# Geochemical response of the mid-depth Northeast Atlantic Ocean to freshwater input during Heinrich events 1 to 4

Anya Crocker, Thomas Chalk, Ian Bailey, Megan Spencer, Marcus Gutjahr, Gavin Foster, Paul Wilson

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To cite this version:

Anya Crocker, Thomas Chalk, Ian Bailey, Megan Spencer, Marcus Gutjahr, et al.. Geochemical response of the mid-depth Northeast Atlantic Ocean to freshwater input during Heinrich events 1 to 4. Quaternary Science Reviews, 2016, 151, pp.236-254. 10.1016/j.quascirev.2016.08.035. hal-04707794

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# Geochemical Response of the Mid-Depth Northeast Atlantic Ocean to Freshwater Input During Heinrich Events 1 to 4

Anya J. Crocker¹,²*, Thomas B. Chalk¹,³, Ian Bailey¹,⁴, Megan R. Spencer¹, Marcus Gutjahr¹,⁵, Gavin L. Foster¹, Paul A. Wilson¹

¹National Oceanography Centre Southampton, University of Southampton, Waterfront Campus, Southampton SO14 3ZH, UK

²Department of Animal and Plant Science, University of Sheffield, Western Bank, Sheffield, S10 2TN, UK

³Department of Physical Oceanography, Woods Hole Oceanographic Institution, Woods Hole, Massachusetts, USA

⁴Now at: Camborne School of Mines & Environment and Sustainability Institute, College of Engineering, Mathematics & Physical Sciences, University of Exeter, Penryn Campus, Treliever Road, Penryn Cornwall, TR10 9FE, UK

⁵Now at: GEOMAR Helmholtz Centre for Ocean Research Kiel, Wischofstraße 1 – 3, 24148 Kiel, Germany

*Corresponding author

Correspondence to: Anya J. Crocker, anya.crocker@noc.soton.ac.uk

1. Heinrich events are intervals of rapid iceberg-sourced freshwater release to the high

2. latitude North Atlantic Ocean that punctuate late Pleistocene glacials. Delivery of

3. fresh water to the main North Atlantic sites of deep water formation during

4. Heinrich events may result in major disruption to the Atlantic Meridional
Overturning Circulation (AMOC), however, the simple concept of an AMOC shutdown in response to each freshwater input has recently been shown to be overly simplistic. Here we present a new multi-proxy dataset spanning the last 40,000 years that resolves four Heinrich events at a classic mid-depth North Atlantic drill site, employing four independent geochemical tracers of water mass properties: boron/calcium, carbon and oxygen isotopes in foraminiferal calcite and neodymium isotopes in multiple substrates. We also report rare earth element distributions to investigate the fidelity by which neodymium isotopes record changes in water mass distribution in the northeast Atlantic. Our data reveal distinct geochemical signatures for each Heinrich event, suggesting that the sites of fresh water delivery and/or rates of input played at least as important a role as the stage of the glacial cycle in which the fresh water was released. At no time during the last 40 kyr was the mid-depth northeast North Atlantic dominantly ventilated by southern-sourced water. Instead, we document persistent ventilation by Glacial North Atlantic Intermediate Water (GNAIW), albeit with variable properties signifying changes in supply from multiple contributing northern sources.

**Keywords:** Heinrich events; North Atlantic palaeoceanography; Last glacial period; Neodymium isotopes; B/Ca; Carbon and oxygen isotopes; Ice-rafted debris; ODP Site 980.

# 1. Introduction
The climate of the last glacial period was punctuated by a number of pronounced events with near global impacts, known as Heinrich (or H-) events (e.g. Heinrich, 1988; Hemming, 2004). During these events, transient catastrophic collapses of the North American Laurentide Ice Sheet (LIS), centered over the Canadian Hudson Bay, produced armadas of icebergs that travelled through the Hudson Straight into the North Atlantic Ocean, eventually depositing large volumes of ice-rafted debris (IRD), including distinctive detrital limestone clasts, in a belt across the North Atlantic (Ruddiman, 1977). LIS surging, together with iceberg calving from other circum-North Atlantic and Arctic ice sheets during H-events added large volumes of fresh (low density) water to the (sub)polar oceans with important implications for northern-sourced deep water formation and the Atlantic Meridional Overturning Circulation (AMOC). The AMOC exerts a strong control on regional and global climate on orbital to suborbital timescales, influencing latitudinal heat budgets (e.g. Boyle and Keigwin, 1987; Broecker and Denton, 1989) and partitioning of carbon between the atmosphere and deep ocean (e.g. Adkins, 2013; Sigman and Boyle, 2000; Sigman et al., 2010). Understanding the response of AMOC to fresh water addition is therefore crucial, particularly in light of recently accelerating mass loss from the Greenland Ice Sheet and other circum-Atlantic ice masses (Gierz et al., 2015; Vaughan et al., 2013).

Previously, it was suggested that fresh water inputs at high latitudes during H-events caused a cessation of deepwater formation in the North Atlantic, resulting in a shoaling of the northern component water overturning cell by ~1 km (e.g. Alley et al., 1999; Sarnthein et al., 1994; Swingedouw et al., 2009). Support for this concept of a near complete shutdown of AMOC during H-events came from two main lines of evidence: (i)
a dramatic drop in circulation vigour (e.g. McCave et al., 1995a; McManus et al., 2004) and (ii) water mass provenance reconstructions (including those based on δ¹³C, Cd/Ca and ¹⁴C) indicating an increased presence of nutrient-rich southern-source waters (SSW) below ~2 – 2.5 km depth in the North Atlantic Ocean (e.g. Keigwin et al., 1991; Robinson et al., 2005; Stern and Lisiecki, 2013; Vidal et al., 1997). More recently, however, both of these arguments have been questioned. ²³¹Pa/²³⁰Thₓs records from sites across a wide range of water depths suggests that overturning persisted at shallower depths during H1 (Bradtmiller et al., 2014; Gherardi et al., 2009; Lippold et al., 2016), with values compatible with a near complete shutdown only identified in the SSW cell and during H-events close to glacial maxima (Böhm et al., 2015; Lippold et al., 2009; McManus et al., 2004). A further recent development is the documentation of a poorly ventilated water mass in the Nordic Seas during the last glacial period that overflowed the Greenland-Scotland Ridge (GSR) into the Atlantic basin during the deglaciation (Thornalley et al., 2015). This discovery means that the presence of nutrient-rich waters in the North Atlantic basin may no longer be simply attributed to the incursion of waters from the south. Indeed, excursions to low oxygen and sometimes also low carbon isotopic signatures of benthic foraminifera in the North Atlantic during H-events have been interpreted to suggest that there may also have been overflow of waters from the Nordic Seas at these times (e.g. Meland et al., 2008; Thornalley et al., 2010; Vidal et al., 1998). In addition, bulk sediment leachate εₙd values from the northeast Atlantic (Crockett et al., 2011) are argued to support a persistent presence of overflow waters from the Nordic Seas in the North Atlantic throughout most, if not all, of the last glacial cycle, providing a northern source of nutrient-rich waters to the northeast Atlantic Ocean.
Detailed palaeoceanographic reconstructions show that the classic concept of a simple repeated response of AMOC to freshwater addition during each H-event may be overly simplistic. Not all H-events show a clear perturbation in every oceanographic reconstruction, and there is no consensus on which H-events involve the largest disruption of the ocean-atmosphere system. There is also little agreement about the factors that control the amplitude of AMOC response to freshwater input and considerable debate over the importance of H-event timing relative to the last glacial cycle (Böhm et al., 2015; Lynch-Stieglitz et al., 2014). New high-resolution data sets spanning multiple H-events therefore are needed to help us to gain a better understanding of the range of associated circulation changes in the Atlantic. Records from sites proximal to the GSR are particularly valuable because this region is especially sensitive to changes in northern deep water formation and therefore can help to better constrain variations in the influence of Nordic Seas overflow waters (NSOW).

Bulk sediment leachate Nd isotope data from northeast Atlantic Ocean Drilling Program (ODP) Site 980 (55°29.1′N, 14°42.1′W; 2170 m water depth; location shown in Fig. 1) have been interpreted to suggest that overflow waters crossing the Wyville-Thomson Ridge (WTR) from the Nordic Seas were supplied to the Feni Basin at intermediate depths throughout the past 40 kyr (Crockett et al., 2011). This interpretation invokes either concentrated overflow waters without substantial entrainment of North Atlantic waters or SSW at this site during H-events 1 – 3. Yet, the potential for overprinting of water mass Nd isotope signatures in this region (Lacan and Jeandel, 2004a; Roberts and Piotrowski, 2015), concerns about the fidelity of bulk sediment leachate Nd isotope records (Elmore et al., 2011; Wilson et al., 2013) and the range of
water masses influencing the intermediate-depth northeast Atlantic (Lacan and Jeandel, 2005) are all sources of uncertainty demanding further careful assessment of the problem.

Additional independent proxies of water mass provenance are therefore required to help reconstruct past vertical water mass structure from the Feni Drift.

To tackle these gaps in our knowledge, we report the results of a new multi-proxy reconstruction of bottom water chemistry and inferred water mass distribution from ODP Site 980. First, we present neodymium isotope reconstructions from three phases (fish debris, planktonic foraminifera and bulk sediment leachates), together with rare earth element distributions to better understand the controls on neodymium association with foraminifera at Site 980. Then we combine our Nd isotope reconstructions with new, high resolution records of three further proxies for water mass chemistry (B/Ca, δ¹³C and δ¹⁸O in benthic foraminifera) at our study site to shed new light on the response of the AMOC in the northeast Atlantic to freshwater addition during four H-events of the last glacial period.

111
<figure><img src="image_3.png" /><figcaption>Fig. 1: (a) The North Atlantic region showing the location of the main study site (ODP Site 980, in black). Dotted lines mark the approximate maximum spatial extent of continental ice sheets during the last glacial maximum (Clark et al., 2012; Dyke et al., 2002; Funder and Hansen, 1996; Sejrup et al., 2005). Arrows represent modern major intermediate and deep current pathways with Nordic Sea overflows (NSOW, orange, also includes Wyville-Thomson Ridge overflow waters or WTROW), Labrador Sea Water (LSW, purple), Northeast Atlantic Deep Water (NEADW, light blue) and Antarctic Bottom Water (AABW, red), based upon McCartney (1992), Hansen and Østerhus (2000), New and Smythe-Wright (2001) and Lacan and Jeandel (2005). (b) Schematic cross-section of the major water masses in the modern Rockall Trough, based upon McGrath et</figcaption></figure>
al. (2012) and Johnson et al. (2010). Colours as panel (a), with the addition of Subpolar Mode Water (SPMW, dark blue). Numbers give neodymium isotope signature of water masses today (Crocker, 2014; Crocket et al., 2011; Lacan and Jeandel, 2004c; Lacan and Jeandel, 2005; Lacan et al., 2012; Stichel et al., 2012). Note that the ε<sub>Nd</sub> value of AABW shown the AABW endmember recorded in the Atlantic section of the Southern Ocean; this value is likely modified by mixing and boundary exchange as the water mass moves northwards (e.g. Lambelet et al., 2016). (c) North Atlantic region showing the location of the main study site (ODP Site 980, in red) and BOFS sites referred to in the text (black). Grey stripes indicate the main belt of ice-rafted debris deposition in the glacial north Atlantic during the Last Glacial Maximum (Ruddiman, 1977). Dark blue numbers: 1 – Rockall Plateau, 2 – Rockall Trough, 3 – Feni Drift. Last Glacial Maximum limit of perennial sea ice indicated by solid light blue line with dashed light blue line marking extreme limit of winter sea ice (de Vernal et al., 2005; Hillaire-Marcel and de Vernal, 2008).

135

## 2. Background

137

### 2.1 Site Location and Oceanography

ODP Site 980 is situated on the Feni sediment drift, and features high sedimentation rates across the studied interval (mean 20 cm kyr⁻¹, see details of age model in supplementary information). Its position on the northern fringe of the main Atlantic belt of North American-sourced detrital carbonate deposition (Hemming, 2004) results in lithologically distinct and hence unambiguously identifiable IRD layers. Site 980 is known for its benchmark archives of millennial-scale climate variability during the Quaternary (e.g. McManus et al., 1999; Oppo et al., 2003; 2006, Fig. 2). Existing records from this site cover an unusually long interval for suborbitally resolved records, extending back
approximately 500,000 years. Published data from Site 980 spanning the last glacial period are, however, of insufficient resolution to resolve changes in ocean chemistry across H-events clearly.

Today, Site 980 is bathed by a mixture of North East Atlantic Deep Water (NEADW) and Labrador Sea Water (LSW), with only a minimal influence of Wyville Thomson Ridge overflow waters (WTROW) from the Nordic Seas (Ellett and Martin, 1973; McGrath et al., 2012). During the last glacial period, Site 980 lay close to the interpreted depth of the boundary between northern- and southern-sourced waters (e.g. Boyle and Keigwin, 1987; Curry and Oppo, 2005; Oppo and Lehman, 1993), but this depth is thought to have been bathed by Glacial North Atlantic Intermediate Water (GNAIW) during background glacial conditions (Yu et al., 2008). If AMOC shutdown occurred (as is classically suggested for last glacial H-events), model simulations suggest that Site 980 would have been bathed by Glacial Antarctic Bottom Water (GAABW) during these times (Flückiger et al., 2008; Singarayer and Valdes, 2010; Swingedouw et al., 2009).

162
<figure><img src="image_4.png" /><figcaption>Fig. 2: Suborbital variability at ODP Site 980. (a) Data set of McManus et al. (1999), benthic foraminiferal oxygen isotopes in **brown** and IRD concentrations in **pale blue**. (b) **New oxygen isotope record** of benthic foraminiferal calcite (C. wuellerstorfi) in **red** (this study) with data of McManus et al. (1999) in **brown** (McManus data adjusted by -0.3 ‰ due to an inter-lab offset). (c) IRD concentrations, new data (this study) in **dark blue** (150 – 500 µm), data from McManus et al. (1999) in **pale blue** (grain size >150 µm). (d) Percentage of polar species *N. pachyderma* (s.)</figcaption></figure>
as a proportion of the total number of planktonic foraminifera, with high values indicating cold sea surface temperatures (this study). (e) Oxygen isotope values of ice from NGRIP ice core, with less negative values indicating warmer temperatures (North Greenland Ice Core Project Members et al., 2004). NGRIP data plotted on the GICC05 age model (Andersen et al., 2006; Rasmussen et al., 2006; Svensson et al., 2008; Vinther et al., 2006). Grey arrows mark Site 980 recalibrated radiocarbon ages from Oppo et al. (2003) and Benway et al. (2010). Data in panels (b) – (d) are plotted on our new age model (see methods), with intervals of high ice rafting shaded in yellow. IRD peaks associated with H1-4 are labelled.

## 2.2 Neodymium isotopes

Over the last few decades, neodymium isotopes have become an important tool for reconstructing the vertical structure of the water column and the provenance and circulation pathways of water masses. A major strength of the Nd isotope technique is that, unlike many commonly used proxies for water mass chemistry (e.g. $\delta^{13}$C), neodymium isotopes are not influenced by biological processes. The estimated residence time of Nd in the ocean (200 – 1000 years (Arsouze et al., 2009; Tachikawa et al., 2003; Tachikawa et al., 1999)) is shorter than the modern oceanic mixing time of 1000 – 1600 years (e.g. Broecker and Peng, 1982; Garrison, 2011; Sarmiento and Gruber, 2004), giving rise to spatial variation in the isotopic signature of different water masses. Rocks exhibit a wide range of neodymium isotopic compositions, depending upon both their age and initial Sm/Nd ratios. These Nd isotope signatures are typically expressed in epsilon notation or $\epsilon_{Nd}$ (representing the $^{143}$Nd/$^{144}$Nd deviation of a sample relative to a chondrite uniform reservoir in parts per 10,000) with values ranging from $\epsilon_{Nd}$ = -56 for old granitic cratons to +12 for young mid-ocean ridge basalts (Jeandel et al., 2007; Lacan et al., 2012;
Sarbas and Nohl, 2008, and references therein). Neodymium from this range of sources is transferred from the continents to the ocean via both riverine and aeolian inputs, resulting in different water masses acquiring distinct neodymium isotopic signatures (Goldstein et al., 1984; Grousset et al., 1992; Mearns, 1988; Tachikawa et al., 1999). Addition and exchange of neodymium at the sediment-bottom water interface along continental margins also modifies the neodymium isotopic signature of bottom waters through the process of boundary exchange (Lacan and Jeandel, 2004a; Rickli et al., 2014; Wilson et al., 2012).

Analysis of Fe-Mn oxyhydroxides within marine sediments extracted by a leaching procedure (e.g. Gutjahr et al., 2007; Rutberg et al., 2000) allows the production of relatively rapid and high resolution records of bottom water ε<sub>Nd</sub>. This technique has now become widely used (e.g. Böhm et al., 2015; Jonkers et al., 2015; Wei et al., 2015; Wilson et al., 2015). A wide range of other phases have also been proposed to record and preserve bottom water ε<sub>Nd</sub>, including corals (e.g. Copard et al., 2010; van de Flierdt et al., 2006), ferromanganese nodules and crusts (e.g. O'Nions et al., 1978; Piepgras et al., 1979), fish debris (e.g. Lang et al., 2016; Martin and Haley, 2000; Staudigel et al., 1985) and foraminifera both with and without authigenic Fe-Mn oxide coatings (e.g. Elmore et al., 2011; Palmer and Elderfield, 1985).

The ε<sub>Nd</sub> signatures recorded by different substrates do not always agree with one another (e.g. Elmore et al., 2011). For example, a pronounced difference between the glacial and Holocene ε<sub>Nd</sub> values of uncleaned foraminifera is seen along a depth transect of Biogeochemical Ocean Flux Study (BOFS) cores on and to the south of the Rockall Plateau (Fig. 1), but this difference is muted or absent in co-registered bulk sediment
217 leachate ε<sub>Nd</sub> (Piotrowski et al., 2012). The leaching of bulk sediment can give a
218 neodymium isotope signature influenced by reactive fine grained sediment components
219 such as volcanic ash (which has a radiogenic signature) and detrital carbonate (which is
220 unradiogenic) (Elmore et al., 2011; Roberts et al., 2010), and thus bulk sediment
221 leachates may not accurately record bottom water chemistry at certain locations. The
222 neodymium isotope signature extracted by bulk sediment leaching is also highly sensitive
223 to the methodology used to separate the authigenic coating signal from the primary
224 sediment (Wilson et al., 2013).

225 The mostly constant value of ε<sub>Nd</sub> = -10 recorded by bulk sediment leachates from
226 Site 980 has been interpreted to demonstrate a continuous presence of WTROW in the
227 Rockall Trough for most, if not all, of the last 40 kyrs due to the similarity of this value to
228 the estimated signature of overflow waters (Crockett et al., 2011). In contrast, modern
229 oceanographic observations indicate only a minimal influence of WTROW at Site 980
today (Johnson et al., 2010; McGrath et al., 2012). Short-lived radiogenic Nd isotopic
230 excursions identified in the Site 980 record with ages estimated as approximately
231 concurrent to H-events 1 – 4 are inferred to record times of either increased influence of
232 SSW during AMOC shutdowns or a more concentrated presence of overflow waters
233 (Crockett et al., 2011). However, the limited resolution of the Site 980 bulk sediment
234 leachate ε<sub>Nd</sub> dataset and the absence of direct identification of H-layers in the site
235 stratigraphy make these conclusions tentative (see also supplementary information). In
236 addition, there are a number of other processes with the potential to alter bottom water
237 ε<sub>Nd</sub> (e.g. variable composition of NEADW or influence of LSW in the Rockall Trough,
238 modification of bottom water isotopic signature by interaction with lithogenic material).
239
The circulation history of the northeast Atlantic during the last glacial period therefore requires further investigation.

A decoupling between $\epsilon_{Nd}$ values and other proxies used as water mass tracers (δ$^{13}$C, B/Ca (Yu et al., 2008)) suggests that the glacial-interglacial variability in foraminiferal $\epsilon_{Nd}$ on and to the south of the Rockall Plateau may not be a result of changing bottom water provenance (Roberts and Piotrowski, 2015). Instead, reactive IRD grains may modify water mass $\epsilon_{Nd}$ in this region as they sink through the water column, with large inputs of volcanic glass ($\epsilon_{Nd}$ = +4.3 ± 0.1) to the Rockall Plateau and its southern flank appearing to modify bottom water towards more radiogenic values during the Last Glacial Maximum (Roberts and Piotrowski, 2015). Unradiogenic labelling of the water column by Hudson Bay-derived detrital limestone grains ($\epsilon_{Nd}$ = -18.6 ± 0.2) may also have occurred during Heinrich events (Roberts and Piotrowski, 2015). These observations raise questions over the temporal and spatial extent of water mass $\epsilon_{Nd}$ relabelling.

To better understand the meaning of bulk sediment $\epsilon_{Nd}$ data from Site 980 and to investigate whether issues of water column relabelling by IRD identified on the Rockall Plateau and its southern margin are applicable to the entire northeast Atlantic, we therefore need higher resolution $\epsilon_{Nd}$ records from multiple phases at Site 980 supplemented by a record of concentrations and lithologies of IRD. Furthermore, independent proxies of bottom water chemistry are also required to improve our understanding of water mass structure of the northeast Atlantic during the last 40,000 years at Site 980, therefore we also present co-registered benthic δ$^{13}$C, δ$^{18}$O, B/Ca records.
263

264 **3. Materials and Methods**

265

266 **3.1 Stratigraphy**

267 We sampled Holes 980B (0.05 – 4.12 meters below seafloor (mbsf) at 8 cm resolution)

268 and 980A (0.32 – 4.08 mbsf at 2 cm resolution). We sampled the same core sections as

269 Crocket et al. (2011) to allow direct comparison with their data. Our Hole A samples are

270 not part of the original primary shipboard splice (Shipboard Scientific Party, 1996),

271 therefore, to assign composite depths (metres composite depth, mcd) to these samples, we

272 correlated Hole A mbsf to the primary splice using shipboard-derived whole core

273 volume-specific magnetic susceptibility data measured shipboard (see supplementary

274 information) (Shipboard Scientific Party, 1996). Our study interval corresponds to 0.06 –

275 8.12 mcd.

276

277 **3.2 IRD and % N. pachyderma (s.)**

278 Sediment samples were sieved at 150–500 µm, and then split until approximately 300

279 lithic grains remained, which were identified and counted using a binocular microscope,

280 applying the categorizations of Hall et al. (2011). Pumice grains are assumed to have a

281 volcanic ash origin, and so are not included in IRD counts. Reproducibility of IRD counts

282 determined by repeat counts is ± 1%. IRD fluxes were estimated following Peck et al.

283 (2007), with bulk density values corrected for a systematic error in the ODP Leg 162

284 shipboard gamma-ray attenuation porosity evaluator (GRAPE) measurements (Jansen et

285 al., 2000). Samples prepared for IRD counts were then split again until approximately
286 300 foraminifera remained. Counts were performed of the number of specimens of
287 *Neogloboquadrina pachyderma* (sinistral) in these splits and the value expressed as a
288 percentage of the total number of planktonic foraminifera. Reproducibility of % N.
289 *pachyderma* (s.) values determined by repeat counts is ± 3%.

290

291 3.3 Stable isotopes

292 To generate benthic stable oxygen and carbon isotope reconstructions for our target
293 interval, specimens of *Cibicides wuellerstorfi* (typically 2 – 4) were picked from the >212
294 µm size fraction (total carbonate mass 60 – 100 µg). Oxygen and carbon isotopic
295 compositions of this carbonate were analysed using a Europa Geo 20-20 stable isotope
296 ratio mass spectrometer at Southampton. All sample values are expressed in delta
297 notation, relative to the Vienna Peedee Belemnite standard (VPBD). External
298 reproducibility is better than 0.053 ‰ for δ¹⁸O and 0.027 ‰ for δ¹³C. Species-specific
299 disequilibrium from seawater δ¹⁸O values was removed by applying an adjustment of
300 +0.64 ‰ (Shackleton and Opdyke, 1973). *C. wuellerstorfi* δ¹⁸O values were adjusted for
301 global sea level change (δ¹⁸Oₗc) to allow us to understand temporal changes in regional
302 water mass chemistry, following the method of Meland et al. (2008) and using the sea
303 level record shown in Grant et al. (2012) with a maximum glacial-interglacial sea level
304 change of 110 m.

305

306 3.4 Neodymium isotope ratios, rare earth elements and trace elements

307 3.4.1 Sample preparation

308 To generate neodymium isotope and rare earth element profiles from foraminifera, a
large number (600 – 1600) of planktic foraminifera of mixed species were picked in the >212 µm size fraction from each sample. These were broken open then sonicated in ELGA water and methanol to remove clays, before being dissolved in 1.75 M HCl. Ferromanganese coatings were not removed from the foraminifera, as these are rich in pore water-derived Nd and have been shown to record a bottom water signature (Elmore et al., 2011; Roberts et al., 2012).

Fish debris $\varepsilon_{Nd}$ is considered to be a faithful recorder of the neodymium isotopic composition of bottom waters (e.g. Grandjean et al., 1987; Martin and Haley, 2000; Staudigel et al., 1985; Stille, 1992; Stille and Fischer, 1990). To provide an additional check on the fidelity of our new foraminiferal $\varepsilon_{Nd}$ record, we therefore picked 1 – 15 pieces of fish debris from the >125 µm size fraction of a subset of samples (n = 8). Adhering clays were removed by sonication in methanol and ELGA water. All samples were oxidatively cleaned based upon the methods of Boyle and Keigwin (1985), although a reductive cleaning step was not necessary (Martin and Haley, 2000). An aliquot of 10 µl was extracted from each of the dissolved foraminifera and fish debris samples for rare earth element (REE) and other trace element analysis, with the remainder of the sample used for neodymium isotope analysis.

### 3.4.2 Rare and trace elements

The mechanism by which Nd becomes associated with foraminifera strongly influences whether the isotopic signature preserved is representative of the bottom water signature (Roberts et al., 2012; Tachikawa et al., 2014). The extent of remobilization of Nd in sediment pore waters is of particular interest, and can be investigated using REE
concentrations and distributions (Roberts et al., 2012). Subtle differences in the behaviour of the various REEs result from a decrease in atomic radius as mass increases, leading to differences in their speciation in seawater (Elderfield and Greaves, 1982; Goldberg et al., 1963). The behaviour of cerium is also distinct from the other elements because it can be oxidized from a soluble (3+) to an insoluble (4+) state (de Baar et al., 1985; Elderfield et al., 1988). This distinction can be expressed by the cerium anomaly (Ce/Ce*), which is the difference in the shale-normalised abundance of cerium compared to the expected value calculated from the nearby light rare earth (LREE) elements (Ce/Ce*=3Ce<sub>n</sub>/(2La<sub>n</sub>+Nd<sub>n</sub>), where n indicates concentrations relative to the Post-Archean Australian Shale, PAAS (McLennan, 1989; Taylor and McLennan, 1985)). Negative cerium anomalies (values <1) indicate a depletion of cerium relative to the other REEs (de Baar et al., 1988).

Rare earth element samples were diluted with 3% HNO<sub>3</sub> spiked with In, Re and Be.
Calcium and strontium concentrations were measured with a Quadrupole ICP-MS:
Thermo X-Series 2 at Southampton. The remaining sample solutions were analysed for
REE and trace element concentrations by High Resolution ICP-MS: Thermo ELEMENT
2XR at Southampton. All samples were corrected for matrix effects and instrument drift
using the In, Re and Be standard incorporated into each sample. A blank correction was
then applied and rare earth element standards used to correct samples for oxide
formation. External reproducibility is estimated as 4–5 %, with internal reproducibility
much less than this for the majority of samples. Rare earth element concentrations of the
samples (REE<sub>sample</sub>) were expressed relative to the REE signature of the PAAS (REE<sub>PAAS</sub>)
(McLennan, 1989; Taylor and McLennan, 1985).
355

356 3.4.3 Neodymium isotope ratios

357 Neodymium was purified from the dissolved samples for isotope analysis using standard
358 procedure column chemistry, based upon the methods of Cohen et al. (1988). Cation
359 columns were used to strip iron and titanium from the samples. The remaining material
360 was then run through LN SpecTM columns to purify and concentrate neodymium (Pin and
361 Zalduegui, 1997). Samples were analysed by Multi-collector ICP-MS: Thermo
362 NEPTUNE at Southampton. Instrumental mass bias ratios were corrected using the
363 procedure of Vance and Thirlwall (2002), adjusting to a ¹⁴⁶Nd/¹⁴⁴Nd of 0.7219 and using
364 cerium-doped standards to correct for interference of ¹⁴²Ce on ¹⁴²Nd. All values were
365 normalised to the JNdi-1 Standard (¹⁴³Nd/¹⁴⁴Nd = 0.512115±7) (Tanaka et al., 2000), and
366 isotopic signatures are expressed in epsilon notation, using a chondritic uniform reservoir
367 value of 0.512638 (Jacobsen and Wasserburg, 1980). Replicate measurements of the
368 JNdi-1 standard across the three analytical sessions gave an external reproducibility 2
369 s.d. = 0.2 εNd units for replicate analyses of Nd standard solutions (n = 25). Error bars
370 plotted show either this external reproducibility or the sample internal reproducibility,
371 whichever is larger. We compared our Nd data to co-measured concentrations of
372 lithophilic elements (e.g. Al, Ti, Zr, Pb) and volcanic grains, and found no evidence to
373 suggest a strong influence of volcanic ash, clay contamination or other terrestrial material
374 on foraminiferal εNd values at Site 980 (see supplementary information).

375

376 3.5 Boron/calcium ratios
Between 8 and 12 *C. wuellerstorfi* tests (212 – 500 µm) were cleaned for trace metal analysis following Barker et al. (2003). Foraminiferal tests were cracked and ultrasonicated in Milli-Q 18.2 MΩcm water to remove clay materials. Samples were then oxidatively cleaned in a 1% hydrogen peroxide solution before a weak acid leach is applied to remove adsorbed cations. Finally, samples were dissolved in ~ 0.075M nitric acid. The above steps were all undertaken in a specialised boron free clean laboratory at the University of Southampton. Element ratios were measured using a ThermoFisher Scientific Element 2XR-ICPMS at Southampton, using the protocol described by Rae et al. (2011) employing matrix-matched in house standards and a variety of consistency standards to ensure reproducibility. Samples were screened for clay contamination using Al/Ca and other contaminant ratios (e.g. Ba/Ca, Fe/Ca), with samples Al/Ca > 100 µmol/mol excluded from this study (n = 2). No anomalous values of other elements were identified. Near complete clay removal is assumed for the remaining samples. Long-term reproducibility for B/Ca during this study was within 4 % at 2 s.d. (~ 8 µmol/mol). Carbonate ion concentrations were calculated using a sensitivity of Δ[CO₃²⁻] on B/Ca of 1.14 ± 0.048 for *C. wuellerstorfi* (Yu and Elderfield, 2007), and a [CO₃²⁻]sat calculated from modern local pressure, temperature and salinity.

## 3.6 Age model generation

A new age model was developed for Site 980 by combining previously published radiocarbon ages (Benway et al., 2010; Oppo et al., 2003) with a new record of the abundance of planktonic foraminifera *N. pachyderma* (s.) correlated to Greenland proxy records (Fig. 2 and supplementary information). An average sedimentation rate of ~ 0.2
400 mm/yr means that samples with 2 cm spacing give an age resolution of approximately
401 100 years.

402 Radiocarbon ages were updated using the Marine13 calibration curve (Reimer et
403 al., 2013) in conjunction with Calib 7.1 software (Stuiver and Reimer, 1993; Stuiver et
404 al., 2005). A constant reservoir age of 400 ± 100 years was assumed, with the exception
405 of the Younger Dryas (800 ± 300 years) and H1 (1600 ± 1000) (e.g. Bard, 1988; Bard et
406 al., 1994; Bondevik et al., 2006; Stanford et al., 2011; Stern and Lisiecki, 2013;
407 Waelbroeck et al., 2001). Additional age constraints were provided by correlating the Site
408 980 % N. pachyderma (s.) record (a proxy for upper ocean temperature when sea surface
409 temperatures are between 4°C and 10°C (Darling et al., 2006)) to North Greenland Ice
410 Core Project (NGRIP) δ¹⁸Oice (a proxy for atmospheric temperatures) (North Greenland
411 Ice Core Project Members et al., 2004) on the GICC05 age model (Andersen et al., 2006;
412 Rasmussen et al., 2006; Svensson et al., 2008; Vinther et al., 2006). Visual correlation
413 was carried out using the Analyseries software (Paillard et al., 1996), and is based on the
414 assumption that temperature changes in the upper North Atlantic Ocean are synchronous
415 with Greenland air temperatures. Distinctive peaks in detrital carbonate fluxes were used
416 to confirm the position of H4, H2 and H1.

417

4. **Results and Discussion**

419

420 4.1 Assessing neodymium isotopes as a water mass tracer at ODP Site 980

421

422 4.1.1 Exploring inter-substrate differences in neodymium isotope signatures
New records of REE distributions and Nd isotope signatures of mixed planktonic foraminifera and fish debris from Site 980 are shown together with IRD concentrations in Fig. 3. Throughout our study interval, there is close agreement between new data from fish debris and planktonic foraminifera εNd data. Glacial εNd values recorded by fish and foraminifera (–9 to –10) are more radiogenic than Holocene values (~ –11.5), with suborbital unradiogenic excursions from baseline values during H4 (~ 38 ka) and the early Holocene (~ 5 – 8 ka), and a strong radiogenic excursion across H2 (~ 23 – 26 ka) (Fig. 3a). Previously published sediment leachate εNd data from Site 980 (Crockett et al., 2011) generally show a systematic offset towards more radiogenic values than our new foraminifera and fish debris values by 1 – 2 εNd units over the last ~11.5 ka (Fig. 3a). This systematic offset is not present in the samples of peak glacial age, where agreement between the two substrates is often much closer, although there are some sporadic disagreements between the two substrates.

Modern oceanographic observations show that the mid-depth western Rockall Trough is predominantly bathed by NEADW (εNd = –12.8 ± 0.2) and LSW (–13.9 ± 0.4), with a small influence of overflow waters from the Nordic Seas (–8.2 ± 0.6) (Ellett and Martin, 1973; Johnson et al., 2010; Lacan and Jeandel, 2005; McGrath et al., 2012; Olsen and Ninnemann, 2010). This suggests that the youngest foraminiferal εNd values at Site 980 (–11.7 ± 0.3) capture bottom water chemistry more accurately than bulk sediment leachate εNd (–10.2 ± 0.3). A stronger overflow signature than indicated by modern observations (Johnson et al., 2010) or local modification of bottom waters towards more radiogenic values is therefore required if the youngest leachate εNd signatures at our study site accurately represent recent bottom water chemistry. Alternative explanations include
446 analytical issues offsetting the bulk sediment leachates to more radiogenic values than the
447 fish debris and foraminifera, and the influence of an additional radiogenic phase is
448 shifting bulk sediment leachates towards more radiogenic values in the Holocene.

449 Bulk sediment leachates can provide a relatively rapid way to reconstruct
450 seawater chemistry, but a number of studies have highlighted the sensitivity of the
451 technique to the precise method used. Wilson et al. (2013) and Blaser et al. (2016) have
452 recently shown that carbonates play an important role as a buffer when extracting the
453 ferromanganese coatings from sediments, preventing the acid-induced mobilization of Fe
454 oxides and volcanigenic material (which has the potential to shift the measured εNd
455 towards more radiogenic values). Crocket et al. (2011) followed the earlier procedure of
456 Gutjahr et al.(2007) which included a prior decarbonation step. Leaching time, sample
457 size and pH have also been shown to influence the bulk sediment leachate εNd signature
458 (Blaser et al., 2016; Wilson et al., 2013)

459 Radiogenic offsets of bulk sediment leachate εNd values from foraminiferal Nd
460 signatures have also been documented in core top samples at other sites in the North
461 Atlantic and attributed to the influence of volcanic ash that is easily leached during the
462 extraction process (Elmore et al., 2011). There is no evidence, however, for increased or
463 sustained volcanic activity throughout the Holocene that would increase the accumulation
464 rate of ash in the sediment at Site 980 (see supplementary information). Instead, North
465 Atlantic ash deposition is much more frequent during the last deglaciation (e.g. Abbott
466 and Davies, 2012; Davies et al., 2012; Lowe et al., 2008), when very good agreement
467 exists between foraminiferal and bulk sediment leachate εNd at Site 980. Airborne
468 volcanic ash therefore cannot easily be invoked to explain the observed offset between
469 substrates during the Holocene.

470 Variable transport of silt-sized titanomagnetite grains from the GSR by bottom currents has been documented at a number of sediment drift sites that lie in the path of overflow waters. Higher accumulation of titanomagnetite is documented associated with higher current velocities under interstadial conditions compared to stadials, with grains likely originating from the young basaltic (and radiogenic) rocks of the Nordic Basaltic Province (Ballini et al., 2006; Kissel, 2005; Kissel et al., 1999). Bottom currents were much stronger in the Holocene than the glacial in the North Atlantic (Innocent et al., 1997; Manighetti and McCave, 1995; McCave et al., 1995a; McIntyre and Howe, 2009; McManus et al., 2004; Thornalley et al., 2013), therefore, increased accumulation of leachable fine material at the Feni Drift during the Holocene represents a possible alternative explanation for the more radiogenic signature recorded by Site 980 sediment leachates. There is a possibility that certain other drift deposits in the region may be susceptible to a similar process because the spatial distribution of sites showing unreasonably radiogenic core top sediment leachate εNd values (as identified by Elmore et al. (2011)) agrees well with the major pathways of modern Nordic Overflow waters (e.g. Hansen and Østerhus, 2000), and bulk sediment leachate εNd records showing the largest radiogenic offset compared to other substrates during the Holocene have been reported at some other sites in the North Atlantic (Piotrowski et al., 2012; Roberts et al., 2010).

488 We conclude that the influence of bottom current transported, silt-sized radiogenic particles on sediment leachate εNd is the most likely explanation for the offset between the Nd isotope substrates during the Holocene at our study site, with leachate methods also contributing to the variable offsets throughout the record. This implies that published
bulk sediment leachate values are not an accurate record of bottom water chemistry at Site 980, at least in the uppermost sedimentary column (above 2.9 mcd). As a consequence, we only consider the Nd isotopic signatures of foraminifera and fish debris as potential tracers of water mass provenance at Site 980.

496

<figure><img src="image_5.png" /></figure>

497
Fig. 3: ODP Site 980 neodymium isotope and rare earth element chemistry. (a) Neodymium isotope composition (and 2σ error) of mixed planktonic foraminifera (blue), fish debris (red), sediment leachates (grey, Crocket et al., 2011)) and the detrital fraction (black, Crocket et al., 2011)). (b) and (c) Nd/Ca and Nd/Mn ratios of mixed planktonic foraminifera. (d), (e) and (f) Ratios of mixed planktonic foraminifera PAAS-normalised REEs (Xn) with (d) Ce/Ce* = 3Ce_n/(2La_n+Nd_n), (e) HREE/LREE = (Er_n+Yb_n+Lu_n)/(La_n+Pr_n+Nd_n), (f) MREE/MREE* = Gd_n+Tb_n+Dy_n/0.5(La_n+Pr_n+Nd_n+Er_n+Yb_n+Lu_n). Fish Ce/Ce* ratios are shown by crosses in panel (d). (g) Fluxes of ice-rafted debris (black) and detrital carbonate (red) grains from the 150 – 500 µm size fraction. Purple shading marks the timing of Heinrich-layer deposition (identified from the IRD and detrital carbonate records). Yellow shading highlights the upper section of the core in which Nd isotope ratios derived from sediment leachates are offset to more radiogenic values than the records of planktonic foraminifera and fish debris (named zone I). The deeper section of the core with higher Nd/Ca values and no systematic inter-substrate εNd offset is named zone II. Note that depth within sediment core (in metres composite depth) is plotted on a non-linear scale.

### 4.1.2 Controls on downcore rare earth element distributions

The downcore REE signatures of planktonic foraminifera (shown in Fig. 3b–e) are characterized by upper (zone I) and lower (zone II) intervals with a distinct transition at 2.9 mcd (11.5 ka). Nd/Ca values (Fig. 3b) are consistently relatively low throughout zone I (~700 – 800 nmol/mol), with higher and more variable values in zone II (900 – 3000 nmol/mol). Nd/Mn values are also more variable in zone II, with the highest values typically associated with (or close to) H-events. Below the transition at 2.9 mcd there is less enrichment of heavy REEs over light REEs (Fig. 3d) and cerium anomaly values are closer to 1 (Fig. 3c). A systematic offset between the sediment leachate εNd values and the other substrates is also recorded in zone I but is not clearly present in zone II (Fig. 3a).
This observation raises the possibility of a link between the processes controlling REE distributions across 2.9 mcd and inter-substrate offsets in ε<sub>Nd</sub> values. We explore three potential explanations for the distinct difference in behavior of REEs in the upper (zone I) and lower (zone II) sections of the studied interval. These are (i) exchange between detrital and authigenic phases, (ii) an active redox front in the sediment and (iii) preserved differences between glacial and interglacial bottom water chemistry.

The REE signature of bulk sediment can co-vary with lithology (e.g. Sholkovitz, 1988). The detrital fraction, however, is typically unreactive, as evidenced by its extraction procedure (Bayon et al., 2002; Jones et al., 1994), and REE concentrations are much lower than in the Fe-Mn oxyhydroxide fraction (Gutjahr et al., 2007). The ε<sub>Nd</sub> values of bulk sediment at Site 980 (Crockett et al., 2011) are consistently offset towards less radiogenic values than any of the substrates discussed here (Fig. 3a). There is also no clear relationship between the relative proportions of radiogenic and unradiogenic IRD and the co-registered foraminiferal ε<sub>Nd</sub> (see supplementary information). These observations suggest that significant exchange between the detrital and authigenic phases is not the main control on foraminiferal REE distributions.

Many of the phases which have been shown to host REEs associated with foraminifera (Fe-Mn (hydr)oxides, organic matter and Mn carbonates) are sensitive to redox conditions (Tachikawa et al., 2014). Therefore, the presence of an active redox front at 2.9 m depth could result in a remobilisation and redistribution of foraminiferal REEs. The similarity of shift in REE distribution recorded by both foraminifera and fish debris, however, argues against this explanation at Site 980. The REE signal of fish debris is acquired at shallow depths in the sediment and is relatively resistant to
diagenesis, with substitution of rare earth elements into the crystal lattice occurring only under more extreme conditions and resulting in a distinctive bell-shaped REE profile (Martin and Scher, 2004; Reynard et al., 1999; Staudigel et al., 1985). This profile is not seen in any of the samples presented here (Fig. 4). In addition, strong negative Ce/Ce* values are preserved in Miocene fish teeth at nearby Site 982 (Martin et al., 2010), similar to that preserved by Holocene (but not glacial) samples at Site 980, suggesting that REE distributions are not reset below the top few metres of sediment. Therefore, it is unlikely that diagenetic redistribution explains the shift in REE distribution in either uncleaned foraminifera or fish debris, *supporting our use of fish debris and foraminiferal REE chemistry at Site 980 to reconstruct past oceanographic conditions*.

The cerium anomaly (Ce/Ce*) is used to track oxidation because cerium is readily converted from the insoluble Ce (IV) to the soluble Ce (III) phase when water column oxygenation decreases (Elderfield et al., 1988). It should, however, be noted that Ce/Ce* preserved by sediment components is not an infallible recorder of local bottom water oxidation, with questions raised regarding the incorporation of REEs into various marine substrates (e.g. German and Elderfield, 1990; Holser, 1997; MacLeod and Irving, 1996). In addition, lower Ce/Ce* ratios are recorded in the modern Pacific and Indian Oceans compared to the North Atlantic Ocean; a result which is attributed to a progressive removal of Ce as the age of a bottom water mass increases (German and Elderfield, 1990). The higher Ce/Ce* values at Site 980 during the glacial could therefore indicate a more vigorous overturning circulation than during the Holocene, with several studies supporting vigorous renewal of intermediate waters in the North Atlantic under glacial conditions (e.g. Gherardi et al., 2009; McCave et al., 1995a). An alternative explanation
is that the higher Ce/Ce* values in zone II (close to 1; Fig. 3c) in our record indicate that glacial bottom waters at Site 980 had a relatively low oxygen content compared with conditions in the Holocene. A similar increase in both Ce/Ce* and Nd/Ca of uncleaned planktonic foraminifera during the transition into the Holocene has been recorded at Bermuda Rise (Roberts et al., 2012). Glacial-interglacial differences in the degree of oxygenation of shallowly buried sediments have been documented in a number of studies around the world (Jaccard and Galbraith, 2012), with decreased oxygenation of glacial bottom waters reported in the intermediate depth northeast Atlantic (Baas et al., 1998; Schönfeld et al., 2003). The broad coincidence of the boundary between zone I and II with the climatic transition into the Holocene therefore raises the possibility that the differences in REE behaviour at Site 980 are the product of differences in the degree of pore water oxidation between glacial and interglacial conditions.
<figure><img src="image_6.png" /><figcaption>Fig. 4: Comparison of REE profiles of ODP Site 980 fish debris (dotted lines) and mixed
planktonic foraminifera (solid lines). Values are normalised to PAAS and La = 1, with samples
from zone I in panel (a) and zone II in panel (b). Sample ages given in figure legend.</figcaption></figure>

Assessing the influence of water mass labelling by IRD at Site 980
Foraminiferal $\epsilon_{Nd}$ values during the late Holocene ($-11.7 \pm 0.3$) at Site 980 are consistent with the representative signatures of the water masses bathing this site today but determining whether older $\epsilon_{Nd}$ signatures are also representative of regional water mass chemistry requires further evaluation. The radiogenic signal ($-7.9$ to $-4.2$) in foraminiferal $\epsilon_{Nd}$ of the nearby BOFS cores (Fig. 1) during the Last Glacial Maximum has been interpreted to reflect labelling of the bottom waters bathing these sites by IRD raining through the water column (Roberts and Piotrowski, 2015). Therefore, before we can use foraminifera-based $\epsilon_{Nd}$ data from Site 980 to determine past variations in water mass sourcing, first we must explore the possibility that exchange with lithogenic material has influenced glacial $\epsilon_{Nd}$ values in our datasets.

Three lines of evidence suggest that there was a smaller influence of local modification of water mass chemistry by IRD at ODP Site 980 than the BOFS sites. First, the radiogenic $\epsilon_{Nd}$ values ($\ge -8$) reported by Roberts and Piotrowski (2015) during MIS 2 are not recorded at Site 980 (Fig. 5a). Second, we observe $\epsilon_{Nd}$ excursions of opposing signs during H2 and H4, despite lithologically similar IRD assemblages (Fig. 3a and supplementary information). Third, our new IRD records demonstrate that IRD inputs during the last glacial period at Site 980 (typically 0 – 650 grains/g, peaking at 1450 grains/g during H2) are lower than those recorded at the BOFS sites (400 – 4700 grains/g, (Roberts and Piotrowski, 2015)) (Fig. 5c and d). The same result is also true of the concentrations of volcanic grains, highlighting that the Feni Drift was likely subject to a different IRD depositional regime than the one influencing the BOFS sites during the last glacial period. Although we cannot discount the influence at Site 980 of water masses modified by IRD upstream of our site, and no direct correlation was found between
volcanic IRD concentration and εNd values by Roberts and Piotrowski (2015), we suggest
that the absence of large volumes of reactive material sinking through the water column
at Site 980 during the last glacial period reduced the potential for local modification of
water mass chemistry. We therefore suggest that IRD water mass relabelling is not the
dominant control on the down core variability in our εNd data. Hence, modification of the
water column εNd by IRD (Roberts and Piotrowski, 2015) may not be significantly
problematic in all regions of IRD deposition.

617
<figure><img src="image_7.png" /><figcaption>Fig. 5: Comparison of palaeoclimate records from Site 980, Feni Drift (black, this study and Crocket et al. (2011)) and BOFS sites (colour) on and along the southern flank of the Rockall</figcaption></figure>

618

619

620
Plateau (Roberts and Piotrowski, 2015) (water depths in legend, site locations shown in Fig. 1).
(a) Uncleaned foraminifera (solid circles and lines) and fish debris $\epsilon_{Nd}$ (crosses). (b) Bulk sediment leachate $\epsilon_{Nd}$ values. (c) Concentration of IRD grains 150 – 500 µm (Site 980) and >150 µm (BOFS cores (Roberts and Piotrowski, 2015)). (d) Concentration of volcanic IRD (including basalt and mafic and felsic volcanic glasses) and pumice clasts with grain size 150 – 500 µm (Site 980) and >150 µm (BOFS sites, Roberts and Piotrowski (2015)). (e) Sedimentation rates (cm/kyr) based on published age models (BOFS sites, Roberts and Piotrowski (2015)) and this study (Site 980).

629

4.2 Water mass changes

631

Having established that fish debris and uncleaned planktic foraminifera provide a reliable record of bottom water Nd isotope composition at ODP Site 980, next we combine these $\epsilon_{Nd}$ data with three other proxies for water mass chemistry (B/Ca, $\delta^{13}$C and $\delta^{18}$O) to reconstruct oceanographic variability in the mid-depth northeast Atlantic over the past 40,000 years.

637
<figure><img src="image_8.png" /><figcaption>Fig. 6: Bottom water chemistry shifts at ODP Site 980 during H-events. Shaded bands mark the position of the H-events identified by IRD abundance at Site 980. (a) Bottom water carbonate ion</figcaption></figure>
concentrations reconstructed from B/Ca ratios of benthic foraminifera C. wuellerstorfi, with error bars indicating analytical uncertainty (calibration uncertainty ± 10µmol/kg (Yu and Elderfield, 2007)). (b) Neodymium isotope ratios of planktonic foraminifera with ferromanganese coatings not removed (in blue) and fish debris (red) (error bars ±2σ). (c) δ¹³C of C. wuellerstorfi, with darker line colour marking the three-point running mean. (d) δ¹⁸O of C. wuellerstorfi, adjusted for global ice volume changes, with darker line colour marking the three-point running mean. (e) Fluxes of ice-rafted debris (black) and detrital carbonate clasts (red) in grains 150 – 500 µm cm⁻² kyr⁻¹.

## 4.2.1 Glacial-interglacial variability

In Fig. 6, we present Site 980 fish debris and foraminiferal ε$_{Nd}$ alongside benthic δ¹³C and B/Ca ratios for the past 41 kyrs. A distinct difference in the properties of the water mass bathing Site 980 can be seen between the glacial and Holocene, with glacial waters having a more radiogenic ε$_{Nd}$ signature (Fig. 6b), slightly lower carbon isotope values (Fig. 6c) and higher [CO₃²⁻] (indicated by higher benthic foraminiferal B/Ca ratios, Fig. 6a). These observations strongly suggest that the interpretation of Crocket et al. (2011) invoking the persistent dominance of WTROW throughout the last 40 kyr at this site cannot be correct without a significant shift in the properties of WTROW.

Marked similarities exist between the long-term evolution of ε$_{Nd}$ at Site 980 and deep water sites at the Bermuda Rise (Böhm et al., 2015; Gutjahr and Lippold, 2011; Roberts et al., 2010) and South Atlantic Cape Basin (Piotrowski et al., 2012; Piotrowski et al., 2005; Rutberg et al., 2000) (Fig. 7). The Bermuda Rise and Cape Basin records are primarily controlled by mixing between unradiogenic northern- and radiogenic southern-sourced waters. An increased contribution of radiogenic southern-sourced waters mixing
with northern-sourced waters in the glacial northeast North Atlantic cannot be ruled out
on the basis of Nd isotopes alone, but B/Ca-based reconstructions of carbonate ion
concentrations argue against this interpretation. Both modern and glacial SSW typically
have low [CO₃²⁻] (75 – 95 µmol/mol), whereas modern North Atlantic Deep Water and
Nordic Sea overflows are much richer in [CO₃²⁻] (>100 µmol/mol) (Key et al., 2004; Yu
et al., 2014; Yu et al., 2008). Reconstructed glacial carbonate ion concentrations (~130 –
150 µmol/mol) at Site 980 are higher than those in the Holocene (~115 – 135 µmol/mol)
(Fig. 6), the opposite direction of change as would be expected if an increased glacial
presence of southern-sourced waters is the dominant driver of the difference in bottom
water chemistry observed at Site 980 between the Holocene and last glacial period.

A reduction in the influence of unradiogenic LSW under glacial conditions
(Cottet-Puinel et al., 2004; Hillaire-Marcel et al., 2001) could contribute to the more
radiogenic glacial εNd signature at Site 980, however, as modern LSW has high [CO₃²⁻]
(Key et al., 2004; Yu et al., 2008), a decrease in glacial LSW influence alone cannot
explain the observed bottom water chemistry changes at Site 980. A decreased export of
Mediterranean outflow water (MOW) has been documented under glacial conditions
(Zahn et al., 1997), but as modern MOW is moderately radiogenic (–9.4 ± 0.5 (Spivack
and Wasserburg, 1988)), a reduction in MOW cannot explain the glacial shift towards
more radiogenic bottom waters at Site 980. Alternatively, Rogerson et al. (2005) suggest
that the core of MOW deepened which could increase its influence at Site 980, and as
MOW has a high modern [CO₃²⁻], potentially contribute to the observed shift in glacial-
interglacial bottom water chemistry. An increased influence of relatively radiogenic and
high [CO₃²⁻] Nordic Sea overflow waters at Site 980 during the last glacial period as
suggested by Yu et al. (2008) can also explain the more radiogenic and [CO₃²⁻]-rich glacial bottom water signature at Site 980. It should be noted that an increased influence of overflow waters on bottom water chemistry at Site 980 does not require an increase in overflow current vigour, particularly in light of the decreased contribution of LSW (Cottet-Puinel et al., 2004; Hillaire-Marcel et al., 2001) to the main Atlantic basin during the last glacial period.

<figure><img src="image_9.png" /><figcaption>Fig. 7: Comparison of Site 980 uncleaned foraminiferal $\epsilon_{Nd}$ records (black) with deep sites at the Bermuda Rise (ODP 1063/OCE326-GGC6 in red/orange (Böhm et al., 2015; Gutjahr and Lippold, 2011; Roberts et al., 2010)) and South Atlantic (TNO57-21/RC11-83 in blue/purple (Piotrowski et al., 2012; Piotrowski et al., 2004, 2005)). Foraminiferal $\epsilon_{Nd}$ values are shown with solid lines and filled circles, while bulk sediment leachates are shown with dotted lines and</figcaption></figure>
hollow circles. Note that GGC6 bulk sediment leachate values are not shown as these were deemed not to be representative of bottom water chemistry by Roberts et al. (2010).

### 4.2.2 Heinrich events

Our multi-proxy reconstruction of bottom water properties at Site 980 shows prominent excursions associated with increased concentrations of IRD during each of H-events 4 – 1 (Fig. 6). Large shifts in the $ε_{Nd}$ of bottom waters are documented for both H4 ($ε_{Nd}$ = –13.6) and H2 (–7.1), but these excursions are of opposing signs. A small (~ 20 µmol/kg) reduction in bottom water [CO$_3$ $^{2-}$] is recorded during H2 and just prior to H3, but there is no significant change during H4, and at no point in the record do the values get close to SSW values (75 – 95 µmol/kg (Key et al., 2004; Yu et al., 2014; Yu et al., 2008)). Excursions towards low benthic $δ^{13}$C are recorded during each of the H-events, although typically lagging slightly behind the main IRD peak.

Arguably, the most notable feature is the unique expression of each event in our proxy records of bottom water chemistry, with differences in the sign of change, as well as the magnitude of excursions. This simple observation shows that circulation changes in the mid-depth glacial northeast Atlantic are much more complex than simple two component northern- and southern-sourced endmember mixing model. Instead, either three or more water masses must have influenced Site 980 during the last glacial period, or alternatively, shifts in surface ocean properties (particularly in deep water formation regions) or circulation vigour significantly modified the properties of water masses.

The surface of the North Atlantic Ocean is thought to have been a highly dynamic environment during the last glacial period, with temporally variable freshwater input
from both icebergs and continental sources (Heinrich, 1988; Hemming, 2004; Lekens et al., 2006; Stanford et al., 2011). Fluctuations in sea surface temperature, salinity, sea ice cover and productivity have all been documented in the high northern latitudes (e.g. Dokken et al., 2013; Maslin et al., 1995; van Kreveld, 1996). An increase in sea ice cover has the potential to reduce air-sea exchange in the deep water formation regions, with CO₂ invasion and evasion acting upon both the δ¹³C and carbonate ion concentration of a water mass in a specific ratio (Fig. 8) (Yu et al., 2008). The variability in our data between H-events and non-Heinrich glacial intervals does not follow this trend, therefore we infer that variable sea-ice cover in deep water formation regions is not a major contributor to the variability in bottom water chemistry recorded at Site 980 during H-events. Instead, there is a closer fit to the “biology” slope, which represents the influence of biological regeneration on seawater chemistry and incorporates the effects of organic tissue degradation and CaCO₃ remineralisation (Yu et al., 2008). The influence of a water mass with increased organic matter remineralisation may therefore explain part of the H-event signals in our proxy records. This signature could reach the site either by increased surface productivity or increased age of the bottom waters reflecting a longer time interval since the water mass was exposed to the atmosphere, either as a result of a rearrangement of circulation patterns, a shift in water mass boundaries or a decrease in circulation vigour. Additional independent proxy records for current vigour such as such as sortable silt (McCave et al., 1995b) or ²³¹Pa/²³⁰Th (Yu et al., 1996) could therefore prove valuable in better understanding the origin of the observed excursions of bottom water chemistry in the northeast Atlantic during H-events.
<figure><img src="image_10.png" /><figcaption>Fig. 8: Comparison of Site 980 bottom water chemistry to regional water mass signatures. (a) Cross-plot of co-measured oxygen and carbon isotopes of *C. wuellerstorfi* from ODP Site 980. All values are smoothed (3-point running mean), with δ¹⁸O values adjusted for global ice volume changes. Estimated water mass compositions shown in grey, based upon Bertram et al. (1995), Voelker et al. (2006), Meland et al. (2008) and Thornalley et al. (2010). Water masses abbreviations: GAABW: Glacial Antarctic Bottom Water; GAAIW: Glacial Antarctic Intermediate Water; LSW: Labrador Sea Water (Holocene value); GNAIW: Glacial North Atlantic Intermediate Water; GNSDW: Glacial Norwegian Sea Deep Water. Note that the isotopic composition of brines has high uncertainty (e.g. Meland et al., 2008; Thornalley et al., 2010). (b) Cross-plot of δ¹³C *C. wuellerstorfi* and estimated bottom water [CO₃²⁻] calculated from *C. wuellerstorfi* B/Ca ratios (see methods). Also shown are lines indicating the calculated slope of co-variance in δ¹³C and [CO₃²⁻] generated by biological regeneration and CO₂ invasion/evasion (Yu et al., 2008). Reconstructed water mass signatures are shown in grey (data from Yu et al., 2014; Yu et al., 2008). NEAP 8K is thought to record a strong influence of overflow waters from the Nordic Seas at the Last Glacial Maximum (Yu et al., 2008).</figcaption></figure>

4.2.2.1 Heinrich event 4
<figure><img src="image_11.png" /><figcaption>Fig. 9: Reconstruction of changes in surface and deep ocean during H4. (a) Bottom water carbonate ion concentrations reconstructed from B/Ca ratios of benthic foraminifera *C. wuellerstori* with error bars indicating analytical uncertainty (calibration uncertainty ± 10</figcaption></figure>

765

766

767

768
769 µmol/kg (Yu and Elderfield, 2007)). (b) Neodymium isotope ratios of planktonic foraminifera
770 with ferromanganese coatings not removed (in blue) and fish debris (red). Error bars indicate
771 ±2σ. (c) δ¹³C of C. wuellerstorfi, with darker line colour marking the three-point running mean.
772 (d) δ¹⁸O of C. wuellerstorfi, corrected for global ice volume changes, with darker line colour
773 marking the three-point running mean. (e) Percentage of polar species N. pachyderma (s.) as a
774 proportion of the total number of planktonic foraminifera. (f) Fluxes of ice-rafted debris (black)
775 and detrital carbonate (red) grains from the 150 – 500 µm size fraction of Site 980 samples.
776

776
777 Heinrich event 4 is distinct from the other H-events in our record because it is
778 characterised by an excursion towards unradiogenic εNd values and the most significant
779 decrease in benthic foraminiferal δ¹⁸O in our record (Figs. 6d and 9). Similar δ¹⁸O
780 excursions have been linked to episodes of brine formation in the Nordic Seas (Dokken
781 and Jansen, 1999; Meland et al., 2008; Vidal et al., 1998), although the origin of the low
782 δ¹⁸O signature is debated (e.g. Bagniewski et al., 2015; Bauch and Bauch, 2001;
783 Rasmussen and Thomsen, 2009; Stanford et al., 2011). Regardless, a Nordic Sea origin
784 for the waters bathing Site 980 during H4 can likely be discounted because modern
785 Nordic Seas Overflows have radiogenic signatures of εNd = -8.2 ± 0.6) (Lacan and
786 Jeandel, 2004b) while our data record a shift of ~ 2 εNd units towards more unradiogenic
787 values. A very large shift in the neodymium isotope composition of the Nordic Seas (for
788 example, due to overprinting by unradiogenic IRD from Greenland) would be required
789 for the shift in bottom water chemistry recorded at Site 980 to be attributed to increased
790 presence of Nordic Seas Overflows at the site. The main source of unradiogenic waters in
791 the North Atlantic today is the Labrador Sea. LSW generation is thought to have been
792 weaker during the last glacial period, although with some evidence of sinking of brine-
rich waters during Heinrich events (Cottet-Puinel et al., 2004; Hillaire-Marcel and de Vernal, 2008; Hillaire-Marcel et al., 2001; Nuttin et al., 2015; Weber et al., 2001).

Alternatively, an increase in deep water formation south of the GSR could explain the $\epsilon_{Nd}$ signal at Site 980 during H4 (Duplessy et al., 1980; Labeyrie et al., 1992). Surface waters today in the Northeast Atlantic have $\epsilon_{Nd}$ values of -13 to -14.8 (Lacan and Jeandel, 2004c), although southward migration of the polar front during H-events (e.g. Eynaud et al., 2009) may have resulted in more radiogenic surface water signatures. Reconstructions of surface water properties in the open ocean suggest that they may have been insufficiently dense to sink (Maslin et al., 1995). Alternatively, brine formation, could have occurred on the European and/or Icelandic margins, Rockall Plateau or GSR (Meland et al., 2008; Thornalley et al., 2010), with a possible additional input of waters with low $\delta^{18}$O and $\delta^{13}$C signatures from the European continent (e.g. Eynaud et al., 2007; Toucanne et al., 2009).

### 4.2.2.2 Heinrich event 2

Heinrich event 2 shows a very different signal to H4 in our multi-proxy records. Unlike H4, there is no $\delta^{18}$O excursion (Fig. 6d), and $\epsilon_{Nd}$ values shift ~2 units towards more radiogenic values (> -8; Fig. 6b). A combination of radiogenic $\epsilon_{Nd}$ and low $\delta^{13}$C with no concurrent excursion in $\delta^{18}$O in the Atlantic basin are commonly associated with SSW in the Atlantic (e.g. Jonkers et al., 2015; Piotrowski et al., 2008; Rutberg et al., 2000; Thornalley et al., 2010). We also see a small decrease in carbonate ion concentrations (~20 µmol/kg), which supports an increase in the contribution of southern-sourced waters to the northeast North Atlantic. Our reconstructed [CO$_3$$^{2-}$] of > 120 µmol/kg at Site 980
during H2 is, however, well above modern Antarctic bottom and intermediate waters (~85 µmol/kg) (Key et al., 2004; Yu et al., 2008), and even more discrepant with values of ~80 µmol/kg reported for deep SSW during the Last Glacial Maximum (Yu et al., 2014), so any southern-sourced water would need to be highly diluted by mixing with northern-sourced endmembers. A alternative source of radiogenic waters with high [CO3<sup>2-</sup>] is the Nordic Seas (Lacan and Jeandel, 2004b; Lacan et al., 2012), a possibility consistent with proposed strengthening of GSR overflows into the North Atlantic during H-events (Crocket et al., 2011; Meland et al., 2008; Thornalley et al., 2010).

A clear temporal lag (~ 2 kyr, represented by 30 cm of core) exists between the onset of excursions in ε<sub>Nd</sub> and δ<sup>13</sup>C during H2 (Fig. 10). The initial shift in ε<sub>Nd</sub> values is approximately concurrent with increasing % N. pachyderma (s.) values. There is no distinct offset between % N. pachyderma (s.) and IRD fluxes during either the onset or termination of Heinrich stadial 2, providing independent support of the suggestion that cooling preceded IRD deposition in the North Atlantic during H-events (Barker et al., 2015). Bottom water [CO<sub>3</sub><sup>2-</sup>] values only start to decrease when stadial conditions are well established. The delayed excursion towards low δ<sup>13</sup>C values could be a result of a local or regional pulse in the flux of organic matter to the seafloor (Mackensen et al., 1993), or the flushing of previously poorly ventilated waters once ice-rafting has ceased and circulation vigour is increasing. Alternatively, the observed lag may indicate the presence of multiple phases of AMOC reorganization (Wilson et al., 2014), involving the existence of short-term circulation modes, rather than a simple switch between two distinct AMOC states ('glacial' and 'off') (Rahmstorf, 2002).
<figure><img src="image_12.png" /><figcaption>Fig. 10: Reconstruction of changes in surface and deep ocean during H2. Symbols as in Fig. 9.</figcaption></figure>
840

841

4.2.2.3 Exploration of the differences between H-events

842

843

844

845

846

847

848

849

850

851

852

853

854

855

856

857

858

859

Comparison between multiple H-events allows greater insight into the processes governing AMOC response to fresh water additions. The results of two recent studies (Böhm et al., 2015; Lynch-Stieglitz et al., 2014) suggest that an important role in determining the magnitude of oceanographic response to freshwater input is played by the stage in the glacial cycle at which freshwater was released. One of these studies, based on reconstruction of flow through the Florida Straits, suggests that AMOC was least sensitive to freshwater forcing during full last glacial conditions when stratification of the Atlantic Ocean was at its greatest (Lynch-Stieglitz et al., 2014). The other, a study of water chemistry in the deep North Atlantic on Bermuda Rise, suggests that during the last glacial cycle, AMOC shutdown only occurs when a threshold in continental ice sheet size is crossed, with ice volumes close to their maximum (Böhm et al., 2015). While the magnitude of AMOC perturbation cannot be simply gauged by the amplitude of the excursion in our proxy records, the variability across our data sets strongly suggests a discernible reorganization in ocean structure associated with both H2 (which occurs at close to peak glacial conditions) and H4 (when ice sheets were much smaller (e.g. Siddall et al., 2003)). Our results therefore offer a different perspective, suggesting that the stage of the glacial cycle occupied by H-events may not be the principle factor governing the response of the mid-depth North Atlantic to freshwater release.

860

861

862

The unique chemical signature of intermediate waters during each of the H-events examined (H4 to H1) at Site 980 therefore requires an alternative explanation to changes in global ice volume. Differences in the location, duration and magnitude of melt water
inputs have been shown to exert a clear influence over the response of AMOC in
modelling experiments (Bigg et al., 2011; Otto-Bliesner and Brady, 2010; Smith and
Gregory, 2009). Differences in the IRD assemblages have been documented between H-
events (e.g. Grousset et al., 1993; Hall et al., 2011; Peck et al., 2007; Snoeckx et al.,
1999), most notably a decreased flux of Hudson Bay-sourced detrital carbonates reaching
the North Atlantic during H3 and H6 (e.g. Bond et al., 1992; Hemming, 2004; Hodell and
Curtis, 2008), suggesting differences in the relative fresh water contributions from the
circum-Atlantic ice sheets. Each H-event is therefore likely associated with different
volumes, rates and/or delivery routes of freshwater to the North Atlantic Ocean. These
differences are likely to have resulted in shifts in the relative strength of surface water
subduction in a number of different potential deep water formation areas. This
interpretation of events also offers a potential explanation as to why different events have
a different prominence in different locations, and therefore why records from different
study regions point to different relationships between global ice volume and AMOC
stability (Böhm et al., 2015; Lynch-Stieglitz et al., 2014). AMOC strength has also been
shown to exert a strong control over continental hydroclimate (e.g. Burckel et al., 2015;
Mulitza et al., 2008), therefore, variations in meltwater inputs between H-events and the
resulting AMOC perturbations may have important implications for terrestrial climate.

The unique response of intermediate water to fresh water input during each H-
event examined (4 to 1) also raises the possibility of variations in the properties of
GNAIW during H-events; a result that may have implications for downstream
reconstructions of the strength of exported water from the high northern latitudes at these
885 times. Glacial bottom water chemistry at Site 980 however appears relatively stable
886 between H-events.

887 Our records indicate that a dominance of northern-sourced waters in the mid-
888 depth northeast North Atlantic persisted throughout the last glacial period, but the
889 absence of a strong influence of SSW at our intermediate depth study site does not
890 preclude a dramatic slowdown of the deep ocean during certain H-events (Burckel et al.,
891 2015; Mulitza et al., 2008). In fact, a shoaling of the overturning circulation cell is a
892 feature common to many simulations of freshwater input to the North Atlantic where
893 AMOC shutdown is incomplete (Böhm et al., 2015).

894

895 **5. Summary and Conclusions**
896

896

897 We investigate the extent to which uncleaned foraminifera, fish debris and bulk sediment
898 leachates can record regionally representative bottom water neodymium isotope
899 compositions in the northeast North Atlantic by examining the isotopic signature of
900 multiple substrates and associated rare earth element distributions. We find that neither
901 local modification of bottom water chemistry by IRD labelling or redistribution of REE
902 at depth within the sediment column exert a strong control over the REE and εNd
903 signatures preserved by planktonic foraminifera without their Fe-Mn oxide coatings
904 removed at Site 980, making them suitable substrates for palaeoceanographic
905 reconstructions. Foraminiferal and fish debris εNd values show good agreement
906 throughout the last 40,000 years, however, bulk sediment leachate values are offset
907 towards more radiogenic values by 1–2 εNd units throughout the Holocene, which we
attribute to the increased influence of fine-grained radiogenic material transported by bottom currents on the sediment leachate values at this time. This result calls for careful evaluation of extracted leachate neodymium isotope compositions for leaching artifacts at North Atlantic drift sites.

By combining our new Nd isotope records with three additional proxies for water mass chemistry (B/Ca, δ¹³C and δ¹⁸O signatures of *C. wuellerstorfi*), we demonstrate that there is an increased influence of Nordic Seas overflow waters (and not southern-sourced waters) at 2.2 km depth in the Rockall Trough during the late glacial, possibly related in part to a reduction in LSW generation. A dominant presence of SSW at Site 980 at any time within the past 40 kyr is ruled out by reconstructed [CO₃²⁻] greater than 110 µmol/kg throughout our record. We find that all four H-events within our study interval have different geochemical signatures, with H4 and H2 marked by εNd excursions in opposing directions. Unradiogenic εNd values and light oxygen and carbon isotope values during H4 could indicate an increased contribution of waters from either the Labrador Sea or deep waters forming south of the Greenland-Scotland Ridge, while the radiogenic εNd signature recorded during H2 is more likely explained by an increased presence of overflow waters from the Nordic Seas. We find that the stage of the glacial cycle occupied by H-events may not be the principle factor governing the response of the mid-depth North Atlantic to freshwater release. Instead, we suggest that this heterogeneity in the intermediate depth northeast North Atlantic most likely arises due to differences in fresh water input locations, magnitudes and fluxes amongst H-events. The balance of contributions from different source regions to the northern intermediate/deep water endmember does not appear to remain constant, with important implications for the
931 chemical signature of GNAIW that is transported downstream *during millennial-scale climate events*.

933

934 Acknowledgements

935 This research used samples provided by the Integrated Ocean Drilling (Discovery)
936 Program IODP, which is sponsored by the US National Science Foundation and
937 participating countries under management of Joint Oceanographic Institutions, Inc. We
938 thank Walter Hale and Alex Wülbers for help with sampling and Matt Cooper, Andy
939 Milton, Mike Bolshaw and Dave Spanner for analytical support. Heiko Pälike, David
940 Thornalley and Rachel Mills are thanked for productive discussions and comments on
941 earlier versions of this work. We also thank three anonymous reviewers for their
942 constructive feedback, which greatly improved the manuscript. Funding for this project
943 was provided by NERC studentships to A.J.C. (grant NE/D005728/2) and T.B.C.
944 (NE/I528626/1), with additional funding support from a Royal Society Wolfson Research
945 Merit Award and NERC grants NE/F00141X/1 and NE/I006168/1 to P.A.W. and
946 NE/D00876X/2 to G.L.F.

947

948 Competing financial interests

949 The authors declare no competing financial interests.

950

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1525
# Geochemical Response of the Mid-Depth Northeast Atlantic Ocean to Freshwater Input During Heinrich Events 1 to 4

Anya J. Crocker¹,²*, Thomas B. Chalk¹,³, Ian Bailey¹,⁴, Megan R. Spencer¹, Marcus Gutjahr¹,⁵, Gavin L. Foster¹, Paul A. Wilson¹

* We examine H1 to H4 in the mid-depth NE Atlantic (ODP 980) using ε$_{Nd}$, B/Ca, δ¹³C and δ¹⁸O.

* Each H-event shows a unique, prominent geochemical excursion with no major influence from southern sourced water.

* We infer the imprint of water masses sourced from different northern hemisphere sites of subduction.

* Our findings may help explain why some H events show bigger climatic perturbations than others.