr ratio shows inverse variations. It increases from 0.718528 to 0.724201 at the base of the Turonian (684 to 636 m depth, CC10b) then decreases to 0.714922 (576 m depth, CC11) and stabilizes around 0.716308 until the beginning of the Middle/Late Turonian (~472 m depth, CC12; Fig. 4). The $^{87}$Sr/$^{86}$Sr ratio then increases from 0.716745 to 0.723844 during the Middle/Late Turonian (472 to 408 m depth, CC12) before decreasing again to 0.721024 at the end of the Turonian (224 m depth, CC13). The $^{87}$Sr/$^{86}$Sr ratio then increases up to 0.726730 from the end of the Turonian to the Coniacian (120 m depth, CC14) and then decreases slightly to 0.721789 in the Santonian (64 m depth, CC16).

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## 5. DISCUSSION

### 5.1 Authigenic and diagenetic processes on clay mineral assemblages

In marine sedimentary record, clay minerals may have different origins (Chamley, 1989). Even if they are usually considered as having a primary origin associated to detrital input, a secondary authigenic origin is possible. This contribution has to be estimated before any interpretation. Indeed, formation of authigenic clay minerals, especially smectite, in marine sediments can occur during early diagenesis, hydrothermal and hydrogenous processes. Authigenic smectites in Mesozoic deposits are usually associated with (i) low sedimentation rates, (ii) submarine alteration of volcanic basalts and (iii) transformation of volcanic glass (Jeans, 1982; Chamley, 1989; Weaver, 1989; Velde, 1995; Fagel, 2007). They can also grow in the pore space of siliceous microfossils tests (Cole and Show, 1983).

At Site U1512, the high sedimentation rate estimated within the Turonian (120 m/Myr) even though it decreases during the Coniacian and the Santonian (25 m/Myr, MacLeod et al., 2020) is not favourable with the formation of authigenic smectites. The development of the mid-oceanic ridge in the Australo-Antarctic Gulf occurs during Late Santonian/Early Campanian, therefore not at the same time as Turonian-Santonian sedimentary record of U1512 (Sayers et al., 2001). Thus, smectites in U1512 deposits are likely not related to submarine weathering of volcanic basalts. About the possible formation of authigenic smectites from alteration of volcanic glass, the association of smectites with clinoptilolites and opal-CT has been proposed a possible clue for this (Pomerol and Aubry, 1977).
However, the occurrence of opal-CT and clinoptilolites could be also associated to enhanced productivity of siliceous organisms. Karpoff et al. (2007) demonstrated that a high productivity of siliceous organisms and the dissolution of their skeleton leads to a high supply of Si, favourable to the formation of clinoptilolites. The occurrence of opal-CT, clinoptilolites, associated with traces of siliceous organisms, has already been observed in Turonian sediments of the western north Atlantic Basin (DSDP Site 387) and in Egypt (Abu Roash Formation, Riech and von Rad, 1979; Hussein and Abd El Rahman, 2020). In Site U1512, similar record, confirmed by the evolution of Ba/Al and to a lesser extent Si/Al and P/Al (Fig. 8), is observed during a 125 m-thick Middle/Late Turonian interval (CC12-CC13, 425 to 300 m depth). So, in Site U1512, clinoptilolites and opal-CT associations have likely a biogenic origin and the smectites would essentially have a detrital origin. This is confirmed by SEM and TEM observations, in which abundant detrital fleecy smectites, which are not associated with authigenic opal-CT and clinoptilolites expanding on detrital particles and filling the porosity, are observed (Fig. 5A & B). This is also confirmed by εNd(0) and ⁸⁷Sr/⁸⁶Sr ratios, which exhibit values higher than those obtained from the scarce isotopic data available for Mesozoic volcanism in Australia (⁸⁷Sr/⁸⁶Sr ratio < 0.705000; εNd(0) values > 1.5) in the Bowen Basin (Allen et al., 1997). As SEM and MET observations, isotopic measurement excludes a dominantly volcanic origin of clay minerals that would involve more radiogenic isotopic values for Nd and less radiogenic values for Sr. However, a possible but limited volcanic-derived origin of some smectites cannot be excluded, because TEM observations also show few folded smectites, typical of the influence of volcanism and the weathering of volcanic ash (Chamley et al., 1985). These smectites are present in Late Turonian to Santonian deposits but are not associated with an increase in smectite proportions and more radiogenic Nd isotope compositions and less radiogenic Sr isotope compositions. The formation of authigenic kaolinites, characterized by well-crystallized booklet morphology (Worden and Morad, 1999), is also unlikely in Site U1512. Despite a higher proportion of kaolinites (> 10%) in some coarser beds in Coniacian sediments (Fig. 4), SEM observations do not reveal any authigenic morphology for kaolinites. This absence of authigenic kaolinites is also confirmed by broad peaks on diffractograms indicating poor crystallinity, typical of detrital kaolinites (Fig. 3).

Lastly, the impact of burial diagenesis on the clay assemblage seems to be limited. This has been confirmed by previous studies, based on the organic fraction and realized in the Bight Basin (Totterdell et al., 2008) and more recently in the Site U1512 (Wainman et al., 2020). Based on low Tmax values (406 to 426°C, average 418°C) obtained by Rock-Eval pyrolysis, the organic matter is considered as immature for the Site U1512 (Wainman et al., 2020). Burial temperature thus probably never exceeded ~60°C (Dellisanti et al., 2010). This is in good agreement with the high proportion of smectites and the absence of illite rich illite/smectite mixed-layers (I-S R1) in clay assemblage. The progressive transformation of smectites layers into R1-type illite/smectite mixed-layers (composed by 50 to 80% of illite layers) known as the illitisation process usually starts at a temperature of about 60-
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70°C (Šucha et al., 1993; Lanson et al., 2009). Based on clay mineral composition, this diagenetic
process thus has been limited and the burial temperatures must not have exceeded ~60°C. We infer
that clay minerals from Site U1512 have mainly a detrital origin and their variations are thus linked to
the weathering impacts, induced by climatic and environmental conditions, on adjacent landmasses.

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**5.2 Origin of clay minerals**

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To determine more precisely the origin of clay minerals, we have compared the ε<sub>Nd(0)</sub> and
<sup>87</sup>Sr/<sup>86</sup>Sr values measured from decarbonated clay-sized fraction of Site U1512 with those obtained for
adjacent landmasses (Fig. 9). In this IODP site, ε<sub>Nd(0)</sub> and <sup>87</sup>Sr/<sup>86</sup>Sr values range from -15.8 to -5.9
(average of -11.5) and 0.714922 to 0.726730 (average of 0.721131) respectively. These values differ
from those recorded in potential sources in the nearest South Australian Craton including Archean to
Proterozoic magmatic and metamorphic rocks of Gawler Craton, Adelaide Fold Belt, Stuart Shelf and
Curnamona Basin (Fig 1B; Parker, 1993; Daly et al. 1998; Fairclough et al., 2003). In these four
provinces, ε<sub>Nd(0)</sub> values are lower than -20 in average and <sup>87</sup>Sr/<sup>86</sup>Sr ratio are much higher with values
close to 0.736250 in average (Fig. 9). Only the Adelaide Fold Belt presents ε<sub>Nd(0)</sub> values (~ -14.5) close
to clay minerals values of Site U1512. These values exclude a main contribution of these nearest
massifs of the South Australian craton to Site U1512 deposits. The potential sediment sources located
in the north of the Ceduna Sub-basin (Officer Basin, Musgrave Province, Amadeus Basin and Arunta
Inlier) as well as those located in the west (Yilgarn Craton; Albany-Fraser Orogen) exhibit also
isotopic values that differ significantly from those obtained from the clay fraction of Site U1512 to
represent a potential source of terrigenous sediments (Fig. 9). In contrast, the compilation of isotopic
values of the geological domains of the Eastern Australian Highlands presents average values for
Paleozoic magmatic-derived rocks, which are close to those recorded in clay fraction of Site U1512,
with ε<sub>Nd(0)</sub> of -12.3 and <sup>87</sup>Sr/<sup>86</sup>Sr ratio of 0.726010 for the Lachlan Fold Belt, ε<sub>Nd(0)</sub> of -9.1 and <sup>87</sup>Sr/<sup>86</sup>Sr
ratio of 0.711850 for the New England Orogen. Only the Nd and Sr isotopic values obtained for
Mesozoic sedimentary and volcanic rocks of the Bowen Basin (ε<sub>Nd(0)</sub> ~ 3.5 and <sup>87</sup>Sr/<sup>86</sup>Sr ratio ~
0.7063101; Fig. 9) do not correspond to the value range of Site U1512. Although these reliefs are
distant from Bight Basin, the Lachlan Fold Belt and New England Orogen seem thus to be the main
sources of terrigenous clay fractions recorded in the Ceduna Sub-basin. Three samples located
between 576 and 472 m depth record more radiogenic values in ε<sub>Nd(0)</sub> (-7.4 to -5.9) and less radiogenic
<sup>87</sup>Sr/<sup>86</sup>Sr (0.714922 to 0.716745). These values are associated with a decrease in Si/Al (3.9 to 2.0) and
a slight increase in Zr/Al (15.1 to 18.9) ratios measured in the clay fraction. These data suggest an
increased contribution of rocks, likely derived from Whitsunday siliceous large igneous province (e.g.,
Bowen Basin, Eromanga Basin) in the formation of clay minerals in the Lower Turonian without
change of clay mineralogy. These rocks would be characterized by a more radiogenic signal for Sm-
Nd isotope systems and a less radiogenic signal for the Rb-Sr isotope system, respectively.
Tectonic and geomorphological processes can explain this distant source for the clay minerals recorded at Site U1512. The Cenomanian (100.5-94 Ma) is characterized in this region by the interruption of the subduction on the East Australian margin and the opening of the Tasman Sea (Jones and Veevers, 1983; Veevers, 2000; Gallagher et al., 1994; Russell and Gurnis, 1994; Pain et al., 2012). This resulted in the uplift of Eastern Australian Highlands, initiated at the end of Cenomanian (~95 Ma) and the rearrangement of the drainage system from east to west, exit in the Ceduna Sub-basin (Bight Basin, Fig. 1B, Bishop, 1985; Raza et al., 2009; Waschbusch et al., 1999; Veevers, 2000; Bradshaw et al., 2012). These conditions had probably led to the weathering of Paleozoic granitoids, volcanic rocks and fine-grained sediments from Eastern Australian Highlands, which were then transported by a SW-flowing rivers system to the Ceduna delta (Taylor, 1994; Veevers, 2000; Totterdell and Krassay, 2003; King and Mee, 2004). This interpretation is in agreement with Hf isotopes measurements performed on detrital zircon from the Gnarlyknots-1 well, (Lloyd et al., 2016), which highlight that sediments from Ceduna delta were largely sourced from Lachlan Fold Belt and New England Orogen rocks. Some studies have highlighted the presence of zircon associated with the pre-Cenomanian volcanism of the Whitsunday siliceous large igneous province (MacDonald et al., 2013; Lloyd et al., 2016; Foley et al., 2021). These data thus indicated that a portion of the sediments recorded in the Ceduna sub-basin originate from remobilization of Early Cretaceous volcaniclastic sediments from the Bowen and Eromanga basins. This is confirmed by isotopic data obtained for Site U1512, which indicate a minor input of the volcaniclastic sediments. However, this minor contribution may explain the more radiogenic values for Nd and less radiogenic values for Sr observed in Lower Turonian (Fig. 4). Previous studies have also suggested that sediments were partly derived from eastern Musgrave Province weathering (MacDonald et al., 2013; Lloyd et al., 2016; Foley et al., 2021, Fig. 10). Thus, clay minerals recorded in Site U1512 sediments reflect weathering of Eastern Australian Highlands and their transport into the Ceduna sub-basin with minor input from Precambrian rocks (e.g. Musgrave Province, Yilgarn and Gawler craton) and pre-Cenomanian volcaniclastic sediments (Fig. 10).

In this case, the high and nearly constant smectite proportions could correspond to a detrital influx, associated to magmatic rocks continental weathering (e.g., granites, basalts) in poorly drained plains located downstream of theses highlands. In addition, the short-term kaolinite-enriched interval, recorded during the Coniacian (180-120 m depth), would indicate a change of sources and/or topography. The kaolinite increase may thus result from either pedogenesis processes under well-drained conditions or reworking of sediments (e.g., claystones, sandstones, Hurst, 1985; Deconinck and Vanderaveroet, 1996). In contrast to kaolinites formed in soils during pedogenesis, reworked kaolinites from ancient sedimentary rocks exhibit covariations with illites because, in this case, they derive both from the same process (Corentin et al., 2020). During the short-term Coniacian kaolinite-enriched interval, a significant covariation ($r = 0.86, n = 16$) between kaolinite and illite is recorded.
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This would indicate a remobilization of these minerals from older sediments. However, the absence of
correlation between the isotopic measurements and the proportions of kaolinites (r = 0.34 and -0,15 for
εNd(t0) and for ⁸⁷Sr/⁸⁶Sr, respectively, n = 12) excludes a change of source to explain a change in clay
mineralogy.

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**5.3 Evolution of climate and weathering conditions**

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In Site U1512, despite rare volcanogenic contributions, the clay minerals are mainly of
terrigenous origin and are thus related to weathering conditions on continental domains of Southern
Australia. The high proportions of smectites (~92%) recorded on U1512 site suggest intermediate
hydrolysis conditions under a warm/temperate and seasonally humid climate for most part of the
sedimentary record, confirmed by absence of chlorite and low proportions of illites, easily weathered
under humid conditions. These intermediate chemical weathering conditions are also confirmed by the
modest values of CIA (~75) obtained for whole-rock analyses. Some previous palynological studies in
northeast Australia (Queensland) mentioned warm climate with seasonally contrast in humidity
confirmed by growth rings on fossil woods (Hallam, 1985; Dettmann, 1994; Macphail, 2007; Fletcher
et al., 2015). Other studies, in the south of Australia, reveal the presence of a hygrophic flora (fern) as
*Cyathidites minor*, *Gleicheniidites* spp., *Dictyophyllidites* spp. and *Microcachryidites antarcticus*
(Dettmann et al., 1992). Recent data from Site U1512 show same assemblages from a humid climate
(Wainman et al., 2020). These spores and pollen grains, commonly observed during Late Cretaceous
in southern Australia, are typical of rainforest assemblages and are thus interpreted as markers of a
humid climate from Turonian to Santonian (Dettmann et al., 1992; Fletcher et al., 2014; Wainman et
al., 2020). Dettmann (1994) and Macphail (2007) believed that these wet conditions were prevalent
along coastline and eastern highlands, but dry seasons persisted inland and on plains. It appears that
the Late Cretaceous forests in south-eastern Australia are similar to the modern forests currently found
in north-western Australia where the mean average temperature is around 18°C (Specht et al., 1992;
Dettmann, 1994). Although Specht et al. (1992) estimated a mean average temperature between 16.5
and 19.5°C during the Santonian, Pirrie and Marshall (1990) found lower temperatures in Antarctica
(13.6°C). While northern palynological (Arnhem Land, Northern Territory) and inland data show a
warm and seasonally humid climate, it appears that southern data reflect more humid and slightly
cooler conditions (Macphail, 2007). Data obtained on clay minerals in Site U1512 seem to confirm the
presence of a warm/temperate and sub-humid climate in southern Australia. These data are consistent
with previous modelling studies determined by lithological and palynological data (Chumakov et al.,
1995; Hay and Floegel, 2012).

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These long-term weathering conditions may have changed during two intervals of kaolinite
increase (respectively up to 14 and 19%), observed during Early Turonian (~600 to 504 m, CC11) and
Coniacian (~204 and 100 m, CC14). Comparable tendencies have been also observed in DSDP Site 361 from Upper Turonian to Lower Santonian deposits (Maillot and Robert, 1980, Cape Basin, South Africa), located at the same paleolatitude. These increases in kaolinite proportions, coupled to highest CIA values (>75) may reflect an increase of chemical weathering conditions at high latitudes in the Lower Turonian and Coniacian times. However, the covariation recorded between kaolinite and illite in Site U1512 during the Coniacian would invalidate this hypothesis for this stage. Watkins and Guerra (2020) observed, at the same time of the kaolinite increase in the Coniacian, an increase of cool taxa of calcareous nannofossils in Site U1512. Several studies have also highlighted a cooler interval at southern high latitudes in western Australia from the end of Turonian to the end of Coniacian followed by a slight increase of sea surface temperatures in the Early Santonian (Clarke and Jenkyns, 1999; Huber et al., 2002; Falzoni et al., 2016; O'Brien et al., 2017; Huber et al., 2018). The increase of kaolinites in a context of cooling does not seem consistent with their formation conditions, which confirms that short-term clay mineral variations are climate-independent.

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## 5.4 Eustatic influence on weathering conditions

The Turonian-Santonian interval is characterized by a sea-level 140 to 250 m higher than the current level (Haq, 2014). This has led to a decrease in continental surfaces and a relative decrease in topography. Detrital clay minerals also inform of drainage conditions on lands (Chamley, 1989). Smectites reflect poor drainage conditions and the preservation of some cations during weathering processes ($Ca^{2+}$, $Na^{+}$) while kaolinites indicate well-drained conditions. Previous studies on clay mineral assemblages performed on several Upper Cretaceous boreholes of the Atlantic Ocean and outcrops of Tethyan domain demonstrate a global domination of smectites in oceanic deposits (Brosse, 1982; Robert and Maillot, 1983; Chamley and Debrabant, 1984; Jacquin, 1987; Daoudi et al., 1989; Deconinck et al., 1989, 2005; Deconinck and Bernouilli, 1991; Debrabant et al., 1992; Thiry and Jacquin, 1993; El Albani et al., 1999). Although these minerals seem to reflect climatic conditions, the high proportions of smectites are interpreted as a marker of flat continents with poor drainage conditions (Chamley, 1979; Chamley et al., 1979; Robert et al., 1979; Enu, 1986; Daoudi and Deconinck, 1994).

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Offshore clay mineral sedimentation is also affected by differential settling process (Gibbs, 1977). Because of their different sizes, clay minerals are not deposited uniformly in a sedimentary basin. Coarser clays minerals such as illites and kaolinites are mainly deposited in proximal environments while smectites are transported over longer distances (Gibbs, 1977; Godet et al., 2008). As for many continental masses during Cretaceous, the low elevation of the Australian continent led to poor drainage conditions. Site U1512 is located at the basin depth below the continental slope. Despite the absence of a reliable paleographic map, the borehole appears to be quite far (probably few
hundreds of kilometers) from the Turonian shoreline. Clay mineral assemblages recorded in Site U1512 could had been affected by differential settling processes, accentuated by the high sea-level during the Late Cretaceous (Chamley, 1989). The poor drainage conditions result in the formation of smectites in soils while the differential settling, emphasized by the high sea-level, accentuates the sedimentation of these smectites in the distal environments. This process leads to the trapping of the coarser clay minerals (illites, kaolinites) on the continental shelf and the deposition of the finer clay minerals (smectites) in the basin which led to a possible over-representation of smectites in Turonian to Santonian deposits in Site U1512. The influence of differential settling process of clay minerals in Site U1512 is confirmed by the increase in coarser clay minerals such as illite and kaolinite during the Coniacian, a period characterized by a slight decrease in sea level (Haq, 2014).

The influence of sea-level variations on the grain size of bulk sediments is illustrated by the increase in Ti/Al and Zr/Al ratios from the late Turonian to the Santonian which indicates an increase of coarse elements. This increase reflects a higher depositional energy, also characterized by an increase in the K/Al ratio in the Santonian which would indicate a higher proportion of K-bearing minerals, like feldspars (orthoclase). The increase of kaolinite proportion in the Coniacian are coeval with the increase of Ti/Al and Zr/Al ratios. These variations would reflect an increase in depositional energy and terrestrial dominance (fluvial runoff) in the basin, associated with the decrease of sea-level and the distance of Site U1512 from the shoreline and characterized by higher content of heavy minerals and coarser clay minerals (kaolinites and illites).

In IODP Site U1512, clay mineral assemblages appear to be mainly related to eustatism and climate conditions. The dominant content of smectites, associated to the presence in low to moderate proportions of kaolinites and the presence in trace of illites reflect moderate hydrolysis processes under a warm/temperate and sub-humid climate but also poor drainage conditions. These conditions are supported by high sea surface temperatures observed at southern high latitudes during Turonian (Huber et al., 1995, 2002) and spore-pollens species that indicate prevailing wet and warm/temperate conditions in the southern part of the Australia during Turonian to Santonian times (Dettmann et al., 1992; Dettmann, 1994; Macphail, 2007; Wainman et al., 2020). However, the influence of eustatism by the process of differential sedimentation of clays complicates the paleoclimatic interpretation. Some clays, especially illites and kaolinites can be, by this process, underrepresented in the distal domain.

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## 6. CONCLUSION

Our study highlights that the Turonian-Santonian succession of the IODP U1512 borehole is mainly composed of terrigenous sediments, including clay minerals for the < 2 µm fraction. As many other Cretaceous DSDP/ODP/IODP sites, the clay minerals assemblages show few variations during
the studied interval and are dominated by R0-type illite/smectite mixed-layers (smectites), followed in small proportions by kaolinites and to a lesser extent illites. These minerals are mostly of detrital in origin. Sr and Nd isotopes, measured on the clay fraction, confirms the hypothesis advanced by some authors of a distant origin for sediments deposited in the Ceduna Sub-basin. Although a small part of smectites appears to be derived from volcanic ashes weathering, the clay minerals for this IODP site would dominantly resulted from the weathering of the Eastern Australian Highlands and uplifted Mesozoic sedimentary basins. The dominance of smectites, coupled to low proportions of illites and the absence of chlorite suggests long-term intermediate hydrolysis and poorly drained conditions on a relative flat continent under warm to temperate and sub-humid climate during a 10-Myr interval, in agreement with palynological data. However, the high sea-level may have influenced the clay mineral assemblages. The influence of eustatism, highlighted by the differential settling process of clay minerals, is supported by the coeval increase of illite and kaolinite proportions during the Coniacian sea-level lowstand. The increase of coarser deposits during this sea-level fall is also confirmed by the gradual increase of Ti/Al and Zr/Al that would reflect an increased influence of continental input in sedimentation.

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## 615 ACKNOWLEDGEMENTS

We thank Ludovic Bruneau (Biogéosciences, Université de Bourgogne) for his help in XRD analysis and Rémi Chassagnon (ICP, Université de Bourgogne) for TEM observation. We also thank the International Ocean Discovery Program (IODP) and the European Consortium for Ocean Research Drilling (ECORD), which enabled Laurent Riquier and Hans-Jürgen Brumsack to participate on Expedition 369, IODP-France and INSU-Tellus-Syster program for financial support for XRD analysis, isotopic analysis and SEM observations.

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## 623 DATA AVAILABILITY

Data are available on supplementary material.

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## 626 REFERENCES

Allen, C. M., Wooden, J. L., Chappell, B. W., 1997. Late Paleozoic crustal history of central coastal Queensland interpreted from geochemistry of Mesozoic plutons: The effects of continental rifting. Lithos, 42(1-2), 67-88. https://doi.org/10.1016/S0024-4937(97)00037-6.

Allen, C. M., 2000. Evolution of a post-batholith dike swarm in central coastal Queensland, Australia: arc-front to backarc. Lithos, 51(4), 331-349. https://doi.org/10.1016/S0024-4937(99)00068-7.

Arthur, M. A., Dean, W. E., Schlanger, S. O., 1985. Variations in the global carbon cycle during the Cretaceous related to climate, volcanism, and changes in atmospheric CO₂. The carbon cycle
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1172 Turner, S., Foden, J., Sandiford, M., Bruce, D., 1993a. Sm-Nd isotopic evidence for the provenance of sediments from the Adelaide Fold Belt and southeastern Australia with implications for episodic crustal addition. Geochimica et Cosmochimica Acta, 57(8), 1837-1856.
1175 https://doi.org/10.1016/0016-7037(93)90116-E.

1176 Uysal, I. T., Zhao, J. X., Golding, S. D., Lawrence, M. G., Glikson, M., Collerson, K. D., 2007. Sm-Nd dating and rare-earth element tracing of calcite: implications for fluid-flow events in the Bowen Basin, Australia. Chemical Geology, 238(1-2), 63-71.
1179 https://doi.org/10.1016/j.chemgeo.2006.10.014.

1180 Veevers, J. J., Powell, C. M., Roots, S. R., 1991. Review of seafloor spreading around Australia. I. Synthesis of the patterns of spreading. Australian journal of earth sciences, 38(4), 373-389.
1182 https://doi.org/10.1080/08120099108727979.

1183 Veevers, J. J., 2000. Change of tectono-stratigraphic regime in the Australian plate during the 99 Ma (mid-Cretaceous) and 43 Ma (mid-Eocene) swerves of the Pacific. Geology, 28(1), 47-50.
1185 https://doi.org/10.1130/0091-7613(2000)28%3C47:COTRIT%3E2.0.CO;2.

1186 Veevers, J. J., 2006. Updated Gondwana (Permian–Cretaceous) earth history of Australia. Gondwana Research, 9(3), 231-260. https://doi.org/10.1016/j.gr.2005.11.005.

1188 Velde, B., 1995. Origin and mineralogy of clays. Clays and the environment. Springer-Verlag, Berlin, 356 pp.

1190 Wade, B. P., Hand, M., Barovich, K. M., 2005. Nd isotopic and geochemical constraints on provenance of sedimentary rocks in the eastern Officer Basin, Australia: implications for the duration of the intracratonic Petermann Orogeny. Journal of the Geological Society, 162(3), 513-530. https://doi.org/10.1144/0016-764904-001

1194 Wade, B. P., Barovich, K. M., Hand, M., Scrimgeour, I. R., Close, D. F., 2006. Evidence for early Mesoproterozoic arc magmatism in the Musgrave Block, central Australia: implications for Proterozoic crustal growth and tectonic reconstructions of Australia. The Journal of geology, 114(1), 43-63. https://doi.org/10.1086/498099.

1198 Wade, C. E., Reid, A. J., Wingate, M. T., Jagodzinski, E. A., Barovich, K., 2012. Geochemistry and geochronology of the c. 1585 Ma Benagerie Volcanic Suite, southern Australia: Relationship to the Gawler Range Volcanics and implications for the petrogenesis of a Mesoproterozoic silicic large igneous province. Precambrian Research, 206, 17-35.
1202 https://doi.org/10.1016/j.precamres.2012.02.020.

1203 Waight, T. E., Dean, A. A., Maas, R., Nicholls, I. A., 2000. Sr and Nd isotopic investigations towards the origin of feldspar megacrysts in microgranular enclaves in two I-type plutons of the Lachlan Fold Belt, southeast Australia. Australian Journal of Earth Sciences, 47(6), 1105-1112.
1206 https://doi.org/10.1046/j.1440-0952.2000.00831.x.

1207 Wainman, C. C., Tagliaro, G., Jones, M. M., Charles, A. J., Hall, T., White, L. T., Bogus, K. A., Wolfgring, E., O'Connor, L. K., McCabe, P. J., Holford, S. P., 2020. The sedimentological
1209 evolution and petroleum potential of a very thick Upper Cretaceous marine mudstone
1210 succession from the southern high latitudes—a case study from the Bight Basin, Australia.
1211 Marine and Petroleum Geology, 118, 104441. https://doi.org/10.1016/j.marpetgeo.2020.104441.

1212 Walter, M. R., Veevers, J. J., Calver, C. R., Gorjan, P., Hill, A. C., 2000. Dating the 840–544 Ma
1213 Neoproterozoic interval by isotopes of strontium, carbon, and sulfur in seawater, and some
1214 interpretative models. Precambrian Research, 100(1-3), 371-433. https://doi.org/10.1016/S0301-
1215 9268(99)00082-0.

1216 Waschbusch, P., Beaumont, C., Korsch, R. J., 1999. Geodynamic modelling of aspects of the New
1217 England Orogen and adjacent Bowen, Gunnedah and Surat basins. New England orogen:
1218 regional geology, tectonics and metallogenesis: Armidale, Australia, Earth Sciences, University
1219 of New England, 203-210.

1220 Watkins, K. P., Hickman, A. H., 1990. Geological evolution and mineralization of the Murchison
1221 Province, Western Australia (Vol. 1). Geological Survey of Western Australia.

1222 Watkins, D. K., Guerra, R. M., 2020. Calcareous nannofossils from the Great Australian Bight (IODP
1223 Site U1512) as a record of the peak and early decline of the mid-Cretaceous super greenhouse.
1224 Marine Micropaleontology, 158, 101893. https://doi.org/10.1016/j.marmicro.2020.101893

1225 Weaver, C. E., 1989. Clays, muds, and shales. Elsevier, Amsterdam. Developments in Sedimentology,
1226 44, 818 pp.

1227 Whelan, J., Hergt, J., Woodhead, J., 2007. Granite–greenstone connection in western Victoria: an
1228 example from the Bushy Creek Igneous Complex. Australian Journal of Earth Sciences, 54(7),
1229 975-990. https://doi.org/10.1080/08120090701488305.

1230 White, L. T., Gibson, G. M., Lister, G. S., 2013. A reassessment of paleogeographic reconstructions of
1231 eastern Gondwana: Bringing geology back into the equation. Gondwana Research, 24(3-4),
1232 984-998. https://doi.org/10.1016/j.gr.2013.06.009.

1233 Wiese, F., Voigt, S., 2002. Late Turonian (Cretaceous) climate cooling in Europe: faunal response and
1234 possible causes. Geobios, 35(1), 65-77. https://doi.org/10.1016/S0016-6995(02)00010-4.

1235 Windrim, D. P., McCulloch, M. T., 1986. Nd and Sr isotopic systematics of central Australian
1236 granulites: chronology of crustal development and constraints on the evolution of lower
1237 continental crust. Contributions to Mineralogy and Petrology, 94(3), 289-303.
1238 https://doi.org/10.1007/BF00371438.

1239 Worden, R. H., Morad, S., 1999. Clay minerals in sandstones: controls on formation, distribution and
1240 evolution. Clay mineral cements in sandstones, 1-41. https://doi.org/10.1002/9781444304336.ch1.
1241 Zhao, J. X., McCulloch, M. T., Bennett, V. C., 1992. Sm-Nd and U-Pb zircon isotopic constraints on the provenance of sediments from the
1242 Amadeus Basin, central Australia: Evidence for REE fractionation. Geochimica et
1243 Cosmochimica Acta, 56(3), 921-940. https://doi.org/10.1016/0016-7037(92)90037-J.
1244

1245 Zhao, J. X., McCulloch, M. T., 1993. Sm-Nd mineral isochron ages of Late Proterozoic dyke swarms
1246 in Australia: evidence for two distinctive events of mafic magmatism and crustal extension.
1247 Chemical geology, 109(1-4), 341-354. https://doi.org/10.1016/0009-2541(93)90079-X.

1248 Zhao, J. X., McCulloch, M. T., 1995. Geochemical and Nd isotopic systematics of granites from the
1249 Arunta Inlier, central Australia: implications for Proterozoic crustal evolution. Precambrian
1250 Research, 71(1-4), 265-299. https://doi.org/10.1016/0301-9268(94)00065-Y.

1252
1253 TABLE
1254
<table><caption>Table 1: r-coefficient between elements and aluminium</caption><thead><tr><th>Element</th><th>As</th><th>Ba</th><th>Ca</th><th>Co</th><th>Cr</th><th>Cu</th><th>Fe</th><th>Ga</th><th>K</th><th>Mg</th></tr></thead><tbody><tr><td>r-coefficient</td><td>0.58</td><td>-0.03</td><td>0.01</td><td>0.86</td><td>0.41</td><td>0.31</td><td>0.78</td><td>0.97</td><td>0.67</td><td>0.72</td></tr><tr><td>p-value</td><td>&lt;0.01</td><td>0.77</td><td>0.9</td><td>&lt;0.01</td><td>&lt;0.01</td><td>&lt;0.01</td><td>&lt;0.01</td><td>&lt;0.01</td><td>&lt;0.01</td><td>&lt;0.01</td></tr><tr><td>Element</td><td>Mn</td><td>Mo</td><td>Na</td><td>Nb</td><td>Ni</td><td>P</td><td>Pb</td><td>Rb</td><td>Si</td><td>Sr</td></tr><tr><td>r-coefficient</td><td>0.45</td><td>-0.04</td><td>0.5</td><td>0.66</td><td>0.82</td><td>0.5</td><td>0.93</td><td>0.75</td><td>-0.92</td><td>0.38</td></tr><tr><td>p-value</td><td>&lt;0.01</td><td>0.71</td><td>&lt;0.01</td><td>&lt;0.01</td><td>&lt;0.01</td><td>&lt;0.01</td><td>&lt;0.01</td><td>&lt;0.01</td><td>&lt;0.01</td><td>&lt;0.01</td></tr><tr><td>Element</td><td>Th</td><td>Ti</td><td>U</td><td>V</td><td>Y</td><td>Zn</td><td>Zr</td><td></td><td></td><td></td></tr><tr><td>r-coefficient</td><td>0.5</td><td>0.86</td><td>0.25</td><td>0.82</td><td>0.78</td><td>0.86</td><td>0.26</td><td></td><td></td><td></td></tr><tr><td>p-value</td><td>&lt;0.01</td><td>&lt;0.01</td><td>0.04</td><td>&lt;0.01</td><td>&lt;0.01</td><td>&lt;0.01</td><td>0.03</td><td></td><td></td><td></td></tr></tbody></table>

1255

1256

Table 1: r-coefficient between elements and aluminium.

1257
<figure><img src="image_3.png" /></figure>

1259
1260

1261 Figure 1: A) Location of the Ceduna Sub-basin, Site U1512 and industry wells including Jerboa-1 and Gnarlyknots 1/1A in the Bight Basin (modified from Bradshaw et al., 2003; Totterdell et al., 2008 and Wainman et al., 2020). B) Turonian (~ 90 Ma) palaeogeographical map of the Australo-Antarctic Gulf (modified from MacLeod et al., 2020). Abbreviations: 1 – Adelaide Fold Belt, 2 – Curnamona Province, 3 – Stuart Shelf.
<figure><img src="image_4.png" /></figure>

1266

1267 Figure 2: Lithology and biostratigraphy of the IODP Site U1512 and position of studied samples.

1268 Lithology modified from Huber et al. (2019), nannofossils zonation from Huber et al. (2019) and

1269 Watkins and Guerra (2020). Abbreviations: Calcareous Nan. Zonation – Calcareous Nannofossils

1270 Zonation, Santo. – Santonian.
<figure><img src="image_5.png" /></figure>

<figure><img src="image_6.png" /><figcaption>Figure 3: Representative XRD diagrams of the clay fraction of Turonian to Santonian sediments of IODP Site U1512. d values are given in Å. Abbreviations: Clino. – Clinoptilolites, I-S R0 – R0-type illite/smectite mixed layers, Kaol. – Kaolinite, Qz. – Quartz.</figcaption></figure>

1271

1272

1273

1274

1275

1276

1277

1278

1279
<figure><img src="image_7.png" /><figcaption>Figure 4: Stratigraphic distribution of clay minerals (illite, smectite and kaolinite) proportions, other minerals observed in the clay-fraction (< 2µm) and Sr and Nd isotopes measured in the clay fraction for the Turonian-Santonian sedimentary succession of the IODP Site U1512. Abbreviations: Clinoptilo. – Clinoptilolites, Nanno. Zone – Nannofossils Zone, Santo. – Santonian.</figcaption></figure>

1280
1281
1282
1283
1284
1285
1286
1287
1288
1289
1290
1291
<figure><img src="image_8.png" /><figcaption>Figure 5: Scanning electron microscopy (A to D) and Transmission electron microscopy (E and F) images of some representative samples of the Site U1512. A) Sample 14R-3W 80-82 (128 m depth, Coniacian) reveals detrital kaolinites. B) Sample 20R-5W 88-90 (188 m, Coniacian) shows detrital clay minerals. C) Sample 68R-4W 82-84 (648 m, Lower Turonian) shows detrital clays and micas. D) Sample 20R-5W 88-90 (188 m, Coniacian) shows opal-CT lepispheres and clinoptilolites filling the porosity. E) Sample 33R-6W 10-12 (313.6 m, Middle/Upper Turonian) reveals authigenic opal-CT and clinoptilolites. F) Sample 36R-5W 100-102 (342.35 m, Middle/Upper Turonian) shows authigenic opal-CT filling gastropod dissolved shell. G & H) Sample 8R-4W 59-61 (72 m, Santonian) shows detrital fleecy smectites (I-S R0), illites characterized by sharp outlines and opal-CT. This sample also present probably volcanogenic folded smectite. I) Sample 35R-2W 15-17 (328 m, Middle/Upper Turonian) reveals fleecy smectites and kaolinites.</figcaption></figure>

1292

1293

1294

1295

1296

1297

1298

1299

1300

1301

1302

1303

1304

1305
<figure><img src="image_9.png" /><figcaption>Figure 6: Ternary plot of the relative concentration of Al₂O₃ (multiplied by 5), SiO₂ and CaO (multiplied by 2). Composition of PAAS and UCC (Taylor and McLennan, 1985) and clay minerals (Weaver, 1989) are reported.</figcaption></figure>

1306

1307

1308

1309

1310
<figure><img src="image_10.png" /></figure>

Figure 7: Stratigraphic distribution of Ti/Al, K/Al and Zr/Al ratios and CIA for the Turonian - Santonian sedimentary succession of the IODP Site U1512. Abbreviations: Nanno. Zone – Nannofossils Zone, Santo. – Santonian.

1311

1312

1313

1314

1315
<figure><img src="image_11.png" /></figure>

Figure 8: Stratigraphic distribution of TOC, Si/Al, P/Al and Ba/Al ratios for the Turonian- Santonian sedimentary succession of the IODP Site U1512. Abbreviations: Nanno. Zone – Nannofossils Zone, Santo. – Santonian.

1316

1317

1318

1319
<figure><img src="image_12.png" /></figure>

Figure 9: Nd and Sr isotopes composition of clay fraction from Site U1512 compared to whole-rock isotopes composition of potential eroded landmasses (eastern landmasses data from McCulloch and Chappell, 1982; Gray, 1984, 1990; Nelson et al., 1984; Hensel et al., 1985; Eberz et al., 1989; Creaser and Gray, 1992; Turner et al., 1993; Creaser, 1995; Johnson and McCulloch, 1995; Allen et al., 1997; Bryant et al., 1997; Keay et al., 1997; Maas et al., 1997; Allen, 2000; Barovich and Foden, 2000; Waight et al., 2000; Foden et al., 2002; Elburg et al., 2003, 2012; Healy et al., 2004; Swain et al., 2005, 2008; Payne et al., 2006, 2010; Budd and Skirrow, 2007; Hergt et al., 2007; Skirrow et al., 2007; Uysal et al., 2007; Whelan et al., 2007; Barovich and Hand, 2008; Raveggi et al., 2008; Haines et al., 2009; Kemp et al., 2009; Fraser et al., 2010; Howard et al., 2011; Wade et al., 2012; Sun et al., 2020; northern landmasses data from Iyer et al., 1976; Black and McCulloch, 1984; Windrim and McCulloch, 1986; Nelson et al., 1989; Zhao et al., 1992; Zhao and McCulloch, 1993, 1995; Foden et al., 1995; Sun et al., 1995; Glikson et al., 1996; Barovich and Foden, 2000; Walter et al., 2000; Edgoose et al., 2004; Hoatson et al., 2005; Wade et al., 2005, 2006; Smithies et al., 2010; Love, 2017, western landmasses data from McCulloch and Compson, 1981; Fletcher et al., 1983, 1984, 1985, 1991; Claoué-Long et al., 1984; Chauvel et al., 1985; Watkins and Hickman, 1990; Maas et al., 1991; Nutman et al., 1993; Turner et al., 1993; Nelson et al., 1995; Champion and Sheraton, 1997; Bateman et al., 2001; Fletcher and McNaughton, 2002; Morris and Kirkland, 2014). White dots correspond to

1320

1321

1322

1323

1324

1325

1326

1327

1328

1329

1330

1331

1332

1333

1334

1335

1336

1337
1338 the mean, thick black bars to the median, the boxes to 50% of the data (1st and 3rd quartiles) and the
1339 side bars to the total data.

1340

1341

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

1342

1343 Figure 10: Possible catchment area and sedimentary routage deduced from εNd(o) values measured on
1344 the clay fraction of Site U1512 compared to εNd(o) of Australian geological units. Catchment area
1345 modified from Veevers (1991), Lloyd et al. (2016) and Foley et al. (2021). Map modified from
1346 Blewett et al. (2012). Abbreviations 1 – Curnamona Province, 2 – Gawler Craton, 3 – Stuart Shelf.