# Constraining oceanic carbonate chemistry evolution during the Cretaceous-Paleogene transition: Combined benthic and planktonic calcium isotope records from the equatorial Pacific Ocean

Arbia Jouini, Guillaume Paris, Guillaume Caro, Annachiara Bartolini, Sivia Gardin

► To cite this version:

Arbia Jouini, Guillaume Paris, Guillaume Caro, Annachiara Bartolini, Sivia Gardin. Constraining oceanic carbonate chemistry evolution during the Cretaceous-Paleogene transition: Combined benthic and planktonic calcium isotope records from the equatorial Pacific Ocean. Earth and Planetary Science Letters, 2023, 619, pp.118305. 10.1016/j.epsl.2023.118305. insu-04198247

HAL Id: insu-04198247

https://insu.hal.science/insu-04198247v1

Submitted on 23 Jul 2025

HAL is a multi-disciplinary open access
archive for the deposit and dissemination of sci-
entific research documents, whether they are pub-
lished or not. The documents may come from
teaching and research institutions in France or
abroad, or from public or private research centers.

L'archive ouverte pluridisciplinaire HAL, est
destinée au dépôt et à la diffusion de documents
scientifiques de niveau recherche, publiés ou non,
émanant des établissements d'enseignement et de
recherche français ou étrangers, des laboratoires
publics ou privés.

<figure><img src="image_1.png" /></figure>

Distributed under a Creative Commons Attribution 4.0 International License
1

Constraining oceanic carbonate chemistry evolution during the Cretaceous-Paleogene

2

transition: combined benthic and planktonic calcium isotope records from the equatorial

3

Pacific Ocean

4

5

Arbia JOUINI¹*, Guillaume PARIS¹, Guillaume CARO¹, Annachiara BARTOLINI², Sivia

6

GARDIN²

7

¹ Université de Lorraine - CNRS, CRPG, F-54000 Nancy, France 54000, France²

8

² CR2P, CNRS - MNHN - Sorbonne Université, 8 rue Buffon CP38, 75005 Paris, France

9

*Corresponding author E-mail address: arbia.jouini@univ-lorraine.fr

10

Abstract

11

12

13

14

15

16

17

18

19

20

21

22

23

24

The Mesozoic-Cenozoic transition is a period of biogeochemical cycle perturbations. The strongest of them is the Cretaceous-Paleogene boundary (K-Pg) crisis, characterized by one of the most important extinctions of planktonic marine calcifiers in Earth's history. One of the primary drivers of this biocalcification crisis is thought to be the increase in atmospheric CO₂ concentration and ocean acidification triggered by the Chicxulub Impact, and/or Deccan volcanism. Because it reflects changes of the calcium cycle and/or depends on parameters of the carbonate system, the Ca isotope composition of carbonate minerals precipitated from seawater (δ⁴⁴/⁴⁰Ca) offers the potential to reconstruct some of the environmental changes that occurred. Here we present new high-resolution planktonic and benthic foraminiferal δ⁴⁴/⁴⁰Ca, δ¹⁸O, δ¹³C, and Sr/Ca records across the K-Pg transition from Shatsky rise (Leg 198; ODP Site 1209, Hole C). The δ⁴⁴/⁴⁰Ca record displays a succession of rapid shifts of ca. -0.4 ‰ across the K-Pg transition. They are similar though not identical between the planktonic and benthic records. These shifts took place on a timescale significantly shorter than the residence time of Ca in the oceans and are therefore unlikely to result from global
disequilibrium in the oceanic Ca budget. Instead, changes in the fractionation factor between carbonate minerals and seawater in response to changes in precipitation rates may explain the observed δ<sup>44/40</sup>Ca and Sr/Ca record. The benthic and planktonic δ<sup>44/40</sup>Ca records show a late Maastrichtian and an early Danian negative excursions best explained by a succession of episodes of ocean alkalinity increase related to increased continental weathering caused by CO<sub>2</sub> emissions from Deccan volcanism and the aftermath of the K-Pg biocalcification crisis. Carbonate compensation via carbonate sediment dissolution, biological carbonate compensation via reduction of biocalcification, and/or an increase in continental weathering must have occurred to offset the excess CO<sub>2</sub>, ultimately resulting in rapid changes in ocean carbonate chemistry, in combination with reduced surface alkalinity export in response to the early Paleogene planktonic biomineralization crisis.

**Keywords: Calcium isotopes, K-Pg transition, planktonic and benthic foraminifera, Deccan**

# 1. Introduction

The transition from the late Maastrichtian to the Paleogene is a period of major environmental changes. During the late Maastrichtian, a period of global warming known as the Late Maastrichtian Warming Event (LMWE, e.g. Li and Keller, 1998) and a transient period of ocean acidification are observed and attributed to the onset of Deccan volcanism (Dameron et al., 2017; Henehan et al., 2016). These environmental disruptions precede the Cretaceous-Paleogene (K-Pg) boundary crisis (-66 Ma, Renne et al., 2013), one of the most important mass extinctions of the Phanerozoic. This extinction is reflected in the oceanic system by a carbonate production crisis that preferentially impacted planktonic biocalcifiers (foraminifera and calcareous nannofossils), reducing the exportation of pelagic calcite and organic matter to the deep ocean (Alegret et al., 2021). Two major geological events are often considered as triggers to the mass extinction: the Chixculub impact or a paroxysmal
peak of activity of the Deccan (Alvarez et al., 1980; Arenillas et al., 2016; Keller et al., 2020). 52

However, the exact trigger and timing of events of the K-Pg crisis remain heavily disputed 53
(e.g. Hull et al., 2020; Sprain et al., 2019; Schoene et al., 2021, 2019). 54

Both events released large quantities of CO₂ and SO₂ into the ocean-atmosphere system 55
(Beerling et al., 2002). The bolide impact created a dust cloud that blocked photosynthesis 56
and instantaneously vaporized the surrounding sediments and released vast amounts of CO₂ 57
and SO₂ (Pope et al., 1997). In contrast, the Deccan volcanism emitted CO₂ and SO₂ in 58
amounts orders of magnitude above those of the impact, but with a duration of 100-150 kyr 59
of thousands of years. Accordingly, volcanic emissions probably had a longer-term influence 60
on oceanic chemistry (Gilabert et al., 2021; Hernandez Nava et al., 2021), although their 61
ability to generate an environmental stress depends on the specific timing of the pulses (e.g. 62
Keller et al., 2020; Paris et al., 2016). Additional constraints on the evolution of 63
biogeochemical cycles and oceanic chemistry across the K-Pg transition are thus needed to 64
help understand the broad paleoenvironmental sequence around this emblematic paleo- 65
environmental crisis. 66

The calcium isotopic composition of biogenic carbonates can be a valuable tool to 67
reconstruct past environmental changes because of its sensitivity to different aspects of 68
seawater chemistry and the carbonate system. Conceptually, the evolution of the calcium 69
isotopic composition of marine carbonates can reflect either changes in seawater 70
composition and/or changes in the isotopic fractionation factor associated with calcium 71
carbonate precipitation. The isotopic fractionation between carbonate minerals and seawater 72
is believed to be kinetically controlled, with CaCO₃ preferentially incorporating light calcium 73
isotopes (Skulan et al., 1997). As a result, seawater is isotopically heavier than marine 74
carbonates, and its isotopic composition (reported in ‰ as δ<sup>44/40</sup>Ca<sub>sw</sub>, defined as 75
[(<sup>44</sup>Ca/<sup>40</sup>Ca)<sub>SW</sub>/(<sup>44</sup>Ca/<sup>40</sup>Ca)<sub>NIST SRM915a</sub>]-1; e.g. Hippler et al., 2003) can change in response to 76
an imbalance between continental and hydrothermal inputs to the ocean and precipitation of 77
calcium carbonate (e.g. Blättler et al., 2011). Furthermore, the isotopic fractionation factor 78
can vary in response to changes in biological activity (e.g. Gussone et al., 2005), 79
temperature (e.g. Marriott et al., 2004), precipitation rate (e.g. Lemarchand et al., 2004; Tang et al., 2008), and seawater carbonate chemistry (Ca/CO₃ ratio; Nielsen et al., 2012). Despite the relatively long residence time of Ca in the ocean (ca. 1 Myr), calcium isotopes can thus be a valuable tool for tracking short-lived environmental changes such as ocean acidification or alkalinization. Investigating high-resolution δ<sup>44/40</sup>Ca records at periods of significant environmental changes, such as the K-Pg boundary, may thus aid in understanding how these factors influence the isotopic composition of marine carbonates and inform environmental changes during such events (e.g. Du Vivier et al., 2015; Linzmeier et al., 2020).

The δ<sup>44/40</sup>Ca composition of carbonates has been investigated for various critical events such as the end-Permian (e.g. Hinojosa et al., 2012), the end-Triassic (e.g. Jost et al., 2017), and the K-Pg (Linzmeier et al., 2020) mass extinctions, as well as for hyperthermal (e.g. PETM; Kitch et al., 2021) and different Oceanic Anoxia Events (OAEs, Blättler et al., 2011; Du Vivier et al., 2015). However, the interpretation of the Ca isotopic composition of bulk sediments can be complicated by the presence of authigenic carbonates (Fantle and Ridgwell, 2020). Records based on single-species biogenic carbonates are thus preferable because these authigenic phases can be eliminated. The only published calcium isotope record spanning the transition from the Maastrichtian to early Danian is based on aragonitic bivalve and gastropod mollusk shells from the López de Bertodano Formation on Seymour Island, Antarctica (Linzmeier et al. 2020). This record corresponds to a nearshore shallow water environment and shows two negative excursions from each side of the K-Pg boundary that are interpreted in terms of local changes in the seawater saturation state, separated by δ<sup>44/40</sup>Ca positive excursion interpreted as episodes of ocean acidification. Additional δ<sup>44/40</sup>Ca data from diverse locations, ideally with a higher resolution, could thus contribute to a better understanding of the consequences of the K-Pg crisis on marine chemistry.

Because of their abundance in Phanerozoic marine carbonates, foraminifera have often been used as an archive for reconstructing past changes in the Ca isotopic composition of
seawater and/or use calcium isotopes from foraminiferal calcite as a way to reconstruct of
seawater chemistry (Hippler et al., 2006; Kitch et al., 2021, 2022; Sime et al., 2005, 2007).
To establish a high-resolution record spanning the K-Pg transition in the open ocean, we
analyzed the Ca isotope composition of monospecific foraminifers spanning the
Maastrichtian-Danian transition from ODP Site 1209, Hole C.(Shatsky Rise) in the central
North Pacific Ocean (Nakanishi et al., 1999). Because surface and deep waters can be
impacted by different processes, we report δ<sup>44/40</sup>Ca, δ<sup>13</sup>C, δ<sup>18</sup>O, and Sr/Ca data for both
planktonic and benthic species. This multiproxy record is interpreted in terms of the evolution
of the carbonate system across the K-Pg transition, shedding new light on
paleoenvironmental conditions of the deep and surface water masses of the Equatorial
Pacific during the K-Pg transition.

118

119

120 **2. Materials and methods**

121 **2.1. Study sites**

122 The multiproxy geochemical record described in this study was obtained from the equatorial
123 Pacific deep water ODP Site 1209, Hole C on the Southern High of Shatsky Rise (Figure 1).
124 Shatsky Rise is a basaltic oceanic plateau (Nakanishi et al., 1999), with an estimated paleo-
125 depth of 2387 m (Bralower et al., 2006), located near the equator in the Early Cretaceous.
126 We selected the more abundant and better-preserved foraminiferal species for geochemical
127 analysis based on previously established biostratigraphic patterns of ODP Site 1209, Hole
128 C (Premoli Silva et al., 2005). This location is characterized by the presence of a dissolution
129 interval during the last Maastrichtian, from ~246.3 to ~236.3 mbsf (meters below seafloor), as
130 defined by Dameron et al. (2017) using the fragmentation index and the planktonic to benthic
131 foraminifer ratio. However, the boundaries of this interval can be further narrowed down
132 based on the steady qualitative deterioration of taxa induced by dissolution (Fig. A shows
133

scanning electron microscope (SEM) views of the samples). Qualitative reexamination of
uppermost Maastrichtian and Danian assemblages and preservation state were therefore
realized from section 1209C_16H_1W to section 1209C_15H_1W, allowing a more precise
definition of the stratigraphic level where dissolution is most severe, from 242 to 238 mbsf
(Figure 3). These new qualitative biostratigraphic observations are detailed in Appendix B.

134

135

136

137

138

139

## 2.2. Foraminifera species

140

To obtain enough tests for geochemical analyses (δ<sup>44/40</sup>Ca, Me/Ca), we selected the most
abundant species through different sections of the core (Fig. A). In the uppermost
Maastrichtian (from 249.73 to 239.73 mbsf), we selected the planktonic species
*Pseudoguembelina costulata* (250-125 µm size fraction) to reconstruct surface seawater
δ<sup>44/40</sup>Ca. *P. costulata* has been described as an upper mixed layer species, recording
environmental conditions of near-surface water masses (Abramovich and Keller, 2003).

141

142

143

144

145

146

Between 238.98 and 234.68 mbsf, planktonic foraminifer assemblages are characterized by
low diversity and the test size is generally small. In this interval, we selected tests of the
biserial taxa *Woodringina claytonensis* from the 125-63 µm fraction, which also occupies
near-surface habitats (MacLeod, 1993). The stratigraphic distribution of this species in ODP
site ODP Site 1209, Hole C is further discussed in Appendix B. Further up in the section,
the mixed layer species *Chiloguembelina morsei* (MacLeod, 1993) collected from a larger
size fraction (250-125 µm) was used up to 231.23 mbsf. Finally, the uppermost Danian (from
230.48 to 227.73 mbsf) is marked by an increase in trochospiral taxa. We selected the mixed
layer species *Praemurica inconstans* (Norris, 1996) from 250-125 µm size fraction, but a
minor contribution from *P. pseudoinconstans* is possible as these two species can be easily
confused under a binocular microscope. For benthic foraminifera, we used exclusively the
shallow infaunal taxa *Oridorsalis umbonatus* (Rathburn and Corliss, 1994; >250 µm size

147

148

149

150

151

152

153

154

155

156

157
158

fraction), which lives at or near the water/sediment interface and is thus a good indicator of
the characteristics of deep water masses.

159

160

For carbon and oxygen isotope records, additional tests were picked from the same mixed
layer species (i.e *Woodringina claytonensis*, *Chiloguembelina morsei* and *Praemurica inconstans*) around the K-Pg transition and the Danian (from 238.23 to 227.73 mbsf).

161

162

163

However, for the base of the Maastrichtian record (from 249.73 to 242.24 mbsf), the samples
did not contain enough specimens of *Pseudoguembelina costulata*. Therefore, another
Cretaceous species (*Pseudoguembelina excolata*) was used to produce the carbon and
oxygen isotope records.

164

165

166

167

168

## 2.3. Sample preparation

169

After picking, the selected carbonate shells were gently crushed between two glass slides
under the binocular microscope and mixed to homogenize the sample. Samples that
contained secondary calcite fillings were discarded. The shells were then cleaned using the
trace element cleaning protocol developed by (Boyle, 1981) and slightly modified by (Rennie
et al., 2018), a suitable process for calcium isotope analyses (Gussone et al., 2009).

170

171

172

173

174

Subsequently, the samples were dissolved in 500 µl of 1% HCl. After complete dissolution,
500 µl of ultra-pure water was added to obtain a total volume of 1 ml. Two sub-samples of 20
µl each were aliquoted for the determination of calcium and trace elements concentrations
(Mg, Mn, Sr, and Fe). The remaining solution was mixed with a ⁴²Ca-⁴³Ca double spike for
Ca isotope analysis. The spike-sample mixtures were equilibrated at 60°C on a hotplate
overnight and then dried down in a laminar flow hood. The spiked samples were finally
dissolved with ultra-pure water and passed through a cation exchange resin AG1X8 (Bio-
Rad) to separate the cation and anion fractions. The cation fraction (referred to as samples
residue) was dried down and stored in a Teflon beaker for Ca isotope analysis and the anion
fraction was saved for later analyses. A blank aliquot spiked in ⁴²Ca was processed with each
batch of samples (typically 7 to 10 samples per batch). Total procedural blanks estimated by

175

176

177

178

179

180

181

182

183

184
185

isotope dilution yielded values between 20 and 30 ng, representing less than 0.5% of the
total Ca. For carbon and oxygen measurements, foraminifera shells were crushed and then
cleaned by adding 10 drops of methanol, ultrasonicated for 30 seconds, and finally dried
down in a laminar flow hood.

186

187

188

189

190

**2.4. Analytical methods**

191

**2.4.1. Calcium isotopes**

192

Ca isotope analyses were performed on a Thermo Scientific Triton Plus Thermal Ionization
Mass Spectrometer using a conventional double $^{42}$Ca-$^{43}$Ca spike technique to correct for
instrumental mass fractionation. The double spike solution was prepared from two individual
single spikes (Oak Ridge National Laboratory) to obtain a $^{43}$Ca/$^{42}$Ca ratio close to the natural
ratio. The solution was then calibrated by repeated measurements of mixed standard-spike
solutions using the NIST SRM 915a standard, yielding a $^{43}$Ca/$^{42}$Ca ratio of 0.216380.

193

194

195

196

197

198

The spiked samples (5 to 7 µg of Ca) were loaded onto zone-refined Rhenium filament in 1 µL of concentrated nitric acid, dried out at an applied current of 0.9 A, and flashed at 2 A for a few seconds before being introduced in the thermal ionization source in a double filament assembly. Samples loaded into the mass spectrometer were heated using an automatic heating sequence. The ionization filament current was first increased to approximately 3000-3500 mA at a rate of 300 mA/min until a 7V $^{40}$Ca signal over a 10$^{11}$ Ω amplifier was obtained.

199

200

201

202

203

204

During analysis, the evaporation filament was kept at 0 mA as the heat generated by the ionization filament was generally sufficient to induce the evaporation of Ca, most likely as a high-mass Ca-rhenate (Davenport et al., 2022). The beam was then centered and focused and intensity was adjusted until a stable beam was obtained. The data were collected in static mode, with $^{40}$Ca, $^{42}$Ca, $^{43}$Ca, and $^{44}$Ca measured on the L2, C, H1, and H3 cups respectively, allowing simultaneous acquisition of $^{40}$Ca/$^{42}$Ca, $^{44}$Ca/$^{42}$Ca, and $^{43}$Ca/$^{42}$Ca. Each measurement consisted of 7 blocks of 10 cycles each with an integration time of 1s and 50

205

206

207

208

209

210
integrations per cycle. At the beginning of each block, a 60 seconds baseline measurement was performed. Lens focus and peak centering were performed every two blocks. Amplifier rotation was performed between each block, allowing all isotopes to be measured on the same "virtual" amplifier after a complete rotation cycle. The isobaric interference of <sup>40</sup>K on <sup>40</sup>Ca was monitored using the <sup>41</sup>K isotope and corrected online using <sup>40</sup>K/<sup>41</sup>K = 1.7384×10<sup>-3</sup>. However, these corrections were always negligible (<sup>40</sup>K/<sup>40</sup>Ca < 10<sup>-6</sup>).

The iterative correction algorithm of (Heuser et al., 2002) was used to correct <sup>44</sup>Ca/<sup>40</sup>Ca samples ratio from instrumental mass fractionation. Corrected <sup>44</sup>Ca/<sup>40</sup>Ca ratios are reported here in the conventional delta notation in permil (‰) relative to the NIST SRM 915a standard.

The instrumental reproducibility during a session was estimated from repeated measurements of the NIST SRM 915a standard during each analytical session, yielding values between ±0.09 to ± 0.11 ‰ (2 S.D).

Finally, the accuracy of the measurements (Figure 2) is assessed by repeat measurements (n = 3) of the seawater standard (OSIL). The average value obtained for all the measurements of the δ<sup>44/40</sup>Ca of seawater is 1.84 ± 0.01 ‰ (2 S.D.). This value is in agreement with the values reported for seawater by TIMS analyzes (e.g. Farkaš et al., 2007).

### 2.4.2. Carbon and oxygen isotopes

Carbon and oxygen isotope analyses were performed at the SSMIM (Isotope Mass Spectrometry Service) of the National Museum of Natural History, Paris, France. Between 40 and 50 µg of carbonate fragments were analyzed using a Delta V Advantage isotope ratio gas mass spectrometer (Thermo Fisher Scientific) directly coupled to Kiel-IV carbonate preparation device (reaction at 70 °C under vacuum). The external precision for oxygen data carbon was better than 0.05 ‰ and 0.03 ‰ respectively. The external precision was based upon repeated measurements (n = 16) of internal lab standard (Carrara marble) calibrated via NIST 19 to the V-PDB (Vienna Pee Dee Belemnite) scale. The external precision represents Carrara marble measurement conducted during our measurement session. The data are reported in delta notation (δ<sup>13</sup>C and δ<sup>18</sup>O) relative to V-PDB.
239

**2.4.3. Major and trace element concentration**

240

241

242

243

244

245

246

247

248

249

250

251

252

For the determination of Ca, Mg, Mn, Sr and Fe concentrations, samples aliquots were diluted with 2ml of 2 % HNO<sub>3</sub>, and elemental concentrations were measured using a Thermo Fisher Scientific Xseries II quadrupole ICP-MS at CRPG Nancy coupled to a Cetac ASX-110 autosampler. The reproducibility of Mg/Ca, Mn/Ca, Sr/Ca and Fe/Ca measurements was determined by repeated measurement of an in-house multi-elemental standard prepared from mono-elemental standards (SCP Science) to fit the typical composition of foraminiferal shells. Calibrations were run with each session to ensure the accuracy of both the Me/Ca ratios and the Ca concentration measurement. The relative standard deviations of Mg/Ca, Mn/Ca, Sr/Ca, and Fe/Ca ratios range from ~3 to ~19 %, ~1 to ~14 %, ~1 to ~10 %, and ~2 to 15%, respectively, and were found to be primarily dependent on sample size. The relative standard deviations were calculated based on repetitive measurement of our in-house standard.
253 **2.5. Age model**

254 The age model applied to the Danian interval and the K-Pg boundary of the Hole 1209 C is
255 based on age model of (Dinarès-Turell et al., 2014) already established by integrated
256 stratigraphy (magnetostratigraphy, biostratigraphy) and astronomical calibration
257 approach. For the Maastrichtian interval of this study, cyclostratigraphy and
258 magnetostratigraphy approach were not possible. Indeed, magnetization intensities for the
259 upper Maastrichtian at Shatsky Rise is weak due to the relatively high carbonate content. In
260 addition, sediment at Shatsky Rise for this time interval are poorly consolidated hampering
261 then a well preserved palaeomagnetic signal (Westerhold et al., 2008). Consequently, the
262 age model for the Maastrichtian interval is based on planktonic foraminiferal biostratigraphy
263 from ODP sites 1209 Hole C (Dameron et al., 2017b; Premoli Silva et al., 2005). The
264 absolute age of the planktonic foraminifer bio-events was estimated based on the
265 Geomagnetic Polarity and Astronomical Time Scale calibrations of Cretaceous and Cenozoic
266 planktonic foraminifer biostratigraphic zonal schemes (Coccioni and Premoli Silva, 2015).
267 Based on the known ages of the different bio-events and the average depth at which each
268 event occurs, a constant sedimentation rate was assumed to obtain a relative age for each
269 sample (Table 1).

270 **3. Results**

271 **3.1. Calcium isotopes**

272 The calcium isotopic results for planktonic and benthic species are listed in Appendix C.1
273 and C.2 and illustrated in Figure 3. The offset in δ<sup>44/40</sup>Ca between different planktonic species
274 has been corrected based on different samples where the two different species could be
275 used to generate δ<sup>44/40</sup>Ca values. Over the analyzed core section, δ<sup>44/40</sup>Ca varies between
276 0.37 and 0.80 ± 0.1 ‰ (2 S.D.) and between 0.25 and 0.82 ± 0.1 ‰ for species-specific
277 planktonic (δ<sup>44/40</sup>Ca<sub>P</sub>) and benthic (δ<sup>44/40</sup>Ca<sub>B</sub>) foraminifera, respectively. From the base of the
278 δ<sup>44/40</sup>Ca record until ~241 mbsf, δ<sup>44/40</sup>Ca<sub>B</sub> remains nearly constant, at ~0.75 ± 0.1 ‰.
279 $\delta^{44/40}\text{Ca}_P$ is also constant but with a slightly lighter average value of ~0.60 ± 0.1 ‰. Based on
280 the $\delta^{44/40}\text{Ca}_B$ record, two isotopic intervals can be distinguished across the uppermost
281 Maastrichtian and lowermost Danian, between ~239 and ~234 mbsf. These $\delta^{44/40}\text{Ca}$ intervals
282 are defined as shift from a reference value than a return to that same value: Interval I is from
283 ~238 to 236.5 mbsf and interval II is from 235.5 to 234 mbsf. These two isotopic intervals are
284 characterized by two consecutive negative shifts where $\delta^{44/40}\text{Ca}_B$ decreases by ~0.4 ‰. The
285 two negative excursions are separated by a positive excursion in $\delta^{44/40}\text{Ca}$ and the K-Pg
286 boundary is located at the base of interval II and is characterized by transient heavier
287 $\delta^{44/40}\text{Ca}$ values. Both the benthic and planktonic records display the largest amplitude
288 fluctuations across the K-Pg transition, although the timing of the negative excursions is
289 slightly asynchronous. Finally, for the last three meters of the record, both $\delta^{44/40}\text{Ca}_P$ and
290 $\delta^{44/40}\text{Ca}_B$ show a decrease of approximately 0.3 ‰. Overall, the benthic $\delta^{44/40}\text{Ca}_B$ record
291 seems to show larger variations (~0.6 ‰) than the noisier planktonic record which seems to
292 follow a long-term decrease of ~0.4 ‰. For these last three meters of the record, the
293 $\delta^{44/40}\text{Ca}_P$ value present in Figure 3 are corrected for isotopic offset.

294 Nevertheless, during the studied period, the two curves exhibit broadly similar trends,
295 implying similar first-order controls on the $\delta^{44/40}\text{Ca}$ composition of foraminifers in deep and
296 shallow environments.

297 **3.2. Carbon and oxygen isotopes**

298 Carbon and oxygen isotopic composition for planktonic species are shown in Figure 3 and
299 summarized in Appendix C.1. Over the analyzed core section, $\delta^{13}\text{C}$ varies between 1 and 3 ±
300 0.03 ‰ (2 S.D), and $\delta^{18}\text{O}$ varies -0.25 and -1.75 ± 0.05 ‰ (2 S.D). From the base until 242
301 mbsf, both $\delta^{13}\text{C}$ and $\delta^{18}\text{O}$ show heavier values compared to the rest of the record. $\delta^{13}\text{C}$
302 ranges from 3 to 2.5 ± 0.03 ‰ and $\delta^{18}\text{O}$ between -1 and -0.6 ± 0.05 ‰. During interval I, $\delta^{13}\text{C}$
303 and $\delta^{18}\text{O}$ are relatively uniform, respectively at 1.25 ± 0.03 ‰ and -1.25 ± 0.05 ‰. This
304 interval is characterized by the lowest carbon isotope values of the studied period. During
305 interval II, the $\delta^{13}\text{C}$ record shows a rapid increase and reaches its maximum value of 2.25 ±
0.03 ‰ at 234.6 mbsf. Concurrently, the δ¹⁸O record decreases to reach a minimum value of -1.5 ± 0.05 ‰ at 235 mbsf. Higher up in the section, the δ¹³C record is marked by a return to lighter compositions 1.5 ± 0.03 ‰ and δ¹⁸O decreases from -1.25 to -0.6 ± 0.05 ‰.

### 3.3. Trace elements

Trace element data are summarised in Appendix C.1 and C.2. For planktonic foraminifera, Mg/Ca ranges from 4 to 19 mmol/mol, with most samples lying below 10 mmol/mol. Mn/Ca ranges from 7×10⁻³ to 2×10⁻¹ mmol/mol (except one sample at 2.03 mmol/mol), and Sr/Ca from 0.6 to 1.2 mmol/mol. For benthic foraminifera, Mg/Ca ranges from 3 to 19.79 mmol/mol (excluding one sample at 47 mmol/mol), with a majority of samples falling below 6 mmol/mol. Mn/Ca ranges from 4×10⁻³ to 2×10⁻¹ mmol/mol, Sr/Ca from 0.8 to 1.4 mmol/mol, and Fe/Ca from 6.6×10⁻⁴ to 3.5×10⁻² mmol/mol.

Crossplots of trace element ratios with δ⁴⁴/⁴⁰Ca and different Me/Ca were used to evaluate the effect of diagenesis on the isotopic signals. As shown in Figure 4, no obvious correlation was found between isotopic compositions and trace element ratios diagnostic of diagenesis, with R² always < 0.09. The scatter plot of δ⁴⁴/⁴⁰Ca vs Sr/Ca is the only noticeable exception, with R² > 0.3.

Figure 4 shows that planktic and benthic Sr/Ca records (Sr/Caₚ and Sr/Ca₉) exhibit significant fluctuations, the most notable being a negative excursion of 0.3 mmol/mol in Sr/Caₚ during the dissolution interval. Furthermore, the Sr/Ca record show a coupled evolution with δ⁴⁴/⁴⁰Caₚ and δ⁴⁴/⁴⁰Ca₉ in some stratigraphic intervals. In particular, during interval II, a clear positive excursion in Sr/Ca₉ and Sr/Caₚ is observed together with the second negative excursion in δ⁴⁴/⁴⁰Ca₉. A similar positive excursion in Sr/Ca₉ is also observed during interval I, although this excursion is of lesser magnitude and constrained by a single point.
332 **4 Discussion**

333 **4.1 Diagenetic overprint: open vs. closed system conditions**

334 Secondary processes such as authigenic carbonate precipitation and/or isotopic equilibration
335 during recrystallization can lead to modifications of the primary geochemical signal. These
336 diagenetic effects must therefore be assessed before exploring the paleoenvironmental
337 implications of the calcium isotope record. As our samples are hand-picked foraminifera, we
338 likely avoided the substantial contribution of authigenic calcite (Fantle and Ridgwell, 2020).
339 The examination of test preservation using SEM images revealed signs of nano-scale
340 cemented overgrowths over the tests as well as the presence of coccoliths attached to
341 planktonic foraminifera (Appendix A). However, this non-foraminiferal calcite was likely
342 removed during the cleaning steps. In some rare cases, we also observed secondary calcite
343 overgrowth filling the tests of some benthic foraminifera and these tests were eliminated
344 during picking. Lastly, SEM observations revealed signs of recrystallization typical of pre-
345 Quaternary samples. However, the benthic and planktonic signals are not strictly identical at
346 a given stratigraphic level, arguing against complete equilibration during recrystallization,
347 despite possible fast reequilibration of calcium isotopes in biogenic calcite (Chanda et al.,
348 2019).

349 The impact of early marine diagenesis on the chemical and isotopic compositions of
350 carbonate sediments depends on the extent to which diagenesis is dominated by the
351 chemistry of the fluid (fluid-buffered) or the chemistry of the sediments (sediment-buffered).
352 In an open system, under fluid-buffered conditions, if alteration involves seawater or
353 seawater-derived fluids, the recrystallization of carbonate minerals is expected to produce an
354 increase in δ<sup>44/40</sup>Ca (Higgins et al., 2018). This is because recrystallization likely occurs at
355 isotopic equilibrium with the diagenetic fluid, with little to no Ca isotopic fractionation
356 (DePaolo, 2011; Fantle and DePaolo, 2007; Higgins et al., 2018). Overall, the current dataset
357 reveals significantly lower δ<sup>44/40</sup>Ca than values measured from carbonate minerals interpreted
to have formed under fluid-buffered condition (with δ<sup>44/40</sup>Ca values altered by up to +1.5 ‰;

Ahm et al., 2018) and show a good agreement with calcite δ<sup>44/40</sup>Ca values reported through

the Neogene (from 0.3 to ~1 ± 0.1 ‰; Farkaš et al., 2007 and reference therein) and in the

late Paleocene *Morozovella* specie (Kitch et al., 2021).

Carbonate sediments that undergo open-system diagenetic alteration tend to be characterized by Sr/Ca ratios lower than primary calcite (e.g. Baker et al., 1982; Kozdon et al., 2013; Regenberg, 2007). During calcite dissolution in the sediment column, Sr²⁺ is released from marine carbonates into pore fluids and then excluded during subsequent inorganic calcite re-precipitation from the same pore fluids (Edgar et al., 2015). This leads to progressively lower Sr/Ca ratios in marine carbonates with increasing diagenetic alteration.

Despite being slightly lower than existing K-Pg data (Henehan et al., 2019), the measured Sr/Ca<sub>B</sub> and Sr/Ca<sub>P</sub> are in most cases higher than 1 mmol/mol, suggesting little to no diagenetic overprint (Edgar et al., 2015; Kozdon et al., 2013). This with the exception of the dissolution interval where planktonic foraminifers show a Sr/Ca ratio lower than 1 mmol/mol (Figure 4). Alteration under fluid-buffered conditions is also expected to produce lower δ¹³C, as well as high Mg/Ca (e.g. Dutton et al., 2005), Mn/Ca (e.g. Brand and Veizer, 1980), and lower δ¹⁸O (e.g. Regenberg, 2007). However, with the exception of a few isolated samples with anomalously high Mn/Ca and/or high Mg/Ca ratios that were discarded from the dataset, crossplots of δ⁴⁴/⁴⁰Ca with isotopic/elemental tracers (δ¹³C, δ¹⁸O, Sr, Mg, Mn) did not reveal any prominent diagenetic signal (e.g. Lau et al., 2017). The measured Mg/Ca ratios are also in agreement with existing data and show no evidence for a diagenetic overprint (e.g. Henehan et al., 2019;

Figure 4; Appendix C.1 and C.2). Collectively, these observations suggest that if the tests recrystallized, they did so under sediment buffered conditions in deep sea carbonate-rich sediments rather than under open system fluid-buffered conditions (Fantle and DePaolo, 2007). Alteration under such conditions likely preserved the primary geochemical signal as
384

the diagenetic phase is expected to inherit the isotopic composition of the primary sediment

385

(Higgins et al., 2018). As a result, we conclude that the observed trends in δ<sup>44/40</sup>Ca reflect

386

primary calcite signals.

387

## 4.2 Seawater calcium isotope changes across the K-Pg transition?

388

In the discussion section, we first focus on the benthic signal as the uppermost Maastrichtian

389

section of the planktic signal can be affected by a possible Danian contamination at the ODP

390

Site 1209, Hole C (Dameron et al., 2017). The second reason is that the benthic signal is

391

more likely to be representative of broader scale processes than the planktonic signal

392

because planktonic signal is more influenced by ocean-atmosphere exchange as well as

393

photosynthesis and respiration.

394

The most prominent feature of our benthic record consists of a succession of two rapid (ca.

395

100 to 400 kyr long) negative excursions in δ<sup>44/40</sup>Ca<sub>B</sub> (intervals I and II) from the late

396

Maastrichtian to the early Paleogene, while the K-Pg transition itself is characterized by

397

transient heavier values of δ<sup>44/40</sup>Ca<sub>B</sub> (Figure 7). These excursions can be viewed as reflecting

398

either the variations of δ<sup>44/40</sup>Ca<sub>sw</sub> or changes in the Ca isotopic fractionation factor between

399

calcite and seawater. Using the δ<sup>44/40</sup>Ca record as a proxy for δ<sup>44/40</sup>Ca<sub>sw</sub> to constrain changes

400

in the oceanic Ca cycle is, however, challenging on short timescales (≤ 100 ka) because of

401

the long residence time of Ca in the ocean (~1Ma; Broecker and Peng, 1982). As a result,

402

calcium isotopes are unlikely to respond to transient disequilibrium in the oceanic Ca budget

403

unless the forcing is large enough to produce extreme variations in the Ca concentration of

404

seawater (Du Vivier et al., 2015; Hinojosa et al., 2012; Jost et al., 2017; Kitch et al., 2021;

405

Linzmeier et al., 2020). To explore the possibility that the δ<sup>44/40</sup>Ca record reflects changes in

406

the calcium isotopic composition of seawater, we use a simplified box model of the global

407

calcium cycle. In this model, calcium inputs to the ocean are combined as a single flux (F<sub>in</sub>)

408

and deposition of biologically precipitated calcium carbonates in sediments is considered the
sole output (Fout). The size of the calcium reservoir (Nca) and calcium isotopic composition (δ$^{44/40}$Ca$_{sw}$) change with time according to the following equations:

$$411 \quad \frac{dN_{Ca}}{dt} = F_{in} - F_{out} \quad \text{Eq. (1)}$$

$$412 \quad \frac{d(N_{Ca}\delta^{44/40}Ca_{SW})}{dt} = F_{in}\delta^{44/40}Ca_{W} - F_{carb}(\delta^{44/40}Ca_{SW} + \Delta_{carb}) \quad \text{Eq. (2)}$$

Following (Sime et al., 2007), these equations can also be rewritten as:

$$415 \quad \frac{dN'_{Ca}}{dt} = \frac{-1}{\Delta_{carb}} \left( \frac{\delta^{44/40}Ca_{W} - \delta^{44/40}Ca_{SW} - \Delta_{carb}}{\tau} \right) + \left( \frac{N'_{Ca}}{\Delta_{carb}} \right) \left( \frac{d\delta^{44/40}Ca_{SW}}{dt} \right) \quad \text{Eq. (3)}$$

$$416 \quad \frac{F_{out}}{F_{in}} = 1 - \tau \left( \frac{dN'_{Ca}}{dt} \right) \quad \text{Eq. (4)}$$

where N'ca is the oceanic calcium reservoir relative to the initial state (N'ca = Nca(t)/Nca(t0)), Δcarb represents the isotopic fractionation between seawater and precipitated carbonate, τ is the residence time of Ca in the ocean (τ = Nca/Fout) and δ$^{44/40}$Ca$_{W}$ is the isotopic composition of the weathering flux. The calcium isotopic composition of seawater (δ$^{44/40}$Ca$_{SW}$) is given by:

$$421 \quad \delta^{44/40}Ca_{SW} = \delta^{44/40}Ca_{B} + \Delta_{carb} \quad \text{Eq. (5)}$$

A smoothed version of the data, produced using a Monte Carlo approach, was used in the model calculations. δ$^{44/40}$Ca$_{W}$ was considered constant and equal to the mean of the δ$^{44/40}$Ca$_{B}$ fitted curve between 240 and 223 mbsf (i.e. δ$^{44/40}$Ca$_{W}$ = 0.54 ‰). We further assume that the carbonate δ$^{44/40}$Ca record is representative of the long-term average isotopic composition of carbonate sediment during the late Cretaceous-early Paleogene. Finally, Δcarb was set at 1.12 ‰ (e.g. Du Vivier et al., 2015).

Equations (3) and (4) were solved numerically using a 4th-order Runge Kutta method and the results are shown in Figure 5. Using the lowest possible value of τ (0.7 Ma; see Table 2), our
430

results show that a fluctuation of at least 40 % in the calcium concentration of seawater
would be needed to explain the rapid excursions in the δ<sup>44/40</sup>Ca signal around the K-Pg
boundary. This change in concentration requires major fluctuations of the F<sub>out</sub>/F<sub>in</sub> ratio, from -
0.5 to +2, which is well beyond any plausible variation of either F<sub>in</sub> or F<sub>out</sub>. Assuming that F<sub>in</sub>
remained constant (even within 50%) throughout the studied period, such a change would
require an inversion of F<sub>out</sub>, which is to say a massive worldwide dissolution event. Only a
residence time shorter than 300 ky produces reasonable F<sub>out</sub> fluctuations, but this is much
lower than the proposed range of τ values and therefore oceanic calcium concentration
(Table 2).

431

432

433

434

435

436

437

438

439

Remains the possibility that δ<sup>44/40</sup>Ca<sub>w</sub> changed across the K-Pg boundary. The strongest argument comes from Li isotopes that were interpreted as reflecting increased denudation rates due to acid rains on land (Misra and Froelich, 2012). This would increase the F<sub>in</sub> flux and could modify the average δ<sup>44/40</sup>Ca<sub>w</sub> value. However, the δ<sup>44/40</sup>Ca<sub>w</sub> value of the model (0.54 ‰) is close to the carbonate endmember. In modern rivers, the carbonate and silicate endmembers are evaluated at 0.60 ‰ and 0.94 ‰, respectively (Fantle and Tipper, 2014).
As a result, acid-rain that would accelerate carbonate dissolution would be unlikely to change the isotopic ratio of calcium delivered with the weathering flux. In addition, the step observed in the Li record is very different from the oscillations of the δ<sup>44/40</sup>Ca record at Shatsky Rise.
For those reasons, we do not explore further the possibility of an evolving δ<sup>44/40</sup>Ca<sub>w</sub>.

440

441

442

443

444

445

446

447

448

449

We conclude that the δ<sup>44/40</sup>Ca record cannot be explained by global changes in the oceanic
Ca budget.

450

451

## 4.3 Factors influencing the calcium isotopic fractionation factor between seawater and calcite

452

453

Calcium isotopes were initially suggested to be a paleotemperature proxy based on
paleoenvironmental investigations of glacial-interglacial cycles (Hippler et al., 2006). The link
between temperature and δ<sup>44/40</sup>Ca was observed in synthetic calcium carbonates (Gussone

454

455
456

et al., 2003; Marriott et al., 2004), culture experiments (e.g. Gussone et al., 2003), and in core top calibrations (Hippler et al., 2009).

457

458

Two ranges of calcium isotope sensitivity to temperature are documented in the literature. The first one is documented for some planktonic foraminifera species and is around ~0.02 ‰ per °C (e.g. *Orbulina universa*; Gussone et al., 2003). However, a much higher sensitivity up to 0.24 ‰ per °C has been documented for the planktonic foraminifera species *Globigerinoides sacculifer* (Gussone et al., 2003; Nägler et al., 2000). According to Gussone et al. (2003), such a difference in calcium isotope sensitivity to temperature between *Orbulina universa* and *Globigerinoides sacculifer* can be related the fact that *Globigerinoides sacculifer* actively dehydrates the Ca²⁺-aquocomplex before calcification takes place. Because no of calcium isotope sensitivity to temperature exists for the studied foraminiferal species here, we test the two range of temperature sensitivity.

459

460

461

462

463

464

465

466

467

468

469

470

471

472

473

474

Assuming a temperature effect of ~0.02 ‰ per °C (Gussone et al., 2003), a δ<sup>44/40</sup>Ca decrease of 0.4 ‰ would require a temperature decrease of 20°C, which is inconsistent with paleotemperature estimates during the studied period (e.g. Barnet et al., 2018; Li and Keller, 1998). Using a temperature dependence of 0.24 ‰ per °C (Gussone et al., 2003), a decrease of 0.4 ‰ in δ<sup>44/40</sup>Ca could be explained by a temperature decrease of about 1.7 °C in the tropical Pacific ocean, which suggests that this higher sensitivity of Ca isotopes to temperature could produce the observed δ<sup>44/40</sup>Ca decrease of 0.4 ‰.

475

476

477

478

479

480

481

Furthermore, δ<sup>44/40</sup>Ca is positively correlated with temperature, while δ<sup>18</sup>O is negatively correlated with temperature (Gussone et al., 2003) so that an anti-correlation of their values should be expected in ODP 1209. However, comparison of our δ<sup>44/40</sup>Ca<sub>B</sub> values with published δ<sup>18</sup>O benthic values (Dameron et al., 2017; Westerhold et al., 2011, see Figure 6 for the expected theoretical relationship between δ<sup>44/40</sup>Ca, δ<sup>18</sup>O and temperature) does not always show such anti-correlation. In conclusion, the observed δ<sup>44/40</sup>Ca<sub>B</sub> variations cannot be explained by temperature variations regardless of which slope of calcium isotopes sensitivity
482

to temperate we use. Finally, more recent culture experiments suggest that the temperature
dependence of δ<sup>44/40</sup>Ca is not always observed (Kısakürek et al., 2011). We therefore
conclude that the observed δ<sup>44/40</sup>Ca variations do not result from temperature changes.

483

484

485

A second parameter that could explain a change in δ<sup>44/40</sup>Ca values is the presence of different vital effects. Vital effects include all the kinetic and physiological processes associated with biomineralization. As a result, different organisms can produce calcium carbonates with different δ<sup>44/40</sup>Ca values from a homogeneous seawater reservoir (Farkaš et al., 2007; Gussone et al., 2005; Kısakürek et al., 2011). A control on δ<sup>44/40</sup>Ca exists depending on whether organisms produce calcite or aragonite, and a wide range of values has been reported for foraminifera. Our planktonic record consists of calcitic monospecific assemblages. As a result, the present record cannot be impacted by changes in carbonate mineralogy (calcite vs. aragonite). In addition, vital effects can only affect this δ<sup>44/40</sup>Ca record if each species fractionates calcium isotopes differently, an effect that we have corrected for when possible for the planktonic record. The offset between *Chiloguembelina morsei* and *Praemurica inconstans* was corrected based on paired δ<sup>44/40</sup>Ca measurements (n = 2), yielding an offset value of 0.3 ± 0.1 ‰. However, no suitable sample was found to correct for possible offsets neither between *Pseudoguembelina costulata* and *Woodringina claytonensis* nor between *Woodringina claytonensis* and *Chiloguembelina morsei* on the other hand. As a result, it is not entirely ruled out that some variations in the planktonic signal might be caused by vital effects. In contrast, because the benthic record is composed of the same species throughout, vital effects cannot account for changes in δ<sup>44/40</sup>Ca<sub>B</sub>.

486

487

488

489

490

491

492

493

494

495

496

497

498

499

500

501

502

503

504

505

506

507

508

If temperature and vital effects do not explain the observed changes in δ<sup>44/40</sup>Ca values, it can prove useful to believe in different investigations of calcium isotope fractionations during inorganic and biogenic carbonate precipitation. The relationship between Δ<sub>carb</sub> and precipitation rate has been widely studied through inorganic calcite precipitation experiments (e.g. Lemarchand et al., 2004; Tang et al., 2008) and modeling (DePaolo, 2011). Though not all experiments agree, there tends to be an agreement on the lack of significant isotopic
fractionation at equilibrium (e.g. Harrison et al., 2023). When the precipitation rate increases, however, the system approaches a kinetic limit where the magnitude of $\Delta_{\text{carb}}$ increases (DePaolo, 2011). Under such conditions, both inorganic and biogenic calcite exhibit negatively correlated $\delta^{44/40}\text{Ca}$ and Sr/Ca (Kısakürek et al., 2011; Tang et al., 2008), which is consistent with the coupled Sr/Ca - $\delta^{44/40}\text{Ca}$ excursions observed during interval I and II.

The $\delta^{44/40}\text{Ca}$ value of calcite is also sensitive to changes in carbonate ion concentration, though the respective contributions of precipitation rates vs. carbonate ion concentration can be difficult to evaluate. In most experiments, the rate is estimated based on $[\text{CO}_3^{2-}]$ and both parameters are correlated in inorganic carbonate precipitation experiments (e.g. Lemarchand et al., 2004, Tang et al., 2008). Inorganic experimental studies also show a direct control of the $[\text{CO}_3^{2-}]$ concentration on $\delta^{44/40}\text{Ca}$, with higher fractionation occurring at higher $\text{CO}_3/\text{Ca}$ ratios (e.g. Nielsen et al., 2012). In biogenic benthic foraminiferal calcite, Sr/Ca$_B$ has also been shown to be sensitive to carbonate ion concentration (Yu et al., 2014). Collectively, these observations suggest that fluctuations of the $\delta^{44/40}\text{Ca}$ records can plausibly relate to carbonate ion concentration and/or precipitation rate and thus to the parameters controlling the saturation state of the solution. These parameters share the same effects on calcium isotopes and are likely to be correlated in the ocean carbonate system. Such an interpretation is also consistent with previous interpretations of the marine calcium isotopic records during biocalcification crises (Du Vivier et al., 2015; Kitch et al., 2022, 2021; Linzmeier et al., 2020).

## 4.4 Evolution of the seawater chemistry across The K-Pg transition

As shown in Figure 7, two negative excursions are observed in the benthic record across the K-Pg transition. Interval I, just before the boundary, is characterized by a negative $\delta^{44/40}\text{Ca}_B$ excursion, possibly associated with a positive peak in Sr/Ca. This interval is also distinguished by a local minimum in the foraminifer fragmentation index, which indicates a decrease in planktonic carbonate dissolution. Together with the negative $\delta^{44/40}\text{Ca}_B$ excursion,
the data suggest an increase in deep water carbonate ion concentrations. Because the lysocline is defined by the depth at which the saturation state of calcite equals 1, it is logical to interpret such negative excursion of $\delta^{44/40}\text{Ca}_\text{B}$ as a deepening of the lysocline, and thus of the CCD.. Though the two limits do not necessarily evolve simultaneously, they evolve similarly on timescales of more than 10 ky (Zachos et al., 2005). The end of this negative excursion coincides with an increase in the fragmentation index, as the deep water saturation state returns to the level observed during the Maastrichtian dissolution interval. Despite the respective approximations in the age models, we note a first order agreement between the calcium isotope records from Shatsky Rise and Seymour Island (Figure 8, Linzmeier et al., 2020). It should also be noted that the $\delta^{44/40}\text{Ca}$ positive shift between intervals I and II, recorded by both planktonic and benthic foraminifera could correspond to the ocean acidification at the end of the Cretaceous (Figure 8, Henehan et al., 2019).

During interval II, the fragmentation index decreases again and a clear negative excursion in $\delta^{44/40}\text{Ca}_\text{B}$ coupled with a major positive excursion in Sr/Ca is observed while a long-term decrease is observed in $\delta^{44/40}\text{Ca}_\text{P}$. Such a relationship between $\delta^{44/40}\text{Ca}$ and Sr/Ca is consistent with a change in precipitation rates (Kısakürek et al., 2011; Tang et al., 2008) or carbonate concentration (Yu et al., 2014), suggesting that, like interval I, the negative $\delta^{44/40}\text{Ca}$ excursion of interval II may reflect an increase in the saturation state of seawater.

Based on our age model, the timing of $\delta^{44/40}\text{Ca}_\text{B}$ and Sr/Ca$_\text{B}$ excursions (400 kyr before and after the K-Pg boundary) is consistent with a scenario controlled by the Deccan pulses through the Cretaceous-Paleogene transition (Figure 9, Schoene et al., 2019; 2021). This, in turn, suggests that the Deccan trap eruptions played a central role in controlling the marine carbonate chemistry during this period of transition between the Mesozoic and the Cenozoic.

It could appear counterintuitive to interpret episodes of alkalinization as a consequence of increased volcanism. On geologically rapid time scales (< 10 Kyr), injection of atmospheric $\text{CO}_2$ into seawater lead to a decrease seawater pH and $\text{CO}_3^{2-}$ concentrations (e.g. Hinojosa
et al., 2012; Zeebe, 2012). An transient episode of oceanic acidification should thus be
expected during each Deccan volcanic eruption, as CO₂ is actively emitted into the
atmosphere. However, as trap emplacement likely occurred as discrete rapid pulses (Chenet
et al., 2007; Keller et al., 2020), the acidification itself would likely be short-lived, lasting only
about 10 ky. We argue that such short-term acidification events, if they occurred, cannot be
resolved in our record. This appears similar for exemple to what is modeled for seawater
chemistry changes in response to the CAMP eruptions at the Triassic-Jurassic boundary
(Paris et al. 2016). The acidification itself only lasts for the first few thousand years, ,
sustained by higher atmospheric pCO₂ before weathering feedbacks. The consequent
increase in temperature and precipitation compensate for the drop in pH and saturation state,
leading thus to an alkalinization of the ocean. Thus, we suggest that the δ⁴⁴/⁴⁰Ca record
reflects alkalinization events of a few hundred of thousand years that followed major volcanic
pulses. On such time scales, the atmosphere would be left with a higher level of pCO₂
(Zachos et al., 2008) promoting weathering of carbonate and silicate rocks on land (e.g.
Zeebe and Zachos, 2013). The latter must have induced an increase in weathering influx of
Ca²⁺ and CO₃²⁻ into seawater leading to a period of ocean alkalinization, the combination of
which resulted in an increase of the saturation state (Paris et al., 2016; Penman and Zachos,
2018). As a result, deep water masses are expected to become more saturated with respect
to calcite. This would likely induce a deepening of the lysocline, consistent with the evolution
of the fragmentation index and the benthic δ⁴⁴/⁴⁰CaB and Sr/CaB records at Shatsky Rise. An
increase in the calcite saturation state can then lead to changes in the calcium fractionation
factor as a response to changes in precipitation rate or carbonate ion concentration and
ultimately produce the observed negative shifts in the δ⁴⁴/⁴⁰Ca record.

It is also interesting to note that, if we follow Schoene et al. (2021) scenario, the only Deccan
peak associated with an increase in δ⁴⁴/⁴⁰CaB record is the peak occurring very close to the
K-Pg boundary. This positive δ⁴⁴/⁴⁰CaB is also concomitant with the decrease in pH
587

suggested by the boron isotope composite record of Henehan et al. (2019) (Figure 8), that occurs at the K-Pg boundary and thus as the meteorite impact occurs.

588

589

Although the planktonic signal is noisier than the benthic, $\delta^{44/40}\text{Ca}_P$ and Sr/Ca$_P$ are overall similar to the benthic records, the Maastrichtian dissolution interval asides. The most notable Sr/Ca$_P$ feature is a negative excursion during the dissolution interval, consistent with a poorer preservation state and high fragmentation index. Thus, Sr/Ca$_P$ possibly reflects a decrease in the surface water saturation state and/or increased dissolution of planktonic foraminifera as they sink to the bottom of the ocean. The latter is consistent with the $\delta^{44/40}\text{Ca}_B$ record, which suggests that the saturation state of the deep ocean remains relatively low at the end of the dissolution interval. During the end-Cretaceous (interval I), the $\delta^{44/40}\text{Ca}$ negative excursion occurs in the surface ocean before the deep ocean. This desynchronization may be explained by a faster response of the surface ocean to the environmental perturbations and its lower stability compared to the deep ocean, combined with the recovery from the oceanic changes at the origin of the late Maastrichtian dissolution interval (Dameron et al. 2017). In addition, surface processes associated with the Deccan and the Chixculub impact might preferentially affect planktonic foraminifera (Hernandez Nava et al., 2021; Pope et al., 1997).

590

591

592

593

594

595

596

597

598

599

600

601

602

603

604

605

606

607

608

609

610

611

612

613

During the Lower-Paleocene (post-K-Pg extinction interval, i.e. interval II), the long-term decrease in $\delta^{44/40}\text{Ca}_P$ could reflect a rebound in the surface seawater super-saturation state following surface ocean acidification and extinction amongst marine calcifiers (Figure 8, Henehan et al., 2019). This increase in surface ocean saturation state is thus be a consequence of the K-Pg calcareous planktonic mass extinction which induce a deacrease in the uptake of $\text{Ca}^{2+}$ and $\text{CO}_3^{2-}$ from the surface ocean (e.g. Caldeira et al., 1990), a process also termed biological compensation (Boudreau et al., 2018). Thus, a long-term alkalinity accumulation could result from the pelagic calcifier extinction, which would sharply decrease the $\text{CaCO}_3/\text{C}_{\text{org}}$ rain ratio to the deep sea in combination with increasing river alkalinity delivery (Bralower et al., 2020) during periodes of Deccan $\text{CO}_2$ emissions. Although biological compensation is relatively instantaneous by comparison to continental weathering,
614
615

both compensation mechanism act in the same direction in order to counterbalance ocean acidification.

616
617
618
619
620
621

Going beyond these different scenarios, the consistent evolutions in δ<sup>44/40</sup>Ca observed in benthic and planktonic foraminifera suggest that the shallow and deep oceans were impacted by alkalinization events. Volcanic pulses associated with the emplacement of the Deccan traps and the role of the carbonate pump in exporting alkalinity from the surface to the deep ocean thus appear to be a major cause of the observed calcium isotope changes during the end-Cretaceous and the Paleocene post-extinction at Shatsky Rise.

622

# 5 Conclusions

623
624
625
626
627
628
629
630
631
632
633
634
635
636
637
638
639

Coupled planktonic and benthic δ<sup>44/40</sup>Ca data document clear isotopic trends across the K-Pg boundary that are not the result of a strong diagenetic alteration. Numerical calculations suggest that the observed changes are too rapid to be interpreted in terms of changes in calcium fluxes. Instead, our results are better interpreted in terms of seawater chemistry, more specifically of the carbonate system, as suggested by laboratory experiments. The data presented in this study suggest a sequence of deep ocean alkalinization events induced by Deccan volcanism and during which the calcium fractionation factor changed in response to carbonate chemistry (CO<sub>3</sub><sup>2-</sup> concentration and/or precipitation rates). The Deccan CO<sub>2</sub> emissions right before and after the K-Pg likely triggered episodes of deep ocean alkalinization with associated changes in the CCD due to increased continental weathering on timescales longer than 100 ky. Following the K-Pg mass extinction of the calcareous plankton, the delayed recovery of the biological carbonate pump, along with the enhanced continental weathering flux produced by the Deccan CO<sub>2</sub> emissions, likely caused a rise in ocean alkalinity. This scenario, combining biotic and environmental changes along the Cretaceous-Paleogene transition, explains the similarities and differences of the benthic and planktonic calcium isotope records and provides further knowledge on the consequences of the K-Pg boundary events.
640
641
642 Acknowledgements: Financial support for this work came from an CNRS-INSU Syster
643 program (grant to G.P. and A.J.) as well as from the OTELo jeunes chercheurs and IODP-
644 France programs (grants to G.P.). At CRPG, the authors would like to thank Aimeryc
645 Schumacher, Christiane Parmentier and Catherine Zimmerman for their active analytical help
646 with the X-series and the TIMS, respectively and their good care of the daily life in the IRISS
647 platform. The instruments used at CRPG are part of the RéGEF network. Maxime Piralla
648 provided invaluable assistance with the modelling, thank you Maxime. The authors woud
649 also like to thank Raphaël Pik who provided useful feedbacks throughout this project and
650 Francesca Falzoni for taxonomic clarifications on the species used in this work. The authors
651 acknowledge help from Sylvain Pont (MNHN IMPMC) and Alexandre Lethiers (SU CR2P) who
652 performed the SEM views and the plate view, respectively, as well as from Sooriadeo Dalliah
653 (MNHN CR2P) who helped processing the ODP samples and Denis Fiorillo (MNHN SSMIM)
654 for the acquisition of the carbon and oxygen isotope data. The authors would like to thank
655 the editor and the reviewers whose careful reading and constructive criticism helped to
656 improve the article. This is CRPG contribution # 2841. Data are available on ORDaR -
657 OTELo Research Data Repository, DOI is : 10.24396/ORDAR-122.

658
659 Abramovich, S., Keller, G., 2003. Planktonic foraminiferal response to the latest Maastrichtian abrupt
660 warm event: a case study from South Atlantic DSDP Site 525A. Marine Micropaleontology 48,
661 225-249. https://doi.org/10.1016/S0377-8398(03)00021-5

662 Ahm, A.-S.C., Bjerrum, C.J., Blättler, C.L., Swart, P.K., Higgins, J.A., 2018. Quantifying early marine
diagenesis in shallow-water carbonate sediments. Geochimica et Cosmochimica Acta 236,
663 140-159. https://doi.org/10.1016/j.gca.2018.02.042

665 Alegret, L., Arreguín-Rodríguez, G.J., Trasviña-Moreno, C.A., Thomas, E., 2021. Turnover and stability
666 in the deep sea: Benthic foraminifera as tracers of Paleogene global change. Global and
667 Planetary Change 196, 103372. https://doi.org/10.1016/j.gloplacha.2020.103372

668 Alvarez, L.W., Alvarez, W., Asaro, F., Michel, H.V., 1980. Extraterrestrial Cause for the Cretaceous-
669 Tertiary Extinction. Science 208, 1095-1108. https://doi.org/10.1126/science.208.4448.1095

670 Arenillas, I., Arz, J.A., Grajales-Nishimura, J.M., Meléndez, A., Rojas-Consuegra, R., 2016. The
671 Chicxulub impact is synchronous with the planktonic foraminifera mass extinction at the
672 Cretaceous/Paleogene boundary: new evidence from the Moncada section, Cuba. Geologica
673 Acta 17.

674 Baker, P.A., Gieskes, J.M., Elderfield, H., 1982. Diagenesis of Carbonates in Deep-Sea Sediments--
675 Evidence From SR/CA Ratios and Interstitial Dissolved SR2+ Data 12.

676 Barnet, J.S.K., Littler, K., Kroon, D., Leng, M.J., Westerhold, T., Röhl, U., Zachos, J.C., 2018. A new
677 high-resolution chronology for the late Maastrichtian warming event: Establishing robust
678 temporal links with the onset of Deccan volcanism. Geology 46, 147-150.
679 https://doi.org/10.1130/G39771.1

680 Beerling, D.J., Lomax, B.H., Royer, D.L., Upchurch, G.R., Kump, L.R., 2002. An atmospheric pCO2
681 reconstruction across the Cretaceous-Tertiary boundary from leaf megafossils. Proceedings
682 of the National Academy of Sciences 99, 7836-7840.
683 https://doi.org/10.1073/pnas.122573099

684 Blättler, C.L., Jenkyns, H.C., Reynard, L.M., Henderson, G.M., 2011. Significant increases in global
685 weathering during Oceanic Anoxic Events 1a and 2 indicated by calcium isotopes. Earth and
686 Planetary Science Letters 309, 77-88. https://doi.org/10.1016/j.epsl.2011.06.029

687 Boudreau, B.P., Middelburg, J.J., Luo, Y., 2018. The role of calcification in carbonate compensation.
688 Nature Geosci 11, 894-900. https://doi.org/10.1038/s41561-018-0259-5

689 Boyle, E.A., 1981. Cadmium, zinc, copper, and barium in foraminifera tests. Earth and Planetary
690 Science Letters 53, 11-35. https://doi.org/10.1016/0012-821X(81)90022-4

691 Bralower, T.J., Cosmidis, J., Heaney, P.J., Kump, L.R., Morgan, J.V., Harper, D.T., Lyons, S.L., Freeman,
692 K.H., Grice, K., Wendler, J.E., Zachos, J.C., Artemieva, N., Chen, S.A., Gulick, S.P.S., House,
693 C.H., Jones, H.L., Lowery, C.M., Nims, C., Schaefer, B., Thomas, E., Vajda, V., 2020. Origin of a
694 global carbonate layer deposited in the aftermath of the Cretaceous-Paleogene boundary
695 impact. Earth and Planetary Science Letters 548, 116476.
696 https://doi.org/10.1016/j.epsl.2020.116476

697 Bralower, T.J., Premoli Silva, I., Malone, M.J. (Eds.), 2006. Proceedings of the Ocean Drilling Program,
698 198 Scientific Results, Proceedings of the Ocean Drilling Program. Ocean Drilling Program.
699 https://doi.org/10.2973/odp.proc.sr.198.2006

700 Brand, U., Veizer, J., 1980. Chemical Diagenesis of a Multicomponent Carbonate System--1: Trace
701 Elements. SEPM JSR Vol. 50. https://doi.org/10.1306/212F7BB7-2B24-11D7-
702 8648000102C1865D

703 Broecker, W.S., Peng, T.H., 1982. Tracers in the Sea 705.

704 Caldeira, K., Rampino, M.R., Volk, T., Zachos, J.C., 1990. Biogeochemical modeling at mass extinction
705 boundaries: Atmospheric carbon dioxide and ocean alkalinity at the K/T boundary, in:
706 Kauffman, E.G., Walliser, O.H. (Eds.), Extinction Events in Earth History, Lecture Notes in
707 Earth Sciences. Springer Berlin Heidelberg, Berlin, Heidelberg, pp. 333-345.
708 https://doi.org/10.1007/BFb0011156

709 Chanda, P., Gorski, C.A., Oakes, R.L., Fantle, M.S., 2019. Low temperature stable mineral
710 recrystallization of foraminiferal tests and implications for the fidelity of geochemical proxies.
711 Earth and Planetary Science Letters 506, 428-440.

712 https://doi.org/10.1016/j.epsl.2018.11.011

713 Chenet, A., Quidelleur, X., Fluteau, F., Courtillot, V., Bajpai, S., 2007. 40K-40Ar dating of the Main

714 Deccan large igneous province: Further evidence of KTB age and short duration. Earth and

715 Planetary Science Letters 263, 1-15. https://doi.org/10.1016/j.epsl.2007.07.011

716 Dameron, S.N., Leckie, R.M., Clark, K., MacLeod, K.G., Thomas, D.J., Lees, J.A., 2017. Extinction,

717 dissolution, and possible ocean acidification prior to the Cretaceous/Paleogene (K/Pg)

718 boundary in the tropical Pacific. Palaeogeography, Palaeoclimatology, Palaeoecology 485,

719 433-454. https://doi.org/10.1016/j.palaeo.2017.06.032

720 Davenport, J., Caro, G., France-Lanord, C., 2022. Decoupling of physical and chemical erosion in the

721 Himalayas revealed by radiogenic Ca isotopes. Geochimica et Cosmochimica Acta 338, 199-

722 219. https://doi.org/10.1016/j.gca.2022.10.031

723 DePaolo, D.J., 2011. Surface kinetic model for isotopic and trace element fractionation during

724 precipitation of calcite from aqueous solutions. Geochimica et Cosmochimica Acta 75, 1039-

725 1056. https://doi.org/10.1016/j.gca.2010.11.020

726 Du Vivier, A.D.C., Jacobson, A.D., Lehn, G.O., Selby, D., Hurtgen, M.T., Sageman, B.B., 2015. Ca

727 isotope stratigraphy across the Cenomanian-Turonian OAE 2: Links between volcanism,

728 seawater geochemistry, and the carbonate fractionation factor. Earth and Planetary Science

729 Letters 416, 121-131. https://doi.org/10.1016/j.epsl.2015.02.001

730 Dutton, A., Lohmann, K.C., Leckie, R.M., 2005. Insights from the Paleogene tropical Pacific:

731 Foraminiferal stable isotope and elemental results from Site 1209, Shatsky Rise: SITE 1209-

732 STABLE ISOTOPES AND Mg/Ca. Paleooceanography 20, n/a-n/a.

733 https://doi.org/10.1029/2004PA001098

734 Edgar, K.M., Anagnostou, E., Pearson, P.N., Foster, G.L., 2015. Assessing the impact of diagenesis on

735 δ11B, δ13C, δ18O, Sr/Ca and B/Ca values in fossil planktic foraminiferal calcite. Geochimica

736 et Cosmochimica Acta 166, 189-209. https://doi.org/10.1016/j.gca.2015.06.018

737 Fantle, M.S., DePaolo, D.J., 2007. Ca isotopes in carbonate sediment and pore fluid from ODP Site

738 807A: The Ca2+(aq)-calcite equilibrium fractionation factor and calcite recrystallization rates

739 in Pleistocene sediments. Geochimica et Cosmochimica Acta 71, 2524-2546.

740 https://doi.org/10.1016/j.gca.2007.03.006

741 Fantle, M.S., Ridgwell, A., 2020. Towards an understanding of the Ca isotopic signal related to ocean

742 acidification and alkalinity overshoots in the rock record. Chemical Geology 547, 119672.

743 https://doi.org/10.1016/j.chemgeo.2020.119672

744 Fantle, M.S., Tipper, E.T., 2014. Calcium isotopes in the global biogeochemical Ca cycle: Implications

745 for development of a Ca isotope proxy. Earth-Science Reviews 129, 148-177.

746 https://doi.org/10.1016/j.earscirev.2013.10.004

747 Farkaš, J., Böhm, F., Wallmann, K., Blenkinsop, J., Eisenhauer, A., van Geldern, R., Munnecke, A.,

748 Voigt, S., Veizer, J., 2007. Calcium isotope record of Phanerozoic oceans: Implications for

749 chemical evolution of seawater and its causative mechanisms. Geochimica et Cosmochimica

750 Acta 71, 5117-5134. https://doi.org/10.1016/j.gca.2007.09.004

751 Gilabert, V., Arz, J.A., Arenillas, I., Robinson, S.A., Ferrer, D., 2021. Influence of the Latest

752 Maastrichtian Warming Event on planktic foraminiferal assemblages and ocean carbonate

753 saturation at Caravaca, Spain. Cretaceous Research 125, 104844.

754 https://doi.org/10.1016/j.cretres.2021.104844

755 Gussone, N., Böhm, F., Eisenhauer, A., Dietzel, M., Heuser, A., Teichert, B.M.A., Reitner, J., Wörheide,

756 G., Dullo, W.-C., 2005. Calcium isotope fractionation in calcite and aragonite. Geochimica et

757 Cosmochimica Acta 69, 4485-4494. https://doi.org/10.1016/j.gca.2005.06.003

758 Gussone, N., Eisenhauer, A., Heuser, A., Dietzel, M., Bock, B., Böhm, F., Spero, H.J., Lea, D.W., Bijma,

759 J., Nägler, T.F., 2003. Model for kinetic effects on calcium isotope fractionation (δ44Ca) in

760 inorganic aragonite and cultured planktonic foraminifera. Geochimica et Cosmochimica Acta

761 67, 1375-1382. https://doi.org/10.1016/S0016-7037(02)01296-6
762 Gussone, N., Hönisch, B., Heuser, A., Eisenhauer, A., Spindler, M., Hemleben, C., 2009. A critical

763 evaluation of calcium isotope ratios in tests of planktonic foraminifers. Geochimica et

764 Cosmochimica Acta 73, 7241–7255. https://doi.org/10.1016/j.gca.2009.08.035

765 Hardie, L.A., 1996. Secular variation in seawater chemistry: An explanation for the coupled secular

766 variation in the mineralogies of marine limestones and potash evaporites over the past 600

767 m.y. Geol 24, 279. https://doi.org/10.1130/0091-7613(1996)024<0279:SVISCA>2.3.CO;2

768 Harrison, A.L., Heuser, A., Liebetrau, V., Eisenhauer, A., Schott, J., Mavromatis, V., 2023. Equilibrium

769 Ca isotope fractionation and the rates of isotope exchange in the calcite-fluid and aragonite-

770 fluid systems at 25 °C. Earth and Planetary Science Letters 603, 117985.

771 https://doi.org/10.1016/j.epsl.2022.117985

772 Henehan, M.J., Hull, P.M., Penman, D.E., Rae, J.W.B., Schmidt, D.N., 2016. Biogeochemical

773 significance of pelagic ecosystem function: an end-Cretaceous case study. Phil. Trans. R. Soc.

774 B 371, 20150510. https://doi.org/10.1098/rstb.2015.0510

775 Henehan, M.J., Ridgwell, A., Thomas, E., Zhang, S., Alegret, L., Schmidt, D.N., Rae, J.W.B., Witts, J.D.,

776 Landman, N.H., Greene, S.E., Huber, B.T., Super, J.R., Planavsky, N.J., Hull, P.M., 2019. Rapid

777 ocean acidification and protracted Earth system recovery followed the end-Cretaceous

778 Chicxulub impact. Proc Natl Acad Sci USA 116, 22500–22504.

779 https://doi.org/10.1073/pnas.1905989116

780 Hernandez Nava, A., Black, B.A., Gibson, S.A., Bodnar, R.J., Renne, P.R., Vanderkluysen, L., 2021.

781 Reconciling early Deccan Traps CO₂ outgassing and pre-KPB global climate. Proc. Natl. Acad.

782 Sci. U.S.A. 118, e2007797118. https://doi.org/10.1073/pnas.2007797118

783 Heuser, A., Eisenhauer, A., Gussone, N., Bock, B., Hansen, B.T., Nägler, T.F., 2002. Measurement of

784 calcium isotopes (δ44Ca) using a multicollector TIMS technique. International Journal of

785 Mass Spectrometry 13.

786 Higgins, J.A., Blättler, C.L., Lundstrom, E.A., Santiago-Ramos, D.P., Akhtar, A.A., Crüger Ahm, A.-S.,

787 Bialik, O., Holmden, C., Bradbury, H., Murray, S.T., Swart, P.K., 2018. Mineralogy, early

788 marine diagenesis, and the chemistry of shallow-water carbonate sediments. Geochimica et

789 Cosmochimica Acta 220, 512–534. https://doi.org/10.1016/j.gca.2017.09.046

790 Hinojosa, J.L., Brown, S.T., Chen, J., DePaolo, D.J., Paytan, A., Shen, S., Payne, J.L., 2012. Evidence for

791 end-Permian ocean acidification from calcium isotopes in biogenic apatite. Geology 40, 743–

792 746. https://doi.org/10.1130/G33048.1

793 Hippler, D., Eisenhauer, A., Nägler, T.F., 2006a. Tropical Atlantic SST history inferred from Ca isotope

794 thermometry over the last 140ka. Geochimica et Cosmochimica Acta 70, 90–100.

795 https://doi.org/10.1016/j.gca.2005.07.022

796 Hippler, D., Eisenhauer, A., Nägler, T.F., 2006b. Tropical Atlantic SST history inferred from Ca isotope

797 thermometry over the last 140ka. Geochimica et Cosmochimica Acta 70, 90–100.

798 https://doi.org/10.1016/j.gca.2005.07.022

799 Hippler, D., Kozdon, R., Darling, K.F., Eisenhauer, A., Nagler, T.F., 2009. Calcium isotopic composition

800 of high-latitude proxy carrier Neogloboquadrina pachyderma (sin.) 14.

801 Hippler, D., Schmitt, A.-D., Gussone, N., Heuser, A., Stille, P., Eisenhauer, A., Nägler, T.F., 2003.

802 Calcium Isotopic Composition of Various Reference Materials and Seawater. Geostandards

803 and Geoanalytical Research 27, 13–19. https://doi.org/10.1111/j.1751-908X.2003.tb00709.x

804 Horita, J., Zimmermann, H., Holland, H.D., 2002. Chemical evolution of seawater during the

805 Phanerozoic: Implications from the record of marine evaporites 24.

806 Hull, P.M., Bornemann, A., Penman, D.E., Henehan, M.J., Norris, R.D., Wilson, P.A., Blum, P., Alegret,

807 L., Batenburg, S.J., Bown, P.R., Bralower, T.J., Cournede, C., Deutsch, A., Donner, B., Friedrich,

808 O., Jehle, S., Kim, H., Kroon, D., Lippert, P.C., Loroch, D., Moebius, I., Moriya, K., Peppe, D.J.,

809 Ravizza, G.E., Röhl, U., Schueth, J.D., Sepúlveda, J., Sexton, P.F., Sibert, E.C., Śliwińska, K.K.,

810 Summons, R.E., Thomas, E., Westerhold, T., Whiteside, J.H., Yamaguchi, T., Zachos, J.C., 2020.

811 On impact and volcanism across the Cretaceous-Paleogene boundary. Science 367, 266–272.

812 https://doi.org/10.1126/science.aay5055
813 Jost, A.B., Bachan, A., van de Schootbrugge, B., Brown, S.T., DePaolo, D.J., Payne, J.L., 2017. Additive effects of acidification and mineralogy on calcium isotopes in Triassic/Jurassic boundary limestones: TRIASSIC/JURASSIC CALCIUM ISOTOPES. Geochem. Geophys. Geosyst. 18, 113–124. https://doi.org/10.1002/2016GC006724

817 Keller, G., Mateo, P., Monkenbusch, J., Thibault, N., Punekar, J., Spangenberg, J.E., Abramovich, S., Ashckenazi-Polivoda, S., Schoene, B., Eddy, M.P., Samperton, K.M., Khadri, S.F.R., Adatte, T., 2020. Mercury linked to Deccan Traps volcanism, climate change and the end-Cretaceous mass extinction. Global and Planetary Change 194, 103312. https://doi.org/10.1016/j.gloplacha.2020.103312

822 Kitch, G.D., Jacobson, A.D., Harper, D.T., Hurtgen, M.T., Sageman, B.B., Zachos, J.C., 2021. Calcium isotope composition of *Morozovella* over the late Paleocene–early Eocene. Geology 49, 723–727. https://doi.org/10.1130/G48619.1

825 Kitch, G.D., Jacobson, A.D., Sageman, B.B., Coccioni, R., Chung-Swanson, T., Ankney, M.E., Hurtgen, M.T., 2022. Calcium isotope ratios of malformed foraminifera reveal biocalcification stress preceded Oceanic Anoxic Event 2. Commun Earth Environ 3, 315. https://doi.org/10.1038/s43247-022-00641-0

829 Kısakürek, B., Eisenhauer, A., Böhm, F., Hathorne, E.C., Erez, J., 2011. Controls on calcium isotope fractionation in cultured planktic foraminifera, Globigerinoides ruber and Globigerinella siphonifera. Geochimica et Cosmochimica Acta 75, 427–443. https://doi.org/10.1016/j.gca.2010.10.015

833 Kozdon, R., Kelly, D.C., Kitajima, K., Strickland, A., Fournelle, J.H., Valley, J.W., 2013. In situ δ¹⁸O and Mg/Ca analyses of diagenetic and planktic foraminiferal calcite preserved in a deep-sea record of the Paleocene–Eocene thermal maximum: DIAGENETIC CEMENTS. Paleceanography 28, 517–528. https://doi.org/10.1002/palo.20048

837 Lasaga, A.C., Berner, R.A., Garrels, R.M., 1985. An Improved Geochemical Model of Atmospheric CO₂ Fluctuations Over the Past 100 Million Years, in: Sundquist, E.T., Broecker, W.S. (Eds.), Geophysical Monograph Series. American Geophysical Union, Washington, D. C., pp. 397–411. https://doi.org/10.1029/GM032p0397

841 Lau, K.V., Maher, K., Brown, S.T., Jost, A.B., Altiner, D., DePaolo, D.J., Eisenhauer, A., Kelley, B.M., Lehrmann, D.J., Paytan, A., Yu, M., Silva-Tamayo, J.C., Payne, J.L., 2017. The influence of seawater carbonate chemistry, mineralogy, and diagenesis on calcium isotope variations in Lower-Middle Triassic carbonate rocks. Chemical Geology 471, 13–37. https://doi.org/10.1016/j.chemgeo.2017.09.006

846 Lemarchand, D., Wasserburg, G.J., Papanastassiou, D.A., 2004. Rate-controlled calcium isotope fractionation in synthetic calcite. Geochimica et Cosmochimica Acta 68, 4665–4678. https://doi.org/10.1016/j.gca.2004.05.029

849 Li, L., Keller, G., 1998. Abrupt deep-sea warming at the end of the Cretaceous 5.

850 Linzmeier, B.J., Jacobson, A.D., Sageman, B.B., Hurtgen, M.T., Ankney, M.E., Petersen, S.V., Tobin, T.S., Kitch, G.D., Wang, J., 2020. Calcium isotope evidence for environmental variability before and across the Cretaceous-Paleogene mass extinction. Geology 48, 34–38. https://doi.org/10.1130/G46431.1

854 MacLeod, N., 1993. The Maastrichtian-Danian radiation of triserial and biserial planktic foraminifera: Testing phylogenetic and adaptational hypotheses in the (micro)fossil record. Marine Micropaleontology 21, 47–100. https://doi.org/10.1016/0377-8398(93)90011-L

857 Marriott, C.S., Henderson, G.M., Belshaw, N.S., Tudhope, A.W., 2004. Temperature dependence of δ 7 Li, δ 44 Ca and Li/Ca during growth of calcium carbonate. Earth and Planetary Science Letters 222, 615–624. https://doi.org/10.1016/j.epsl.2004.02.031

860 Misra, S., Froelich, P.N., 2012. Lithium Isotope History of Cenozoic Seawater: Changes in Silicate Weathering and Reverse Weathering. Science 335, 818–823. https://doi.org/10.1126/science.1214697
863 Molina, E., Arenillas, I., Arz, J.A., 1996. The Cretaceous/Tertiary boundary mass extinction in planktic
864 foraminifera at Agost, (Spain). Revue de Micropaléontologie 39, 225-243.

865 https://doi.org/10.1016/S0035-1598(96)90075-6

866 Nägler, T.F., Eisenhauer, A., Müller, A., Hemleben, C., Kramers, J., 2000. The δ<sup>44</sup>Ca-temperature
867 calibration on fossil and cultured *Globigerinoides sacculifer*: New tool for reconstruction of
868 past sea surface temperatures. Geochem. Geophys. Geosyst. 1, n/a-n/a.

869 https://doi.org/10.1029/2000GC000091

870 Nakanishi, M., Sager, W.W., Klaus, A., 1999. Magnetic lineations within Shatsky Rise, northwest
871 Pacific Ocean: Implications for hot spot-triple junction interaction and oceanic plateau
872 formation. J. Geophys. Res. 104, 7539-7556. https://doi.org/10.1029/1999JB900002

873 Nielsen, L.C., DePaolo, D., Yoreo, J.D., 2012. Self-consistent ion-by-ion growth model for kinetic
874 isotopic fractionation during calcite precipitation.

875 https://doi.org/10.1016/J.GCA.2012.02.009

876 Norris, R.D., 1996. Symbiosis as an evolutionary innovation in the radiation of Paleocene planktic
877 foraminifera. Paleobiology 22, 461-480. https://doi.org/10.1017/S0094837300016468

878 Paris, G., Donnadieu, Y., Beaumont, V., Fluteau, F., Goddéris, Y., 2016. Geochemical consequences of
879 intense pulse-like degassing during the onset of the Central Atlantic Magmatic Province.

880 Palaeogeography, Palaeoclimatology, Palaeoecology 441, 74-82.

881 https://doi.org/10.1016/j.palaeo.2015.04.011

882 Penman, D.E., Zachos, J.C., 2018. New constraints on massive carbon release and recovery processes
883 during the Paleocene-Eocene Thermal Maximum. Environ. Res. Lett. 13, 105008.

884 https://doi.org/10.1088/1748-9326/aae285

885 Pope, K.O., Baines, K.H., Ocampo, A.C., Ivanov, B.A., 1997a. Energy, volatile production, and climatic
886 effects of the Chicxulub Cretaceous/Tertiary impact. Journal of Geophysical Research:

887 Planets (1991-2012) 102, 21645-21664. https://doi.org/10.1029/97JE01743

888 Pope, K.O., Baines, K.H., Ocampo, A.C., Ivanov, B.A., 1997b. Energy, volatile production, and climatic
889 effects of the Chicxulub Cretaceous/Tertiary impact. J. Geophys. Res. 102, 21645-21664.

890 https://doi.org/10.1029/97JE01743

891 Premoli Silva, I., Petrizzo, M.R., Melloni, D. (Eds.), 2005. DATA REPORT: PLANKTONIC FORAMINIFERAL
892 BIOSTRATIGRAPHY ACROSS THE CRETACEOUS/PALEOCENE BOUNDARY AT SHATSKY RISE

893 (ODP LEG 198, NORTHWEST PACIFIC), Proceedings of the Ocean Drilling Program, Scientific
894 Results, Volume 198. Ocean Drilling Program.

895 Rathburn, A.E., Corliss, B.H., 1994. The ecology of living (stained) deep-sea benthic foraminifera from
896 the Sulu Sea. Paleooceanography 9, 87-150. https://doi.org/10.1029/93PA02327

897 Regenberg, M., 2007. Early diagenetic overprint in Caribbean sediment cores and its effect on the
898 geochemical composition of planktonic foraminifera 17.

899 Renne, P.R., Deino, A.L., Hilgen, F.J., Kuiper, K.F., Mark, D.F., Mitchell, W.S., Morgan, L.E., Mundil, R.,
900 Smit, J., 2013. Time Scales of Critical Events Around the Cretaceous-Paleogene Boundary.

901 Science 339, 684. https://doi.org/10.1126/science.1230492

902 Rennie, V.C.F., Paris, G., Sessions, A.L., Abramovich, S., Turchyn, A.V., Adkins, J.F., 2018. Cenozoic
903 record of δ34S in foraminiferal calcite implies an early Eocene shift to deep-ocean sulfide
904 burial. Nature Geosci 11, 761-765. https://doi.org/10.1038/s41561-018-0200-y

905 Schoene, B., Eddy, M.P., Keller, C.B., Samperton, K.M., 2021. An evaluation of Deccan Traps eruption
906 rates using geochronologic data. Geochronology 3, 181-198.

907 https://doi.org/10.5194/gchron-3-181-2021

908 Schoene, B., Eddy, M.P., Samperton, K.M., Keller, C.B., Keller, G., Adatte, T., Khadri, S.F.R., 2019. U-Pb
909 constraints on pulsed eruption of the Deccan Traps across the end-Cretaceous mass

910 extinction. Science 363, 862-866. https://doi.org/10.1126/science.aau2422

911 Sime, N.G., De La Rocha, C.L., Galy, A., 2005. Negligible temperature dependence of calcium isotope
912 fractionation in 12 species of planktonic foraminifera. Earth and Planetary Science Letters
913 232, 51-66. https://doi.org/10.1016/j.epsl.2005.01.011
914 Sime, N.G., De La Rocha, C.L., Tipper, E.T., Tripati, A., Galy, A., Bickle, M.J., 2007. Interpreting the Ca
915 isotope record of marine biogenic carbonates. Geochimica et Cosmochimica Acta 71, 3979-
916 3989. https://doi.org/10.1016/j.gca.2007.06.009

917 Skulan, J., DePaolo, D.J., Owens, T.L., 1997. Biological control of calcium isotopic abundances in the
918 global calcium cycle. Geochimica et Cosmochimica Acta 61, 2505-2510.
919 https://doi.org/10.1016/S0016-7037(97)00047-1

920 Sprain, C.J., Renne, P.R., Vanderkluysen, L., Pande, K., Self, S., Mittal, T., 2019. The eruptive tempo of
921 Deccan volcanism in relation to the Cretaceous-Paleogene boundary. Science 363, 866-870.
922 https://doi.org/10.1126/science.aav1446

923 Stanley, S.M., Hardie, L.A., 1998. Secular oscillations in the carbonate mineralogy of reef-building and
924 sediment-producing organisms driven by tectonically forced shifts in seawater chemistry.
925 Palaeogeography, Palaeoclimatology, Palaeoecology 144, 3-19.
926 https://doi.org/10.1016/S0031-0182(98)00109-6

927 Tang, J., Dietzel, M., Böhm, F., Köhler, S.J., Eisenhauer, A., 2008. Sr2+/Ca2+ and 44Ca/40Ca
928 fractionation during inorganic calcite formation: II. Ca isotopes. Geochimica et Cosmochimica
929 Acta 72, 3733-3745. https://doi.org/10.1016/j.gca.2008.05.033

930 Yu, J., Elderfield, H., Jin, Z., Tomascak, P., Rohling, E.J., 2014. Controls on Sr/Ca in benthic
931 foraminifera and implications for seawater Sr/Ca during the late Pleistocene. Quaternary
932 Science Reviews 98, 1-6. https://doi.org/10.1016/j.quascirev.2014.05.018

933 Zachos, J.C., Dickens, G.R., Zeebe, R.E., 2008. An early Cenozoic perspective on greenhouse warming
934 and carbon-cycle dynamics. Nature 451, 279-283. https://doi.org/10.1038/nature06588

935 Zachos, J.C., Röhl, U., Schellenberg, S.A., Sluijs, A., Hodell, D.A., Kelly, D.C., Thomas, E., Nicolo, M.,
936 Raffi, I., Lourens, L.J., McCarren, H., Kroon, D., 2005. Rapid Acidification of the Ocean During
937 the Paleocene-Eocene Thermal Maximum. Science 308, 1611-1615.
938 https://doi.org/10.1126/science.1109004

939 Zeebe, R.E., 2012. LOSCAR: Long-term Ocean-atmosphere-Sediment Carbon cycle Reservoir Model
940 v2.0.4. Geosci. Model Dev. 5, 149-166. https://doi.org/10.5194/gmd-5-149-2012

941 Zeebe, R.E., Zachos, J.C., 2013. Long-term legacy of massive carbon input to the Earth system:
942 Anthropocene versus Eocene. Phil. Trans. R. Soc. A. 371, 20120006.
943 https://doi.org/10.1098/rsta.2012.0006

945
<figure><img src="image_2.png" /></figure>

Figure 1: (a)Paleogeographic map of Shatsky Rise with present-day continental configuration. (b) Bathymetric map of Shatsky Rise showing in red the positon of ODP Site 1209 in the Southern High (modified from Bown, 2005).

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

Figure 2: The accuracy of the measurements is assessed by repetitive measurements (n =3) of the seawater standard (OSIL)
<figure><img src="image_4.png" /><figcaption>Figure 3: The δ<sup>44/40</sup>Ca<sub>P</sub> and δ<sup>44/40</sup>Ca<sub>B</sub> records along with Sr/Ca<sub>P</sub>, Sr/Ca<sub>B</sub>, δ<sup>13</sup>C and δ<sup>18</sup>O records of planktic foraminefera. The major Ca isotopic shifts are delimited by colored bands (Intervals I and II). Different foraminiferal species are represented by different symbols. Dark blue samples on the δ<sup>44/40</sup>Ca<sub>P</sub> and δ<sup>44/40</sup>Ca<sub>B</sub> records represent samples suspected to have been diagenetically overprinted (i.e high Mg/Ca or/and high Mn/ Ca or/and Sr/Ca ratios, see Appendix C.1 and C.2). Here, we present the Cretaceous-Paleogene transition at the basis of the first occurrence (FO) of *Parvularugoglobigerina eugubina* (Premoli Silva et al., 2005), marker of the lowermost Danian. However this bioevent is slightly younger than the K-Pg boundary (e.g. Molina et al., 1996).</figcaption></figure>
<figure><img src="image_5.png" /></figure>

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

<figure><img src="image_7.png" /></figure>
964

outlier Mg/Ca or/and Mn/Ca or/and Sr/Ca) are not plotted here. See Appendix C.1 and C.2
and appendix D for all data values.

965

966

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

967

968

969

970

971

972

Figure 5: The δ<sup>44/40</sup>Ca<sub>B</sub> record across the K-Pg transition together with the modeled evolution of the oceanic Ca reservoir (N'Ca) and carbonate sedimentation flux (i.e. F<sub>carb</sub>/F<sub>w</sub>) assuming a Ca residence time of 0.7 My. Here, the K-Pg boundary is illustrated with a dotted line to indicate that this boundary is not well defined, as the primary markers (i.e. the clay and iridium level) are missing in Shatsky rise.

973
<figure><img src="image_9.png" /></figure>
<figure><img src="image_10.png" /><figcaption>Figure 6: Comparison between our δ<sup>44/40</sup>Ca, δ<sup>13</sup>C<sub>P</sub> and δ<sup>18</sup>O<sub>P</sub> with δ<sup>13</sup>C<sub>B</sub> and δ<sup>18</sup>O<sub>B</sub> data from Dameron et al. (2017) and Westerhold et al. (2011). Different foraminiferal species are represented by different symbols. Dark blue samples on the δ<sup>44/40</sup>Ca<sub>P</sub> and δ<sup>44/40</sup>Ca<sub>B</sub> records represent samples suspected to have been diagenetically overprinted (i.e high Mg/Ca or/and high Mn/Ca or/and Sr/Ca ratios, see Appendix C.1 and C.2). Here, the K-Pg boundary is illustrated with a dotted line to indicate that this boundary is not well defined, as the primary markers (i.e. the clay and iridium level) are missing in Shatsky rise.</figcaption></figure>
Figure 7: The δ<sup>44/40</sup>Ca<sub>P</sub> and δ<sup>44/40</sup>Ca<sub>B</sub> records along with Sr/Ca records and fragmentation index (Dameron et al., 2017) across the K-Pg transition. The main shifts in δ<sup>44/40</sup>Ca are delimited by colored bands and are numbered from interval I to II. Different foraminiferal species are represented by different symbols. Dark Blue samples on the δ<sup>44/40</sup>Ca<sub>P</sub> and δ<sup>44/40</sup>Ca<sub>B</sub> records represent samples suspected to have been diagenetically overprinted (i.e high Mg/Ca or/and high Mn/Ca or/and lower Sr/Ca ratios, see Appendix C.1 and C.2). Here, the K-Pg boundary is illustrated with a dotted line to indicate that this boundary is not well defined, as the primary markers (i.e. the clay and iridium level) are missing in Shatsky rise.

<figure><img src="image_11.png" /></figure>
<figure><img src="image_12.png" /><figcaption>Figure 8: Comparison between our δ<sup>44/40</sup>Ca, δ<sup>44/40</sup>Ca (Linzmeier et al., 2020) and δ<sup>11</sup>B data (Henehan et al., 2019). Different foraminiferal species are represented by different symbols. Dark blue samples on the δ<sup>44/40</sup>Ca<sub>P</sub> and δ<sup>44/40</sup>Ca<sub>B</sub> records represent samples suspected to have been diagenetically overprinted (i.e high Mg/Ca or/and high Mn/Ca ratios or/and Sr/Ca ratios, see Appendix C.1 and C.2). Here the K-Pg boundary is illustrated with a dotted line to indicate that this boundary is not well defined, as the primary markers (i.e. the clay and iridium level) are missing in Shatsky rise.</figcaption></figure>
Figure 9: Correlation of our δ<sup>44/40</sup>Ca, Sr/Ca data and timing of Deccan eruptions from (Schoene et al., 2019). Different foraminiferal species are represented by different symbols. Dark blue samples on the δ<sup>44/40</sup>Ca<sub>P</sub> and δ<sup>44/40</sup>Ca<sub>B</sub> records represent samples suspected to have been diagenetically overprinted (i.e high Mg/Ca or/and high Mn/Ca lower or/and Sr/Ca ratios, see Appendix C.1 and C.2). Here the K-Pg boundary is illustrated a dotted line to indicate that this boundary is not well defined, as the primary markers (i.e. the clay and iridium level) are missing in Shatsky rise.

998

999

1000

1001

1002

1003

1004

1005

1006

1007

1008

1009

1010

1011
<table><caption>Table 1: Age model for the Maastrichtian interval of this study</caption><thead><tr><th>Bioevents</th><th>Site</th><th>Depth (rmcd)</th><th>Age (Ma)</th><th>Age references</th><th>Bioevent references</th></tr></thead><tbody><tr><td>LO Gansserina gansseri</td><td>198_1209C</td><td>266.05</td><td>66.53</td><td>Coccionia et Premoli Silva. 2015</td><td>Dameron et al. 2017</td></tr><tr><td>FO A. mayaroensis</td><td>198_1209C</td><td>298.135</td><td>69.33</td><td>Coccionia et Premoli Silva. 2015</td><td>Dameron et al. 2017 and Premolisilva et al. 2005</td></tr></tbody></table>

<table><caption>Table 2: Oceanic calcium residence time estimation</caption><thead><tr><th>Ca Reservoir size (mol)</th><th>F<sub>out</sub></th><th>Estimated resident time of Ca in the ocean (Ma) for different Ca Reservoir size</th></tr></thead><tbody><tr><td>4.08E+19 (Horita et al., 2002)</td><td rowspan="5">1.5E+13 mol/yr (Mackenzie and Andersson, 2013)</td><td>2.7</td></tr><tr><td>2.16E+19 (Lasaga et al., 1985)</td><td>1.3</td></tr><tr><td>4.80E+19 (Stanley and Hardie, 1998)</td><td>3.3</td></tr><tr><td>7.20E+19 (Hardie, 1996)</td><td>4.7</td></tr><tr><td>1.44E+19 (Du Vivier et al., 2015)</td><td>0.7</td></tr></tbody></table>