# On climate and abyssal circulation in the Atlantic Ocean during late Pliocene marine isotope stage M2, ~3.3 million years ago

Nicola Kirby, Ian Bailey, David Lang, Anieke Brombacher, Thomas Chalk, Rebecca Parker, Anya Crocker, Victoria Taylor, James Milton, Gavin Foster, et al.

► To cite this version:

Nicola Kirby, Ian Bailey, David Lang, Anieke Brombacher, Thomas Chalk, et al. On climate and abyssal circulation in the Atlantic Ocean during late Pliocene marine isotope stage M2, ~3.3 million years ago. Quaternary Science Reviews, 2020, 250, pp.106644. 10.1016/j.quascirev.2020.106644. hal-04707843

HAL Id: hal-04707843

https://hal.science/hal-04707843v1

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1 On climate and abyssal circulation in the Atlantic Ocean

2 during late Pliocene marine isotope stage M2, ~3.3 million
3 years ago

4

5 Nicola Kirby¹,², Ian Bailey*³, David C. Lang¹, Anieke Brombacher¹, Thomas B.

6 Chalk¹, Rebecca L. Parker³, Anya J. Crocker¹, Victoria E. Taylor¹, James A.

7 Milton¹, Gavin L. Foster¹, Maureen E. Raymo⁴, Dick Kroon⁵, David B. Bell⁵,
8 Paul A. Wilson¹

9 1. School of Ocean and Earth Science, University of Southampton, Waterfront Campus, National Oceanography Centre Southampton, European Way, Southampton SO14 3ZH, UK.

10

11 2. Now at: Earth Sciences, School of Geography, Earth and Environmental Sciences, University of
12 Birmingham B15 2TT, UK.

11

12

13 3. Camborne School of Mines, College of Engineering, Mathematics & Physical Sciences, University of Exeter, Penryn Campus, Treliever Road, Penryn TR10 9FE, UK.

13

14

15 4. Columbia University, Lamont-Doherty Earth Observatory, P.O. Box 1000, 61 Route 9W, Palisades, NY 10964, USA.

15

16

17 5. School of GeoSciences, University of Edinburgh, Edinburgh EH9 3JW, UK

17

18 6. Centre for Environmental Policy, Weeks Building, 16-18 Prince's Gardens, Imperial College London South Kensington, London SW7 1NE, UK.

18

19

20

21 *corresponding author: njk883@student.bham.ac.uk

22

23 Abstract: Marine Isotope Stage (MIS) M2, 3.3 Ma, is an isolated cold stage

24 punctuating the benthic oxygen isotope (δ¹⁸O) stratigraphy of the warm Piacenzian

25 interval of the late Pliocene Epoch. The prominent (~0.65‰) δ¹⁸O increase that

26 defines MIS M2 has prompted debate over the extent to which it signals an early
prelude to the rhythmic extensive glaciations of the northern hemisphere that
characterise the Quaternary and raised questions about the forcing mechanisms
responsible. Recent work suggests that CO₂ storage in the deep Atlantic Ocean played
an important role in these events but detailed reconstructions of deep ocean chemical
stratification are needed to test this idea and competing hypotheses. Here we present
new records of the Nd isotope composition of fish debris and δ¹³C and B/Ca ratios of
benthic foraminifera from the northwest and southeast Atlantic Ocean. Our novel
geochemical data show that, in contrast to major Quaternary glaciations such as MIS
2 (~21 kyr) and MIS 100 (~2.52 Ma), the deep North Atlantic Ocean was weakly
chemically stratified during MIS M2. We show that Southern Component Water
incursion into the Atlantic Ocean was limited to the deep South Atlantic basin during
MIS M2 and peaked well before (~10-15-kyr) the atmospheric CO₂ minimum. Our
findings imply that the deep Atlantic Ocean was not the principle sink of CO₂
sequestered from the atmosphere during MIS M2, implicating a different CO₂ storage
deep-water reservoir mechanism, presumably Southern Component Water incursion
into the Pacific Ocean. Weak chemical stratification in the deep Atlantic Ocean
during MIS M2 relative to MIS 100 and 2 suggests comparatively active Atlantic
meridional overturning circulation. That suggestion is consistent with the warmth of
the high latitude North Atlantic during MIS M2 – surface water temperatures cooled
during M2 but only to Holocene values. Our findings may help to explain the paucity
of evidence for extensive early glaciation of the northern hemisphere during M2 but
leave open the possibility of ice sheet advance on Antarctica.

49

50 **Key words:** AABW; NADW; Atlantic circulation; Last Glacial; MIS 100; MIS M2;
51 AMOC
52

## 1. Introduction

Much of the Pliocene Epoch is considered to have been warmer than the pre-industrial with peak atmospheric carbon dioxide ($pCO_2$) levels about as high as present (413 ppm in 2019; Seki et al., 2010; Dowsett et al., 2012; Haywood et al., 2013; Martinez-Boti et al., 2015; de la Vega et al., 2020). Many studies of Pliocene warmth over the past two decades have focused on the mid-Piacenzian warm period (mPWP, 3.264 to 3.025 Ma), the most recent interval of sustained warmth during this epoch (McClymont et al., 2020). However, just prior to this time there occurred a transient interval of high latitude cooling beginning around 3.5 Ma during Marine Isotope Stage (MIS) MG5 that culminated during MIS M2, 3.295 Ma (Prell, 1984; Keigwin, 1987; Haug and Tiedemann, 1998; Jansen et al., 2000; Lisiecki and Raymo, 2005).

One interpretation of MIS M2 is that it signals early extensive northern hemisphere (and therefore) bi-polar glaciation (De Schepper et al, 2009; De Schepper et al., 2013). Yet the magnitude and location of land ice growth during this cold stage is poorly documented – if there was a major glacial advance it is not even known whether it was concentrated in the northern or southern hemisphere (e.g., Prell, 1984; Maslin et al., 1998; Dwyer and Chandler, 2009; De Schepper et al., 2009; Naish et al., 2009; McKay et al., 2012; Gao et al., 2012; De Schepper et al., 2013; De Schepper et al., 2014; Brigham-Grette et al., 2013; Dolan et al., 2015). Estimates of global sea-level fall associated with full glacial conditions of MIS M2 range from ~ 10 m lower than the present day (Naish et al., 1997; Miller et al., 2012) to as much as 65 m lower (Dwyer and Chandler, 2009). If the larger sea-level fall estimates relative to modern are substantiated, MIS M2 may mark the first time during the Cenozoic when the seawater component of benthic $\delta^{18}$O increase was controlled by ice sheet growth in
the northern hemisphere. Yet while there is firm evidence for glacial expansion on Greenland, Svalbard and Iceland during M2, the same is not true of North America or mainland Eurasia (De Schepper et al., 2014). Furthermore, numerical ice-sheet modelling experiments, simulate only small ice caps in the northern hemisphere, equivalent to ~12 m fall in sea level (msle) (Tan et al., 2017), twice the seawater equivalent of the modern Greenland Ice Sheet but an order of magnitude smaller than the fall associated with the major glaciation of the Last Glacial Maximum (Lambeck et al., 2014).

A small M2 ice-sheet budget for the northern hemisphere is also implied by a new detailed record of atmospheric CO₂ for late Pliocene which shows a marked decline during MIS M2 (from ~420 to 310 ppm) but to concentrations unlikely to have fallen below the threshold suggested (DeConto et al., 2008) to trigger extensive northern hemisphere glaciation (de la Vega et al., 2020). Atmospheric CO₂ decline during MIS M2 lags benthic δ¹⁸O by ~10 kyr but is in-phase with possible evidence for Southern Component Water (SCW) incursion into the deep mid-latitude North Atlantic Ocean (Lang et al., 2016; de la Vega et al., 2020). These observations imply that expansion of a CO₂-charged SCW reservoir in the deep Atlantic may have been active more than 600-kyr before the onset of significant northern hemisphere glaciation ~2.72 Ma (de la Vega et al., 2020). Yet, while Nd isotope and δ¹³C records from Integrated Ocean Drilling Program (IODP) Site U1313 (Fig. 1) suggest that major incursions of SCW of last glacial-magnitude occurred at this site during large Quaternary glacials at least from MIS 100, 2.52 Ma (Lang et al., 2014; 2016), those records are insufficiently resolved for MIS M2 to provide a clear test of whether the same is true for this older more enigmatic cold stage (Fig. 1a,b).

To shed new light on the role of the Atlantic Ocean in CO₂ decline during MIS
M2, we present a new high-resolution (suborbital-scale) assessment of changes in water-mass distribution at deep Northwest Atlantic Ocean IODP Site U1313 (Fig. 2) and deep Southeast Atlantic Ocean Drilling Program (ODP) Site 1267 (~4350 m depth; ~29°S, Fig. 2). We apply the same multi-proxy approach shown to be effective for tracking water-mass distribution during the last glacial (e.g., Yu et al., 2008; Chalk et al., 2019), combining neodymium-isotope composition of fish debris as well as the carbon isotope and B/Ca composition of benthic foraminifera to trace water-mass properties. Our new records for MIS M2 show contrasting structure to that seen for MIS 100 and MIS 2, strongly suggesting that the water-mass structure of the deep Atlantic Ocean during MIS M2 was quite different to that of large Quaternary glacials. We compare our data to published records for MIS M2 and discuss the implications of our new findings for our understanding of late Pliocene climate.

## 2. Methods

### 2.1. Study sites, sampling and chronology

We document change in three water-mass tracers ($\delta^{13}$C and B/Ca in benthic foraminiferal calcite and Nd isotopes in fish debris) for MIS M2 from the deep western North Atlantic (Site U1313; Fig. 2) and deep eastern South Atlantic (Site 1267; Fig. 2) and compare their amplitude of change to new and published records for MIS 100 (2.52 Ma) and MIS 2 (21 ka), large Quaternary glacials for which major SCW incursion into the deep North Atlantic is documented by characteristic $\delta^{13}$C-$CO_3^{2-}$Nd isotope relationships (e.g., Oppo and Fairbanks, 1987; Raymo et al., 1987, 1990, 1992; Hodell and Venz, 1992; Yu et al., 2008; Lang et al., 2016; Chalk et al., 2019). Given the caveats that come with water-mass provenance proxies, we follow Lang et al. (2016) in placing most confidence in signals of change documented in all
three water-mass-proxy records. To produce our new data sets we: 1) increased the resolution of the Site U1313 fish debris Nd isotope record published by Lang et al. (2016) for MIS M2 and MIS 100 (from ~6-7 kyr to ~3.5 and 1 kyr, respectively), and 2) developed new suborbitally resolved records of fish debris Nd isotopes and benthic foraminiferal B/Ca ratios and δ¹³C for MIS M2 (3.24-3.33 Ma), 100 (2.5-2.55 Ma) and the last glacial-Holocene (0-40 ka) from Site 1267.

## 2.1.1. IODP Site U1313

IODP Site U1313 (41°N, 32.4°W, 3426 m depth, Fig. 2) constitutes a re-occupation of Deep Sea Drilling Project (DSDP) Site 607, a benchmark site for monitoring North Atlantic Deep Water (NADW) production during the Plio-Pleistocene (e.g., Ruddiman et al., 1989; Raymo et al., 1990, 1992; Raymo et al., 2004). Today, Site U1313 is bathed by 100% lower NADW. To increase the resolution of the previously published fish debris Nd isotope record for Site U1313 spanning MIS M2 and MIS 100, we analysed 14 additional samples from the shipboard primary splice (between 152.48-154.38 mcd; Expedition 306 Scientists, 2006) and 42 samples from the shipboard secondary splice (between 118.74-121.39 mcd; Expedition 306 Scientists, 2006). The age model for our MIS M2 record is based on retuning the Site U1313 benthic δ¹⁸O stratigraphy (Bolton et al., 2010) to the LR04 stack (Lisiecki and Raymo, 2005) (Supplementary Information Section S2). To improve the age model for the Site U1313 stratigraphy spanning MIS 100 (which was previously based on assigning ages from the primary splice to the secondary splice stratigraphy by correlating between the two using shipboard-derived sediment colour reflectance data; Lang et al., 2016), we generated a new suborbitally resolved benthic foraminiferal δ¹⁸O record (at on average ~1 kyr) between 118.64-122.44 mcd from the secondary
<table>
  <tr>
    <td>152</td>
    <td>splice, which we subsequently tuned to the LR04 stack (Supplementary Information Section S2).</td>
  </tr>
  <tr>
    <td>153</td>
    <td></td>
  </tr>
  <tr>
    <td>154</td>
    <td></td>
  </tr>
  <tr>
    <td>155</td>
    <td>2.1.2. ODP Site 1267</td>
  </tr>
  <tr>
    <td>156</td>
    <td>ODP Site 1267 (29°S, 2°E at 4350 m depth; Fig. 2) sits in the Angola Basin in the Southeast Atlantic and monitors bottom-water conditions at the base of the northern face of Walvis Ridge. Today Site 1267 is bathed by Lower Deep Water (LDW), a mixture of 70-80% North Atlantic Deep Water (NADW) and 20-30% Antarctic Bottom Water (AABW) sourced from the western Atlantic via the equatorial Romanche and Chain Fracture zones and mixed vertically in the Angola Basin downstream of these entry points (Fig. 2; Mercier and Morin, 1997; Arhan et al., 2003). The Rio de Janeiro and Rio Grande Fracture zones, which cut the Mid-Atlantic Ridge at 4000 m or greater depth at 21°S and 26°S, respectively, also funnel AABW from the western South Atlantic Brazil Basin into the Angola Basin (Fig. 2; Mercier et al., 2000). Together with two other passages below 4000 m in the southern part of the Walvis Ridge near 36°S 7°W (the Walvis Passage) and 30°S 2°W (Connary and Ewing, 1974), these fracture zones are the only pathways for AABW to reach Site 1267 (Fig. 2). During late Pleistocene glacials this site was more influenced by AABW (with its relatively low δ<sup>13</sup>C, radiogenic Nd isotope and low carbonate ion signature) because the interface between AABW and overlying NADW shoaled and moved equatorward allowing more southern source deep water to enter the basin via these fracture zones (Yu et al., 2008; Chalk et al., 2019; Farmer et al., 2019; Pöppelmeier et al., 2020). The more southerly location (by ~70° of latitude) and greater depth (by ~900 m) of Site 1267 relative to Site U1313 allows us to investigate</td>
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    <td>157</td>
    <td></td>
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    <td>158</td>
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    <td>159</td>
    <td></td>
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    <td>160</td>
    <td></td>
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    <td>161</td>
    <td></td>
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    <td>162</td>
    <td></td>
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    <td>163</td>
    <td></td>
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    <td>164</td>
    <td></td>
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    <td>165</td>
    <td></td>
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    <td>166</td>
    <td></td>
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    <td>167</td>
    <td></td>
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    <td>168</td>
    <td></td>
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    <td>169</td>
    <td></td>
  </tr>
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    <td>170</td>
    <td></td>
  </tr>
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    <td>171</td>
    <td></td>
  </tr>
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    <td>172</td>
    <td></td>
  </tr>
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    <td>173</td>
    <td></td>
  </tr>
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    <td>174</td>
    <td></td>
  </tr>
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    <td>175</td>
    <td></td>
  </tr>
</table>
changes in the extent of SCW incursion northwards and shoaling of its interface with
Northern Component Water (NCW) during MIS 2, 100 and M2.

To develop new records of Nd isotopes from fish debris and benthic
foraminiferal B/Ca ratios from Site 1267 for MIS M2, MIS 100 and the last glacial-
Holocene we analysed 89 samples taken from the shipboard primary splice between
0-0.34 mcd, 29.4-29.95 mcd and 40.27-41.42 mcd (41 of which were originally used
to generate the published benthic δ¹⁸O and δ¹³C stratigraphies for Site 1267; Bell et
al., 2014). The published age model for this site is based on alignment of a stack of
Site 1267 and 1264 benthic δ¹⁸O data for the past ~3.6 Ma to the LR04 (Bell et al.,
2014). To improve this age model for MIS M2 and the last glacial-Holocene, we
supplemented the previously published benthic δ¹⁸O stratigraphy for Site 1267 by
generating new δ¹⁸O data between 0-0.34 mcd (n = 16) and 40.27-41.22 mcd (n = 50)
and tuned the new resultant stratigraphies to the LR04 (Supplementary Information
Section S2).

2.2. Stable isotope preparation and analysis

To generate new benthic foraminiferal stable isotope data (δ¹⁸O and δ¹³C) for our
target intervals, individual tests of the epifaunal species *Cibicidoides wuellerstorfi*
(typically n>8) were picked from the 150-400 µm size fraction in 66 samples from
Site 1267 (the same samples used to generate Nd isotope data for MIS M2, MIS 100
and the last glacial-Holocene) and from the 212-500 µm fraction in 73 samples from
the secondary splice of Site U1313 (typically >8 tests per sample, the same samples
used to generate Nd isotope data for MIS 100). After drying, picked individuals were
cracked open and ultrasonically cleaned in methanol. This was done to remove
adhering clays and any low δ¹⁸O clay infilling before a homogenised aliquot was
201 transferred to reaction vials and analysed using a Thermo Finnigan MAT 253 dual
202 inlet mass spectrometer equipped with a Kiel IV Carbonate Device at NOCS,
203 University of Southampton. Data were calibrated using a two-point calibration to
204 standards NBS 18 and 19. We also measured an internal standard (GS1) in each run as
205 a quality control measure that yielded an analytical precision of 0.08‰ for δ¹⁸O and
206 0.02‰ for δ¹³C (at 2 s.d.) during the study period. Results are reported relative to the
207 Vienna Peedee belemnite (VPDB) standard. Where the samples picked for stable
208 isotope analysis were also used to generate benthic B/Ca data (e.g., from Site 1267, n
209 = 56), the tests of the picked foraminifera were gently broken open and mixed before
210 being split into two aliquots.

211

212 2.3. Nd-isotope preparation and analysis

213 The local Nd-isotope composition of bottom-water is reliably preserved in fossilised
214 fish tooth, bone and scale hydroxyfluorapatite (henceforth fish debris) during early
215 stage diagenesis, and is resistant to alteration on multi-million year timescales (Martin
216 and Scher, 2004). To develop new Nd isotope datasets, we hand-picked ~50-200 mg
217 of fish debris (10 to 30 pieces of bones and teeth) from the >150 µm fraction of Site
218 U1313 samples and from the >38 µm fraction of Site 1267 samples. Fish debris was
219 cleaned with two steps of ultrasonication in methanol, followed by three steps in
220 Milli-Q 18.2 MΩ·cm water. Organic material was removed by bathing each sample in
221 a buffered oxidising solution (1% H₂O₂ – 0.1M NH₄OH) at 80°C for 30 minutes, and
222 sonicated for 30 seconds every 10 minutes. Finally, a weak-acid (0.001M HNO₃)
223 polish was applied to ensure removal of any reabsorbed contaminants before samples
224 were dissolved in 0.1ml 1.75M HCl.
Pure samples of Nd were extracted from our dissolved samples using a two-stage column procedure. First the rare earth element (REE) fraction was separated using Eichrom TRUSpec resin, then Nd was separated from the REE fraction using Eichrom LNSpec resin. The ¹⁴³Nd/¹⁴⁴Nd ratio was determined at the University of Southampton using a multi-collector inductively coupled plasma mass spectrometer (MC-ICP-MS, Thermo Scientific Neptune). Neodymium isotopic compositions were obtained using the method of Vance and Thirlwall (2002) through adjustment to a ¹⁴⁶Nd/¹⁴⁴Nd ratio of 0.7219 and a secondary normalization to ¹⁴²Nd/¹⁴⁴Nd = 1.141876. Mass-bias corrected ratios were normalised to the given ¹⁴³Nd/¹⁴⁴Nd of the standard JNdi-1 (0.512115 (Tanaka et al., 2000)). Total procedural blanks averaged 35 pg during the analysis of our U1313 fish debris. Long-term replication of the JNdi-1 standard during this time yielded ¹⁴³Nd/¹⁴⁴Nd = 0.512118 ± 0.000006 (2 s.d.), giving an external reproducibility of 0.13 εNd (2 s.d). Total procedural blanks averaged 39 pg during the analysis of our 1267 fish debris. Long-term replication of the JNdi-1 standard during this time yielded ¹⁴³Nd/¹⁴⁴Nd = 0.512116 ± 0.00001 (2 s.d.), giving an external reproducibility of 0.18 εNd (2 s.d). In all cases we report the external uncertainty, unless the internal error is larger than the external error when we plot a combined error calculated as √(external error² + internal error²). All Nd isotope ratios are reported in epsilon notation (DePaolo and Wasserburg, 1976) as:

$$ \epsilon_{Nd} = \left[ \frac{^{143}_{Nd}/^{144}_{Nd}_{sample}}{^{143}_{Nd}/^{144}_{Nd}_{CHUR}} - 1 \right]^{244} \times 10^{4} $$

Radiogenic in-growth of ¹⁴³Nd is corrected following (DePaolo and Wasserburg, 1976) using a crustal ¹⁴⁷Sm/¹⁴⁴Nd value of 0.125 (MacLeod et al., 2008). Such age-corrections are minor (<0.05 εNd for our late Pliocene-earliest Pleistocene data, and
250 <0.0004 εNd for last glacial-Holocene data), reflecting the young age of our samples
251 relative to the half-life of ¹⁴⁷Sm. We note that fish debris and authigenic Fe-Mn
252 oxyhydroxide ¹⁴⁷Sm/¹⁴⁴Nd typically varies within a narrow range (e.g., 0.121-0.136
253 (Martin et al., 2010)) and that our results are insensitive to the choice of value. The
254 sensitivity of our stated values to application of the range of ¹⁴⁷Sm/¹⁴⁴Nd = 0.121-
255 0.136 (Martin et al., 2010) is always <0.01 εNd, which is negligible.
256
257 2.4 Fresh volcanic glass counts
258 The εNd signal recorded by fish debris can be modified by partial dissolution of fresh
259 radiogenic basaltic volcanic glass by sediment pore waters post deposition (Du et al.,
260 2016). It has been suggested that young volcanic glass sourced from the nearby
261 Azores volcanic island eruptions may have influenced the εNd signal recorded at
262 U1313 during the Holocene (Blaser et al., 2016). To evaluate whether porewater
263 volcanic glass dissolution may have played a role in the trends documented in our Site
264 U1313 εNd data we determined the concentration of fresh volcanic glass (>63 µm g⁻¹)
265 in 53 of the Site U1313 samples analysed in this study and by Lang et al. (2016) for
266 fish debris εNd using a standard method (e.g., Bond and Lotti, 1995).

267 It has been suggested that increased inputs of detrital carbonate sourced from
268 the Canadian Shield (Hudson Bay) to Site U1313 during Heinrich events are
269 responsible for millennial-scale unradiogenic excursions in authigenic εNd at this site
270 (e.g., Lang et al., 2016; Du et al., 2020). Yet, Heinrich layer deposition at Site U1313
271 was restricted to the past ~640 kyr (Naafs et al., 2014). Furthermore, IRD
272 accumulation rates at U1313 are extremely low during ambient glacial conditions
273 over the past ~3.3 Myr (Lang et al., 2014; Lang et al., 2016).
274
<table>
  <tr>
    <td>275</td>
    <td>2.5 Benthic foraminiferal boron/calcium ratios</td>
  </tr>
  <tr>
    <td>276</td>
    <td>Between 8 and 12 C. wuellerstorfi tests (212-500 µm) were cleaned for trace metal</td>
  </tr>
  <tr>
    <td>277</td>
    <td>analysis from 56 samples from Site 1267 spanning MIS M2, MIS 100 and the last</td>
  </tr>
  <tr>
    <td>278</td>
    <td>glacial-Holocene following Chalk et al. (2019). Foraminiferal tests were cracked open</td>
  </tr>
  <tr>
    <td>279</td>
    <td>and ultrasonicated several times in Milli-Q 18.2 MΩ·cm water to remove clay</td>
  </tr>
  <tr>
    <td>280</td>
    <td>materials. Samples were then oxidatively cleaned in a NH3OH buffered 1% hydrogen</td>
  </tr>
  <tr>
    <td>281</td>
    <td>peroxide solution before a 0.0005M weak acid leach was applied to remove adsorbed</td>
  </tr>
  <tr>
    <td>282</td>
    <td>cations. Finally, samples were dissolved in ~0.075 M nitric acid. The above steps</td>
  </tr>
  <tr>
    <td>283</td>
    <td>were all undertaken in the boron-free clean laboratory at the University of</td>
  </tr>
  <tr>
    <td>284</td>
    <td>Southampton. Element ratios were measured using a ThermoFisher Scientific Element</td>
  </tr>
  <tr>
    <td>285</td>
    <td>2XR-ICPMS at Southampton, using the protocol described by Rae et al. (2011)</td>
  </tr>
  <tr>
    <td>286</td>
    <td>employing matrix-matched in-house standards and a variety of consistency standards</td>
  </tr>
  <tr>
    <td>287</td>
    <td>to ensure reproducibility. Samples were screened for clay contamination using Al/Ca</td>
  </tr>
  <tr>
    <td>288</td>
    <td>and other contaminant ratios (e.g., Ba/Ca, Fe/Ca), with samples Al/Ca > 200</td>
  </tr>
  <tr>
    <td>289</td>
    <td>µmol/mol excluded from this study (n = 4 out of 56). Long-term reproducibility for</td>
  </tr>
  <tr>
    <td>290</td>
    <td>B/Ca during this study was evaluated using consistency standards analysed during</td>
  </tr>
  <tr>
    <td>291</td>
    <td>each analytical session and within 4% at 2 s.d. (~2 mmol/mol). Carbonate ion</td>
  </tr>
  <tr>
    <td>292</td>
    <td>concentrations were calculated using a sensitivity of Δ[CO2<sup>3-</sup>] on B/Ca of 1.14 ±</td>
  </tr>
  <tr>
    <td>293</td>
    <td>0.048 for C. wuellerstorfi (Yu and Elderfield, 2007), and a [CO2<sup>3-</sup>]<sub>sat</sub> calculated from</td>
  </tr>
  <tr>
    <td>294</td>
    <td>modern local pressure, temperature and salinity (see Supplementary Information</td>
  </tr>
  <tr>
    <td>295</td>
    <td>Section S3). The influence of this assumption on the calculated Δ[CO2<sup>3-</sup>] is very</td>
  </tr>
  <tr>
    <td>296</td>
    <td>small (&lt;0.5 umol/kg; Yu and Elderfield, 2007).</td>
  </tr>
  <tr>
    <td>297</td>
    <td></td>
  </tr>
  <tr>
    <td>298</td>
    <td>2.6 Water-mass geochemical end-members</td>
  </tr>
  <tr>
    <td>299</td>
    <td>We focus on relative change in water-mass composition between different glacials</td>
  </tr>
</table>
rather than absolute changes but, nevertheless, consider the potential for change in end-member compositions of NCW and SCW. To monitor the δ¹³C of NCW during our three target intervals, we use benthic δ¹³C data from shallow northeast ODP Site 982 (1145 m water depth; Venz et al., 1999). The question of how best to monitor the δ¹³C of SCW is more contentious. Sequences recovered from areas ideally located to monitor the end-member composition of SCW (e.g., in the Atlantic sector of the Southern Ocean) are carbonate-poor. Because of this issue, one pioneering study used benthic δ¹³C from mid-latitude South Atlantic ODP Site 1090 (~3700 m water depth) for this purpose for the Plio-Pleistocene (Venz and Hodell, 2002). We now know, however that, today and in the past, Site 1090 is partially bathed by NCW (Pena and Goldstein, 2014) making it ill-suited for this purpose. We therefore follow Raymo et al. (1990) and Oppo et al. (1995) and use the δ¹³C of Pacific deep water (equatorial Pacific ODP Site 849 (3851 m water depth; Mix et al., 1995) as a conservative, albeit aged, record of the end-member composition of SCW. However, we also compare benthic δ¹³C data from our study sites to the benthic δ¹³C stratigraphy of Southwest Pacific ODP Site 1123 (3290 m water depth, Fig. S1; Patterson et al., 2018), which is bathed today by a mixture of Antarctic Bottom water and Lower Circumpolar Deep Water. While its Plio-Pleistocene benthic δ¹³C stratigraphy shows that the composition of SCW bathing the deep Atlantic has changed over the period of our study interval, it also confirms that marked differences persisted in the δ¹³C of NCW and SCW during cold stages over the past 3.6 Ma. The δ¹³C of Site 1123 is thought to have decreased from 3.6 Ma in response to increased carbon remineralisation in SCW caused by Southern Ocean sea-ice expansion and a reduction in deep ocean ventilation (Patterson et al., 2018).

Knowledge of the εNd of NCW and SCW endmembers during the Plio-
325 Pleistocene, as well as the validity of using fish debris Nd isotope records to 326 reconstruct water-mass mixing in the deep Atlantic over this time, is comprehensively 327 discussed in Lang et al. (2016). Briefly, despite their extremely slow precipitation 328 rates, our best source of information on the evolution of the Nd isotope composition 329 of NCW and SCW during the Plio-Pleistocene arguably remains the εNd of Fe-Mn 330 crusts. To estimate the εNd of these water-masses during MIS M2 and MIS 100, we 331 therefore followed Lang et al. (2016) and took the average value of all published 332 northwest Atlantic Fe-Mn crust εNd data and their 2 s.d. that span ~3.3-2.4 Ma (from 333 BM1969.05, ALV539 2-1 and TR079-D14; see supplementary Table S7). For the last 334 glacial-Holocene, following the approach of Piotrowski et al. (2005), NCW εNd was 335 defined as the average of previously published Holocene data (<14 ka, MIS 2/1 336 boundary as defined by Lisiecki and Raymo (2005)) and their 2 s.d. for Site U1313 337 (Lang et al., 2016). We also define the SCW end-member εNd for all our time target 338 intervals using the εNd crust record D18-1 from the Indian Ocean Sector of the 339 Southern Ocean (see supplementary Table S7).

340 Traditionally, near-consistency in the NCW and SCW end members on orbital 341 timescales is widely assumed in Nd isotope studies of water-mass mixing (e.g., 342 Piotrowski et al., 2005; Böhm et al., 2014; Howe et al., 2016; Lippold et al., 2016). 343 Yet, recently published records (Wilson et al., 2014; Roberts and Piotrowski, 2015; 344 Hu et al., 2016; Pöppelmeier et al., 2018; 2019a; Pöppelmeier et al., 2019b; Zhao et 345 al., 2019; Huang et al., 2020) and numerical modelling studies (e.g., Gu et al., 2019; 346 Du et al., 2020), raise the possibility that NCW and SCW εNd may have changed, at 347 least, between the Last Glacial Maximum and the Holocene. In particular, the εNd of 348 cores from the New England slope at ~1800 m depth (KNR198), a region that should 349 have been bathed in NCW throughout the Last Glacial Maximum-Holocene, suggest
that the preformed composition of NCW may have been up to ~4 ε-units more
positive than today during the Last Glacial Maximum (Zhao et al., 2019). In this
revised framework, radiogenic excursions in εNd in the deep North Atlantic during the
Last Glacial Maximum may reflect changes in the end-member composition of NCW
rather than SCW incursion (Zhao et al., 2019; Pöpplemeier et al., 2020; Du et al.,
2020). One interpretation of these signals is that higher εNd in the deep North Atlantic
during the Last Glacial Maximum (by ~3.5 ε-units, e.g., Böhm et al., 2015; Lang et
al., 2016) reflects a reduction in chemical weathering of old unradiogenic cratonic
rocks on Canada and Greenland under extensive northern hemisphere ice-sheets and
that, when end-member changes are taken into account, the deep North Atlantic
remained bathed in near 100% NCW with northward expansion of SCW limited to the
deep South Atlantic (Zhao et al., 2019; Pöppelmeier et al., 2020).

The plausibility of temporal variations in NCW εNd being driven by chemical weathering has recently been challenged by Du et al. (2020) from a tracer budget perspective. The notion that the extent of chemical weathering on North America controls the εNd of NCW is also challenged by North Atlantic crust records which demonstrate that, as ice-sheet coverage of high northern latitude Archean and Proterozoic cratons increased over the past 3.3 Myr, the long-term Nd isotope composition of NCW has decreased (Burton et al., 1999; Vance and Burton, 1999; Reynolds et al., 2004). A dominant role for chemical weathering is also hard to reconcile with the fact that Site U1313 εNd appears to vary in phase with southern (not northern) hemisphere insolation, especially prior to 2.7 Ma (Supplementary Figure S6). An alternative interpretation of the high εNd values in the North Atlantic during the Last Glacial Maximum is that they reflect a reduction in the contribution to NCW of unradiogenic Labrador Seawater (Hillaire-Marcel et al., 2001a; 2001b) and a shift
in the location and mode of NCW formation to south of Iceland (Sarnthein et al., 1994) where its preformed composition is made more positive by 'bottom-up' non-conservative processes (i.e., partial dissolution of radiogenic authigenic (volcanic) phases) enhanced by a more sluggish glacial circulation (Du et al., 2020). Given these uncertainties, we proceed with caution herein, assuming near-consistency in Nd end-members on orbital timescales (the traditional interpretive framework), but also discuss the implications of this revised framework.

382

## 3 Results and Discussion

### 3.1. Nd isotopes and B/Ca ratios

In Figures 3 and 4 we integrate new and published data to present high resolution records of fish debris-derived $\varepsilon_{Nd}$ and B/Ca and $\delta^{13}$C in benthic foraminiferal calcite for the deep Atlantic Ocean during MIS M2, ~3.3 Ma, MIS 100, ~2.52 Ma, and the last glacial-Holocene, ~40-0 ka. Our reconstruction of $\varepsilon_{Nd}$ and $[CO_3]^{2-}$ in bottomwaters bathing Site 1267 during the latest Holocene (averages = ~ -10.7±0.2 and 93±10 µmol kg$^{-1}$) are strikingly similar to those reported for LDW bathing Walvis Ridge in the Angola Basin today ($\varepsilon_{Nd}$ = -10.6±0.7; 30°S 1°W, 4675 m water depth (Jeandel, 1993) and $[CO_3]^{2-}$ = ~100 µmol kg$^{-1}$ (Lauvset et al., 2016)) strongly suggesting that our data from this site faithfully track changes in deep water-mass mixing in the Angola Basin (also see Supplementary Information Section 5).

The magnitude of peak interglacial-full glacial change in $[CO_3]^{2-}$, $\varepsilon_{Nd}$ and benthic $\delta^{13}$C at Site 1267 is comparable between all three of our studied time windows (with average IG-G change of ~10-22 µmol kg$^{-1}$, ~0.7-1.9 $\varepsilon$-units and ~0.4-0.5 ‰, respectively; data in red in Fig. 3d-l; Tab. 1). The new highly resolved Nd isotope record from Site U1313 for MIS 100 confirms the last glacial-magnitude shift
in $\epsilon_{Nd}$ (~1.85 for peak interglacial to full glacial; data in black and grey in Fig. 3h; Tab. 1) inferred from a less well resolved record (data in grey in Fig. 3h; Lang et al., 2016). The new high-resolution record for MIS M2 at Site U1313 also shows that, while there is a prominent short-term maximum in $\epsilon_{Nd}$ during the early glacial phase of M2 (~3.305 Ma, labelled EG in Fig. 3i), by the time full glacial (FG) conditions are reached at this site, $\epsilon_{Nd}$ is only ~0.7 $\epsilon$-units higher on average than during the preceding interglacial MIS MG1 (Tab. 1). No analogous early glacial short-term structure in $\epsilon_{Nd}$ is documented in the corresponding benthic $\delta^{13}$C record for MIS M2 (compare data in grey in Fig. 3i to 3l). The magnitude of change in benthic $\delta^{13}$C at Site U1313 is small throughout MIS M2 (average IG-G change of only ~0.2 ‰) relative to both MIS 100 (average IG-G change of ~0.75 ‰) and the last glacial-Holocene (average IG-G change of ~0.55 ‰).

412

## 3.2. Deep Atlantic Ocean structure during the Last Glacial-Holocene

Our data reveal distinct geochemical differences between bottom waters at Site 1267 (today bathed in 70-80% lower NADW) and Site U1313 (today bathed in 100% lower NADW) in $\epsilon_{Nd}$, benthic $\delta^{13}$C and $[CO_3]^{2-}$ over the past ~10 kyr (>+2 $\epsilon$-units, -0.3 ‰ and ~-20 $\mu$mol kg$^{-1}$, respectively) (Fig. 4a, d, j). These observations demonstrate that SCW influenced bottom-waters bathing the northern flank of Walvis Ridge throughout the Holocene. Our $[CO_3]^{2-}$ estimates for MIS 2 at Site 1267 (data in red in Fig. 4a) are comparable to those reported for the deep northeast North Atlantic (e.g., IODP Site U1308 at 51°N and ~3800 m depth, data in dark blue in Fig. 4a) but they are ~80 $\mu$mol kg$^{-1}$ lower than those estimated for the mid-depth northeast North Atlantic (e.g., ODP Site 980 at 55°N and ~2170 m depth, data in green in Fig. 4a). The new $[CO_3]^{2-}$ estimates from Site 1267 therefore also show that the pronounced
carbonate-system depth-stratification reconstructed (e.g., by Yu et al., 2008; Chalk et al., 2019) for the eastern Atlantic basin during the last glacial (compare [CO3<sup>2-</sup>] 980 data in green with U1308 and 1267 data in dark blue and red, respectively, in Fig. 4a) probably extended southwards to Walvis Ridge (at ~30°S). The new data are also consistent with the suggestion (Chalk et al., 2019) that the dissolved inorganic carbon (DIC) contents of the deep eastern Atlantic, at depths >3800 m, may be larger than those of contemporaneous deep waters in the western Atlantic (compare data in red, dark blue and black in Fig. 4a).

The simplest explanation for the radiogenic εNd, low benthic δ<sup>13</sup>C and low B/Ca ratios in our MIS 2 datasets for Site 1267 is that they reflect an increased influence of SCW on deep-waters bathing Walvis Ridge in the Angola Basin relative to the Holocene. This explanation holds true even when recent assumptions about potential changes in NCW and SCW during the Last Glacial Maximum (e.g., Zhao et al., 2019; Pöppelmeier et al., 2020) are considered because our Site 1267 εNd are still more radiogenic than those recorded at Site KNR198, a recently proposed record of preformed NCW for the Last Glacial Maximum (Zhao et al., 2019); Fig. 4d). In this explanation, the pathway for SCW to the Angola Basin is the Romanche Fracture zone, which today provides an equatorial passageway for SCW to this basin down to 4350 m (labelled RFZ in Fig. 2). This fracture zone was recently invoked as a route for SCW during MIS 2 and used to explain low [CO3<sup>2-</sup>] in bottom waters bathing the northeast North Atlantic (Site U1308 at 51°N) during this time (Chalk et al., 2019). As in the modern, the Vema Fracture Zone (VFZ), situated further north but in even deeper water (~5000 m) (Fig. 2; Demidov et al., 2006) provides an even more effective conduit for SCW flow to reach the northeast North Atlantic during Plio-Pleistocene glacials. The deeper VFZ may help explain why bottom waters bathing
450 Site U1308 during the Last Glacial Maximum have lower $[CO_3^{2-}]$ than those bathing
451 Site 1267.

452 Today, the $[CO_3^{2-}]$ of deep-waters can be used as a quasi-conservative tracer of
453 water-mass source, but this may not have been the case during MIS 2. A variety of
454 processes unrelated to water-mass mixing can modify the $[CO_3^{2-}]$ of the ocean (e.g.,
455 air-sea exchange, alkalinity change and biological processes), but given that $[CO_3^{2-}]$
456 will react differently to these processes compared to $\delta^{13}C$, paired $[CO_3^{2-}]$-$\delta^{13}C$ data
457 can be used to determine whether changes in the former reflect changes in water-mass
458 structure or non-conservative processes (e.g., Yu et al. 2008; Chalk et al., 2019). The
459 trend of paired $[CO_3^{2-}]$-$\delta^{13}C$ data from Site U1308 for the last glacial and Holocene
460 approximates best to a trend line indicative of $CO_2$ influx (e.g., of ocean uptake of
461 carbon), indicating carbonate system proxy data can be used at this site as a quasi-
462 conservative tracer of SCW-NCW mixing (small black filled and hollow circles in
463 Fig. 5a). By contrast, the gradients of paired $[CO_3^{2-}]$-$\delta^{13}C$ data from sites 1267 and
464 U1313 for the same time interval in $\delta^{13}C$-$[CO_3^{2-}]$ space is significantly steeper (solid
465 and hollow green and red circles in Fig. 5a) and better approximate the slope of the
466 ‘biology line’ of Yu et al. (2008). This observation indicates that carbon
467 remineralization may have contributed to the paired low $\delta^{13}C$ and $[CO_3^{2-}]$ values at
468 sites 1267 and U1313 during MIS 2 (by decreasing the $\delta^{13}C$ of DIC bathing 1267 and
469 U1313 beyond that anticipated from linear mixing of SCW-NCW alone).
470 Nevertheless, an increase in the contribution of SCW to these two sites relative to the
471 Holocene is indicated by our $\epsilon_{Nd}$ records (Fig. 4d).

472 Dense Norwegian-Greenland Sea (NGS) overflows are proposed to contribute
473 to the deep waters bathing Walvis Ridge throughout much of the Pleistocene (Bell et
474 al., 2014; 2015). This suggestion invokes a deep Eastern Atlantic route for this water-
mass to the Angola Basin. This possibility cannot be ruled out for our study intervals based on our εNd data alone because today both Nordic Sea overflows and AABW share a radiogenic εNd signature (of ~-7 to -9; Crocker et al., 2016). Yet [CO₃²⁻] is markedly higher for modern NGS overflow water (>115 µmol kg⁻¹) than for SCW (~85 µmol kg⁻¹) and this contrast would have been greater during the Last Glacial Maximum than today (Yu et al., 2008; Crocker et al., 2016). Thus, it is unlikely that the radiogenic (εNd = -9, Fig. 4d-e) and [CO₃²⁻] low (= ~80±10 µmol kg⁻¹, Fig. 4a-b) deep-waters bathing Site 1267 during the MIS 2 and MIS 100 were composed of dense overflow waters sourced directly from the NGS via an abyssal eastern basin route. This is a robust interpretation even if their composition was overprinted en route by extensive biological remineralization of carbon.

486

### 3.3. Deep Atlantic Ocean structure during MIS M2 and 100

The consistency in the amplitude of IG-G change in all three water-mass proxies at Site 1267 during all three of our target time slices (Tab. 1) suggests that the oceanographic response to MIS M2, MIS 100 and MIS 2 in the deep South Atlantic was similar across all three glacials, i.e., all proxies reflect an increase in the influence of SCW. In contrast, IG-G changes in δ¹³C and εNd at Site U1313 suggest a smaller contribution of SCW to deep waters bathing this shallower more northerly situated site during MIS M2 than during MIS 100 and MIS 2 (Tab. 1).

The validity of this conclusion depends on there not being major changes in the offset between the end-member compositions of NCW and SCW during these times. A last-glacial-Holocene-like change in εNd at Site 1267 during MIS M2 and MIS 100 could be artificially produced by a more positive SCW εNd during these cold stages relative to MIS 2. Because SCW is a mixture of relatively unradiogenic Atlantic and
radiogenic Pacific waters (Stichel et al., 2012) a more positive SCW could arise in two ways. First, if the composition and mixing contribution of the Pacific end-member to SCW remained similar during our three studied glacials, a more positive NCW would lead to a more positive SCW when it mixes with Pacific deep waters in the Southern Ocean. Second, SCW could become more positive if NCW makes a smaller contribution to SCW due to a reduction in the southward advection rate of North Atlantic deep waters. Based on the available crust data (Frank et al., 2002), it is reasonable to assume that the $ε_{Nd}$ composition of the Pacific end-member remained stable during our study intervals, but the Nd isotope composition of NCW did not (compare red and blue horizontal bars in Fig. 4d-f). Both the crust records and Site U1313 $ε_{Nd}$ illustrate that NCW was ~1 ε-unit more radiogenic during the late Pliocene than it was during the late Pleistocene (Fig. 4d-f). Yet, a relatively radiogenic NCW during MIS 100 and MIS M2 would not have changed the gradient in $ε_{Nd}$ between NCW-SCW during these cold stages relative to MIS 2 because it would also have made SCW more radiogenic. We can also rule out the possibility that a reduction in the southward advection rate of NCW relative to MIS 2 could explain the existence of a last-glacial-Holocene-like change in $ε_{Nd}$ at Site 1267 during MIS M2 and MIS 100 because it is unlikely that Atlantic Meridional Overturning Circulation (AMOC) weakening was greater during late Pliocene and earliest Pleistocene cold stages than during MIS 2.

Our interpretations of relative changes in our water-mass proxy data are supported by a comparison between the benthic $δ^{13}$C data from our study sites and the well-constrained $δ^{13}$C-based estimates of NCW and SCW end members (Fig. 4j-l). A comparison of benthic $δ^{13}$C values at Site 1267 to those at Site 982 (the NCW end-member) reveals that there was an increased contribution of SCW bathing Walvis
Ridge for most of early- to mid-glacial M2 until ~3295 ka (compare data in red and purple in Fig. 4l). While our knowledge of NCW and SCW $\epsilon_{Nd}$ during MIS M2 is less certain, our Nd isotope data from this site record the same history of SCW incursion during this cold stage as inferred from benthic $\delta^{13}$C (compare data in red and orange in Fig. 6d). In contrast, at Site U1313, benthic $\delta^{13}$C values are comparable to those of NCW throughout MIS M2 (compare data in black and purple in Fig. 4l). We infer that, while SCW contributed significantly to the deep waters bathing Site 1267 during this glacial at least until ~3295 ka, Site U1313 was mainly bathed by NCW during MIS M2. We note the radiogenic spike in Nd during early MIS M2 (star in Fig. 4f) but find no supporting evidence that it arises from water-mass mixing (see supplementary information Section S6).

### 3.3.1 What would be the implication of changing $\epsilon_{Nd}$ end members for our interpretation of MIS M2?

We have interpreted the history of deep Atlantic Ocean structure using the traditional framework wherein the end-member $\epsilon_{Nd}$ compositions of NCW and SCW did not change appreciably on glacial-interglacial timescales (e.g., Foster et al., 2007; Piotrowski et al., 2005; Böhm et al., 2015; Howe et al., 2016). If major end-member change in NCW occurred over the last glacial cycle (of ~+4 $\epsilon$-units) then it is possible that the deep North Atlantic remained bathed in near 100% NCW during the Last Glacial Maximum and northward expansion of SCW was limited to the deep South Atlantic (Zhao et al., 2019; Pöppelmeier et al., 2020). However, an absence of SCW in the deep North Atlantic during the Last Glacial Cycle is hard to reconcile with the fact that benthic $\delta^{18}$O from, e.g., the Iberian Margin correlates so well with Antarctic temperature (Shackleton et al., 2000). Nevertheless, in this scenario, the large
negative excursion recorded in North Atlantic benthic δ¹³C records during the Last
Glacial Maximum is proposed to reflect enhanced remineralisation of carbon in deep-
waters (Howe et al., 2016) under a more sluggish (i.e., diminished AMOC) glacial
deep Atlantic circulation relative to modern (Böhm et al., 2015; Ng et al., 2018).

What would be the implication for our interpretation of MIS 100 and M2 if this
alternative framework is correct? There are two suggested mechanisms for NCW end-
member change. In the first case, change is driven by chemical weathering extent on
North America (Zhao et al., 2019; Pöppelmeier et al., 2020) such that deep North
Atlantic εNd acts as a proxy of Plio-Pleistocene North American Ice Sheet extent. In
this case, the large radiogenic εNd excursions during MIS 2 and MIS 100 at Site
U1313 (Lang et al., 2016; this study; Fig. 3g-h; Tab. 1) signal comparable reductions
in chemical weathering on North America and Greenland due to a similar spatial
configuration of ice-sheets on Canada and Greenland whereas the relatively
unradiogenic εNd values during MIS M2 point to more limited glaciation of North
American and Greenland (Lang et al., 2016; this study; Fig. 3g-i; Tab. 1). The second
mechanism for large radiogenic εNd excursions involves relabelling of NCW via
partial dissolution of east Greenlandic- and Icelandic-sourced radiogenic volcanic
phases deposited in the Iceland basin, following a shift in the location of the main site
of NCW formation from the sea ice-covered Nordic Seas to south of Iceland, and a
reduction in unradiogenic Labrador Seawater production (Du et al., 2020). In this
case, our data for MIS 2 and MIS 100 signal a slowdown in AMOC (reduced bottom
current speeds in the Iceland Basin promotes labelling via partial dissolution of
volcanic phases because the longer benthic exposure time that results allows seawater
and authigenic εNd in this region to converge toward detrital sediment εNd (Du et al.,
2020)), whereas the muted IG-G shift in εNd at Site U1313 during MIS M2 (Tab. 1)
signals vigorous formation of NCW in largely ice-free Nordic and Labrador seas.

Either way, if radiogenic excursions in the deep North Atlantic during glacials reflect end-member changes in NCW εNd, the lack of a large negative excursion in Site U1313 benthic δ¹³C during MIS M2 (compare data in red in Fig. 3l and 3g-h; Tab. 1) implies reduced remineralisation of ¹²C in the deep North Atlantic relative to the Last Glacial Maximum and MIS 100 under a more vigorous AMOC (Pöppelmeier et al., 2020). Thus, regardless as to whether our North Atlantic εNd records reflect changes in water-mass provenance or a change in the preformed composition of NCW, they strongly suggest that Atlantic Ocean circulation during MIS M2 was markedly different to that of large Quaternary glaciations with markedly weaker chemical stratification.

### 3.4 Implications for glacial CO₂-storage in the ocean abyss

A ~110 ppm decline in atmospheric CO₂ during MIS M2 (Fig. 6a) strongly implies increased carbon storage in the abyssal ocean (de la Vega et al., 2020). A Southern Ocean-focused driving mechanism is favoured by de la Vega et al. (2020) because the tail end of the CO₂ decline documented appears out of phase with the decline in northern hemisphere insolation and benthic δ¹⁸O but in-phase with both insolation decline at 65°S (compare CO₂ record in Fig. 6a to data in black in 6b,f) and Nd isotope-based evidence for SCW incursion into the North Atlantic (Lang et al. (2016) (compare Fig. 6a to U1313 εNd data points denoted by grey boxes in 6c). Yet, in contrast to previous suggestions (e.g., Lang et al., 2016), we find that SCW expansion during MIS M2 was limited to the deep South Atlantic (as tracked by 1267). Our data also suggest that this modest SCW expansion in the Atlantic during MIS M2 peaked ~10-15-kyr before minimum values in atmospheric CO₂ were reached during this
600 glacial (compare red horizontal bar in Fig. 6d to CO₂ data in 6a). Together, these
601 observations imply that the deep Atlantic was not the ultimate store for the CO₂
602 sequestered from the atmosphere during this cold stage. Future work should focus on
603 the possibility that the deep Pacific played a major CO₂ storage role, as hinted at by
604 the temporal relationship between benthic δ¹³C at Site 1123 and atmospheric CO₂
605 during MIS M2 (compare Fig. 6a to 1123 δ¹³C data in blue in 6d).

606

607 **3.5 How sluggish was AMOC during MIS M2?**

608 The lag between atmospheric CO₂ and benthic δ¹⁸O during MIS M2 (Section 3.4)
609 implies that high latitude cooling and glaciation during this cold stage were not
610 simply controlled by radiative forcing from greenhouse gases. Instead, other
611 mechanisms must have been important. One mechanism invoked to have played at
612 least a part in amplifying glaciation of the high northern latitudes is AMOC
613 weakening (De Schepper et al., 2013).

614 AMOC suppression of a magnitude comparable to that reconstructed for major
615 Quaternary glaciations including the Last Glacial Maximum (e.g., Naafs et al., 2010;
616 Bolton et al., 2018), is invoked for M2 based on evidence from dinoflagellate
617 assemblages for a southward deflection of the North Atlantic Current (NAC) to ~41°N
618 (see panel ‘3’ in Figure 6 of De Schepper et al., 2013) and significant reductions in
619 North Atlantic SSTs (of up to 6°C; Fig. 7f). The sediment abundance of the
620 dinoflagellate *Operculodinium centrocarpum*, which has a strong affinity with NAC
621 waters in the modern (De Schepper et al., 2013), in Site U1308 (50°N) and 610 (53°N)
622 (Fig. 7g) provide evidence for NAC weakening and/or its deflection south of ~50°N
623 throughout MIS M2 (Fig 8c). The abundance of *O. centrocarpum* in Site U1313
624 sediments deposited at 41°N during this cold stage is only high, however, during early
625 glacial conditions (Fig. 7h), indicating that maximum southward deflection of the
626 NAC during this time occurred well before the full glacial MIS M2 (Fig. 8d). An
627 early glacial timing for maximum NAC decline during MIS M2 is consistent with our
628 observation that SCW incursion into the deep Atlantic reached its MIS M2 peak
629 before ~3295 ka (compare Fig. 7b-c to 7h).

630 Palaeoceanographic records suggest that AMOC can remain relatively strong in
631 the face of a significant reduction in the volumetric extent of NCW (e.g., Böhm et al.,
632 2015). Yet, numerical models of ocean circulation indicate that if AMOC vigour
633 during MIS M2 was comparable to the Last Glacial Maximum, it should be associated
634 with a similarly large (i.e., late Pleistocene glacial-like) reduction in the volumetric
635 importance of the NCW overturning cell (e.g., Nikurashin and Vallis, 2011; Kostov et
636 al., 2014; Burke et al., 2015). If changes in εNd at U1313 reflect the magnitude of
637 SCW incursion into the deep North Atlantic during our three target glacials, our new
638 datasets show that M2 was not associated with NCW shoaling on the scale of the last
639 glacial or even MIS 100, 2.52 Ma (Fig. 4d-f, j-l). If the evolution of εNd at U1313
640 instead reflects changes in the preformed composition of NCW (e.g., Zhao et al.,
641 2019; Pöppelmeier et al., 2020), the small amplitude of IG-G change in εNd and
642 benthic δ¹³C at this site during MIS M2 would also be inconsistent with a
643 significantly weaker-than-modern AMOC (see Section 3.3.1). Either way, our new
644 records strongly suggest that AMOC during this cold stage was not as sluggish as
645 during major Quaternary glacials.

646 Our interpretations are consistent with the surface temperature history of the
647 North Atlantic. A ~6°C decline in SSTs in the subpolar North Atlantic during MIS
648 M2 provides clear evidence for a reduction in northward heat transport from a
649 warmer-than-pre-industrial baseline but temperatures remained as warm (Sites 982
and U1313) or warmer (Sites 642 and 610) than the Holocene average (Fig. 7f). These
observations may help us to understand why firm evidence for glacial expansion in
the high northern latitudes ~3.3 Ma is limited to circum-Arctic landmasses,
Greenland, Svalbard and Iceland (De Schepper et al., 2014; Tan et al., 2017). We note
that sediments interpreted as glacial tills from mid-latitude James Bay Lowland,
Canada, have been loosely dated to ~3.5 Ma (Gao et al., 2012), but this interpretation
has been called into question (Clague et al., 2020).

657

658 **4. Conclusions**

We present new Nd isotope data and benthic foraminiferal B/Ca ratios for southeast
Atlantic Site 1267 (29°S, ~4350 m water depth) and northwest Atlantic Site U1313
(41°N, ~3400 m water depth) spanning the last glacial-Holocene, MIS 100, 2.52 Ma,
and MIS M2, ~3.3 Ma, often inferred to signal a prelude to Quaternary extensive
glaciation of the northern hemisphere and therefore a bi-polar Cenozoic icehouse
climate state. We demonstrate that Southern Component Water expanded into the
deep South and North Atlantic during MIS 100. The resulting glacial water-mass
properties at our study sites are comparable to those that we document for the last
glacial. This Quaternary glacial picture contrasts with the one that we document for
MIS M2, in which Northern Component Water persists in the deep North Atlantic.

Our new observations are incompatible with AMOC suppression during MIS
M2 to the scale of MIS 100 or 2 and suggest that the atmospheric CO₂ decline during
this cold stage was not achieved by carbon storage in the deep Atlantic Ocean,
implying storage in deep Pacific Ocean reservoirs. Our findings, together with
Holocene-like surface ocean temperatures in the high latitude North Atlantic, help to
resolve the mystery of why there is no evidence to suggest that the
675 northern hemisphere was extensively glaciated during M2.

676

677 **5. Acknowledgements**

678 This research uses samples provided by the IODP, which is sponsored by the U.S.

679 National Science Foundation and participating countries under management of Joint

680 Oceanographic Institutions, Inc. We thank the shipboard party of IODP Expedition

681 306 and A. Wuelbers and W. Hale for their help at the Bremen Core Repository and

682 Bastian Hambach and Megan Spencer for laboratory support in Southampton. We

683 also thank the editor Antje Voelker and three anonymous reviewers for comments and

684 suggestions that helped to improve this manuscript considerably. I.B. acknowledges

685 support from the EU project MARINEFF (Interreg VA France-(Channel)-England

686 Programme project #162). M.E.R. acknowledges support from the Vetlesen

687 Foundation. P.A.W. and G.L.F. acknowledge funding from Royal Society Wolfson

688 Merit Awards and NERC grants NE/F00141X/1 to P.A.W. and I.B. and NE/T012382

689 to P.A.W. and A.B.

690

691 **Data Availability:** The data presented in this manuscript will be made available upon

692 request. These data will be deposited in PANGAEA upon acceptance on this

693 manuscript for publication.

694

695 **Declaration of interests:** none.

696

697 **Author Contributions:** N.K generated the Site 1267 data and contributed to

698 manuscript writing; I.B. designed the study, supervised N.K. in data generation and

699 wrote the manuscript; A.B generated the Site U1313 stable isotope data for MIS 100
and contributed to manuscript writing; T.B.C. supervised R.L.P in generating the Site
1267 benthic B/Ca data and contributed to the ideas presented and to manuscript
writing; R.L.P generated a subset of the Site 1267 data; A.J.C. supervised N.K. and
R.L.P in data generation, contributed to the ideas presented and to manuscript writing;
V.E.T. supervised N.K. and R.L.P. in data generation, generated the volcanic glass
count data and contributed to manuscript writing; J.A.M helped to generate the Nd
isotope data; G.L.F. contributed to the ideas presented and manuscript writing;
M.E.R. contributed to the ideas presented and manuscript writing; D.K. contributed
samples and to the ideas presented and manuscript writing; D.B.B contributed
samples and to manuscript writing; P.A.W. designed the study, supervised N.K. in
data generation, co-wrote the manuscript with I.B.

711
712

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1139

1140 **Figure Captions**

1141 **Figure 1.** Climate indicators across marine isotope stage (MIS) M2, ~3.3 Ma, and the

1142 Plio-Pleistocene intensification of northern hemisphere glaciation, ~3.5-2.5 Ma

1143 (Mudelsee and Raymo, 2005): (A) Site U1313 εNd composition of fish debris (Lang et

1144 al., 2016). Estimates of the εNd composition of Northern Component Water, NCW,

1145 (red bar) follow Lang et al. (2016) and is taken as the average value of Fe-Mn crust

1146 εNd (and their 2 s.d.) from TR079-D14 (Reynolds et al., 1999), BM1969.05 (Burton et

1147 al., 1997; 1999; on chronology of O’Nions et al., 1998), ALV539 2-1 (Burton et al,

1148 1999; on chronology of O’Nions et al., 1998) spanning 2.4-3.3 Ma. Estimate for
1149 Southern Component Water, SCW, (blue bar) is average value of Fe-Mn crust $\epsilon_{Nd}$
1150 (and their 2 s.d.) from D18-1 (Frank et al., 2002) spanning 2.4-3.3 Ma; (B) Site
1151 U1313 (Lang et al., 2014), Site 1267 (Bell et al., 2014) and ODP Site 982 (Hodell and
1152 Venz-Curtis, 2006) benthic $\delta^{13}C$; (C) a stack of globally representative benthic $\delta^{18}O$
1153 records, the LR04 (grey; Lisiecki and Raymo, 2005) and benthic $\delta^{18}O$ from North
1154 Atlantic Ocean IODP Site U1313 (black; Bolton et al., 2010) and Walvis Ridge (red;
1155 Bell et al., 2014). The mid Piacenzian warm period (mPWP) is labelled. The
1156 black/white bar at the bottom of the figure denotes palaeomagnetic (sub)chronozones
1157 (Cande and Kent, 1995); M = Matuyama, G = Gauss, K = Keana and Ma =
1158 Mammoth. See Figure 2 and Supplementary Figure S1 for site locations.

1159
1160 **Figure 2.** Location of Atlantic Ocean drill sites used in this study (Sites 1267 and
1161 U1313, red circles) and others (black circles) discussed in the main text. Also shown
1162 is a schematic of modern pathways of deep-waters relevant to our study sites: NADW
1163 = North Atlantic Deep Water (dark grey lines), AABW = Antarctic Bottom Water
1164 (light blue lines), LSW = Labrador Sea Water, DSOW = Denmark Strait Overflow
1165 Water, ISOW = Iceland Scotland Overflow Water (ISOW), WTRO = Wyville
1166 Thomson Ridge Overflow, NGS = Norwegian Greenland Seas. The Verma (VFZ),
1167 Romanche (RGZ), Chain (CFZ), Rio de Janeiro (RJFZ) and Rio Grande (RGFZ)
1168 fracture zones and Walvis Passage (WP) are also labelled and represent important
1169 pathways for AABW to the eastern Atlantic (see main text). The Kane Gap (GP) is
1170 also labelled. Depths stated for VFZ, RGZ, RJFZ and RGFZ are known maximum
1171 depths through which SCW can pass through these fracture zones. Bottom-water
1172 pathway reconstructions based on: McCartney et al. (1991); McCartney (1992); Coles
1173 et al. (1996); Smyth-Wright and Boswell (1998); Stramma and England (1999);
spanning 2.4-3.3 Ma. Also shown in (D) is a composite Nd isotope record from the New England Slope in the northwest North Atlantic, considered by Zhao et al. (2019) to record the evolution of NCW over the past ~21 ka. Estimate for the SCW end-member during MIS M2 and MIS 100 is average value of Fe-Mn crust (εNd (and their 2 s.d.) from D18-1 (Frank et al., 2002) spanning 2.4-3.3 Ma. Last glacial-Holocene NADW defined as average (and their 2 s.d.) of Holocene U1313 data (<14 ka; MIS 2/1 boundary as defined by Lisiecki and Raymo, 2005). Last glacial-Holocene SCW end-member estimated using an average value of Fe-Mn crust (εNd (and their 2 s.d.) from D18-1 (Frank et al., 2002) spanning 0-0.04 Ma. Vertical range bars in D show range of modern Upper and Lower North Atlantic Deep Water (UNADW and LNADW, respectively) εNd (Lambelet et al., 2016); (G-I) fresh volcanic glass and ice-rafted debris (IRD) >150 µm per gram of dry sediment at Site U1313. Last glacial, MIS 100 and MIS M2 IRD data from Lang et al. (2014), Lang et al. (2016) and Bolton et al. (2010); (J-L) benthic δ¹³C data (including estimates of the δ¹³C of NCW and SCW); (M-O) benthic δ¹⁸O data. Error bar shown for benthic δ¹³C (δ¹⁸O) signifies that the uncertainty for the different sources of these data plotted is no larger than 0.1 ‰ (0.15 ‰) at 2 s.d. See Figure 2 and Supplementary Figure S1 for site locations.

**Figure 5.** δ¹³C: [CO₃²⁻] cross plot. B/Ca-derived carbonate data and co-occurring δ¹³C data from ODP Site 1267 for Marine Isotope Stage, MIS, 2-Holocene (A), and MIS 99, 100, 101, M1, M2 and MG1 (B). Data from IODP U1313 (hollow green interglacial circles and solid green glacial circles; Chalk et al., 2019), IODP U1308 (hollow black interglacial circles and solid black glacial circles; Chalk et al., 2019), ODP 980 (hollow light brown interglacial circles and solid light brown glacial circles;
Chalk et al., 2019) are also shown in (A). Shown in both panels are shaded endmember compositions for the Southern Ocean (grey oval, Yu et al., 2014) and the Pacific (yellow oval, Kerr et al. (2017), based on Last Glacial Maximum and Holocene data from these regions. The Last Glacial Maximum northern (NADW) endmember is assumed to be denoted by the cluster of glacial ODP 980 points (solid light brown circles). Calculated slopes for CO₂ addition/evasion due to AABW-NADW mixing only (pale blue, dashed) and biological utilization (black, dashed) are shown. Sample gradients for alkalinity and temperature change are also shown, following Yu et al. (2008). See Figure 2 for site locations.

**Figure 6.** Palaeoceanographic proxies spanning marine isotope stage (MIS) M2, ~3.3 Ma: (A) interpolated 1-kyr resolution δ¹¹B-derived CO₂ data from *Globigerinoides ruber* white sensu stricto (with a 6-point running mean) at ODP Site 999 (Martinez-Boti et al., 2015; de la Vega et al., 2020). Shading = 68% (dark red) and 95% (light red) confidence intervals; (B) summer boreal (June 21st) and austral (December 21st) insolation forcing at 65°N and 65°S, respectively (Laskar, 2004); (C) Site U1313 benthic δ¹³C (Lang et al., 2014) and fish debris εNd (thus study; Lang et al., 2016) and estimates of the δ¹³C of NCW (Site 982; Venz et al., 1999) and SCW (Site 849; Mix et al., 1995; Site 1123; Patterson et al., 2018) with data symbols removed for clarity; (D) Site 1267 benthic δ¹³C (Lang et al., 2014) and fish debris εNd (thus study) and estimates of the δ¹³C of NCW (Site 982; Venz et al., 1999) and SCW (Site 849; Mix et al., 1995; Site 1123; Patterson et al., 2018) with data symbols removed to aid plotting clarity; (E) [CO₃²⁻] estimates from Site 1267 based on benthic B/Ca data; (F) Benthic δ¹⁸O data from Site U1313 (Bolton et al., 2010) and Site 1267 (this study).

Horizontal dashed line in D represents atmospheric CO₂ value considered to be an
important threshold for extensive glaciation of continents in the northern hemisphere (De Conto et al., 2008). All records plotted on LR04-based age models. See Figure 2 and Supplementary Figure S1 for site locations.

**Figure 7.** Palaeoceanographic proxies spanning marine isotope stage (MIS) M2, ~3.3 Ma and periods of early and late Pliocene warmth: (A) globally representative benthic δ¹⁸O stack the LR04 (grey; Lisiecki and Raymo, 2005) and benthic δ¹⁸O from North Atlantic Ocean IODP Site U1313 (black; Bolton et al., 2010); (B) fish debris Nd isotope data (εNd) from sites 1267 and U1313 alongside estimates of the εNd composition of NCW and SCW water-mass end-member (see Figure 4 and main text for methods); (C) benthic δ¹³C data from sites 1267 and U1313 and estimates of the δ¹³C of Northern component waters (NCW) and southern component waters (SCW); (D) interpolated 1-kyr resolution δ¹¹B-derived CO₂ data from *Globigerinoides ruber* white sensu stricto (with a 6-point running mean) at ODP Site 999 (Martinez-Boti et al., 2015; de la Vega et al., 2020). Shading = 68% (dark red) and 95% (light red) confidence intervals; (E) summer boreal (June 21st) and austral (December 21st) insolation forcing at 65°N and 65°S, respectively (Laskar, 2004); (F) Alkenone-based reconstructions of sea-surface temperature (SST) from Nordic Seas ODP Hole 642B at 67°N (Bachem et al., 2017), subpolar North Atlantic subpolar North Atlantic ODP Site 982 at 58°N (Lawrence et al., 2009), DSDP Hole 610A at 53°N (Naafs et al., 2020) and IODP Site U1313 at 41°N (Naafs et al., 2010); abundance of dinoflagellate cyst *Operculodinium centrocarpum* (De Schepper et al., 2009; De Schepper et al., 2013), with increased numbers taken to indicate the presence of the NAC over a study site (De Schepper et al., 2013), at (G) DSDP Site 610 and IODP Site U1308, and at (H) Site U1313 (De Schepper et al., 2013). All data plotted on LR04 or respective published age models.
1274 mPWP = mid Piacenzian Warm Period. Horizontal dashed line in D represents atmospheric CO₂ value considered to be an important threshold for extensive glaciation of continents in the northern hemisphere (DeConto et al., 2008). Colour-coded horizontal dashed lines in F are averages of equivalent Holocene-aged SST data from the same site. A modern SST value is highlighted for Hole 610A since no equivalent Holocene-aged alkenone SST data are available for this site. See Figure 2 and Supplementary Figure S1 for site locations.

1281

1282 **Figure 8.** Sub-polar North Atlantic dinoflagellate cyst records and their schematic interpretation in terms of North Atlantic Current (NAC) evolution during MIS M2: (Ai) abundance of dinoflagellate cyst *Operculodinium centrocarpum* (De Schepper et al., 2009; De Schepper et al., 2013), with increase numbers taken to indicate the presence of the NAC over a study site (De Schepper et al., 2013), at DSDP Site 610 and IODP Site U1308, (Aii) and at Site U1313 (De Schepper et al., 2013). (Aii) Globally representative benthic δ¹⁸O stack the LR04 (grey; Lisiecki and Raymo, 2005) and benthic δ¹⁸O from North Atlantic Ocean IODP Site U1313 (black; Bolton et al., 2010). Numbers 1-3 in (A) relate to NAC pathway interpretations illustrated in map shown in B-D. The most southerly position for the NAC during MIS M2 shown in map 2 is inferred from peak abundances of *O. centrocarpum* in U1313 sediments during early glacial conditions. A latitudinal position for the NAC in map 3 that is intermediate between the extremes shown in maps 1 and 2 is based on low abundances of *O. centrocarpum* in both U1313 and 610A/U1308 sediments during full glacial MIS M2.

1297

1298
1299 Table 1. Average values of water mass proxy data from Site 1267 (U1313) during
1300 target glacials (G) and associated interglacials (IG), as well as average IG-G change.

<table><thead><tr><th rowspan="2">MIS</th><th colspan="3">Marine Isotope Stage average*</th><th colspan="3">IG-G difference</th></tr><tr><th>εNd</th><th>δ<sup>13</sup>C</th><th>[CO<sub>3</sub><sup>2</sup>]</th><th></th><th>ΔεNd</th><th>Δδ<sup>13</sup>C</th><th>Δ[CO<sub>3</sub><sup>2</sup>]</th></tr></thead><tbody><tr><td>MG1</td><td>-11.03 (-11.92)</td><td>+0.85 (0.80)</td><td>92.7</td><td>MG1-M2</td><td>+1.09 (+0.70)</td><td>-0.46 (-0.20)</td><td>-0.5</td></tr><tr><td>M2</td><td>-9.95 (-11.22)</td><td>+0.39 (0.59)</td><td>92.1</td><td></td><td></td><td></td><td></td></tr><tr><td>M1</td><td>-11.04 (-11.60)</td><td>+0.79 (0.83)</td><td>103.1</td><td>M1-M2</td><td>-1.10 (-0.37)</td><td>+0.40 (+0.23)</td><td>11</td></tr><tr><td>101</td><td>-11.18 (-12.55)</td><td>+0.55 (0.88)</td><td>108.1</td><td>101-100</td><td>-1.87 (+1.84)</td><td>-0.41 (-0.76)</td><td>22.2</td></tr><tr><td>100</td><td>-9.31 (-10.71)</td><td>+0.14 (0.12)</td><td>85.8</td><td></td><td></td><td></td><td></td></tr><tr><td>99</td><td>-10.04 (-11.88)</td><td>+0.30 (0.89)</td><td>102.1</td><td>99-100</td><td>-0.73 (-1.17)</td><td>+0.16 (+0.76)</td><td>16.2</td></tr><tr><td>MIS 2</td><td>-9.57 (-11.48)</td><td>+0.05 (0.19)</td><td>82.3</td><td></td><td></td><td></td><td></td></tr><tr><td>MIS 1</td><td>-10.74 (-13.98)</td><td>+0.57 (0.90)</td><td>92.9</td><td>1-2</td><td>-1.17 (-1.84)</td><td>+0.52 (+0.55)</td><td>10.6</td></tr></tbody></table>

1301 *determined using LR04 definitions of MIS boundaries (Lisiecki and Raymo, 2005) apart from for

1302 MIS M1 (where we use 3238-3285 ka) and MIS M2 (where we use 3285-3312 ka).
<figure><img src="image_2.png" /><figcaption>Figure 1</figcaption></figure>
<figure><img src="image_3.png" /><figcaption>Figure 2</figcaption></figure>
<figure><img src="image_4.png" /><figcaption>Figure 3</figcaption></figure>
<figure><img src="image_5.png" /><figcaption>Figure 4</figcaption></figure>
<figure><img src="image_6.png" /><figcaption>Figure 5</figcaption></figure>

<figure><img src="image_7.png" /></figure>
<figure><img src="image_8.png" /><figcaption>Figure 6</figcaption></figure>
<figure><img src="image_9.png" /><figcaption>Figure 7</figcaption></figure>
<figure><img src="image_10.png" /><figcaption>Figure 8</figcaption></figure>

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

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

<figure><img src="image_13.png" /></figure>
Kirby et al.: On climate and abyssal circulation in the Atlantic Ocean during late Pliocene marine isotope stage M2, ~3.3 million years ago

**Conflict of interest statement:** The authors declare that they have no competing financial interests or personal relationships that influenced the work reported in this paper
**Author Contributions:** N.K generated the Site 1267 data and contributed to manuscript writing; I.B. designed the study, supervised N.K. in data generation and wrote the manuscript; A.B generated the Site U1313 stable isotope data for MIS 100 and contributed to manuscript writing; T.B.C. supervised R.L.P in generating the Site 1267 benthic B/Ca data and contributed to the ideas presented and to manuscript writing; R.L.P generated a subset of the Site 1267 data; A.J.C. supervised N.K. and R.L.P in data generation, contributed to the ideas presented and to manuscript writing; V.E.T. supervised N.K. and R.L.P. in data generation, generated the volcanic glass count data and contributed to manuscript writing; J.A.M helped to generate the Nd isotope data; G.L.F. contributed to the ideas presented and manuscript writing; M.E.R. contributed to the ideas presented and manuscript writing; D.K. contributed samples and to the ideas presented and manuscript writing; D.B.B contributed samples and to manuscript writing; P.A.W. designed the study, supervised N.K. in data generation, co-wrote the manuscript with I.B.