# Environmental changes during the onset of the Late Pliensbachian Event (Early Jurassic) in the Cardigan Bay Basin, Wales.

Teuntje P. Hollaar, Stephen P. Hesselbo, Jean-François Deconinck, Magret Damaschke, Clemens V. Ullmann, Mengjie Jiang, Claire M. Belcher

► To cite this version:

Teuntje P. Hollaar, Stephen P. Hesselbo, Jean-François Deconinck, Magret Damaschke, Clemens V. Ullmann, et al.. Environmental changes during the onset of the Late Pliensbachian Event (Early Jurassic) in the Cardigan Bay Basin, Wales.. Climate of the Past, 2023, 19 (5), pp.979-997. 10.5194/cp-19-979-2023. hal-04100386

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1 Environmental changes during the onset of the Late Pliensbachian

2 Event (Early Jurassic) in the Mochras Borehole, Cardigan Bay Basin,

3 NW Wales.

4

5 Teuntje P. Hollaar¹,², Stephen P. Hesselbo²,³, Jean-François Deconinck⁴, Magret Damaschke⁵,

6 Clemens V. Ullmann²,³, Mengjie Jiang², Claire M. Belcher¹.

7 1 WildFIRE Lab, Global Systems Institute, University of Exeter, Exeter, EX4 4PS, UK

8 2 Camborne School of Mines, Department of Earth and Environmental Science, University of Exeter, Penryn

9 Campus, Penryn, TR10 9FE, UK

10 3 Environment and Sustainability Institute, University of Exeter, Penryn Campus, Penryn, TR10 9FE, UK

11 4 Biogéosciences, UMR 6282 CNRS, Université de Bourgogne/Franche-Comté, 21000 Dijon, France

12 5 Core Scanning Facility, British Geological Survey, Keyworth, NG12 5GG, UK

13 Correspondence to: Teuntje P. Hollaar (t.p.hollaar@exeter.ac.uk)

14

15 **Abstract.** The Late Pliensbachian Event (LPE), in the Early Jurassic, is associated with a perturbation in the
16 global carbon cycle (positive carbon isotope excursion (CIE) of ~ 2 ‰), cooling of ~5°C, and the deposition of
17 widespread regressive facies. Cooling during the Late Pliensbachian has been linked to enhanced organic matter
18 burial and/or disruption of thermohaline ocean circulation due to North Sea doming. Orbital forcing had a strong
19 influence on the Pliensbachian environments and recent studies show that the terrestrial realm and the marine
20 realm in and around the Cardigan Bay Basin were strongly influenced by orbital climate forcing. In the present
21 study we build on the previously published data for long eccentricity cycle E459 ± 1 and extend the
22 palaeoenvironmental record to include E458 ± 1. We explore the environmental and depositional changes on
23 orbital time scales for the Mochras core during the onset of the LPE. Clay mineralogy, XRF elemental analysis,
24 isotope ratio mass spectrometry, and palynology are combined to resolve systematic changes in erosion,
25 weathering, fire, grain size and riverine influx. Our results indicate distinctively different environments before
26 and after the onset of the LPE positive CIE, and show increased physical erosion relative to chemical
27 weathering. We also identify 5 swings in the climate, in tandem with the 405 kyr eccentricity minima and
28 maxima. Eccentricity maxima are linked to precessionally repeated occurrences of a semi-arid, monsoonal
29 climate with high fire activity and relatively coarser fraction of terrestrial runoff. In contrast, 405 kyr minima in
30 the Mochras core are linked to a more persistent, annually wet climate, low fire activity, and relatively finer
31 grained deposits across multiple precession cycles. The onset of the LPE +ve CIE did not impact the expression
32 of the 405 kyr in the proxy records, however, during the second pulse of lighter carbon (¹²C) enrichment, the
33 clay minerals record a change from dominant chemical weathering to dominant physical erosion.

34

35 **1.1 Introduction**

36 The Early Jurassic is a period marked by large climatic fluctuations and associated carbon-isotope excursions
37 (CIE's) in an overall warm and high pCO₂ world (McElwain et al., 2005; Korte & Hesselbo, 2011). A series of
38 small and medium sized CIE's have recently been documented for the Sinemurian and Pliensbachian, which
39 have mainly been recorded in European records (Korte & Hesselbo, 2011; Franceschi et al., 2014; Korte et al.,
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2015; Price et al., 2016; Hesselbo et al., 2020a; Storm et al., 2020; Silva et al., 2021; Cifer et al., 2022) and recently at the NW end of the Tethys Ocean in Morocco (Mercuzot et al., 2020) and in North America (De Lena et al., 2019). Notable is the pronounced positive CIE in the Late Pliensbachian, which has been called the Late Pliensbachian Event (LPE) and is linked to climatic cooling (Hesselbo & Korte, 2011; Korte et al., 2015) and a supra-regional/global sea level low stand (Hallam, 1981; de Graciansky et al., 1998; Hesselbo & Jenkyns, 1998; Hesselbo, 2008). The LPE has been recognized by a positive shift in benthic marine oxygen-isotopes (~1.5–2 per mil) (Bailey et al., 2003; Rosales et al., 2004, 2006; Suan et al., 2010; Dera et al., 2011a; Korte & Hesselbo, 2011; Gómez et al., 2016; Alberti et al., 2019, 2021), coeval with a positive shift in marine and terrestrial carbon isotopes (~2 per mil) (Jenkyns & Clayton, 1986; McArthur et al., 2000; Morettini et al., 2002; Quesada et al., 2005; Rosales et al., 2006; Suan et al., 2010; Korte & Hesselbo, 2011; Silva et al., 2011; Gómez et al., 2016; De Lena et al., 2019).

A cooler Late Pliensbachian climate has been suggested based on low pCO₂ values inferred by leaf stomatal index data from eastern Australia (Steinthorsdottir & Vajda, 2015), the presence of glendonites in northern Siberia (Kaplan, 1978; Price, 1999; Rogov & Zakharov, 2010), vegetation shifts from a diverse flora of different plant groups to one mainly dominated by bryophytes in Siberia (Ilyina, 1985; Zakharov et al., 2006), and possible ice rafted debris found in Siberia (Price, 1999; Suan et al., 2011). Whilst the presence of ice sheets is strongly debated, a general cooling period (~5°C lower; Korte et al., 2015; Gómez et al., 2016) is evident from several temperature reconstructions of NW Europe. A cooling is hypothesized via enhanced carbon burial in the marine sediments, leading to lower pCO₂ values and initiating cooler climatic conditions (Jenkyns & Clayton, 1986; Suan et al., 2010; Silva et al., 2011; Storm et al., 2020). Direct evidence of large-scale carbon burial in Upper Pliensbachian marine deposits has not yet been documented (Silva et al., 2021). Alternatively, cooling has been suggested to be caused by regional tectonic updoming of the North Sea region, which would have disrupted the ocean circulation in the Laurasian Seaway, reducing poleward heat transport from the tropics (Korte et al., 2015). Disruption of the ocean circulation between the western Tethys and the Boreal realm is supported by marine migration patterns (Schweigert, 2005; Zakharov et al., 2006; Bourillot et al., 2008; Nikitenko, 2008; Dera et al., 2011b; van de Schootbrugge et al., 2019) and numerical models (Bjerrum et al., 2001; Dera & Donnadieu, 2012; Ruvalcaba Baroni et al. 2018); however, the direction of the flows remain debated.

An additional factor at this time is that a strong orbital control exists on the sedimentary successions in the Pliensbachian (Weedon & Jenkyns, 1990; Ruhl et al., 2016; Hinnov et al., 2018; Storm et al., 2020; Hollaar et al., 2021). Previous studies have indicated that sea level changes, possibly coupled to glacio-eustatic rise and fall, occurred during the LPE on a 100 kyr (short eccentricity) time scale (Korte & Hesselbo, 2011). A high-resolution record of charcoal, clay mineralogy, bulk-organic carbon-isotopes, TOC and CaCO₃ encompassing approximately one 405 kyr cycle from the Llanbedr (Mochras Farm) borehole, Cardigan Bay Basin, NW Wales, UK suggested that the long-eccentricity orbital cycle had a significant effect on background climatic and environmental change at this time, particularly affecting the hydrological regime of the region (Hollaar et al., 2021). This previous research focussed on orbital forcing of environmental change for a time lacking any large excursion in δ¹³Cₗ, and so unaffected by perturbations to the global carbon cycle. Here, we expand on the record of Hollaar et al. (2021) to cover two long eccentricity cycles (which we identify as cycle E459 ± 1 and
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79 E458 ± 1 of Laskar *et al.* 2011), where the final parts of E458 and the start of E457 are interrupted by onset of the Late Pliensbachian Event. We find that the long eccentricity forcing continued to dictate the precise timing of major environmental changes in the Cardigan Bay Basin, including the initial step of the positive carbon isotope excursion.

83

84 **1.2 Material**

85 **1.2.1 Palaeo-location and setting**

86 Associated with the break-up of Pangea, connections between oceans via epicontinental seaways were
87 established during the Early Jurassic, such as the Hispanic Corridor, which connected the north-western Tethys
88 and the eastern Panthalassa, and the Viking Corridor which linked the north-western Tethys Ocean to the Boreal
89 Sea (Sellwood & Jenkyns, 1975; Smith et al., 1983; Bjerrum et al., 2001; Damborenea et al., 2012). The linking
90 passage of the NW Tethys Ocean and the Boreal Sea (south of the Viking Corridor) is the palaeogeographical
91 location of the Llanbedr (Mochras Farm) borehole, Cardigan Bay Basin, NW Wales, UK (Fig. 1) – referred to
92 hereafter as Mochras. Due to the location of the Mochras succession during the Late Pliensbachian, it was
93 subject to both polar and equatorial influences allowing the study of variations in the circulation in the N-S
94 Laurasian Seaway (including the Viking Corridor) prior to and across the LPE. Mochras was located at a mid-
95 palaeolatitude of ~35° N (Torsvik & Cocks, 2017).

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<figure><img src="image_9.png" /><figcaption>Figure 1: Palaeolocation of the Mochras borehole in the context of potential North Sea doming. Figure</figcaption></figure>

reprinted and adapted from Korte *et al.* (2015), which is open access

(https://creativecommons.org/licenses/by/4.0/). The Mochras borehole was located at a paleolatitude of ~35° N in the Cardigan Bay Basin (Torsvik & Cocks, 2017). Circulation in the Tethys Ocean and between there and the Boreal region influenced the depositional environment of the Mochras core (Pieńkowski *et al.*, 2021). Late Pliensbachian uplift of the North Sea dome potentially led to occlusion of the Viking Corridor and disrupted circulation in the seaway (Korte *et al.*, 2015).

The depositional environment of Mochras is likely characterized by a rift setting, which is reflected by the relatively open and deep marine facies and the evidence for below storm wave-base and contourite deposition (Pieńkowski *et al.*, 2021), but always with a strong terrestrial influence (van de Schootbrugge *et al.*, 2005; Riding *et al.*, 2013) from the nearby landmasses (Dobson & Whittington, 1987). The Cardigan Bay Basin fill was downthrown against the Early Paleozoic Welsh Massif on the SE side by a major normal fault system, probably comprising the Bala, Mochras and Tonfanau faults at the eastern and south-eastern margins of the basin in Late Paleozoic–Early Mesozoic time (Woodland, 1971; Tappin *et al.*, 1994). The main source of
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113 detrital material is understood to be the Welsh Massif, followed by the Irish Massif (Deconinck et al., 2019).

114 Other Variscan massifs that could have influenced the provenance are the London-Brabant Massif to the south

115 east, and Cornubia to the south (van de Schootbrugge et al., 2005), depending on the marine circulation and

116 sediment transport at the time.

117 **1.2.2 Core location and material**

118 The Llanbedr (Mochras Farm) Borehole was drilled onshore in the Cardigan Bay Basin (52 48' 32"N, 4 08'

119 44"W) in 1967–1969, North Wales (Woodland, 1971; Hesselbo et al., 2013). The borehole recovered a 1300 m

120 thick Early Jurassic sequence (601.83 – 1906.78 metres below surface (mbs)), yielding the most complete and

121 extended Early Jurassic succession in the UK, being double the thickness of same age strata in other UK cores

122 and outcrops (Hesselbo et al., 2013; Ruhl et al., 2016). The Lower Jurassic is biostratigraphically complete at

123 the zonal level (Ivimey-Cook, 1971; Copestake & Johnson, 2014), with the top truncated and unconformably

124 overlain by Cenozoic strata (Woodland, 1971; Dobson & Whittington, 1987; Tappin et al., 1994; Hesselbo et

125 al., 2013). The lithology is dominated by argillaceous sediments, with alternating muddy limestone, marl and

126 mudstone (Woodland, 1971; Sellwood & Jenkyns, 1975).

127 **1.2.3 Pliensbachian**

128 The Pliensbachian Stage in the Mochras borehole occurs between ~865 to ~1250 mbs, with the Margaritatus

129 Zone between ~1013 and 909 mbs (Page in Copestake & Johnson, 2014). The Pliensbachian interval comprises

130 alternations of mudstone (with a relatively moderate total organic carbon [TOC]) and organic poor limestones,

131 with a pronounced cyclicity at ~1 ± 0.5 m wavelength (Ruhl et al., 2016). The Upper Pliensbachian contains

132 intervals that are silty and locally sandy, whilst levels of relative organic enrichment also occur through the

133 Pliensbachian (Ruhl et al., 2016). Overall, the Upper Pliensbachian is relatively rich in carbonate (Ruhl et al.,

134 2016; Ullmann et al., 2022).

135

136 **1.3 Methods**

137 For this study, samples were taken at a ~30 cm resolution from slabbed core from 918–934 mbs for XRD and

138 mass spectrometry, as well as palynofacies and microcharcoal analysis. These new samples complement a set at

139 10 cm resolution from 951–934 mbs (results published in Hollaar et al., 2021).

140 **1.3.1 TOC, $CaCO_3$ and bulk organic carbon isotope mass spectrometry**

141 TOC and $\delta^{13}C_{org}$ were measured to track the changes in the total organic fraction and the bulk organic carbon

142 isotope ratios simultaneously with the other palaeoenvironmental proxy data.

143 Powdered bulk-rock samples (~ 2 g) were decarbonated in 50 ml of 3.3% HCl. After this, the samples were

144 transferred to a hot bath of 79 °C for 1 h to remove siderite and dolomite. Subsequently, the samples were

145 centrifuged and the liquid decanted. The samples were rinsed repeatedly with distilled water to reach neutral pH.

146 After this, the samples were oven-dried at 40 °C, re-powdered and weighed into tin capsules for mass

147 spectrometry with the Sercon Integra 2 stable isotope analyser at the University of Exeter Environmental &

148 Sustainability Institute (ESI), stable isotope facility on the Penryn Campus, Cornwall. Samples were run

149 alongside in house reference material (bovine liver; $\delta^{13}C$ -28.61 and Alanine; $\delta^{13}C$ -19.62) which was used to

150 correct for instrument drift and to determine the $\delta^{13}C$ values of the samples. $\delta^{13}C_{org}$ values reported relative to V-
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151 PDB following a within-run laboratory standard calibration. Total organic carbon was determined using the CO₂
152 beam area relative to the bovine liver standard (%C 47.24). Replicate analysis of the in-house standards gave a
153 precision of ± <0.1 ‰ (2 SD).

154 The carbonate content was measured by the dry weight sample loss before and after decarbonation. The %C
155 content derived from the mass spectrometer was corrected for carbonate loss to derive TOC.

1.3.2 X-Ray Diffraction (XRD) to determine clay mineralogy

157 Clay mineral analysis was performed to gain insight into the relative importance of physical erosion versus
158 chemical weathering and related changes in the hydrological cycle.

159 About 2–3 g of gently powdered bulk-rock was decarbonated with a 0.2 M HCl solution and the clay sized
160 fraction (< 2 µm) extracted and oriented on glass slides for X-ray diffraction analysis (XRD) using a Bruker D4
161 Endeavour diffractometer (Bruker, Billerica, MA, USA) with Cu Kα radiations, LynxEye detector and Ni filter
162 under 40 kV voltage and 25 mA intensity (Biogéosciences Laboratory, Université Bourgogne/Franche-Comté,
163 Dijon). Following Moore & Reynolds (1997), the clay phases were discriminated in three runs per sample: (1)
164 air-drying at room temperature; (2) ethylene-glycol solvation during 24 h under vacuum; (3) heating at 490 °C
165 during 2 h.

166 Identification of the clay minerals was based on their main diffraction peaks and by comparing the three
167 diffractograms obtained. The proportion of each clay mineral on glycolated diffractograms was measured using
168 the MACDIFF 4.2.5. software (Petschick, 2000). Identification of the clay minerals follows the methods in
169 Deconinck et al. (2019) and Moore & Reynolds (1997).

1.3.3 Palynofacies and microcharcoal

171 Palynofacies were examined to explore shifts in the terrestrial versus marine origin of the particulate organic
172 matter. Each ~ 20 g bulk rock sample was split into 0.5 cm³ fragments, minimizing breakage of charcoal and
173 other particles, to optimize the surface area for extraction of organic matter using a palynological acid
174 maceration technique. The samples were first treated with cold hydrochloric acid (10% and 37% HCl) to remove
175 carbonates. Following, hydrofluoric acid (40% HF) was added to the samples to remove silicates. Carbonate
176 precipitation was prevented, by adding cold concentrated HCl (37%) after 48 h. The samples were neutralized
177 via multiple DI water dilution-settling-decanting cycles, after which 5 droplets of the mixed residue were taken
178 for the analysis of palynofacies prior to sieving. The remaining residue was sieved through a 125 µm and 10 µm
179 mesh to extract the microcharcoal fraction.

180 A known quantity (125 µl) of the 10–125 µm sieved residue was mounted onto a palynological slide using
glycerine jelly. This fraction, containing the microscopic charcoal, was analysed and the charcoal particles
counted using an Olympus (BX53) transmitted light microscope (40x10 magnification). For each palynological
slide four transects (two transects in the middle and one on the left and right side of the coverslip) were followed
and the number of charcoal particles determined. Charcoal particles were identified with the following criteria:
opaque and black, often elongated lath-like shape with sharp edges, original anatomy preserved, brittle
appearance with a lustrous shine (Scott, 2010). These data were then scaled up to the known quantity of the
sample according the method of Belcher *et al.* (2005).
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188 Palynofacies were grouped broadly according to Oboh-Ikuenobe *et al.* (2005): sporomorphs, fungal remains,
189 freshwater algae, marine palynomorphs, structured phytoclasts, unstructured phytoclasts, black debris,
190 amorphous organic matter (AOM), and charcoal (further described in Hollaar *et al.*, 2021). The palynofacies
191 were quantified on a palynological slide using the optical light microscope (40x10 magnification) and counting
192 a minimum of 300 particles per slide. Because the samples are AOM-dominated, counting was continued until a
193 minimum of 100 non-AOM particles were observed. We used the percentage of terrestrial phytoclasts, which
194 includes sporomorphs, structured and unstructured phytoclasts, to examine changes in terrestrial organic particle
195 content.

196
197 **1.3.4 X-Ray Fluorescence (XRF) to determine detrital elements**

198 Detrital elemental ratios were examined to analyse changes in relative terrestrial influx and the type of material
199 transported from the land to the marine realm. The slabbed archive halves of the Mochras borehole were
200 scanned via automated X-ray fluorescence (XRF) at a 1 cm resolution for the interval 951 – 918 mbs, with the
201 ITRAX MC at the British Geological Survey Core Scanning Facility (CSF), Keyworth, UK (Damaschke et al.,
202 2021). The measurement window was 10 s and long-term drift in the measurement values was counteracted by
203 regular internal calibration with a glass reference (NIST-610). Duplicate measurements were taken every 5 m
204 for a 50 cm interval to additionally verify the measured results.

205 **1.3.5 Statistical analysis**

206 Principal component analysis (PCA) was performed in the software PAST on the normalized dataset including
207 microcharcoal, TOC, CaCO3, δ13Corg, S/I, K/I, primary clay mineralogy, Si/Al, Zr/Rb. The samples before the
208 +ve CIE (951.0–930.4 mbs) and the samples after the +ve CIE (930.3–918.0 mbs) are grouped to examine a
209 potential difference in the sedimentary composition before and after the +ve CIE.
210 A Pearson’s correlation was executed in Matlab R2017b. The *p* value tests the hypothesis of no correlation
211 against the alternative hypothesis of a positive or negative correlation (significance level at *p* ≤ 0.05).

212
213 **1.4 Results**

214
215 **1.4.1 TOC, CaCO3 and bulk organic carbon isotope ratio mass spectrometry**

216 Alternating TOC-enhanced and Ca-rich lithological couplets occur on a metre scale through the studied interval (r = -0.64, p = 0.001). TOC content fluctuates in the range 0.17–1.72 wt% (mean 0.8 wt%) and the highest fluctuations of TOC content are found from 939–930 mbs. The CaCO3 content fluctuates in opposition of TOC and varies between 14 and 89 %. The studied interval is generally high in CaCO3 (mean 58 %) (Fig. 2). The δ13Corg displays a minor (~0.5 ‰) shift towards more positive values at ~944 mbs (as reported in Storm et al., 2020; Hollaar et al., 2021). At ~930 mbs an abrupt shift of ~1.8 ‰ (Fig. 3 and 4; Storm et al., 2020) indicates the onset of the Late Pliensbachian Event (LPE) in the Mochras core. In agreement with this, the results of the present study show a shift from ~ minus 27 per mil to ~ minus 25.15 per mil between 930.8 and 930.4 mbs (Fig. 3). The δ13Corg data presented here have been divided into three phases: the pre-LPE gradual rise, followed by the +ve CIE, which is subdivided into pulses 1, 2 and 3 (Fig. 4). After the onset of the positive δ13Corg excursion, the TOC content drops to the lowest values (from 0.85 % before and 0.6 % after the +ve CIE on average), but
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227 the 1 metre fluctuations continue (Fig. 2 and Fig. 3). No overall change in the CaCO₃ content is observed

228 through the positive carbon-isotope excursion (Fig. 2).

229

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230

231 **Figure 2: Detrital ratios over the lithological Ca-rich and TOC-enhanced couplets for the studied**

232 **interval.** Overview of Ca (black, derived from Ruhl *et al.* 2016), CaCO₃ (blue), and TOC content of the studied

233 interval 951–918 mbs. The grey shading represents the TOC-enhanced lithological beds and the unshaded bands

234 mark the Ca-rich (limestone) beds. The detrital ratios reflect the silt to fine sand fraction (Si, Zr) versus the clay

235 fraction (Rb, Al, K). Two increasing upward cycles are observed in the Si/Al and Zr/Rb ratios. The pattern

236 observed in all detrital ratios (except the Ti/Al) is similar and likely reflects changes in grain size.

237

238 **1.4.2 Clay minerals**

239 XRD analysis shows that the main clay types found in this interval are illite, random illite-smectite mixed-layers

240 (I-S R0) [hereafter referred to as smectite], and kaolinite. Illite and kaolinite co-fluctuate in the interval studied

241 here, and are directly out of phase with smectite abundance. Chlorite and R1 I-S are present in minor

242 proportions, but reach sporadically higher relative abundance (> 10%) from ~ 932 mbs upwards, with sustained

243 >10% abundance at ~925–918 mbs (Fig. 3 and SI Fig. 1). The relative abundances of smectite and illite and of

244 kaolinite and illite are expressed by the ratio S/I and K/I respectively. These ratios were calculated according to

245 the intensity of the main diffraction peak of each mineral.

246
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**Figure 3: Synthesis diagram showing the climatic swings observed in tandem with the long eccentricity cycle.** The studied interval comprises part of the pre LPE gradual rise, the initiation of the LPE +ve CIE and pulse 1 and 2. Five climatic phases (A–E) are interpreted from the Si/Al, smectite/illite, kaolinite/illite, chlorite and I-S R1 abundance and the microcharcoal abundance. In tandem with the 405 kyr cycle (Storm *et al.*, 2020) climatic state of a year-round wet climate, low fire activity and fine-grained sediments across multiple precession cycles (phase A and C) alternates with a climatic state that includes repeated precessionally driven states that are semi-arid, with high fire activity and coarser sediments (phase B and D). The top of the record (phase E) indicates increased physical erosion (chlorite + I-S R1, kaolinite) relative to chemical weathering.

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266

267 **Figure 4: The δ¹³Corg (Storm et al., 2020) and δ¹⁸Obulk and δ¹⁸Ofossil (Ullmann et al., 2022) from the Late Pliensbachian of the Mochras core.** A pre-LPE gradual rise is recorded in the δ¹³Corg of the Mochras core, followed by the initiation of the LPE +ve CIE, which consists of three pulses. After the LPE +ve CIE, δ¹³Corg values drop and the Spinatum negative CIE is recorded. The δ¹⁸Obulk of the Mochras core (blue) are diagenetically altered and are unlikely to preserve a palaeoclimatic imprint (Ullman et al., 2022); however, peak values in the δ¹⁸Obulk occur during the LPE +ve CIE. Also, the δ¹⁸Ofossil values (red) are slightly more positive during pulse 3 of the LPE +ve CIE.

274

275 **1.4.3 Organic matter**

276 The type of particulate organic matter, and more specifically the abundance in the either marine or terrestrial
277 origin of the particles, fluctuate on a metre scale from 18–42 %. Palynofacies indicate that the type of organic
278 matter does not change in relation to the metre-scale lithological facies cycles (no correlation between
279 percentage terrestrial phytoclasts and TOC or CaCO₃). The proportion of terrestrial phytoclasts increases
280 towards the top of the record and has 4 high phases: between 944.6 and 942.0 mbs, 937.5 and 934.9 mbs, 930.4
281 and 925.4 mbs, and 920.3 and 918.0 mbs (SI Fig. 2). The first and second high phase falls within the + 0.5 ‰
282 positive swing in the δ¹³Corg; the latter two high phases correspond to pulse 1 and pulse 2 in the +ve CIE.
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283 Amorphous organic matter (AOM) is very abundant, followed by unstructured phytoclasts, with lower

284 proportions of structured phytoclasts and charcoal (SI Fig. 3). Microcharcoal particles make up a relatively large

285 proportion of the terrestrial particulate organic matter (~10 % on average) and ~3.5 % on average of the total

286 particulate organic matter fraction (SI Fig. 3). Only sparse marine and terrestrial palynomorphs were observed

287 (SI Fig. 3). No abrupt changes are recorded in the terrestrial/marine proportions, but small long-term

288 fluctuations are observed in the percentage of terrestrial phytoclasts, with three phases of increase noted, of

289 which the overall highest phase occurs after the start of the +ve CIE.

290 To assess the character of the observed fluctuations in charcoal abundance, whether changes in charcoal can be

291 related to enhanced runoff from the land and/or organic preservation, or if the charcoal signifies changes in fire

292 activity on land, the charcoal record was corrected for detrital influx. We adjust the charcoal particle abundances

293 using the XRF elemental record, normalizing to the total terrigenous influx following Daniau *et al.* (2013) and

294 Hollaar *et al.* (2021). The stratigraphic trends in the normalized microcharcoal for E<sub>icer</sub>/Ca, Si/Al, Ti/Al and

295 Fe/Al remain the same (SI Fig. 4). The absolute number of charcoal particles decreases, with raw mean charcoal

296 particles 1.06x10<sup>5</sup> per 10 g and E<sub>icer</sub>/Ca normalized mean 9.7x10<sup>4</sup> n/10g, Ti/Al normalized 6.4x10<sup>4</sup> n/10g, Si/Al

297 normalized 7.7x10<sup>4</sup> n/10g, Fe/Al normalized 9.8x10<sup>4</sup> mean number of microcharcoal particles per 10 g (SI Fig.

298 4). The number of charcoal particles per 10 g processed rock decreases when correcting for terrestrial run-off

299 changes, hinting that perhaps part of the ‘background’ charcoal is related to terrestrial influx; the normalisation

300 also shows that the observed patterns in charcoal abundances are not influenced by changes in terrestrial runoff

301 and taphonomy. Hence, the highs and lows in the charcoal record can be interpreted to represent changes in the

302 fire regime on land. The microcharcoal abundance fluctuates strongly in the record presented here; however, no

303 clear difference in charcoal content has been observed before and after the onset of the +ve CIE.

304

305 **1.4.4 Detrital elemental ratios (XRF)**

306 Strong similarities are observed between the fluctuating ratios of Si/Al, Si/K, Zr/Rb, Zr/Al and Zr/K (Fig. 2).

307 The elements Al, Rb and K sit in the clay fraction (e.g. Calvert & Pederson, 2007), whereas Si and Zr are often

308 found in the coarser fraction related to silt and sand grade quartz and heavy minerals (Calvert & Pederson,

309 2007). The ratios all show clear metre-scale fluctuations, and these are superimposed on two increasing-upward

310 trends observed in both the Si/Al and the Zr/Rb, followed by a drop and rise to peak values in the latest part of

311 phase D and phase E above the onset of the +ve CIE (Fig. 2; Fig. 3). A parallel trend is observed between the

312 clay ratios (XRD) and elemental ratios Si/Al and Zr/Rb (Fig. 2). Phases of high S/I correspond to the peaks in

313 the two coarsening upward sequences, whereas phases of high K/I correspond to the low phases in the two

314 coarsening upward sequences. After the +ve CIE onset (in phase E) this relationship turns around, and an

315 enrichment in the kaolinite/illite ratio corresponds to the elemental ratios, where highest kaolinite relative

316 abundance is observed in parallel with elemental ratios suggesting maximum coarse fraction.

317

318 **1.4.5 PCA analysis**

319 The proxy datasets (δ<sup>13</sup>C<sub>org</sub>, TOC, percentage terrestrial phytoclasts, microcharcoal, smectite/illite,

320 kaolinite/illite, abundance of chlorite and R1 I-S, Si/Al, Zr/Rb, Zr/Al) were normalized between 0-1 and run for
<figure><img src="image_21.png" /></figure>

321 PCA analysis in PAST. Sixty-four percent of the variance is explained by the first three axes (PCA-1 27.7 %,

322 PCA-2 19.7 %, PCA-3 15.3 %) inside the 95 % confidence interval.

323 PC-1 mainly explains the anti-correlation of TOC and CaCO₃. PC-2 shows the anti-correlation of K/I and S/I.

324 Positive loadings were observed for S/I, microcharcoal, macrocharcoal and CaCO3. For PC-2, negative loadings

325 were observed for K/I, abundance of chlorite + I-S R1. PC-3 shows strong positive loadings (> 0.3) for δ¹³Corg,

326 Si/Al and Zr/Al.

327 Plotting PC-1 (y-axis) over PC-3 (x-axis) shows that the samples after the onset of the +ve CIE are grouped to

328 the top of the y-axis (more associated with S/I compared to K/I) and to the right of the x-axis (more associated

329 with primary minerals, phytoclasts, and higher Si/Al, Zr/Rb and Zr/Al) (Fig. 5).

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

330 **Figure 5: PCA-analysis shows a distinctly different depositional signature before and after the onset of**

331 the LPE +ve CIE in the Mochras core. PCA plot of PC-1 and PC-3: all samples before the onset of the LPE

332 +ve CIE are marked in black closed circles and the samples after the onset of the LPE +ve CIE are marked in

333 blue open squares.

334

335 **1.5 Discussion**

336 Figure 2 provides the context for the LPE ‘cooling event’ at Mochras set within the background record. Shifts in

337 bulk δ¹⁸Ocarb are coeval to the δ¹³Org change to heavier isotopic values (~930 mbs) and reach a maximum in the

338 Margaritatus Zone (>1 ‰) (Ullmann *et al.*, 2022). The bulk oxygen-isotope excursions of Mochras are affected

339 by diagenesis and are deemed unlikely to reflect environmental conditions (Ullmann *et al.*, 2022). However,
<figure><img src="image_23.png" /></figure>

340 oxygen isotope data from marine benthic and nektobenthic molluscs and brachiopods show heavier values

341 during the late Margaritatus Zone concurrent with a positive shift in δ¹³Corg, indicating cooling during the LPE

342 in the nearby Cleveland Basin (Robin Hood's Bay and Staithes) (Korte & Hesselbo, 2011) and this trend is also

343 observed in several European sections (e.g. Korte et al., 2015). The duration of the +ve CIE has been estimated

344 as ~0.4–0.6 Myr in the Cardigan Bay Basin (Ruhl et al., 2016; Storm et al., 2020).

345
346 **1.5.1 Background sedimentological and environmental variations**

347 The Mochras succession shows metre-scale alternating TOC-enhanced and Ca-rich lithological couplets

348 (mudstone/limestone; Fig. 2). Previous assessments of the palaeoenvironmental signature of these TOC-

349 enhanced and Ca-rich couplets indicate strongly that the different depositional modes are driven by orbital

350 precession (Ruhl et al., 2016; Hinnov et al., 2018; Storm et al., 2020; Hollaar et al., 2021; Pieńkowski et al.,

351 2021). Precession driven changes in monsoonal strength have been suggested to influence the deposition and

352 preservation of TOC and carbonate in the Cardigan Bay Basin (Ruhl et al., 2016), although the impact may have

353 been expressed, at least partially, by changes in strength of bottom currents in the seaway as a whole

354 (Pieńkowski et al., 2021).

355 The preservation of primary carbonate is poor in the Mochras borehole, making it complex to determine in

356 detail the relative importance of carbonate producers for the bulk carbonate content (Ullmann et al., 2022).

357 However, Early Jurassic, pelagic settings in the Tethys region often received abiotic fine grained carbonate from

358 shallow marine carbonate platforms (Weedon, 1986; Cobianchi & Picotti, 2001) and/or via carbonate producing

359 organisms (such as coccolithophores in zooplankton pellets) (Weedon, 1986; van de Schootbrugge et al., 2005, e.g. Weedon et al., 2018). Coccolithophores are often poorly preserved and recrystallized (Weedon, 1986;

360 Weedon et al., 2018). The organic matter found in the studied section of the Mochras borehole varies between
361 18 and 42% terrestrial phytoclasts (Fig. 3). Phytoclasts are common, but palynomorphs are relatively sparse and

363 poorly preserved. Marine amorphous organic matter is the main constituent in the present study of particulate

364 organic matter in unsieved macerated samples, in the interval studied here (951 – 918 mbs). Examination of

365 variations in the terrestrial/marine proportions of organic matter, shows no correspondence between the type of

366 organic matter and the TOC-enhanced or Ca-rich lithological alternations. However, previous research has

367 indicated that the percentage of terrestrial phytoclasts show precession forcing independent of the lithological

368 couplets (so out of phase with precession scale changes in Ca-TOC content) between 951 – 934 mbs in the

369 Mochras core (Hollaar et al., 2021). Such orbital forcing of the terrestrial vs marine proportions of organic

370 matter were also found in Early Jurassic sediments of Dorset, and were similarly independent of the lithological

371 facies (Waterhouse, 1999). Terrestrial phytoclast content show a weak expression of long-eccentricity driven

372 variations in the section studied (Fig. 3).

373 Fossil charcoal makes up a substantial proportion of the organic fractions (11% of the terrestrial fraction) and

374 has previously been shown to vary considerably over long-eccentricity cycle 459 ± 1 peaking in abundance

375 during the phase of maximum eccentricity (Hollaar et al., 2021). Microcharcoal also appears to be most

376 abundant during the maximum phase of the subsequent long eccentricity cycle 458 ± 1 (Fig. 3). Additionally,

377 K/I and S/I clay mineral ratios appear to alternate in response to long-eccentricity drivers (Fig. 3) up to 931 mbs
<figure><img src="image_24.png" /></figure>

378 where the clay mineral signature changes. Between 951 and 930 mbs high K/I occurs during phases of low long eccentricity suggesting an enhanced hydrological cycle (Hollaar et al., 2021) with more intense weathering, and enhanced fine grained terrestrial runoff to the marine record (Deconinck et al., 2019). In contrast, phases of maximum long-eccentricity appear to be smectite-rich, indicating seasonal rainfall, enhanced fire (Hollaar et al., 2021) and thus periods of droughts, and lower terrestrial runoff and subsequent lower dilution (Deconinck et al., 2019). Detrital elemental ratios increase accordingly during the smectite-rich phases, and are lower during kaolinite-rich phases between 951 and 930 mbs. Detrital elemental ratios can be used to explore changes in sediment composition (e.g. Thibault et al., 2018; Hesselbo et al., 2020b) and the similarity of the long-term trend in Zr/Rb and Si/Al (Fig. 2) indicates that these elemental ratios reflect grainsize. The clay fraction (hosting Al, and Rb (Chen et al., 1999)), diminishes upwards, whereas the coarser silt to sand fraction (associated with Si (Hesselbo et al., 2020b) and Zr (Chen et al., 2006)), increases upward (Fig. 3 and 4). The grainsize changes inferred here reflect two overall coarsening upwards sequences (Fig. 3 and 4). These sequences may reflect changes in clastic transport due to changes in the proximity to the shore/siliciclastic source, changes in runoff due to a changing hydrological cycle, or accelerated bottom currents with greater carrying capacity of coarser sediments.

393

394 **1.5.2 Depositional and environmental changes before and after the LPE +ve CIE**

395 The LPE +ve CIE begins around 930 mbs in the Mochras core and encompasses the remaining part of the studied section (Fig. 3). We contrasted all the pre-CIE sediment signatures with those of the +ve CIE signatures using principal components analysis which indicates a distinctly different sedimentary composition and environmental signature before and after the onset of the +ve CIE in Mochras (Fig. 5).

399 Before the +ve CIE onset, the clay mineral assemblage shows alternating phases of smectite and kaolinite, indicating pedogenic weathering. The relative abundance of the detrital clay types observed in the studied interval have the potential to hold important palaeoclimatic information regarding the hydrological cycle and the relative proportion of chemical weathering and physical erosion. Chemical weathering is enhanced in a high humidity environment with relatively high temperatures and rainfall, when clays are formed in the first stages of soil development. In the modern day, kaolinite is primarily formed in tropical soils, under year-round rainfall and high temperatures (Thiry, 2000). Smectite also occurs in the tropics, but is more common in the subtropical to Mediterranean regions, where humidity is still high, but periods of drought also occur (Thiry, 2000). Hence, smectite forms predominantly in soil profiles under a warm and seasonally dry climate (Chamley, 1989; Raucsik & Varga, 2008), and kaolinite in a year-round humid climate (Chamley, 1989; Ruffell et al., 2002). Similarly, alternating intervals of kaolinite and smectite dominance were observed for the Late Sinemurian (Munier et al., 2021) and the Pliensbachian of Mochras (Deconinck et al., 2019). The predominantly detrital character of these clay minerals has been confirmed by TEM scanning of Pliensbachian smectite minerals, which revealed the fleecy morphology and lack of overgrowth (Deconinck et al., 2019). Therefore, the alternations of smectite and kaolinite are interpreted to reflect palaeoclimatic signatures of a changing hydrological cycle, with a year-round wet climate evidenced by high K/I ratios, and a more monsoon-like climate with seasonal rainfall with high S/I (Deconinck et al., 2019; Hollaar et al., 2021; Munier et al., 2021) (See Fig. 3). The intervals with a signal for

415
<figure><img src="image_25.png" /></figure>

416 weaker seasons appears to correspond to phases of low eccentricity in the 405 kyr cycle, and signals of greater
417 seasonality with periods of high more pronounced eccentricity (Fig. 3) in the 405 kyr cycle.

418 Higher frequency cycles are not observed in the clay mineral ratios, with no precession or obliquity forcing
419 detected in the high-resolution part of the study 951 – 934 mbs (Hollaar et al., 2021) and no expression of the
420 100 kyr cycle in the record presented here. The formation of developed kaolinite-rich, and to a lesser extent
421 smectite-rich soil profiles, requires a steady landscape for many tens of thousands of years, although the ~1 Myr
422 timescale of Thiry (2000) seems excessive in our case given the clear expression of clay mineral changes
423 through long-eccentricity cycles. Also, the transportation and deposition of continental clays will occur after soil
424 formation and add further time between formation and final deposition (Chamley, 1989; Thiry, 2000). Thus,
425 there is likely to be a lag of the climatic signal observed in the marine sediments (Chamley, 1989; Thiry, 2000).
426 However, we note that high frequency climatic swings have been recorded in the clay mineral record in some
427 instances, such as in the Lower Cretaceous in SE Spain (Moiroud et al., 2012). The limestone-marl alternations
428 there are enhanced in smectite versus kaolinite and illite, respectively, reflecting precession scale swings from a
429 semi-arid to a tropical humid climate (Moiroud et al., 2012). Precession and higher frequency shifts in the clay
430 record are likely caused by fluctuations in runoff conditions rather than the formation of soils with a different
431 clay fraction.

432 Directly after the initial +ve CIE shift from 930-924 mbs (Phase 1 of Fig. 3) little seems to change, and the
433 system evidently continued to respond as before to the long eccentricity forcing, despite the predicted cooling
434 (Korte & Hesselbo, 2011; Korte et al., 2015; Gómez et al., 2016). However, from around 924 mbs up to the top
435 of the studied section (Phase 2 of Fig. 3) the clay mineral assemblage displays a distinctly different composition,
436 with kaolinite dominating especially the early part of phase 2 of the LPE (Fig. 3). At the same time there is an
437 enhancement of the primary minerals illite and chlorite, and I-S R1 (Fig. 3 and SI Fig. 1). Although an
438 enhancement in detrital kaolinite indicates an acceleration of the hydrological cycle, detrital kaolinite is dual in
439 origin and can also be derived from reworking of the primary source material (Deconinck et al., 2019). If the
440 climate is cooler, chemical weathering becomes less dominant and physical erosion of the bedrock becomes the
441 main detrital source of clay minerals. In the Cardigan Bay Basin, the bedrock of the surrounding Variscan
442 massifs (such as the Caledonian, Welsh and Irish massifs) were a likely source of these clays. In the Early
443 Jurassic of the NW Tethys region, mica-illite and chlorite bearing Lower Paleozoic mudrock were emergent
444 (Merriman, 2006; Deconinck et al., 2019), hence the enhancement of illite and chlorite likely indicates physical
445 erosion in the region surrounding the study site. Finally, authigenic clay particles could have been formed
446 during burial diagenesis. At temperatures between 60-70 °C smectite illitization occurs and I-S R1 is formed;
447 however, the high abundance of smectite in Mochras indicates limited burial diagenesis in the Mochras core
448 (Deconinck et al., 2019). Weak thermal diagenesis is confirmed for the Pliensbachian of Mochras, with Tmax
449 between 421 °C and 434 °C (van de Schootbrugge et al., 2005). Therefore, I-S R1 in Mochras is interpreted to
450 be derived from chemical weathering of illite (Deconinck et al., 2019). The coeval increase of these primary
451 clay minerals, I-S R1 and kaolinite, indicate that during this period physical erosion dominated over soil
452 chemical weathering (Deconinck et al., 2019; Munier et al., 2021). This is similar to what is observed for the
453 latest Pliensbachian in Mochras (Deconinck et al., 2019).
<figure><img src="image_26.png" /></figure>

454 Erosion of weathering profiles transports clay minerals (including kaolinite and smectite) to the marine realm. In the ocean, the differential settling of kaolinite (near shore) and smectite (more distal) could occur based on the morphology and size of clay particles (Thiry, 2000). However, comparison of long-term inferred regional sea level changes from surrounding UK basins (Hesselbo, 2008) suggests that the relative proportions of smectite and kaolinite are not influenced by changes in relative sea level in the Pliensbachian of Mochras (Deconinck et al., 2019). On the assumption that the coarsening upward sequences at Mochras are indicative of relative sea level change, it can also be argued that the proximity to shore did not impact the proportions of smectite and kaolinite with enhanced smectite during ‘proximal’ deposition and enhanced kaolinite at times of more ‘distal’ deposition (Fig. 3).

463 We suggest that the first phase of the LPE (Fig. 3, phase 1) was characterised by repeated periods of rainfall in a seasonal climate forced by precession in which chemical weathering (smectite formation) dominated the sedimentary signatures. This corresponds to maximum long-eccentricity and shows the same climatic signature as during maximum eccentricity phases before the +ve CIE. This is then followed by a second phase (Fig. 3, phase 2) where the climate is generally cooler, overall potentially more arid, but with rainfall throughout the year over multiple precession cycles. This appears to have favoured deep physical erosion, owing to the abundance of primary clay minerals, kaolinite and I-S R1. This interval corresponds to a minimum phase in the 405 kyr eccentricity based on Storm et al. (2020). This interpretation is further supported by decreasing and then low microcharcoal abundance, pointing to suppression of fire activity at this time.

472 Two coarsening upward cycles that predate the onset of the +ve CIE and continue for a few metres after its initiation, are present in the detrital elemental ratios (best expressed in Si/Al and Zr/Rb records) (Fig. 3 and 4), and indicate a changing sediment influx over the studied interval. Previous study of the lithofacies of the Mochras borehole has also shown the coarsening-upward sequences of 0.5–3 m thickness, which are observed to be followed upwards by a thinner fining-upward succession (Pieńkowski et al., 2021). This reported fining-upward part is not reflected in the elemental ratios of the two sequences shown in this study. Furthermore, the coarsest phases of these sequences are approximately coeval with decreasing trends in the K/I ratio and increasing trends in the S/I. This could indicate that periods of a strong monsoonal/seasonal climate (indicated by S/I) brought coarser grained material to the basin, whereas periods of year-round humidity (K/I) are associated with higher chemical weathering (low Si/Al). Therefore, these two coarsening upwards cycles appear to link to increasing long-eccentricity. A similar mechanism has been inferred for the northern South China Sea region in the Miocene, where coarser grained material is found during periods of a strong summer monsoon and relatively lower chemical weathering (Clift et al., 2014). Present day studies show that bedrock erosion and associated sediment transport is greater in areas with high seasonal contrast (Molnar, 2001; Molnar, 2004). Hence, the Si/Al record also appears to reflect weathering and erosion conditions on land (Clift et al., 2014, 2020), driven by long-eccentricity modulated climate (SI Fig.5). However, other scenarios that would influence the grain size on this time scale cannot be dismissed and include changes in proximity to siliciclastic source, or changes in sediment transport via bottom water currents.

490 Changes in bottom water current strength and direction likely affected the depositional site of the Mochras core (Pieńkowski et al., 2021) although there is as yet no consensus on the processes that likely controlled these palaeoceanographic parameters. An early phase of regional tectonic updoming of the North Sea disrupted the
<figure><img src="image_27.png" /></figure>

493 circulation in the N-S Laurasian Seaway (including the Viking Corridor) and therefore diminished the connectivity between western Tethys and the Boreal realm, hypothetically reducing poleward heat transport from the tropics (Korte et al., 2015). This mechanism has also been argued to explain the later cooling observed in NW Europe during the transition of the warmer Toarcian to the cooler Aalenian and Bajocian (Korte et al., 2015). Late Pliensbachian occlusion of the Viking Corridor is supported by the provincialism of marine faunas at this time, showing a distinct Euro-Boreal province and a Mediterranean province (Dera et al., 2011b). During the Toarcian a northward expansion of invertebrate faunal species has been found (Schweigert, 2005; Zakharov et al., 2006; Bourillot et al., 2008; Nikitenko, 2008), indicating a northward (warmer) flow through the Viking corridor (Korte et al., 2015). More recently, a southward expansion of Arctic dinoflagellates into the Viking Corridor was suggested for the termination of the T-OAE (van de Schootbrugge et al., 2019), which is in agreement with a N to S flow through the Viking Corridor suggested by numerical models (Bjerrum et al., 2001; Dera & Donnadieu, 2012; Ruvalcaba Baroni et al. 2018) and sparse Nd-isotopes (Dera et al., 2009).

505 Over the European Epicontinental Shelf (EES), and the Tethys as a whole, a clockwise circular gyre likely brought oxygenated warm Tethyan waters to the southwest shelf, with a progressively weaker north and eastward flow due to rough bathymetry and substantial islands palaeogeography (Ruvalcaba Baroni et al., 2018). This predominantly surface flow is modelled to have extended to shelfal sea floor depths. Only episodically might nutrient-rich Boreal waters have penetrated south onto the EES in these coupled ocean-atmosphere GCM model scenarios (Dera & Donnadieu, 2012). The modelling also suggests – counter-intuitively – that the clockwise surface gyre of the Tethys extended further northwards and impacted the EES more effectively when the Hispanic corridor was more open.

513 An alternative bottom current configuration was discussed for Mochras specifically wherein changes in north-to-south current strength (cf. Bjerrum et al., 2001) are proposed for the changes in grainsize and siliciclastic versus clay content via contour currents (Pieńkowski et al., 2021). A strong flow from the cooler and shallow boreal waters is hypothesized to have brought a coarser grainsize fraction in suspension and as bedload, which was then deposited in the Cardigan Bay Basin while flowing to the deeper and warmer waters of the peri-Tethys (Pieńkowski et al., 2021). Times of a strong north to south current are proposed to be associated with more oxygenated bottom waters (Pieńkowski et al., 2021). In contrast, when the north to south current became weaker, less coarse material will have been carried in suspension and as bedload, and a relatively higher clay proportion will have been deposited in the Cardigan Bay Basin (Pieńkowski et al., 2021). In this scenario, times of sluggish currents are associated with low bottom water oxygenation (Pieńkowski et al., 2021) and thus climate forcing of current strength could explain the deposition of alternating coarser and finer fractions in the Mochras borehole (Pieńkowski et al., 2021).

525 Our research suggests that orbital cycles both before and during the onset of the +ve CIE have a significant influence on seasonality and hydrology, affecting both fire regimes and sediment depositional character. Further research is required to consider how long-eccentricity and obliquity cycles might interact with north-south flow in the Cardigan Bay Basin and circulation processes. What is clear is that orbital cycles have impact on terrestrial processes in the terrestrial sediment source areas (Hollaar et al., 2021) and led to differences in deposition within the marine sediments in Mochras core (Ruhl et al., 2016; Pieńkowski et al., 2021). Our data indicate that periods of coarser sediment deposition correspond to periods that include more seasonal climates
<figure><img src="image_28.png" /></figure>

532 before the onset of the +ve CIE (low kaolinite), which is in line with the hypothesized grainsize changes caused by contour currents (Pieńkowski *et al.*, 2021). However, after the onset of the +ve CIE, although we suggest that the chemical weathering rate decreased, enhanced runoff and physical erosion are indicated by a peak in primary clay minerals and K/I. Enhanced runoff could be expected to impact the thermohaline contour currents (Dera & Donnadieu, 2012). Simultaneously, an increasingly cold climate (as indicated by enhanced physical erosion over chemical weathering) indicates a boreal influence. It remains to be determined to what extent orbital cycles might have the power to influence ocean circulation in the basin.

539 Relatively coarse sediments in the Late Pliensbachian have also been related to shallower sediment deposition in UK basins (Hesselbo & Jenkyns, 1998; Hesselbo, 2008; Korte & Hesselbo, 2011). These regressive facies may have been caused by an early phase of North Sea doming (Korte & Hesselbo, 2011). Sequence stratigraphy of the Lower Jurassic of the Wessex, Cleveland and Hebrides basins (Hesselbo & Jenkyns, 1998; Hesselbo, 2008) shows relative sea level changes and sand influxes in the late Margaritatus Zone in the studied basins.
541 542 543
544 Noteworthy in the Mochras borehole are phases of low δ¹⁸O of macrofossils which seem to correspond to high phases of macrofossil wood concomitant with low sea level, suggesting a possible control of relative sea level on the oxygen-isotope record and the source of detrital material (Ullmann et al., 2022). The broad spatial distribution of these basins suggests that associated regression and/or sediment influx is of at least regional scale (Hesselbo, 2008). The results presented here fall within this phase of regression (Hesselbo & Jenkyns, 1998; Hesselbo, 2008).
545 546 547 548 549

550 In the context of North Sea doming as a possible cause of the Late Pliensbachian cooling, these facies can be interpreted to represent shallowing upward facies in a shallower system, or deep water system receiving coarser sediment input. The doming is hypothesized to have minimized or prohibited southward flow of cooler waters from the Boreal and northward flow from warmer waters from the Mediterranean area (Korte et al., 2015). The Mochras borehole is situated on the southwestern flank of the dome and would have been cut-off from the northern parts of the Laurasian Seaway, including the Hebrides Basin and Cleveland Basin (Korte et al., 2015).
551 552 553 554 555 556 This change in seaway circulation could have impacted the source area of the detrital sediments in the Mochras borehole and brought the shallow shoreface facies closer to the borehole site.
557

558 Doming of the North Sea area would have led to greater radial spread of nearshore facies; however, owing to the strong eccentricity forcing that we interpret here, an additional factor that is influenced by the seasonal distribution of insolation forced by orbital cyclicity needs to be included. The Cardigan Bay Basin (Mochras) is positioned about 290 km to the SW of the Cleveland Basin and at a similar latitude, but to the W of the Wessex Basin (Ziegler, 1990; Torsvik & Cocks, 2017), and is therefore expected to be impacted by the regional changes in sea level and/or sediment flux. In the Late Pliensbachian of the Cleveland Basin, the detrital ratios of Si/Al, Zr/Al and Zr/Rb show similar coarsening upward sequences, which have been interpreted to reflect changes in riverine transport of siliciclastic grains and grainsize (Thibault et al., 2018). The inferred changes in sea-level in the Cleveland Basin occur at a 100 kyr pacing (Huang & Hesselbo, 2014; Hesselbo et al., 2020b), potentially linking the regression cycles to short eccentricity (Huang et al., 2010 and refs therein) and long-eccentricity (Thibault et al., 2018). This would mean that eccentricity driven changes in inferred sea level change could be linked to glacioeustatic cycles during these times (Brandt, 1986; Suan et al., 2010; Korte & Hesselbo, 2011; Krencker et al., 2019; Ruebsam et al., 2019, 2020b; Ruebsam & Schwark, 2021; Ruebsam & Al-Husseini,
559 560 561 562 563 564 565 566 567 568 569 570
<figure><img src="image_29.png" /></figure>

571 2021). Glacioeustatic sea level changes are discussed for the Early Jurassic and Middle Jurassic (Krencker et al., 2019; Bodin et al., 2020; Ruebsam & Schwark, 2021; Nordt et al., 2022). A recent study on the rapid transgression observed at the Pliensbachian–Toarcian boundary, ruled out other mechanisms that could force sea level at this time scale, such as aquifer-eustacy, and show that glacioeustatic changes in sea level are a likely possibility at times in the Early Jurassic (Krencker et al., 2019). Therefore, our findings overall provide support the episodic occurrence of continental ice at the poles (Brandt, 1986; Price, 1999; Suan et al., 2010; Korte & Hesselbo, 2011; Korte et al., 2015; Bougeault et al., 2017; Krencker et al., 2019; Ruebsam et al., 2019, 2020a, 2020b; Ruebsam & Schwark, 2021; Ruebsam & Al-Husseini, 2021).

579
580 **1.6 Conclusions**

581 The terrestrial environment adjacent to the Cardigan Bay Basin was strongly influenced by both orbitally driven
582 climate forcings and the Late Pliensbachian Cooling Event (LPE). Long-eccentricity forcing remained strong
583 both prior to and during the LPE. Prior to the LPE, eccentricity-driven shifts in maximum seasonality influence
584 the degree of chemical weathering (S/I vs K/I), sediment flux to the basin (Si/Al), and fire activity. As maximum
585 precessional seasonality decreases with reduced 405 kyr eccentricity, the year-round relatively cool and wet
586 climate extended over multiple precession cyles drove significant erosion of bedrock on emergent land surfaces
587 as evidenced by high bedrock-derived mineral content, high K/I and I-S R1. Therefore, both the Milankovitch
588 forcings and larger climatic shifts operate in tandem to drive changes in the terrestrial environment.

589
590 **Data availability:** Supplementary data is available at the National Geoscience Data Centre at Keyworth (NGDC) at (doi to be added) for the interval 934 – 918 mbs. All data presented for the interval 951 – 934 mbs is available at the National Geoscience Data Centre at Keyworth (NGDC) at https://doi.org/10.5285/d6b7c567-49f0-44c7-a94c-e82fa17ff98e (Hollaar et al., 2021). The full Mochras XRF dataset is in Damaschke et al. (2021).

594
595 **Author contribution:** CMB, SPH and TPH designed the research. TPH conducted the laboratory
596 measurements, with JFD contributing to the XRD-measurements and MD, CU and ML to the XRF-
597 measurements. TPH, CMB and SPH wrote the manuscript, with contributions of all authors.

598 **Competing interests:** The authors declare that they have no conflict of interest.

599 **Acknowledgements:** This is a contribution to the JET project funded by the Natural Environment Research Council (NERC) (grant number NE/N018508/1). SPH, CMB, JFD, CU, ML and TPH, acknowledge funding from the International Continental Scientific Drilling Program (ICDP) and TPH acknowledges funding from the University of Exeter. We thank the British Geological Survey (BGS), especially James Riding, Scott Renshaw for facilitating access to the Mochras core. Also, Simon Wylde and Charles Gowing for their contribution to the XRF scanning and discussion on the results. We further thank Chris Mitchell for help with the TOC and δ¹³Corg analyses. Finally, we thank Ludovic Bruneau for technical assistance with the XRD-analysis.

606
<figure><img src="image_30.png" /></figure>

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