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A reconstruction of sea surface warming in the northern North Atlantic duringMIS 3 ice-rafting events

Lukas Jonkers a,*, Matthias Moros b,c, Maarten A. Prins d, Trond Dokken c, Carin Andersson Dahl c,e,Noortje Dijkstra d,1, Kerstin Perner b, Geert-Jan A Brummer aaRoyal Netherlands Institute for Sea Research, Department of Marine Geology. PO Box 59, 1790 AB Den Burg, The Netherlandsb Leibniz Institute for Baltic Sea Research Warnemünde. Seestrabe 15, D-18119 Rostock, GermanycBjerknes Centre for Climate Research. Allégaten 55, NO-5007 Bergen, NorwaydVU University, Faculty of Earth and Life Sciences, Marine Biogeology section. De Boelelaan 1085, 1081 HV Amsterdam, The NetherlandseUniversity of Bergen, Department of Earth Science. Allégaten 41, N-5007 Bergen, Norway

a r t i c l e i n f o

Article history:Received 23 November 2009Received in revised form24 March 2010Accepted 26 March 2010

a b s t r a c t

Marine isotope stage 3 (29e59 kyr BP) is characterised by rapid shifts fromcold stadial towarm interstadialperiods, which may be linked to changes in the vigour of the Atlantic Meridional Overturning Circulationdue to variable freshwater input by melting ice. Here we present two northern North Atlantic multi-proxyrecords of sea surface conditions that indicate warm (near) sea surface conditions during such ice-raftingevents. We infer near surface temperature from planktonic foraminiferal counts, Mg/Ca and oxygenisotopes of left-coiling Neogloboquadrina pachyderma and from calcite content. Temperatures increasedduring ice-rafting and rose rapidly to interstadial values after ice-rafting ceased. This pattern is clearestduring Heinrich Event 4, but also present throughout the other millennial scale ice-rafting events. Itindicates that stadials in the Greenland ice-cores are concurrentwith a (near) surfacewarming in theNorthAtlantic,whichwas probably restricted to the summer, aswinter temperaturesmust have remained low forsediment-laden ice to reach the site. As similar warming during ice-rafting events is seen regionally in thenorthern North Atlantic it cannot be explained by a rerouting of the North Atlantic Current. Rather, weattribute it to a shoaling of a warm subsurface water mass that was formed as a result of decreasedventilation of the upper waters and a continued northward subsurface !ow of warm water. Planktoniccarbon isotopes support this suggestion showing coincident decreased ventilation during deposition of ice-rafted detritus (IRD). The absence of a clear meltwater spike in the d18O records during IRD input suggeststhat besides glacial freshwater, sea ice may have been responsible for the ventilation decrease and asso-ciatednear surface heat built up. The proposed scenario is in agreementwithmodelling studies that requirethe release of heat trapped below the surface to restart the overturning circulation.

! 2010 Elsevier Ltd. All rights reserved.

1. Introduction

Freshwater input into theNorthAtlanticOceanbymelting ice hashad profound impacts on climate during the last glacial, through itsslowing effect on the Atlantic Meriodional Overturning Circulation(AMOC) that affects the global distribution of heat (e.g. Bond et al.,1992; Ganopolski and Rahmstorf, 2001). The Greenland ice-coreshave revealed abrupt climatic changes from cold stadials to warminterstadials and back, now known as DansgaardeOeschger events,duringmarine isotope stage 3 (MI S3; 29e59 kyr BP) (Johnsen et al.,

1992; Dansgaard et al., 1993). Similar climatic variability is alsoobserved inmarine records and it is often associatedwith ice-raftingand hence freshwater input (Bond et al., 1993; Cortijo et al., 2000;Van Kreveld et al., 2000; Elliot et al., 2002; Prins et al., 2002).Freshwater input by melting icebergs left a clear trace in the sedi-ment in the form of layers of ice-rafted detritus (IRD) (e.g. Heinrich,1988; Bondet al.,1992). These IRD layers are generally recognized bytheir coarse-grained nature and are found over a huge area in theNorth Atlantic (Hemming, 2004 and references therein). Majorepisodes of ice-rafting are known as Heinrich events (HEs), whichare interspaced with ice-rafting events on millennial time scales.Many studies have con"rmed the impact of ice-rafting events onpastocean circulation (e.g. Vidal et al.,1997;VanKreveld et al., 2000;Elliot et al., 2002; Prins et al., 2002). Much attention has been givento a glacial origin of IRD and the source areas of IRD are well known

* Corresponding author. Tel.: !31 (0)222 369 343; fax: !31 (0)222 319 674.E-mail address: lukas.jonkers@nioz.nl (L. Jonkers).

1 Present address: Department of Geology, University of Tromsø. Dramsveien 201,9037 Tromsø, Norway.

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(Bond et al., 1992; Grousset et al., 2001; Peck et al., 2007). However,the underlying mechanism that causes the ice-rafting eventsremains subject to debate and discussion about whether HEs andmillennial scale ice-rafting events have a similar mechanistic originpersists (Marshall and Koutnik, 2006; Alvarez-Solas et al., 2010).Internal ice-sheet dynamics have been proposed (Alley andMacAyeal, 1994) as well as (climatic) factors outside the ice-sheets(Heinrich, 1988; Lagerklint and Wright, 1999; Moros et al., 2002;Rasmussen and Thomsen, 2004). Modelling studies showa pronounced slowdown of the AMOC in response to freshwaterforcing, but also suggest that the rapid resumptionwhen freshwaterforcing ceases may be related to the sudden release of subsurfaceheat, built up due to reduced ventilation during freshwater forcing(Kaspi et al., 2004; Shaffer et al., 2004; Mignot et al., 2007). Suchrelease of heat could explain the abrupt change to warm conditionsthat characterises the Greenland interstadials. However, proxyrecords that document such awarming during ice-rafting are scarce(Rasmussen and Thomsen, 2004) and the abrupt temperatureincrease, therefore the onset of circulation, remain not fully under-stood. Hence, high-resolution time-series that document pastoceanic changes during such ice-rafting events are instrumental inincreasing our understanding of the rapid changes in the MI S3climate system.

Here we present data from sediment cores LO09-18 and DS97-2P from the central northern North Atlantic spanning w31e48 kyrBP. This period is characterised by high frequency and amplitudeclimatic changes (cf. Cortijo et al., 2000). We focus on MIS 3because of this pronounced variability and the clarity of thesignals in the two cores. Both records show frequent input of IRDand associated slowdown of deep-circulation (Moros et al., 1997,2002; Prins et al., 2002; Rasmussen et al., 2002; Snowball andMoros, 2003). Through planktonic foraminiferal counts, stableisotope and trace metal analysis as well as bulk mineralogical andgeochemical tracers/proxies we infer upper ocean dynamicsassociated with these ice-rafting events.

2. Sediment cores and regional surface hydrography

Sediment cores LO09-18 and DS97-2P were retrieved fromintermediate water depths on the eastern !ank of the ReykjanesRidge (Fig. 1; Table 1). Surface circulation at the sites is presently

in!uenced by the Irminger Current, a branch of the North AtlanticCurrent that brings warm and salty water to the north. TheIrminger Current is thus a major component of the warm north-ward !ow of the AMOC. The polar front, which separates the coldand fresh Polar waters carrying ice to the south, from the warmand salty Atlantic waters, presently lies just east of the EastGreenland Current (Dickson et al., 1988) (Fig. 1). During the colderperiods of the last glacial however, the polar front shiftedtowards the southeast and repeatedly reached our core locations(Ruddiman and McIntyre, 1981; Bard et al., 1987; Eynaud et al.,2009) and associated IRD deposition occurred. The sites lie northof the main IRD belt (Ruddiman, 1977) and received IRD fromGreenland, Iceland and Eurasian sources, rather than fromLaurentian sources during the last glacial (Van Kreveld et al., 2000;Prins et al., 2002; Moros et al., 2004). During the Last GlacialMaximum (LGM) the cyclonic Central North Atlantic Gyre served asthe major transport belt of icebergs to the North Atlantic (Robinsonet al., 1995; Watkins et al., 2007). The location of our core sites atthe northern edge of the gyre likely led to the overall reduced inputof IRD compared to areas located along the main track of icebergtransport (Robinson et al., 1995). According to the reconstructionsof Sarnthein et al. (2003) our core locations are close to the sea-icelimit in LGM winter.

3. Material & methods

3.1. Foraminiferal and isotope analysis

To complement the Rasmussen et al. (2002) dataset, additionalforaminiferal census counts in core DS97-2Pweremade at irregularintervals on >200 specimen per sample between 150 and 250 mm.The same size fraction is used for stable isotope analyses, whichwere conducted at a resolution of 1 cm. For the intervals580.5e621.5 cm, 672.5e710.5 cm and 784.5e815.5 cm at least fourmeasurements of four specimens for every sample of left-coiling N.pachyderma (N. pachyderma s.) were performed on a Finnigan MAT252 mass spectrometer with a Kiel II device. Reproducibility wasmonitored using an internal standard and amounted to <0.1& ford18O and w0.05 for d13C (1 s.d.). Other samples were measuredusing a Delta Plus mass spectrometer with GASBENCH, usingaliquots of w25 specimens. Reproducibility was determined usingboth an internal standard and NBS20 and was w0.15& andw0.1&for oxygen and carbon isotopes, respectively. Overlapping sectionswere measured using both methods, con"rming that the twodifferent methods had not affected our results.

Paleotemperature estimates for core DS97-2P were obtainedfrom planktonic foraminiferal Mg/Ca. For Mg/Ca analyses approxi-mately 50 shells between 150 and 250 mm of left-coilingN. pachyderma were cleaned following the protocol outlined byBarker et al. (2003). Samples were centrifuged after sample disso-lution and measured on a Varian Vista ICP-AES according to deVilliers et al. (2002). Analytical precision was determined bysimultaneous measurement of a standard solution and amountedto "0.01 mmol/mol (1 s.d.). Contamination by terrigenous materialwas monitored by Fe and Al concentrations; no indications ofcontamination were found. Paleotemperature estimates are basedon the Nürnberg et al. (1996) temperature calibration, which yields

Fig. 1. Simpli"ed modern surface circulation in the northern North Atlantic Ocean(based on McCartney and Talley, 1982). Cores LO09-18 and DS97-2P are indicated bya black dot and a white plus, respectively. The dashed grey line depicts the maximumextent of icebergs in the North Atlantic (www.geus.dk), most of the ice howeverpresently remains within the EGC. NAC: North Atlantic Current, IC: Irminger Currentand EGC: East Greenland Current.

Table 1Core positions and depths.

Core Latitude Longitude Water depth (m)

LO09-18 58# 58.04 N 30# 40.99 W 1471DS97-2P 58# 56.33 N 30# 24.59 W 1685

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temperatures in agreement with previous estimates of summer seasurface temperatures for the same period based on foraminiferalabundance data from LO09-18 and with a second nearby core (VanKreveld et al., 2000). The skeletal Mg/Ca of another foraminiferalspecies was recently shown to be dependent, not only on temper-ature, but also on salinity (Mathien-Blard and Bassinot, 2009). Ifthis would apply to N. pachyderma s. as well then the temperatureincreases during ice-rafting events may be overestimated. Addi-tionally, Mg/Ca temperature calibrations tend to be relativelyinsensitive at low temperatures (e.g. Elder"eld and Ganssen, 2000;Kozdon et al., 2009). The temperature amplitude inferred fromforaminiferal Mg/Ca is thus likely to represent a minimum value.

Foraminiferal census counts of core LO09-18 were performed ata resolution of approximately 1 cm on the >150 mm fraction; witha minimum of 500 individuals counted per sample (see supple-mentary information). Sea surface temperatures (SST) were esti-mated from the foraminiferal assemblages using weightedaveraging partial least squared regression (WA-PLS) (Ter Braak andJuggins, 1993) and the training set of Kucera et al. (2005). The WA-PLS method is relatively well suited for use with (virtually) mono-speci"c samples, which are frequently encountered in the record.Consequently, the method yields more robust/conservative SSTminima when compared to other transfer functions, such as themaximum likelihood method (Telford, pers. comm.). The moderntemperature values for 10 m water depth during summer (July,August, September) for the calibration dataset were taken from theWorld Ocean Atlas version 2 (WOA, 1998).

For stable isotope analyses of core LO09-18 (every 2 cm) at leastthree measurements of seven shells of N. pachyderma s. >150 mmwere performed per sample using a Finnigan MAT 253 coupled toa Kiel IV device. Reproducibility of an external standard (NBS19)amounted to $0.05& and $0.03& (1 s.d.) for d18O and d13C,respectively. The abundance of foraminifrela shells >250 mm wasvery low. Comparison of stable isotope ratios between both cores isthus not likely to be complicated by size-speci"c offsets. All isotopevalues are reported on the PDB scale and samples were ultrasoni-cally cleaned with ethanol prior to measurement.

Studies on the ecology of N. pachyderma s. indicate that its stableisotope composition re!ects conditions around 50 m depth in thewatercolumn (Carstens and Wefer, 1992; Volkmann and Mensch,2001; Jonkers et al., in press). Mg/Ca ratios probably indicatetemperatures at a similar depth. Foraminiferal census counts, onthe other hand, yield estimates of sea surface temperature s.s. Sinceall inferences show similar patterns we have applied the term nearsurface, however (small) discrepancies between census counts andchemical inferences may be due to this depth difference.

3.2. Grain-size measurements

The grain-size record of DS97-2P (Prins et al., 2001, 2002) wasextended at resolution of 1 cm from 650 cm downward usinga Fritsch A22 laser diffraction particle sizer. Precision of the laserparticle sizer has been shown to be better than 0.5% for the meangrain-size (Jonkers et al., 2009). To isolate the lithogenic fraction,samples were treated according to Konert and Vandenberghe(1997). The sand percentage reported here represents the volumefraction >63 mm. The presence of sand grains in North Atlanticsediments has since long been used to infer ice-rafting events, aslarge lithogenic particles are most likely transported by ice. Thesand % thus provides a robust, albeit conservative, indicator of thecontribution of IRD to the sediment.

Previous analyses on these cores include XRF scanning of theDS97-2P core for Ti, K and Ca (Prins et al., 2001, 2002). Ca countshave been converted to carbonate weight % (Prins et al., 2001). TheTi/K ratio of the sediment re!ects the relative contribution of

continental crust vs. mid-oceanic ridge basalt derived material(Richter et al., 2006), thus serving as an indicator of continentallyderived IRD. Calcite, quartz and plagioclase content of core LO09-18were determined by XRD (Moros et al., 2002, 2004). The quartz/plagioclase ratio of North Atlantic sediments re!ects the varyingcontribution of continental derived quartz versus that of basalticplagioclase and can also be used as a proxy for the contribution ofIRD (Moros et al., 2002, 2004). Although calcite and carbonatecontent are not the same, for the sake of simplicity they are here-after referred to as calcite.

3.3. Age model

Both cores show remarkably similar patterns of calcite contentand contribution of IRD, enabling a detailed correlation. Acceler-ated mass spectrometry (AMS) radiocarbon dating was performedon 11mono-speci"c samples of N. pachyderma s. from between 300and 550 cm of core LO09-18 at the Poznan radiocarbon laboratory(Fig. 2; Table 2). After correction for reservoir effects (400 yrs)radiocarbon years were converted to calendar years based on theFairbanks0107 calibration curve (Fairbanks et al., 2005) and an agemodel was developed as outlined by Heegaard et al. (2005). Theactual reservoir ages may have varied considerably over time andthe assumption of a constant reservoir age might be too simple(Waelbroeck et al., 2001; Adkins et al., 2002). However, since wederive proxies for ice-rafting and temperature from the same cores,

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Table 2Radiocarbon dates.

Sample, depth (cm) AMS 14C-0.4a (kyr) Lab code Calibrated age (kyr)

LO09-18 312.5 25.59" 0.21 Poz-8174 30.79" 0.26LO09-18 350 28.04" 0.28 Poz-8176 33.41" 0.33LO09-18 394.5 29.60" 0.30 Poz-29243 35.02" 0.33LO09-18 403 31.10" 0.40 Poz-8171 36.47" 0.42LO09-18 426.5 31.36" 0.28 Poz-29244 36.73" 0.32LO09-18 453 31.90" 0.50 Poz-8182 37.28" 0.53LO09-18 467.5 33.30" 0.40 Poz-29246 38.69" 0.42LO09-18 484.5 35.00" 0.60 Poz-29247 40.34" 0.59LO09-18 503 36.60" 0.80 Poz-8178 41.78" 0.72LO09-18 514.5 39.20" 0.70 Poz-29250 44.03" 0.63LO09-18 552.5 41.20" 0.90 Poz-29252 45.78" 0.81DS97-2P 535 27.86" 0.18 Poz-29234 33.22" 0.24DS97-2P 586 28.80" 0.17 Poz-29235 34.20" 0.23DS97-2P 631.5 31.61" 0.29 Poz-29237 36.99" 0.33DS97-2P 667.5 32.27" 0.32 Poz-29238 37.66" 0.36DS97-2P 701.5 31.09" 0.28 Poz-29239 36.45" 0.31DS97-2P 740 26.42" 0.20 Poz-29240 31.69" 0.26DS97-2P 752 37.40" 0.60 Poz-29810 42.48" 0.55D\7-2P 769.5 29.31" 0.28 Poz-29242 34.73" 0.32a Reservoir effect.

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our inferences are independent of the chosen reservoir age. Theaverage sedimentation rate of sediment core LO09-18 isw16.5 cm kyr%1. Aroundw37 kyr some rapid short term changes insediment accumulation rate occurred, possible related to thesudden input of IRD, which is known to affect sedimentation rates(Francois and Bacon, 1994; Thomson et al., 1995). To illustrate thedifference between atmospheric (Greenland) and sea surfacetemperature patterns, the SST record of LO09-18 was visuallytuned, within the 14C age bounds, to the NGRIP d18O record(NGRIP_Members, 2004) on the GICC05 age scale (Svensson et al.,2008). We assumed that the highest SSTs correspond to thewarm and long interstadial 8.

The age model of DS97-2P has been revised because of thechronological inconsistencies in the glacial part of the record inearlier versions (Prins et al., 2001; Rasmussen et al., 2002).Matching of the calcite and IRD records of both cores using Ana-lySeries 2.0 (Paillard et al., 1996) served as our primary age esti-mate, which was subsequently con"rmed by new AMS radiocarbondates of mono-speci"c samples of N. pachyderma s. (Table 2). Downto w700 cm the new radiocarbon dates agree well with the ageinferred from the correlationwith LO09-18 (Fig. 3). Beyond that, theradiocarbon ages appear younger, although the correlationbetween both cores remains high until at least 820 cm, suggestingthat the stratigraphy of DS97-2P is uncompromised. The proxyrecord of DS97-2P in fact resembles that of a suite of cores from theReykjanes Ridge area (Moros et al., 2002), supporting the strati-graphic robustness of the core and indicating that it representsa signal of regional signi"cance. The average sedimentation rate ofcore DS97-2P is slightly higher, w18 cm kyr%1.

4. Results

Sand percentages and quartz/plagioclase ratios display distinctmillennial scale variability during the glacial period (Fig. 4A).Increases in sand percentages and the quartz/plagioclase ratio areconcurrent with decreasing log (Ti/K) ratios (Fig. 4A), suggestingmajor changes in the provenance and transport mechanism of thesediment. All three parameters point to a continental crust derivedice-rafted origin of the sediment. In core DS97-2P the ice-raftingevents are identi"ed on the basis of the sand content. Since IRD isnot limited to coarse particles only, this yields a more conservativeestimate of IRD content than the geochemical indicators (Lemmen,1990; Nürnberg et al., 1994; Dowdeswell et al., 1998). Regardless ofthe method used to infer ice-rafting, both records show a similarsequence of IRD input events. In none of the cores Heinrich events(HEs) stand out as events of high magnitude or long duration.

The calcite contributions show signi"cant variationwith distinctpeaks between the ice-rafting events (Fig. 4B). Calcite contentstarted to increase after IRD input, but also during IRD input in theolder part of the sequence (HE5 and events k and j). Highest calcitecontent was reached after IRD deposition ceased. Assumingconstant accumulation rates between dating points, the duration ofthese calcite peaks can be constrained to w600e800 years,although the maximum following HE4 may have lasted longer.

In core DS97-2P the carbonate record is broadly mimicked bythe relative abundance of left-coiling N. pachyderma, where thespecies is less abundant during high carbonate excursions (Fig. 4B).The species dominates the sedimentary planktonic foraminiferalassemblage at high Northern latitudes (P!aumann et al., 1996;

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Schröder-Ritzrau et al., 2001) and its relative abundance serves asan indicator of the in!uence of polar waters (Bé and Tolderlund,1971; P!aumann et al., 1996, 2003). Consequently, foraminiferalbased estimates of past sea surface temperatures (SST) at highlatitudes are largely determined by the abundance of N. pachy-derma s. (Kandiano and Bauch, 2003). We therefore use its abun-dance in DS97-2P as a relative paleo (near) SST indicator. Thepresence of Globigerina bulloides on the other hand, is indicative ofwarmer conditions; as is that of Turborotalita quinqueloba, which isalso often associated with oceanic fronts (Bé and Tolderlund, 1971;Schröder-Ritzrau et al., 2001). In LO09-18 the relative abundance ofN. pachyderma s. decreased fromvirtually 100 to around 50% duringthe transition from HE4 to interstadial (IS) 8 (Figs. 4 and 5). At thesame time G. bulloides and T. quinqueloba together increased to over40% (Fig. 5). A similar, but less pronounced, pattern is also visibleduring IRD input associated with HE4 (Figs. 4 and 5), whenG. bulloides and T. quinqueloba together gradually reach up to 18%

before abruptly dropping back to almost 0%. Even the short IS9 ischaracterised by a small, but signi"cant change in the foraminiferalassemblage (Fig. 5 and S2 for absolute abundances).

Foraminiferal census based paleotemperatures in LO09-18 rangebetween 3.8 and 7.9 #C (RMSE & 1.68 #C) (Figs. 4D and 5D). Pre HE4temperatures were slightly below 4 #C, with IS9 marginallywarmer. Only during HE4 SSTs increased considerably to >5 #C.After IRD input ceased temperatures dropped slightly, and subse-quently increased to around 8 #C during IS8 (Fig. 5D). A similartrend is visible in the oxygen isotopes of N. pachyderma s (Fig. 5D).The paleotemperature estimates indicate unambiguously thatduring peak ice-rafting, sea surface temperatures reached (at leastseasonally) over 5 #C.

The Mg/Ca values of N. pachyderma s. obtained from DS97-2Pvary between 0.6 and 1.0 mmol mol%1 and show a pattern verysimilar to the oxygen isotopes (Fig. 4C). Estimated temperaturesvary between 3.0 and 9.6 #C (Fig. 4D). The Mg/Ca derived

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Fig. 4. Proxy records of LO09-18 (left panel) and DS97-2P (right panel). A: IRD contribution. B: Calcite/carbonate content and relative abundance of N. pachyderma s. (note invertedaxis). C: planktonic oxygen isotopes. D: paleotemperature estimates based on Mg/Ca (values in mmol/mol are indicated) for DS97-2P and foraminiferal assemblages for LO09-18.Vertical bars depict temperature error: 1 #C for Mg/Ca (Peck et al., 2008) and 1.68 #C for foraminiferal census data. E: Carbon isotope values of N. pachyderma s.with 5-point runningaverage. Grey bars indicate ice-rafting events.

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temperatures arewithin error of the assemblage based estimates ofLO09-18. In accordance with the record of LO09-18 highertemperatures are reached during the interstadial periods after ice-rafting ceased, but notably HE4 and event h are already charac-terised by relatively warm conditions (Fig. 4D). The use of differentcalibrations to estimate temperature does not qualitatively affectthe reconstructions and con"rms that even during HE4 (near) SSTsreached seasonally up to w5 #C (see supplementary information).The Mg/Ca data thus corroborate the synchronous (near) surfacewarming and IRD deposition.

The observed alternation from periods with ice-rafting to warmintervals resembles the stadial-interstadial pattern of theGreenland ice-cores. However, in contrast to the cold Greenlandstadials, some warming of the surface waters already occurred

during input of IRD. In LO09-18 this is clear during HE4, but thecalcite record indicates that also during events k and j, surfacewater conditions became warmer (Figs. 4 and 5). Surface warmingin DS97-2P may have started during ice-rafting events HE5, HE4, k,j, h and g (Fig. 4). Although the resolution of the foraminiferalcensus record is not very high, its covariance with the calcite recordis clear. Peaks in the calcite content of both cores are thus concur-rent with (near) surface warming, and importantly, support thatsome warming also took place during IRD input.

The d18O of N. pachyderma s. varies between w3 and 4.5&. Theamplitude of the changes in DS97-2P is slightly larger than in LO09-18 (Fig. 4C). Temporal patterns are very similar between the twocores and resemble the evolution of paleotemperature. Highestvalues tend to occur during the start of, or between, IRD events. Thelowest values are associated with the interstadials. The transitionfrom high to low d18O starts in both cores during IRD deposition,except during event i in LO09-18 and event f in DS97-2P (Figs. 4 and5). Throughout the LO09-18 d18O record decreases occur graduallyand are terminated by an abrupt increase back to heavier values;the d18O curve is thus not dissimilar to a saw-tooth pattern. Suchpattern is less clear in DS97-2P, but HE5, HE4 and h have a similarshape (Fig. 4C). In general the d18O record co-varies with the CaCO3content and is inversely related to the abundance of N. pachydermas. Only prior to HE4 and HE5, when d18O already reached lowvalues, the calcite and isotope patterns diverge. Even the smalldecrease in the abundance of N. pachyderma s. during the secondhalf of HE4 is re!ected in the d18O record of LO09-18 (Fig. 5). Sincethe oxygen isotopes, Mg/Ca ratios, calcite content and N. pachy-derma s. abundance co-vary, the d18O values appear to be domi-nated by temperature rather than by changes in the d18O of theseawater. The planktonic d18O data support the onset of a warmingtrend during IRD deposition and the temperature maxima after theevents (Figs. 4 and 5). The minimum in d18O prior to HE3 in DS97-2P seems anomalous, but coincides with increased abundance ofthe infaunal dwelling benthic foraminifer Fursenkoina fusiformisthat is indicative of low oxygen levels possibly related to highnutrient !uxes from the surface (Rasmussen et al., 2002). Theinterval thus seems to re!ect atypical conditions, but the d18Orecord generally con"rms the warming during IRD input.

Stable carbon isotope values of left-coiling N. pachyderma varybetween w0.5 and %0.4& with, again, the amplitude of thechanges greatest in DS97-2P (Fig. 4D). Negative excursions areassociated with IRD-events. In LO09-18 peaks in IRD are concurrentwith d13C minima and the decrease in d13C starts with the onset ofIRD deposition (Fig. 4D). In DS97-2P IRD events also show lowplanktonic d13C values, but the exact timing with respect to the IRDpeak depends on the parameter used to infer IRD (Fig. 4). Whensand is used only HE4 and event h lead the d13C minima, but whenthe Ti/K is used events HE4 to f lead. Note however that d13Cminima always occur before the peak in calcite content (Fig. 4).Controls on the stable carbon isotope ratios of planktonic forami-nifera are poorly constrained; the d13C may amongst others bein!uenced by air-sea exchange, the carbonate ion concentrationand nutrient concentration and respiration (Labeyrie and Duplessy,1985; Lynch-Stieglitz et al., 1995; Bauch et al., 2002). Here weassume that the latter process dominates and we use the d13C toinfer ventilation of the (near) surface waters (cf. Sarnthein et al.,1995). The consistency in the signal suggests that reduced surfaceventilation was a common feature of the IRD-events and warmingonly peaked after mixing of the upper ocean commenced.

5. Discussion

In both cores a similar sequence of IRD-events is present.Paleotemperature estimates based on Mg/Ca of N. pachyderma s. in

36 37 38 39 40 41GICC05 kyr b2k

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Fig. 5. Detailed proxy records of LO09-18 tuned to the NGRIP d18O record(NGRIP_Members, 2004) on the GICC05 age scale (Svensson et al., 2008). A: NGRIPd18O recorded on inverted axis. B: Relative planktonic foraminiferal abundances.C: quartz/plagioclase values as indicator of IRD input. D: Sea surface temperature (errorbars indicated). E: d18O of N. pachyderma s. (error bars are standard deviations on themean of three measurements). F: calcite percentage. The AMS radiocarbon dates ofLO09-18 are indicated. Note that the error on the GICC05 scale is in the order of 1.5 kyr.The shaded areas highlight interstadials (IS) 8 and 9.

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core DS97-2P and foraminiferal census counts in LO09-18 indicatethat many of these IRD-events were characterised by increasing(near) surface temperatures (Figs. 4 and 5). Temperature maximawere reached after IRD deposition stopped and coincide with peaksin calcite content (Figs. 4 and 5). Since detrital carbonates areabsent (Van Kreveld et al., 2000), the calcite content, thus appearsrelated, through changes in productivity, to (near) surfacetemperature (Brummer and van Eijden, 1992; Lackschewitz et al.,1998). The oxygen isotopes of the planktonic foraminiferN. pachyderma s. generally con"rm the temperature record (Figs. 4and 5). Both records show repeatedly that the interstadial warmingthat peaked after the ice-rafting event already started when IRDdeposition was going on. During HE4 and event h warming mayeven have started before or at least simultaneously with ice-rafting(Figs. 4 and 5). The generally cold stadial periods were thus char-acterised by relatively warm (near) surface waters at the ReykjanesRidge sites. In the North Atlantic IRD-events are generally charac-terised by decreased planktonic d18O values associated withdecreased surface salinities (e.g. Bond et al., 1992; Labeyrie et al.,1995; Vidal et al., 1997). However, such changes are not clearlyvisible in the d18O records of the cores presented here, similar tomany sites north of the Heinrich belt (Cortijo et al., 1997). Appar-ently the oxygen isotope signal of N. pachyderma s. at the Reykjanesridge sites was not dominated by the vast amounts of isotopicallydepleted glacier-derived ice that were released into the oceanduring ice-rafting events.

In all events maximum warming is observed following IRDdeposition and after increasing ventilation as indicated by the d13Cof N. pachyderma s. (Fig. 4). The transition from cold to warmconditions however started already during the ice-rafting events(Figs. 4 and 5). This warming was probably a very seasonalphenomenon as the growing season of planktonic foraminera livingat high latitudes is limited to a short period in summer (Schröder-Ritzrau et al., 2001). Nevertheless, it unequivocally indicates thata climatic amelioration already started during ice-rafting. It appearslikely that this warming even contributed to increased melting ofice and IRD deposition. Similar warm conditions during IRDdeposition are also evident from calcite content and foraminiferalassemblages in core SO82-05 on the western !ank of ReykjanesRidge (Van Kreveld et al., 2000). Mg/Ca derived temperatures fromN. pachyderma s. and G. bulloides in a core offshore of southernIreland also show relatively high temperatures during IRD input(Peck et al., 2008), which was suggested to be related to changes inthe pathway of the warm NAC. Furthermore, records from thewestern !ank of Reykjanes Ridge and the Irminger basin alsosupport the inferred warming during IRD input (Elliot et al., 1998;Van Kreveld et al., 2000). These records show shifts to lower d18ON. pachyderma s. values occurring during the IRD input andminimum d18O values being reached afterwards. During someevents the relative abundance of N. pachyderma s. showed a tran-sient decrease (Elliot et al., 1998; Van Kreveld et al., 2000), sug-gesting similar progressively more in!uence of warmer waterduring IRD input as observed in LO09-18 and DS97-2P. Elliot et al.(1998) and Van Kreveld et al. (2000) explained the lag of the d18Oshift with respect to IRD deposition by a decoupling between IRDinput and freshwater forcing, where either exclusively clean icemelted after IRD deposition or the fresh water anomaly wastransported away from the site of ice melting and release of IRD.Such decoupling would however not explain the warming inferredfrom other proxies that accompanied the shift to low d18O values. Itwould also imply advection of the meltwater anomaly over longdistances, which would be possible only if very little mixingoccurred, which is unlikely. We argue therefore that the ReykjanesRidge records of planktonic d18O largely re!ect a seasonal warmingconcurrent with IRD input.

It is important to constrain the depth at which this warmingoccurred, especially since the water column likely exhibited strongstrati"cation (cf. the low d13C values) due to meltwater input. Npachyderma s. is able to tolerate low temperatures and salinities,which is evident from its calci"cation habitat in the upper 100 m ofthe East Greenland Current, characterised by cold and low salinitywater (Kohfeld et al., 1996; Stangeew, 2001; Simstich et al., 2003;Kozdon et al., 2009). The species’ oxygen isotopes re!ect temper-ature and d18Ow of the halocline waters (Carstens and Wefer, 1992;Volkmann and Mensch, 2001; Jonkers et al., in press); the inferredwarming is thus (partly) a subsurface phenomenon. However, theforaminiferal assemblages indicate that sea surface temperaturesalso increased when the halocline waters warmed. Since Mg/Cavalues indicate calci"cation temperature and foraminifera assem-blages yield SST, the mismatch between the Mg/Ca and assemblagebased temperature minima may re!ect depth differences betweenboth proxies. It may also be due to the low sensitivity of the Mg/Cacalibration at low temperatures (Kozdon et al., 2009). The rapidwarming into the interstadials may re!ect the sudden release ofheat from the subsurface after strati"cation, due to the presence ofa meltwater lens or drift ice, broke down (Bakke et al., 2009). Thisimplies continued warm water !ow to the north, beneath thesurface (Rasmussen and Thomsen, 2004), and reduced ventilationof intermediate waters during ice-rafting (Mignot et al., 2007) toallow for the built up of a heat reservoir. Our data indicate thatwarmer waters already reached the surface episodically, when ice-rafting was going on, most likely during warm summers. The datathus hint at the presence of warmer waters, mostly below thesurface, during generally cold conditions and at the formation ofsuch a subsurface heat reservoir. Similar observations further west(Van Kreveld et al., 2000) and off Ireland (Peck et al., 2008) suggestthat the warming was indeed a regional phenomenon caused bya shoaling of a warm subsurface water mass, rather than by a localrerouting of the North Atlantic Current. Warming parallel with theIRD input (e.g. HE4) would support the hypothesis that advection ofwarm waters to the north lead to the destabilisation of tide-waterglaciers and thus triggered the ice-rafting events (Lagerklint andWright, 1999; Moros et al., 2002).

The absence of a clear isotopically depleted glacier-derivedmeltwater signal during IRD input may re!ect a balance betweenmeltwater and temperatureeffects. Anotherpossibilitycouldbe thatpart of the IRD was supplied via sea-ice, since it is isotopicallyidentical to seawater (O’Neil,1968; TanandStrain,1996). Sea-ice cancontain considerable amounts of sediment, even sand-sized(Kempema et al., 1989; Nürnberg et al., 1994; Eicken et al., 2000).Such a scenario would require enhanced transportation of sea-icefrom the north (Duplessy et al., 1996; Bradley and England, 2008;Hillaire-Marcel and de Vernal, 2008). Future geochemical prove-nance and/or particle shape analyses of the IRDmayprovide insightsinto the exact role of sea-ice as a carrier of IRD. However, a sea-icecover would provide a very ef"cient insulator of the ocean and itsdisappearance may play an important role in the rapid climaticswitches during MIS 3 (Gildor and Tziperman, 2000; Kaspi et al.,2004; Li et al., 2005). Whether the implied break-down in strati"-cation during interstadial conditions is due to external factors (e.g.changing wind stress), or due to increased buoyancy derived fromthe subsurface warming, remains dif"cult to infer from the data.However, regardless of its cause, rapid emptyingof a subsurfaceheatreservoir could be instrumental in reinvigorating the AMOC,possibly even cause the overshooting of the circulation, and explainthe abrupt warming associated with DansgaardeOeschger events(Ganopolski and Rahmstorf, 2001; Shaffer et al., 2004; Wang andMysak, 2006; Mignot et al., 2007; Liu et al., 2009). In any case, theReykjanes Ridge records show that MI S3 ice-rafting events,although occurring under generally cold conditions, were

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characterised by relatively high and progressively increasingtemperatures. This warming may even have played a role in thedestabilisation of tide-water glaciers providing a trigger for theonset of ice-rafting.

6. Conclusions

We have presented two high-resolution records of sea surfaceconditions from the northern North Atlantic spanning 31e48 kyrBP. The records show a series of nine ice-rafting events inwhich theHEs do not stand out. Using planktonic foraminiferal census counts,stable carbon and oxygen isotopes and paleothermometry we infersea surface conditions associated with these ice-rafting events. Therecords show relatively high (near) surface temperatures duringstadial ice-rafting events. At the same time d13C values were low,possibly indicating a reduced ventilation of surface waters. Thetemperature increase initiated during ice-rafting events culmi-nated in the following interstadial, immediately after IRD deposi-tion. Subsequently, sea surface temperature decreased, re!ectingtypical stadial conditions before the start of the next ice-raftingevent. In accordance with climate models, we attribute the warmtemperatures during ice-rafting to the presence of a subsurfaceheat reservoir that was formed because of reduced ventilation ofsubsurface waters due to the insulating effect of a meltwater lensand/or a sea-ice layer. Shoaling of this warm water mass duringsummer would explain the observed warming during ice-rafting.Upon a total stop in ice-rafting, the insulating layer at the seasurface disappeared and was followed by rapid emptying of thesubsurface heat reservoir, leading to the abrupt onset of the warminterstadial periods.

Acknowledgements

We thank Luke Skinner and Mervyn Greaves for Mg/Ca analysesand Evaline van Weerlee, Michiel Kienhuis, Simon Jung, Bram vander Kooij and Hubert Vonhof for help with the oxygen isotopemeasurements. We are also grateful for the constructive commentsby two anonymous reviewers on a previous version of this manu-script. LJ was supported through the VAMOC project within RAPID(NWO grant no. 854.00.020).

Appendix. Supplementary data

Supplementary data associated with this article can be found inthe online version at doi: 10.1016/j.quascirev.2010.03.014

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