etters 245 (2006) 39–55www.elsevier.com/locate/epsl

Earth and Planetary Science L

Mediterranean outflow strengthening during northern hemispherecoolings: A salt source for the glacial Atlantic?

Antje H.L. Voelker a,⁎, S.M. Lebreiro a, J. Schönfeld b, I. Cacho c,H. Erlenkeuser d, F. Abrantes a

a INETI, Departamento de Geologia Marinha, Estrada da Portela, Zambujal, 2721-866 Alfragide, Portugalb Leibniz-Institute of Marine Science, IFM-GEOMAR, Dienstgebäude Ostufer, Wischhofstraße 1-3; D-24148 Kiel, Germany

c GRC Geociències Marines, Departamento Estrat. Paleont. i Geoc. Marines, Facultat de Geologia, Universitat de Barcelona,C/Martí Franques s/n, 08028 Barcelona, Spain

d Leibniz Labor für Altersbestimmung und Isotopenforschung, Universität Kiel, Max-Eyth-Straße 11, D-24118 Kiel, Germany

Received 15 June 2005; received in revised form 8 March 2006; accepted 9 March 2006Available online 19 April 2006

Editor: E. Boyle

Abstract

High-resolution grain size and magnetic susceptibility records from the eastern Gulf of Cadiz (site MD99-2339; 1170m waterdepth) reveal contourites formed by the Mediterranean Outflow (MOW) during the last 47kyr BP. Oscillations in the MOW'sintensity occurred in phase with Greenland temperature variations with a stronger outflow during northern hemisphere coolingssuch as Dansgaard-Oeschger stadials, Heinrich events, and the Younger Dryas. Benthic δ13C data implies the WesternMediterranean Deep Water as one of the main sources feeding the outflow current, while differential changes in the properties ofthe Mediterranean source and entrained North Atlantic Central Water largely control the MOW's strength. Detailed studies forHeinrich events 1, 4 and 5 show that the flow strength peaked only when subtropical surface waters prevailed in the eastern Gulf ofCadiz, while incursions of icebergs and subpolar surface water were not favorable for the MOW's intensification. As the MOWwas strong when the Atlantic's thermohaline circulation (THC) was weakened, the heat and salt injected by the MOW into theintermediate North Atlantic waters might have preconditioned the THC to switch from the stadial to the interstadial mode.© 2006 Elsevier B.V. All rights reserved.

Keywords: Mediterranean outflow; contourites; Dansgaard-Oeschger cycles; subtropical water; millennial-scale climate variability

1. Introduction

Millennial-scale Dansgaard-Oeschger (DO) cycles,consisting of a cold stadial or Heinrich event (H) phaseand a warm interstadial phase, were first observed inGreenland ice core records [1,2] and later in climate

⁎ Corresponding author.E-mail address: antje.voelker@ineti.pt (A.H.L. Voelker).

0012-821X/$ - see front matter © 2006 Elsevier B.V. All rights reserved.doi:10.1016/j.epsl.2006.03.014

records worldwide [3]. Surface and deep water recordsfrom the North Atlantic region [4-6] as well as climatemodels [7,8] indicate the thermohaline circulation(THC) as a major force controlling these oscillations.The glacial THC had three modes of overturning: thestadial, basic mode with a reduction in North AtlanticDeep Water (NADW) formation [7]; the Heinrich modewith a collapsed NADW formation; and the interstadialmode, which was similar to the present-day circulation.

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The stadial and Heinrich modes were both forced byfreshwater derived from the calving and meltingnorthern hemisphere ice sheets. The reason for theswitch into the interstadial mode, however, is still underdebate. Possible trigger mechanisms include saltconcentration during sea ice formation in the NorthAtlantic's convection areas [6], southern hemisphereforcing in a bipolar seesaw [8] or an oceanic feedbackinvolving the Mediterranean Outflow (MOW) [9].Changes in the properties of mid-depth water massesin the Pacific and Indian Oceans have been linked tomillennial-scale climate change, e.g. [10-12], butoutside of the Bahamas bank [4,13] little is knownabout the North Atlantic's intermediate water massesthat also include the MOW.

2. Modern hydrography

The MOW's Mediterranean water consists ofchangeable parts of Levantine Intermediate Water(LIW) and Western Mediterranean Deep Water(WMDW), which contributes an estimated 0.2Sv[14,15] to the 1Sv outflow volume [14,16]. Bothwater masses are formed in winter when outbursts of

Fig. 1. Chart of the southern Iberian margin and the northern Balearic Islandscores mentioned in this paper. Depth contours are 200, 500 and 2000m (blackstylize pathways of the Western Mediterranean Deep Water (WMDW). Major[24] with principal pathways indicated by black arrows. Note that the MOWwestern Iberian margin and in the open Atlantic. Thick black dashed lines in tWestern IberianWinter Front in the north (after [28]). During winter surface wand transported eastwards with the Azores Current (AC). The wintertime nortCurrent (IPC) [28].

cold dry air from the European continent induce surfacewater cooling and deep convection in the Gulf of Lions(for WMDW), in the southern Adriatic Sea, in thesouthern Aegean Sea, and in the Levantine Basinbetween Rhodos and Cyprus (all three for LIW) [17].Fluctuations in the physical characteristics and volumeof the MOW result from the interannual to decadalvariability in WMDW and LIW formation [17,18]. Asthe MOW enters the Atlantic through the Strait ofGibraltar (main sill depth of 280m) the dense Mediter-ranean water entrains fresher North Atlantic CentralWater (NACW) increasing the transport volume by afactor of 3, decreasing the density, and turning from adensity driven plume into an intermediate depth jet-stream, the Mediterranean undercurrent [19,20]. Thedownstream modification affects the MOW's hydrody-namic and energetic behavior and thereby the lateralspreading [20]. Possibly along with the variable bottomtopography, it also leads to a splitting of the MOW intotwo stability levels: an upper core centered between 500and 800m and a more saline and denser lower corebetween 1000 and 1400m, whose low velocity marginalflow nowadays influences studied site MD99-2339 (Fig.1) as evidenced by CANIGO CTD stations 27 and 45

(inset) with the position of the deep-sea core MD99-2339 and the other) and in addition 1200m (dark gray) in the Atlantic. Dashed gray arrowsspreading area of the Mediterranean Outflow (MOW) (light gray) afteronly occupies water depths between 500 and 1200–1400m along thehe Atlantic indicate the position of the Azores front in the south and theaters and central waters between the two fronts are of subtropical originhern branch of the Azores Current is referred to as the Iberian Poleward

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[21]. Because of their high density and current speed,these cores act as contour currents shaping the seafloorin the Gulf of Cadiz and inducing the formation ofcontourites and sediment drift bodies [22,23]. Despitethe entrainment, the MOW maintains a surplus ofsalinity and temperature, traceable into the open Atlanticwhere it spreads westward across the Atlantic andnorthward along the European continental margin up tothe Irish Sea providing a source of heat and salt for theNorth Atlantic's subtropical and subpolar gyres [24].The NACWat water depths between 100 and 500m hastwo source regions: the warmer and more salinesubtropical (st) NACW is formed in winter along theAzores front at 35–36°N (Fig. 1) and the subpolar (sp)NACW by winter cooling in the eastern North Atlanticnorth of 46°N [25]. NACWst flows eastwards towardsthe Iberian margin and forms a poleward subsurfaceundercurrent along the Portuguese margin [26]. Under-neath the NACWst, NACWsp can move as far south as30–35°N [25]. As both NACW are formed close-byoxygen draw down related to water mass aging or highorganic matter fluxes in the upwelling zones off westernIberia and Morocco is relatively small and no oxygenminimum zone [27] occurs in the study area. Thesurface water in the Gulf of Cadiz consists of subtropicalwater that in winter is derived from the Azores Current[28]. The main surface water current, however, followsthe northern coastline and transports water from thePortugal Current into the Mediterranean where itbalances the volume lost by the outflow.

3. Material and methods

The MOW's history for the last 47kyr Before thePresent (kyr BP) is studied in Calypso piston coreMD99-2339 (35.88°N, 7.53°W; 1170m water depth;Fig. 1), retrieved from a mud wave field [23] in theeastern Gulf of Cadiz during leg 5 of the IMAGES Vcampaign [29]. MOW related data of core MD99-2339is compared to the Mediterranean Sea records ofCalypso cores MD95-2043 and MD99-2343 (Fig. 1).Core MD99-2343 (40°29.84'N; 4°1.69'E; 2391m) wasrecovered north of Menorca during IMAGES V. SiteMD95-2043 (36°8.6'N; 2°37.3'W) was cored in thewestern Alboran Sea from a water depth of 1841mduring the first IMAGES cruise in 1995 [30].

In core MD99-2339, decadal- to centennial-scalerecords of planktonic and benthic isotopes, ice-rafteddebris (IRD), and sea surface temperature (SST) basedon the plankton foraminifera census counts reflecthydrographic conditions of the surface waters. Grainsize and magnetic susceptibility depict hydrodynamic

changes and sediment composition and thereby addressthe MOW's strength.

The sample series consists of two sample sets: onetaken consistently at 10cm spacing throughout the core(Kiel series) and the second one (DGM series) with 5cmresolution for the upper 10.5m and a resolution of 1–20cm for the rest of the core, with the closer spacedintervals in and around the visible contouritic layers andthe section of Dansgaard-Oeschger interstadial (IS)8 and Heinrich event (H) 4. Bulk sediment samples ofthe Kiel series (ca. 25cm3) were weighted, freeze dried,weighted again and then washed through a 63μm mesh-size sieve. The residue was dried at 40 °C. Bulksediment samples of the DGM series (8–25cm3

volume) were weighted wet, dried at 40 °C andweighted again. Dried sample material was disaggre-gated in 0.033mol sodium hexametaphosphate (Calgon)solution and washed over a 63μm mesh-size sieve andin some cases in addition over a 149µmmesh-size sieve.The fraction >63µm was dried in filter paper at 40 °C.From the fraction <63µm the organic matter wasremoved with hydrogen peroxide, excess reagenteliminated in a 60 °C warm water bath, and the sampleswashed with distilled water through diatom ceramiccandles. Prior to grain size analysis each sample washomogenized by stirring and then measured in theMicrometrics Sedigraph 5100 of the Departamento deGeologia Marinha (DGM) of INETI (Alfragide). Resultsshown present the mean grain size of the total finefraction. The sand fraction of both sample sets was usedfor the isotopic and faunal analyses and lithic grain>250μm counts; these grains are considered ice-rafteddebris (IRD) and are presented as number per gram dryweight of the total sample.

Stable isotope measurements were performed on 4–22 shells of G. bulloides or G. ruber white, 1–4 shells ofCibicidoides pachyderma, Cibicidoides wuellerstorfi orCibicidoides sp., and 1–6 specimens of Uvigerina sp.(mainly U. mediterranaea), all picked from the fraction>250μm. Benthic foraminifera were measured solelywith the Carbo Kiel (Kiel 1) device coupled to aFinnigan MAT 251 mass spectrometer at the Leibniz-Labor für Altersbestimmung und Isotopenforschung atKiel University with a precision of ±0.07‰ for δ18Oand ±0.04‰ for δ13C. Planktonic isotope measurementswere partly performed in Kiel and partly in theDepartment of Geosciences (FB5-Geowissenschaften)at Bremen University using a Finnigan MAT 252 massspectrometer (precision: ±0.07‰ for δ18O and ±0.05‰for δ13C). The benthic δ18O values based on Cibici-doides or Cibicides sp. were corrected to Uvigerinaequivalents by adding 0.64‰ [31]. Wherever possible

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the benthic δ18O record is based on C. pachyderma.Values from other Cibicid(oid)es species or Uvigerinasp. were mainly used when C. pachyderma was notavailable (like at the beginning of IS 8 and in the sectionof IS 12 and H 5). The latter species tend to have slightlylighter δ18O values (Fig. 2), which might be related toslightly different temperature and salinity properties inthe MOW.

Sea surface temperatures (SST) in core MD99-2339were reconstructed from the planktonic foraminiferafaunal composition. The faunal assemblage was countedin splits of the fraction >250μm and 150–250μm,respectively, with the total number of specimenscounted in both fractions exceeding 250, when possible.Temperatures were calculated with the SIMMAXprogram [32] using a modern analog data file, whichfollowing the recommendations of the IMAGESMARGO project [33] is restricted to North Atlanticsurface samples. The analog set includes new surfacesamples from the Iberian margin [34], resulting in a totalof 1020 modern analogs. Modern analog temperatures–for a water depth of 10m–were taken from the WorldOcean Atlas 1998 using the webpage of the MARGOproject (http://www.palmod.uni-bremen.de/~csn/woa-sample.html). Paleo-temperatures were calculated

Fig. 2. Core MD99-2339: benthic isotope data split by species groups and on+0.64‰ [31]. In the late Holocene, Uvigerina spp. shows consistently lighfluctuations towards heavier values.

based on the 10 nearest neighbors. Between 13 and19kyr the temporal resolution of the SST record is 80–150yr, 30–80yr from 35.5 to 40.5kyr, and 20–115yrfrom 44 to 47kyr.

Magnetic susceptibility was measured with a Bar-tington MS2F Point Sensor on the GEOTEK Multi-Sensor Core Logger bench at the Bremen Corerepository. Measurement step was 1cm. Data presentedin this paper is stored at the World Data Centre-MARE:http://www.pangaea.de.

4. Chronostratigraphy

4.1. Primary age model of site MD99-2339

The stratigraphy of core MD99-2339 is based on 20AMS 14C ages and three tie points to the GISP2 icecore chronology [2] via δ18O tuning (Table 1). Thetuning points are within H 2, at the transition from H 5to IS 12 and within IS 13. H 2 is identified by its IRDpeak (J. Schönfeld, unpublished data). As the warmSST at the beginning of IS 12 were not favorable for G.bulloides, the correlation point at the onset of the H/IStransition is also based on the δ18O record of G. ruberwhite (see Fig. 8C).

the primary age scale. Cibicid(oid)es spp. δ18O values were adjusted byter δ18O values than C. pachyderma and does not reflect the strong

Table 1Radiocarbon ages, calendar ages and sedimentation rates for core MD99-2339

Depth(s) cm Dated material Laboratorycode

Sampleweight (mg)

Reservoir corrected14C age±1σ error (yr BP)

Calendar age(yr BP)

Sedimentation rate(cm/kyr)

0.5cm Assumed to be recent 0 –57.5cm G. ruber white and pink OS-37284 2.2 1800±95 1727 33.057.5cm Deep dwelling foram. OS-37236 12.2 1900±35 1847a

157.5cm Surface dwelling foram. OS-37237 10.2 5310±35 5888 24.0180cm G. ruber white and pink OS-39990 13.5 7750±60 8524 8.5217.5cm Surface dwelling foram. OS-37238 5.6 10,350±65 12,214217.5cm G. inflata OS-36650 4.0 10,450±70 12,394a 10.2277.5cm Deep dwelling foram. OS-37239 6.9 12,250±55 14,112 31.6433.5cm Pteropod fragments OS-37240 21.3 15,950±85 19,129 31.1633.5cm 2nd δ18O minimum of H 2 in GISP 2 23,800 42.8813.5cm Deep dwelling foram. OS-37241 4.3 25,000±180 28,700 36.7956.5cm Surface dwelling foram. OS-37242 3.5 28,600±170 33,100 32.5956.5cm Deep dwelling foram. OS-37243 4.6 26,000±250b

1020.5cm Deep dwelling foram. KIA19146 9.2 31,490+400/−380 35,090 32.21055.5+

1056.5cmSurface dwellingforaminifera

KIA19147 5.5 32,410+440/−420 35,810 49.3

1055.5+1056.5cm

Deep dwelling foraminifera KIA19148 4.1 32,450+480/−450 35,850a

1195.5cm Surface dwelling foram. KIA19149 7.5 34,550+590/−550 38,550 50.91250.5+

1251.5cmDeep dwelling foraminifera KIA19150 4.7 36,400+720/−660 40,800c –

1307.5+1308.5cm

Deep dwelling foraminifera KIA19151 7.4 36,870+770/−700 41,020 45.6

1379.5cm Deep dwelling foram. KIA19152 9.7 38,230+930/−830 41,480 155.41541.5+

1542.5cmDeep dwelling foraminifera KIA19153 10.6 41,020+1370/−1170 42,820 121.3

1768cm Transition H 5/IS 12 in GISP 2 45,400 87.71838cm 2nd δ18O minimum in IS 13 in GISP 2 46,940 45.5

Depths are the mean of the respective 1cm thick sampling interval. (a) Not used in age model because reliable surface dweller ages are available forthe same depth; (b) on the basis of planktonic δ18O record assumed to be too young; (c) on the basis of planktonic δ18O record rejected as too old.

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Except for the sample at 433.5cm (pteropod frag-ments), 14C ages were measured on multi-speciesforaminifera samples–picked from the fraction>250µm–either consisting of surface dwelling species(mainly: G. ruber white, G. ruber pink, O. universa, G.bulloides, G. falconensis) or deep dwelling species (G.inflata, G. scitula, G. truncatulinoides). Samples of theupper core section were measured in the National OceanSciences AMS (NOSAMS) facility at WHOI (OS code),those from the lower part in the AMS facility of theLeibniz-Labor für Altersbestimmung und Isotopen-forschung at Kiel University (KIA code). Two datesgave problematic results and were excluded. The deepdweller age at 956.5cm of 26,000 14C-yr appears to betoo young, especially in context of the planktonic δ18Orecord, which closely mimics the MD95-2042 [35] andthe Greenland ice core δ18O records during MIS 3 (seeFig. 4). The 14C age at 1250.5+1251.5cm, on the otherhand, appears to be too old, possibly related to aZoophycus burrow [36] or to redistribution of forami-nifera with the contour current.

At four levels (Table 1) surface and deep dwellingspecies were measured in parallel in order to investigate,if the deep dwelling planktonic foraminifera record alarger reservoir effect. With one exception, age offsetsare within the respective 1σ errors (Table 1), indicatingthat the thermocline was as well ventilated as the surfacewater and that no additional correction for the oceanreservoir effect needs to be applied to the deep dwellerages. Consequently, all 14C ages were corrected by−400 yr for the reservoir effect and then calibratedeither with INTCAL04 [37] or according to [38] (allages >433.5cm).

The resulting calendar age model and sedimentationrates are also listed in Table 1. For most of the sedimentsequence sedimentation rates vary between 30 and50cm/kyr. There are two exceptions, which are alsoclearly visible in the age–depth relationship (Fig. 3A).Between 157.5 and 217.5cm, the sedimentation ratesdecrease to 8.5–10.2cm/kyr (Table 1), indicating a“condensed”core section. In contrast, sedimentationrates doubled to tripled between 1379.5 and 1768cm,

Fig. 3. Age–depth plots for the age models of cores MD99-2339, MD95-2043 and MD99-2343. (A) shows the primary age models for all cores,which for core MD99-2339 (black) is the one listed in Table 1 and for core MD95-2043 (gray), the “secondary” age model of [40]. The age model ofMD99-2343 is derived from [44]. Dots or stars represent 14C age based control points, while correlation points to GISP2 are marked by circles incores MD99-2339 and MD99-2343 and by crosses in core MD95-2043. Comparison between the primary (black) and secondary, GISP2-tuned (gray)age models in the sections >20kyr is shown in (B) for MD99-2339 and in (C) for MD95-2043. Differences between the age models are generallysmall for both cores. The arrow in (B) indicates the level (∼1251cm) of the rejected 14C age in core MD99-2339 resulting in a poorer constraint of thesedimentation rates in the primary age model in this core section and to a larger offset between the two age models.

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i.e. between 41 and 45.4kyr BP. The increase coincideswith the summer insolation minimum at 35–40°Nbetween 40 and 45kyr BP, but the reason for the risein accumulation at site MD99-2339 is not understood,yet.

4.2. Secondary age models for cores MD99-2339 andMD95-2043

For the direct comparison of the benthic δ13Crecords of cores MD99-2339 and MD95-2043, the agemodels of both cores were slightly adjusted tominimize signal shifts due to age model differences.Adjustment was achieved by tuning to the bidecadalGISP2 δ18O record [39] (Fig. 4). The bidecadal GISP2record was chosen because stadial (S)/IS transitionsand in most cases also stadial minima are betterdefined in it than in the lower resolution record of [2].In addition, parts of the MIS 3 record of MD99-2339have a similar temporal resolution. For both sedimentcores, tuning was only performed in the sections older

than 23kyr while the younger 14C ages werecalibrated with INTCAL04 [37]. Re-calibration ofthe AMS 14C ages in core MD95-2043 [40] mainlyaffected the control points at 487, 512, 682, 708, and802cm where ages became about 200 yr older. Forthis core, the alkenone SST record was tuned toGISP2 by aligning the onsets of S/IS or IS/Stransitions and some stadial temperature minima(Fig. 4). This approach is valid as no offsets areobserved between the alkenone and planktonicforaminifera faunal data [40] and a high correlationcoefficient exists between the marine and the Green-land temperature records for the Holocene and the lastglacial/interglacial transition [41].

For core MD99-2339, the δ18O record of G.bulloides was chosen for tuning. The scarcity of G.bulloides in the Holocene and Termination 1B and localsalinity overprints during the deglaciation and the LastGlacial Maximum (LGM) hamper a good correlationbetween the δ18O of G. bulloides and the GISP2 recordin the younger 14C dated section. For the glacial part,

Fig. 4. Correlation of the secondary age models of MD99-2339 (B) and MD95-2043 (D) to the δ18O record (C) of GISP2 [39]. Gray dots in (B) and(D) mark the tuning points and dashed lines connect some of them with the GISP2 record. To highlight that the G. bulloides δ18O record of MD99-2339 is reflecting a more basin-wide signal in the tuned glacial part, the G. bulloides δ18O record of core MD95-2042 [35] (A) with its tuning points(circles) is shown on top. Arrows between (A) and (B) indicate offsets between the chronologies of the two cores, which in the section >24kyr arerelated to the fewer control points in MD95-2042. Numbers above the GISP2 δ18O record indicate some of the Dansgaard-Oeschger interstadials, H1–H 5 below Heinrich events 1–5.

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however, the δ18O record of core MD99-2339 shows agreat similarity to the one of core MD95-2042 ([35];Fig. 4) from an open ocean location off the southwesternPortuguese margin (Fig. 1). Thus, we assume that themillennial-scale oscillations in the G. bulloides record atsite MD99-2339 reflect conditions in the NorthAtlantic's subtropical gyre, in particular in the Portugaland Azores Currents. Because the subtropical gyre is asignificant moisture source for the precipitation aboveGreenland [42,43], changes in the δ18O records ofGISP2 and MD99-2339 should climatically be linked.Like for MD95-2043 onsets of S/IS and IS/S transitionsand some stadial temperature minima were chosen as tiepoints (Fig. 4B). Although MD95-2042 has fewer tiepoints to GISP2, the tuned chronologies of MD99-2339and MD95-2042 agree well. Clear age offsets are onlyvisible at the onset of the H 4 cold interval, whereMD95-2042 has poorer age control, and at the onset of

the H 1 δ18O minimum, which in core MD99-2339 is14C dated (Fig. 4).

For both, MD99-2339 and MD95-2043, the age–depth relationship of the secondary age model does notdeviate much from the one of the primary time scale(Fig. 3B,C), clearly indicating that the tuning does notbias the major deductions of this paper. For core MD99-2339, the age scales mainly differ around 40kyr BP (upto 420yr), i.e. at the base of H 4, where the tuned agemodel implies already higher sedimentation rates (Fig.3B). As IS 9 is so clearly defined in the G. bulloidesδ18O record, the difference is most probably linked tomore variable sedimentation rates in this contouriticsection (Fig. 5) than derived from the linear interpola-tion between two 14C ages. If the 14C age at 1250.5–1251.5cm (arrow in Fig. 3B) was not biased (seeprimary age scale), the 14C based stratigraphy mighthave indicated an earlier change in the sedimentation

Fig. 5. Paleoclimate records of core MD99-2339 from the Gulf of Cadiz (B–F; on primary age scale) in comparison to the GISP2 δ18O temperaturerecord (δ18Oice; A; [2]) over the last 48kyr. The planktonic δ18O record (δ18Opla; B) was measured inG. bulloides and the benthic one (δ18Obenth; E)in Cibicidoides and Uvigerina species (presented as Uvigerina equivalents). The epibenthic δ13C data (F) was derived from C. pachyderma and C.wuellerstorfi. Changes in the grain size and thereby in the strength of the Mediterranean Outflow are revealed by the mean grain size (gs) <63μm (C)and the magnetic susceptibility (D). Dashed lines connect some gs peaks with the GISP2 record. Numbers above the GISP2 δ18O record indicateDansgaard-Oeschger interstadials 1–13, H 1–H 5 below Heinrich event 1–5 (also marked by gray squares). Arrows on top show periods of MarineIsotope Stages (MIS) 2 and 3. Triangles in (F) mark positions of primary age control points with tuning points to GISP2 shown in gray.

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rate as well. In core MD95-2043, the two age scalesmainly differ between 30.5 and 31.5kyr BP (Fig. 3C).

For the comparison between cores MD99-2339 andMD99-2343 no adjustments were necessary as theoriginal age model of MD99-2343 is mainly based ontuning to the GISP2 chronology ([44]; Fig. 3A).

5. Results and discussion

5.1. Millennial-scale variability

The δ18O record of G. bulloides, a species absent orsparse in parts of the Holocene, has values ranging from

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−0.05‰ to 2.87‰ (Fig. 5B). The heaviest planktonicδ18O values coincide with a precursor event to Heinrichevent (H) 2 at 25.6kyr BP and with the beginning of H1, previously denoted as Last Isotopic Maximum (LIM)[45]. In this paper, the H 1 interval is equal to theGreenland Stadial GS-2 [46] and encompasses the H 1ice-rafting events sensu strictu and the remaining coldperiod prior to the Bølling interstadial. The 25.6kyr BPcold event can be traced along the western Iberianmargin (e.g., MD95-2042; Fig. 4A), but is–in contrast toH 2 (J. Schönfeld, unpublished data)–not associatedwith IRD deposition.

Benthic δ18O values vary between 1.9‰ and 4.4‰and reveal a glacial–interglacial shift of −1.6‰ from18.6 to 10.2cal kyr BP (Fig. 5E). This shift exceeds thesea level rise related value by −0.8‰ and thereforeindicates a Holocene MOW that is either about 3°Cwarmer or fresher than the glacial MOW. While largeδ18O excursions are seen in the late Holocene, glacialbenthic δ18O oscillations were mostly small (≤0.3‰).This pattern is similar to the SO75-26KL record [47]further down the MOW stream but is quite differentfrom the pronounced stadial–interstadial cyclicity in theWMDW records [44,48]. The small oxygen isotopevariability in the Atlantic cores is related to the long-term stable temperature and salinity properties of theNACW entrained into the MOW and possibly also toconditions in the LIW, for which, however, a continuousrecord from the western Mediterranean Sea does not yetexist. The benthic δ13C record, on the other hand,reveals Dansgaard-Oeschger- and Bond-cycle typevariability with values ranging from 0.3‰ to 1.5‰(Fig. 5F). In particular, lighter values are related to theHeinrich events and to the early Holocene, whenSapropel 1 was formed in the Mediterranean Sea [49].

The mean grain size of the silt and clay fractions(<63μm) varies between 6 and 23.5μm and shows aclear millennial-scale variability with higher amplitudesof change during MIS 3 and the deglaciation (Fig. 5C).Grain size variations indicate changes in the strength ofthe prevailing bottom current. Magnetic susceptibility,on the other hand, reveals changes in the mineralogicalcomposition, namely the concentration of ferromagneticgrains in the sediment. At site MD99-2339, themagnetic susceptibility data reflects the grain sizerecord, even in the fine structures, indicating thatstronger bottom currents led to increased deposition offerromagnetic and therefore lithic grains. Thus, maximain both records evidence a contourite and thereby anincreased flow strength of the lower MOW core. Thelowest grain size values are recorded during the earlyHolocene. They coincide with the condensed core

section and the timing of Sapropel 1 [49] and indicatea sudden reduction of both flow strength and sedimentdelivery by the lower MOW core to site MD99-2339.Lower magnetic susceptibility values between 16 and17kyr BP are also seen in the shipboard data, but theextreme values in the point sensor data (Fig. 5D) couldbe an artefact, as the thinner and variable core diameterin this section might have prevented precisemeasurements.

Between 26 and ∼18kyr, mean grain size values atsite MD99-2339 are in the range of those fromContourite I [50,51] (2.1 to 4.4kyr BP; Fig. 5D), butlower than during most of the MIS 3 stadials, the late H1 contourite and the Younger Dryas. Increased currentstrength during the LGM is seen in records from thenearby Gil Eanes Drift [52] and from the westernPortuguese margin [53], which indicate a strong, butdeeper (∼2000m) lower MOW core. So even with arestricted exchange through the Strait of Gibraltar due tothe progressive glacial sea level lowering [54,55],Mediterranean water flowed into the Gulf of Cadiz.

Except for the late Holocene, the contourites wereformed at times when temperatures above Greenlandwere cold, that is they are contemporary with the DOstadials, part of the Heinrich phases and the YoungerDryas (Fig. 5). Even less profound cooling events likethe Older Dryas (∼14kyr BP) and the cooling stepswithin DO IS 8 and 12, which fit with the secondary1.47kyr cycle in these long-term interstadials [56,57](Fig. 6), are associated with a flow strengthening. Thecoupling between lower MOW intensification and coldclimatic conditions in the North Atlantic regionindicates that MOW intensity variations are linked tothe climatic conditions during the DO cycles and mighteven play a role in the driving mechanism(s) ofmillennial-scale climate change as implied by [9].

5.2. Benthic δ13C signals and contour current sourcewaters

To verify if MOW is really the water mass formingthe contourites in the Gulf of Cadiz we compared thebenthic δ13C and grain size records of core MD99-2339with the benthic δ13C variations at sites MD95-2043 andMD99-2343 in the deep western Mediterranean Sea(Fig. 7). Heavier benthic δ13C values in the Mediterra-nean Sea are more related to ventilation than to nutrientlevels and therefore indicate a vigorous deep convectionin the Gulf of Lions [44] during cold climate periods.Interstadials, on the other hand, are marked by lowerδ13C values indicating a lesser ventilated WMDW[44,48]. Site MD99-2339 shows similar δ13C

Fig. 6. Comparison of the mean grain size record of MD99-2339 (on its primary age scale) with the trapezoidal 1.47kyr pacing model for the DOcycles in GISP 2 [56], which resembles the filter of the 1.47kyr periodicity [57]. Numbers on top indicate DO interstadials (IS) 1 and 4–13. For IS8 and 12 the number is placed above the first cycle, which coincides with the onset of the warming above Greenland, while the letters a and b indicatethe secondary cycle, respectively.

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oscillations (Fig. 7). Changes at both sides of the Straitof Gibraltar occurred in synchrony with better ventila-tion during stadials and poorer ventilation duringinterstadials (Fig. 7B). So, the benthic δ13C signal atsite MD99-2339 mainly follows the Mediterraneanpattern, which is opposite to the North Atlantic onewith a better ventilation during interstadials [4,5].Unfortunately, a comparison for the period youngerthan 11kyr is not possible as both Mediterranean siteslack epibenthic foraminifera in that interval [44,48]. Thegreat similarity between the MD99-2339 and MD95-2043 δ13C records in terms of both absolute values andpatterns (Fig. 7B) suggests that WMDW contributed asubstantial part to the water leaving the MediterraneanSea and generating the contourites in the Gulf of Cadiz.This is further supported by the synchroneity betweenvigorous deep convection in the Gulf of Lions (highδ13C in MD99-2343) and contourite formation (grainsize maxima) at site MD99-2339 (Fig. 7D). Nowadays,WMDW contributes only 0.2Sv to the MOW [15], butbuoyancy loss due to wintertime cooling can reduce thedensity gradient between the intermediate and deepwaters of the Mediterranean Sea and lead to an increasedoutflow of WMDW [16]. So it seems that the colderclimatic conditions during the glacial also led to a higherbuoyancy loss and consequently to an increased exportof WMDW.

During most of the glacial period, a positive benthicδ13C gradient (δ13CMD95-2043–δ13CMD99-2339) of0.3‰ to 0.8‰ existed across the Strait of Gibraltar (Fig.7C). Lower or even short intervals of negative gradientsare associated with MIS 3 interstadials. For most of thecold periods, including Heinrich events, stadials and theLGM, comparable gradients are observed. The contrarysignal during stadials S6 and S7 is due to local effects in

the Alboran Sea, as WMDW δ13C values at site MD95-2043 are much lower than those at site MD99-2343 (Fig.7). Lower δ13C values in the MOW are probably morerelated to the ventilation status of and nutrient content inthe entrained NACW than to LIW contributions to thedeeper outflow, as LIW, with the exception of early IS8 when a sapropel was formed in the easternMediterranean Sea [58], should also have been wellventilated. Especially during MIS 2, NACW nutrientlevels might have played a role, as nutrients could havebeen advected from the upwelled waters off Portugaland Morocco, both of which experienced an increase inupwelling related productivity [59], or from AntarcticIntermediate Water that penetrated at least as far north as30°N [60]. The 13C depletion between 26.5 and 13kyron the Atlantic side of the Strait of Gibraltar is also seenin the upper MOW core (site MD99-2341; J. Schönfeld,unpublished data) and can be traced further into theeastern North Atlantic [61].

Records from site MD99-2339 show a complexpattern for the Heinrich events, which includes a shortphase of intense δ13C depletion coincident with lowgrain size values (Figs. 7 and 8). This pattern isconsistent with a brief weakening of WMDW formationin the Gulf of Lions as evidenced by the westernMediterranean Sea records [44,48] (Fig. 7). During H 1,2 and 4, minimum δ13C values at site MD99-2339 areequal to those observed at site SO75-26KL [47], forwhich [47,60] postulated a stronger influence of Atlanticwaters during Heinrich events. The H 1 δ13C minimumin core MD99-2339 is 0.2‰ heavier than at site V29-204, which was retrieved from 1849m water depth fromthe western Reykjanes ridge and which reflects condi-tions in the glacial Upper North Atlantic Deep Water(UNADW) [4], and 0.2‰ lighter than the UNADW/

Fig. 7. Comparison between the Mediterranean Sea benthic δ13C records (black; B: MD95-2043 [48]; D: MD99-2343 [44]) with the benthic δ13C (B;magenta) and the mean grain size (D; red) record of core MD99-2339. MD99-2339 and MD95-2043 data is shown on their respective secondary agescale, MD99-2343 on its original time scale [44]. As epibenthic foraminifera are absent in the Mediterranean Sea during the Holocene, data can onlybe compared for the interval older than 11kyr BP. Record (C) indicates the gradient between the benthic δ13C values of MD95-2043 andMD99-2339;calculated for all levels that are contemporary within a ±35-year margin. The GISP2 δ18O record [2] (A) is shown for comparison with some DOinterstadials and Heinrich events (H) 1–5 and some stadials (S) indicated above and below. Nomenclature for stadials follows [46,80].

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Labrador Sea Water signal at the shallower site SO82-5(1416m) [6,62] from the same region. During H 2,MD99-2339 δ13C levels converge with those at the twoUNADW sites, but during early H 4 only with V29-204.So conform with the intermediate water dynamicsproposed for sites above 1500m water depth on thewestern Portuguese margin [47,53,60], the δ13C evi-dence would indicate that UNADW replaced MOW atsite MD99-2339 during Heinrich events. Mg/Ca-basedbottom water temperatures for core MD99-2339 [63],however, only reach minima of 7–8°C during H 1, 2 and3, which excludes pure UNADW as prevailing watermass, but could indicate that a mixture of UNADW andMOW bathed site MD99-2339 during the times of low

WMDW production. The entrainment of progressivelyless ventilated North Atlantic waters [64,65] would alsoexplain the gradual δ13C depletion observed at siteMD99-2339 (Figs. 7 and 8). The ends of H 1, 4 and 5 aremarked by higher δ13C values (Fig. 8), probablyinfluenced by both the well ventilated WMDW and arecovery in NACW ventilation.

The Younger Dryas contourite (Peak II; [22]) is seenin many sediment cores from the Gulf of Cadiz (e.g.,[50,51]), which are either in the pathway of the lower orupper MOW core. The comparably low δ13C values inthe WMDW (Fig. 7; [44,48]) imply a decreased deepconvection in the Gulf of Lions, most probably causedby a reduced density gradient between the upper and

Fig. 8. Close-up studies of the intervals around Heinrich (H) events 1 (A), 4 (B) and 5 (C) in core MD99-2339 on its primary age scale. Shown fromtop to bottom are the δ18O record of G. bulloides, planktonic foraminifera assemblage based sea surface temperatures (SST), the abundance of ice-rafted debris (IRD) in the fraction >250µm, the benthic δ13C data, and the mean grain size <63µm. In (C), the δ18O record of G. ruberwhite (gray) isshown in addition to the one ofG. bulloides as it better displays the transition from H 5 to IS 12. For H 5, the IRD record is negligible as the number ofgrains is within the noise level. Note that in (B) the boundaries of the H 4 bar are defined by the IS 9 and 8 δ18O peaks and that the length of H 4therefore deviates from the one shown for the secondary age scale in Fig. 7. IS refers to DO interstadial, with IS 1 being equal to the Bølling, and S toDO stadial. LIM: Last Isotopic Maximum; OD: Older Dryas.

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lower layer in the Mediterranean basin along with thepostglacial sea level rise [54,55]. So, the outflow fromthe Mediterranean Sea must have been mainly LIW,which fed both MOW cores. As models [54,66] showthat the salt in the deeper layer of the Mediterranean Seahas a residence time in the order of centuries tomillennia, the LIW was salty enough to produce adense MOW in the Gulf of Cadiz.

5.3. Case studies of H 1, H 4 and H 5: surface oceancontrol on MOW intensification

Variations in MOW flow strength were also coupledto hydrographic changes in the North Atlantic, as will bediscussed in detail for the intervals around H 1, H 4 and

H 5 (Fig. 8). The raised MOW intensity during the LGMpersisted until the onset of the LIM at 17.5kyr BP. Thisevent marks the onset of the H 1 interval at the westernIberian margin [45] and is displayed by heavyplanktonic δ18O values in core MD99-2339 (Figs. 5and 8A). SST around 19°C prevailed at the beginning ofthe LIM (until 17.2kyr BP), so that surface watersalinity must have been high. This may have reduced thedensity gradient between Atlantic inflow and Mediter-ranean outflow and decreased the transport rate of theMOW. The maximum current intensity as displayed bythe mean grain size did not recover for more than1000yr at site MD99-2339. During this interval,however, surface water conditions changed drasticallywith an incursion of subpolar waters (17–16.8kyr BP), a

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reappearance of temperate to subtropical waters, andonce again a re-advance of subpolar waters into theeastern Gulf of Cadiz at 16.1kyr BP. The surface waterswere even fresher this time, as indicated by lighter δ18Ovalues of G. bulloides and the presence of IRD, bothderived from melting icebergs (Fig. 8A). While siteMD99-2339 is the easternmost site where IRD has beenfound, subpolar waters penetrated further into theMediterranean Sea [40,44,67,68]. Site MD99-2339recorded, however, only a later stage of the H 1 icebergsurges as the IRD peak lags the onset of IRD depositionin the North Atlantic and at the northwestern Portuguesemargin by 500–700yr [45,69]. Nonetheless, the firstincursion of subpolar waters into the Gulf of Cadizcorresponds with the onset of IRD sedimentation at thewestern Iberian margin. The IRD maximum is coevalwith the benthic δ13C minimum during H 1 and the endof the low current strength interval (Fig. 8A). With thecessation of IRD deposition and rising SST after16.2kyr BP, indicating the renewed northward extensionof the subtropical gyre, a deep contour current existedonce more in the eastern Gulf of Cadiz. The secondgrain size maximum coincides with the end of the abruptSST rise and a better ventilation of the deep water (Figs.5 and 8A). The strengthening of the MOW and thewarming still occurred within the H 1/GS-2 cold interval(Figs. 5 and 8A). The warming is more pronounced inthe Gulf of Cadiz than in the Mediterranean Sea [41].However, a warmer SST does not increase a water mass'density. The most likely reason for the re-invigoration ofthe MOW and contourite formation just at that time andnot earlier was salt. An excess of salt must haveaccumulated in the subtropical gyre during the south-ward penetration of the polar front. Deep-sea recordsshow indeed that the glacial subtropical gyre was muchsaltier than the subpolar one [70] and that off NW-Africasalinities spiked during most Heinrich events [71]. Thesalt was fed into the Mediterranean by the Atlanticinflow, concentrated even more in the LIWand WMDWdue to the dry and windy conditions around theMediterranean Sea (e.g., [3,72-75]), and was finallyentrained into the MOW. In the model of [9] theadvection of MOW leads to increased upper oceansalinity in the North Atlantic [76], but the changingfluxes in the North Atlantic's meridional salinity cycleare difficult to estimate. It is obvious, however, that thisexcess salt would not be available anymore when theAtlantic's THC switched back into the interstadial modeat the beginning of the Bølling interstadial, exactly thetime, when the MOW's strength ceases again(∼14.7kyr BP). Furthermore, the Mediterranean cli-mate turned to warmer and more humid conditions,

which mutually promoted a reduction in the intermedi-ate and deep water formation [44,48].

During H 4 (Fig. 8B), the MOW and surface waterrecords show the same pattern as during H 1: anincreased flow strength and warm SST at the beginningof the cold interval, a current strength reduction whenicebergs and subpolar water masses protruded into theGulf of Cadiz and the western Mediterranean Sea, and are-invigoration of the MOW once the IRD depositionceased. Also, similar rapid shifts in the latitudinalposition of the boundary between subtropical andsubpolar water masses are revealed by a warmingsandwiched in between two short incursions of subpolarwaters (Fig. 8B). The intermittent warming seems tooriginate from the Gulf Stream [77] and was transportedinto the Gulf of Cadiz by the paleo-Azores Current.During H 4 the icebergs arrived at site MD99-2339 withthe first southward advance of the polar front and theassociated meltwater intruded into the western Mediter-ranean Sea [40,44,67,68] where it caused a cessation ofdeep convection in the Gulf of Lions ([44]; Fig. 7D).The SST record of core MD99-2339 suggests that thefreshwater transporting Atlantic current flowed–liketoday–close to the coast and only temporarily affectedsite MD99-2339, in agreement with the scenariodiscussed in [44]. High percentages of G. bulloidesand T. quinqueloba during H 4 (A. Voelker, unpublisheddata) even imply that like during the LGM [78] theAzores front penetrated into the Gulf of Cadiz and waspositioned close to site MD99-2339. The grain sizemaximum at the end of H 4 coincides directly with thetransition to IS 8 and the time when WMDW formationin the Gulf of Lions had resumed again (Fig. 7D). Alsoduring H 5 (Fig. 8C), the coldest SST, indicating theadvance of fresh subpolar waters, are coeval withreduced MOW intensity, a reduced ventilation of thewater mass bathing site MD99-2339 and the cessation ofWMDW formation ([44]; Fig. 7D). Just prior to 46kyrBP, temperate to subtropical surface waters advancedinto the eastern Gulf of Cadiz and MOW intensityincreased, in phase with rising SST (Fig. 4; [40]) andresumed WMDW formation in the western Mediterra-nean Sea [44]. Similar to the onset of H 4, the beginningof H 5 is also marked by strong SST oscillations at siteMD99-2339. Some of these oscillations and thewarming after 46kyr also affected the western Portu-guese margin where they were recorded at site MD95-2040 (40.58°N, 9.86°W) [79].

The repeated patterns observed for the 3 Heinrichevents hint to the hydrographic conditions in thesubtropical gyre as one of the controlling factor on theMOW's strength. The SST record reveals the presence

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of subtropical waters also when contourites were formedduring DO stadials 10 and 8 (Fig. 8B) and during thecooling step within DO IS 8 (∼36.5kyr BP). Asmentioned above, the subtropical gyre was more salineduring the glacial and, although it has so far only beenshown for Heinrich events [71], salinity most probablyalso spiked during the other stadial phases. So, when thepolar front was positioned north off the Strait ofGibraltar, salt enriched subtropical waters, probablyassociated with an eastward extension of the Azoresfront [78], could advance into the Gulf of Cadiz andsubsequently into the western Mediterranean Sea. Herethey increased the surface water density and togetherwith the prevailing atmospheric conditions around theMediterranean facilitated deep convection in the Gulf ofLions. The importance of the source for the inflowingsurface water is indicated by the MD99-2343 records[44]. During Heinrich events climatic conditions on thesurrounding continents did not change, so that ceasing/cessation of WMDW production clearly depended onthe prevailing surface water, i.e. on the advance ofsubpolar/subtropical surface waters into the Mediterra-nean Sea.

6. Conclusions

The records from site MD99-2339 show that theMOW underwent large changes during the last glacial.With the intensified MOW a significant amount of heatand salt was added into the intermediate depth NorthAtlantic during times when the THC was in stadial orHeinrich mode. This is also true if the contouritesformed at site MD99-2339 are caused by a displacementof the lower MOW's flow path rather than a realintensification of the lower core. In the only climatemodel that includes the MOW in the glacial THC [9],heat and salt from the MOW help to graduallyprecondition the North Atlantic until strong, interstadialoverturning resumes. In the model two short-lived, butlarge increases in the MOW transport volume follow thefreshwater perturbation in the North Atlantic by 700 and900yr and precede the onset of strong meridionaloverturning by 300yr [9]. The MOW transport increaseat 700yr fits in timing with the MD99-2339 grain sizerecord. During MIS 3, variations in the grain size dataclosely follow the stadial/interstadial cycles of theGISP2 δ18O record (Fig. 6), which have a periodicityof 1.48 to 1.47yr [2,57]. As vigorous WMDWformation and MOW strengthening is generally coevalwith the stadial half of the cycle (Figs. 5 and 7), MOWvolume increase follows theoretically about 740yr aftera meltwater perturbation in the North Atlantic. Howev-

er, contourite formation at site of MD99-2339 lasted formore than 500yr and directly preceded the interstadialoverturning mode in the Atlantic ocean [5]. So the timeof contourite formation would encompass both peaks inMOW transport volume and the delay time to the THCswitch of the model [9]. As the data implies a continuoustransport of salt and heat into the glacial North Atlanticduring stadials, the time needed to reach a threshold forthe THC's switch was longer either because less salt andheat was exported than injected by the short-termedincreases in the model or the advection into the surfacewaters of the North Atlantic's convection areas [76]took longer. In any case, even if the MOW signal mightturn out to be not strong enough as the ultimate triggerfor the THC's mode switch, the export of salt must haveincreased the density in the North Atlantic's subsurfacewaters thereby bringing the THC one step closer to thethreshold needed for the interstadial overturning toresume.

Acknowledgements

We thank J.C. Faugeres and T. Mulder for thegenerous access to core MD99-2339, Yvon Balut, IPEVand the shipboard party for their support and expertiseduring the IMAGES-GINNA cruise, and the laboratoryteam of the Departamento Geologia Marinha for theirefficient work. A.V., S.L. and F.A. were funded by theFundação para a Ciência e a Tecnologia (FCT) projectMOWFADRI. High-resolution magnetic susceptibilitymeasurements were made possible by an EC Access toResearch Infrastructure-Paleostudies grant to S.L. andA.V. J. S. was supported by the Deutsche Forschungs-gemeinschaft (Grants Za157/16-1 and Ti240/9). I.C.acknowledges a fellowship from the Comer AbruptClimate Change Foundation (USA). The comments ofthe two reviewers significantly benefited this paper.

References

[1] W. Dansgaard, S.J. Johnsen, H.B. Clausen, D. Dahl-Jensen, N.S.Gundestrup, C.U. Hammer, C.S. Hvidberg, J.P. Steffenson, A.E.Sveinbjörnsdottir, J. Jouzel, G. Bond, Evidence for generalinstability of past climate from a 250-kyr ice-core record, Nature364 (1993) 218–220.

[2] P.M. Grootes, M. Stuiver, 180/160 variability in Greenland snowand ice with 10−3 to 105 year time resolution, J. Geophys. Res.102 (1997) 26455–426470.

[3] workshop participantsA.H.L. Voelker, Global distribution ofcentennial-scale records for marine isotope stage (mis) 3: adatabase, Quat. Sci. Rev. 21 (2002) 1185–1214.

[4] W.B. Curry, T.M. Marchitto, J.F. McManus, D.W. Oppo, K.L.Laarkamp, Millennial-scale changes in ventilation of thethermocline, intermediate, and deep waters of the glacial North

53A.H.L. Voelker et al. / Earth and Planetary Science Letters 245 (2006) 39–55

Atlantic, in: P.U. Clark, R.S. Webb, L.D. Keigwin (Eds.),Mechanisms of Global Climate Change At Millennial TimeScales, . Geophys. Monogr., 112, AGU,Washington, D.C., 1999,pp. 59–76.

[5] M. Elliot, L. Labeyrie, J.-C. Duplessy, Changes in North Atlanticdeep-water formation associated with the Dansgaard-Oeschgertemperature oscillations (60-10ka), Quat. Sci. Rev. 21 (2002)1153–1165.

[6] S. van Kreveld, M. Sarnthein, H. Erlenkeuser, P.M. Grootes, S.Jung, M.-J. Nadeau, U. Pflaumann, A. Voelker, Potential linksbetween surging ice sheets, circulation changes, and theDansgaard-Oeschger cycles in the Irminger Sea, 60-18kyr,Paleoceanography 15 (2000) 425–442.

[7] A. Ganopolski, S. Rahmstorf, Rapid changes of glacial climatesimulated in a coupled climate model, Nature 409 (2001)153–158.

[8] R. Knutti, J. Flückiger, T.F. Stocker, A. Timmermann,Strong hemispheric coupling of glacial climate throughfreshwater discharge and ocean circulation, Nature 430(2004) 851–856.

[9] G.R. Bigg, M.R. Wadley, Millennial-scale variability in theoceans: an ocean modelling view, J. Quat. Sci. 16 (2001)309–319.

[10] J.P. Kennett, K.G. Cannariato, I.L. Hendy, R.J. Behl, Carbonisotopic evidence for methane hydrate instability during thequaternary interstadials, Science 288 (2000) 128–133.

[11] K. Pahnke, R. Zahn, Southern hemisphere water mass conversionlinked with North Atlantic climate variability, Science 307 (2005)1741–1746.

[12] M.A. Altabet, M.J. Higginson, D.W. Murray, The effect ofmillennial-scale changes in Arabian Sea denitrification onatmospheric CO2, Nature 415 (2002) 159–162.

[13] T.M. Marchitto, W.B. Curry, D.W. Oppo, Millennial-scalechanges in North Atlantic circulation since the last glaciation,Nature 393 (1998) 557–561.

[14] T.H. Kinder, G. Parrilla, Yes, some of the mediterranean outflowdoes come from great depth, J. Geophys. Res-Oceans. 92 (1987)2901–2906.

[15] M. Tomczak, J.S. Godfrey, Regional Oceanography: AnIntroduction, Pergamon, London, 1994, 422 pp.

[16] H.L. Bryden, H.M. Stommel, Limiting processes that determinebasic features of the circulation in the Mediterranean Sea,Oceanol. Acta 7 (1984) 289–296.

[17] J. Candela, Mediterranean water and global circulation, in: G.Siedler, J. Church, J. Gould (Eds.), Ocean Circulation andClimate—Observing and Modelling the Global Ocean, Int.Geophys. Ser., 77Academic Press, San Diego, London, 2001, pp.419–429.

[18] M. Astraldi, G.P. Gasparini, A. Vetrano, S. Vignudelli,Hydrographic characteristics and interannual variability ofwater masses in the central mediterranean: a sensitivity test forlong-term changes in the Mediterranean Sea, Deep-Sea Res., PartI, Oceanogr. Res. Pap. 49 (2002) 661–680.

[19] I. Ambar, M.R. Howe, Observations of the mediterraneanoutflow: I. Mixing in the mediterranean outflow, Deep-SeaRes. 26A (1979) 535–568.

[20] M.O. Baringer, J.F. Price, Mixing and spreading of themediterranean outflow, J. Phys. Oceanogr. 27 (1997)1654–1677.

[21] I. Ambar, N. Serra, M.J. Brogueira, G. Cabecadas, F. Abrantes, P.Freitas, C. Goncalves, N. Gonzalez, Physical, chemical andsedimentological aspects of the mediterranean outflow off iberia,

Deep-Sea Res., Part II, Top. Stud. Oceanogr. 49 (2002)4163–4177.

[22] J. Faugères, E. Gonthier, D. Stow, Contourite drift molded bydeep mediterranean outflow, Geology 12 (1984) 296–300.

[23] E.L. Habgood, N.H. Kenyon, D.G. Masson, A. Akhmetzhanov,P.P.E. Weaver, J. Gardner, M. T., Deep-water sediment wavefields, bottom current sand channels and gravity flow channel-lobe systems: Gulf of Cadiz, NE Atlantic, Sedimentology 50(2003) 483–510.

[24] M.C. Iorga, M.S. Lozier, Signatures of the Mediterraneanoutflow from a North Atlantic climatology: 1. Salinity anddensity fields, J. Geophys. Res-Oceans. 104 (1999)25985–926009.

[25] M.S. McCartney, L.D. Talley, The sub-polar mode water of thenorth-atlantic ocean, J. Phys. Oceanogr. 12 (1982) 1169–1188.

[26] R. Haynes, E.D. Barton, A poleward flow along the atlantic coastof the Iberian Peninsula, J. Geophys. Res-Oceans. 95 (1990)11425–11441.

[27] R.A. Locarnini, T.D. O'Brien, H.E. Garcia, J.I. Antonov, T.P.Boyer, M.E. Conkright, C. Stephens, World ocean atlas 2001,volume 3: oxygen, in: S. Levitus (Ed.), Noaa Atlas Nesdis 51, U.S. Government Printing Office, Washington, D.C., 2002, p. 286,CD-ROMs.

[28] A. Peliz, J. Dubert, A.M.P. Santos, P.B. Oliveira, B. Le Cann,Winter upper ocean circulation in the western iberian basin—fronts, eddies and poleward flows: an overview, Deep-Sea Res.,Part I, Oceanogr. Res. Pap. 52 (2005) 621.

[29] L. Labeyrie, E. Jansen, E. Cortijo, MD 114/IMAGES V, á bordduMarion Dufresne, Fort de France, 11 juin 1999—Marseille, 20septembre 1999, IPEV Ref: OCE/2003/02 2, Institut PolaireFrancais Paul-Emile Victor, 2003, pp. 381–849.

[30] F. Bassinot, L. Labeyrie, IMAGES MD 101 A coring cruise ofthe R/V Marion Dufresne in the North Atlantic and NorwegianSea, Institut Francais pour la Recherche et la TechnologiePolaires, Plouzane, 1996, p. 217.

[31] N.J. Shackleton, Attainment of isotopic equilibrium betweenocean water and the benthonic foraminifera genus uvigerina:isotopic changes in the ocean during the last glacial, Colloq. Int.Centre Natl. Rech. Sci. 219 (1974) 203–209.

[32] U. Pflaumann, J. Duprat, C. Pujol, L.D. Labeyrie, Simmax: amodern analog technique to deduce atlantic sea surfacetemperatures from planktonic foraminifera in deep-sea sedi-ments, Paleoceanography 11 (1996) 15–36.

[33] M. Kucera, M. Weinelt, T. Kiefer, U. Pflaumann, A. Hayes, M.Weinelt, M.T. Chen, A.C. Mix, T.T. Barrows, E. Cortijo, J.Duprat, S. Juggins, C. Waelbroeck, Reconstruction of sea-surfacetemperatures from assemblages of planktonic foraminifera:multi-technique approach based on geographically constrainedcalibration data sets and its application to glacial atlantic andpacific oceans, Quat. Sci. Rev. 24 (2005) 951–998.

[34] E. Salgueiro, A. Voelker, F. Abrantes, H. Meggers, U.Pflaumann, N. Loncaric, R. Gonzalez-Alvarez, P. Oliveira, H.B. Bartels-Jónsdóttir, J. Moreno, G. Wefer, Planktonic forami-nifera from modern sediments reflect upwelling patterns offIberia. Insights from a regional transfer function, Mar. Micro-paleontol. (in preparation).

[35] N.J. Shackleton, M.A. Hall, E. Vincent, Phase relationshipsbetween millennial-scale events 64,000–24,000 years ago,Paleoceanography 15 (2000) 565–569.

[36] L. Loewemark, F. Werner, Dating errors in high-resolutionstratigraphy: a detailed x-ray radiograph and ams 14c study ofzoophycus burrows, Mar. Geol. 177 (2001) 191–198.

54 A.H.L. Voelker et al. / Earth and Planetary Science Letters 245 (2006) 39–55

[37] P.J. Reimer, M.G.L. Baillie, E. Bard, A. Bayliss, J.W. Beck, C.J.H. Bertrand, P.G. Blackwell, C.E. Buck, G.S. Burr, K.B. Cutler,P.E. Damon, R.L. Edwards, R.G. Fairbanks, M. Friedrich, T.P.Guilderson, A.G. Hogg, K.A. Hughen, B. Kromer, G. McCor-mac, S. Manning, C.B. Ramsey, R.W. Reimer, S. Remmele, J.R.Southon, M. Stuiver, S. Talamo, F.W. Taylor, J. van der Plicht, C.E. Weyhenmeyer, INTCAL04 Terrestrial Radiocarbon Agecalibration, 0–26cal kyr BP, Radiocarbon 46 (2004) 1029–1058.

[38] A.H.L. Voelker, P.M. Grootes, M.-J. Nadeau, M. Sarnthein, 14Clevels in the Iceland Sea from 25–53kyr and their link to theearth’s magnetic field intensity, Radiocarbon 42 (2000) 437–452.

[39] M. Stuiver, P.M. Grootes, Gisp2 oxygen isotope ratios, Quat.Res. 53 (2000) 277–284.

[40] I. Cacho, J.O. Grimalt, C. Pelejero, M. Canals, F.J. Sierro, J.A.Flores, N.J. Shackleton, Dansgaard-Oeschger and Heinrichevents imprints in Alboran Sea paleotemperatures, Paleoceano-graphy 14 (1999) 698–705.

[41] I. Cacho, J.O. Grimalt, M. Canals, L. Sbaffi, N.J. Shackleton, J.Schoenfeld, R. Zahn, Variability of the western MediterraneanSea surface temperature during the last 25,000 years and itsconnection with the northern hemisphere climatic changes,Paleoceanography 16 (2001) 40–52.

[42] C.D. Charles, D. Rind, J. Jouzel, R.D. Koster, R.G. Fairbanks,Glacial–interglacial changes in moisture sources for greenland:influences on the ice core record of climate, Science 263 (1994)508–511.

[43] S.J. Johnsen, W. Dansgaard, J.W.C. White, The origin of arcticprecipitation under present and glacial conditions, Tellus 41B(1989) 452–468.

[44] F.J. Sierro, D.A. Hodell, J.H. Curtis, J.A. Flores, I. Reguera, E.Colmenero-Hidalgo, M.A. Bárcena, J.O. Grimalt, I. Cacho, J.Frigola, M. Canals, Impact of iceberg melting on mediterraneanthermohaline circulation during heinrich events, Paleoceanogra-phy 20 (2005) 1–13.

[45] J. Schönfeld, R. Zahn, L. de Abreu, Surface and deep waterresponse to rapid climate changes at the western iberian margin,Glob. Planet. Change (2003), doi:10.1016/S0921-8181(1002)000197-000192.

[46] J. Lowe, W.Z. Hoek, Inter-regional correlation of palaeoclimaticrecords for the last glacial–interglacial transition: a protocol forimproved precision recommended by the intimate project group,Quat. Sci. Rev. 20 (2001) 1175–1187.

[47] R. Zahn, J. Schönfeld, H. Kudrass, M. Park, H. Erlenkeuser, P.Grootes, Thermohaline instability in the North Atlantic duringmeltwater events: stable isotope and ice-rafted detritus recordsfrom core so75-26kl, Portuguese margin, Paleoceanography 12(1997) 696–710.

[48] I. Cacho, J.O. Grimalt, F.J. Sierro, N. Shackleton, M. Canals,Evidence for enhanced mediterranean thermohaline circulationduring rapid climatic coolings, Earth Planet. Sci. Lett. 183 (2000)417–429.

[49] D. Mercone, J. Thomson, I.W. Croudace, G. Siani, M. Paterne, S.Troelstra, Duration of s1, the most recent sapropel in the easternMediterranean Sea, as indicated by accelerator mass spectrom-etry radiocarbon and geochemical evidence, Paleoceanography15 (2000) 336–347.

[50] C. Vergnaud-Grazzini, M. Caralp, J.-C. Faugères, E. Gonthier, F.Grousset, C. Pujol, J.-F. Saliége, Mediterranean outflow throughthe Strait of Gibraltar since 18 000 years BP, Oceanol. Acta 12(1989) 305–324.

[51] T. Mulder, P. Lecroart, M. Voisset, J. Schönfeld, E. Le Drezen, E.Gonthier, V. Hanquiez, R. Zahn, J.-C. Faugères, E.J. Hernandez-

Molina, scientific party of Cadisar cruise, The Gulf of Cadiz. Akey area for understanding paleoclimate record and oceaniccirculation, EOS Trans. AGU 83 (2002) 481–488.

[52] M. Rogerson, E.J. Rohling, P.P.E. Weaver, J.W. Murray, Glacialto interglacial changes in the settling depth of the mediterraneanoutflow plume, Paleoceanography 20 (2005) 1–12.

[53] J. Schönfeld, R. Zahn, Late glacial to holocene history of themediterranean outflow. Evidence from benthic foraminiferalassemblages and stable isotopes at the portuguese margin,Paleogeography, Palaeoclimatology, Palaeoecology 159 (2000)85–111.

[54] S. Matthiesen, K. Haines, A hydraulic box model study of themediterranean response to postglacial sea-level rise, Paleoceano-graphy 18 (2003 (10-21)) 8-1–8-12.

[55] E.J. Rohling, H.L. Bryden, Estimating past changes in the easternmediterranean fresh-water budget, using reconstructions of sea-level and hydrography, Proceedings of the Koninklijke Neder-landse Akademie Van Wetenschappen-Biological ChemicalGeological Physical and Medical Sciences 97 (1994) 201–217.

[56] M. Schulz, On the 1470-year pacing of Dansgaard-Oeschgerwarm events, Paleoceanography 17 (2002 (05-01)) 4-1–4-9.

[57] M. Schulz, W.H. Berger, M. Sarnthein, P.M. Grootes,Amplitude variations of 1470-year climate oscillations duringthe last 100,000 years linked to fluctuations of continental icemass, Geophys. Res. Lett. 26 (1999) 3385–3388, doi:10.1029/1999GL006069.

[58] M. Geraga, S. Tsaila-Monopolis, C. Ioakim, G. Papatheodorou,G. Ferentinos, Short-term climate changes in the southernAegean Sea over the last 48,000 years, Palaeogeography,Palaeoclimatology, Palaeoecology 220 (2005) 311.

[59] F. Abrantes, Paleoproductivity oscillations during the last 130 kaalong the Portuguese and NW African margins, UpwellingSystems: Evolution Since the Early Miocene, Spec. Publ. —Geol. Soc., 64, 1992, pp. 499–510.

[60] C. Willamowski, R. Zahn, Upper ocean circulation in the glacialNorth Atlantic from benthic foraminiferal isotope and traceelement fingerprinting, Paleoceanography 15 (2000) 515–527.

[61] M. Sarnthein, K. Winn, S. Jung, J. Duplessy, L. Labeyrie, H.Erlenkeuser, G. Ganssen, Changes in east atlantic deepwatercirculation over the last 30,000 years: eight time slicesreconstructions, Paleoceanography 9 (1994) 209–267.

[62] K.S. Lackschewitz, K.-H. Baumann, B. Gehrke, H.-J. Wallrabe-Adams, J. Thiede, G. Bonani, R. Endler, H. Erlenkeuser, J.Heinemeier, North Atlantic ice sheet fluctuations 10,000–70,000yr ago as inferred from deposits on the Reykjanes Ridge,southeast of Greenland, Quat. Res. 49 (1998) 171–182.

[63] P. Martin, A. Voelker, S. Lebreiro, F. Abrantes, Millennial-scalevariability in mediterranean outflow in the Gulf of Cadiz: insightsfrom benthic mg/ca and ba/ca, AGU Fall Meeting Eos Trans.AGU 85 (47) (2004) (Fall Meet. Suppl., AGU, San Francisco,pp. Abstract PP12A-03).

[64] J.H. Baas, J. Schönfeld, R. Zahn, Mid-depth oxygen drawdownduring heinrich events: evidence from benthic foraminiferalcommunity structure, trace-fossil tiering, and benthic δ13c at thePortuguese margin, Mar. Geol. 152 (1998) 25–55.

[65] L. Loewemark, J. Schoenfeld, F. Werner, P. Schaefer, Tracefossils as a paleoceanographic tool: evidence from latequaternary sediments of the southwestern Iberian margin, Mar.Geol. 204 (2004) 27–41.

[66] E.J. Rohling, Review and new aspects concerning theformation of eastern mediterranean sapropels, Mar. Geol. 122(1994) 1–28.

55A.H.L. Voelker et al. / Earth and Planetary Science Letters 245 (2006) 39–55

[67] E. Colmenero-Hidalgo, J.-A. Flores, F.J. Sierro, M.A. Barcena,L. Loewemark, J. Schoenfeld, J.O. Grimalt, Ocean surface waterresponse to short-term climate changes revealed by coccolitho-phores from the Gulf of Cadiz (NE Atlantic) and Alboran Sea (Wmediterranean, Palaeogeography, Palaeoclimatology, Palaeoecol-ogy 205 (2004) 317–336.

[68] M. Perez-Folgado, F.J. Sierro, J.A. Flores, I. Cacho, J.O. Grimalt,R. Zahn, N. Shackleton, Western Mediterranean planktonicforaminifera events and millennial climatic variability during thelast 70kyr, Mar. Micropaleontol. 48 (2003) 49–70.

[69] G. Bond, R. Lotti, Iceberg discharges into the North Atlantic onmillennial time scales during the last glaciation, Science 267(1995) 1005–1010.

[70] M.R. Chapman, M.A. Maslin, Low-latitude forcing of meridi-onal temperature and salinity gradients in the subpolar NorthAtlantic and the growth of glacial ice sheets, Geology 27 (1999)875–878.

[71] T. Kiefer, Produktivität und Temperaturen im subtropischenNordatlantik: zyklische und abrupte Veränderungen im spätenQuartär, Ber. — Rep. Geol. — Pälaontol. Inst., Univ. Kiel 90(1998) (127 pp.).

[72] J.R.M. Allen, B. Huntley, Weichselian palynological recordsfrom southern Europe: correlation and chronology, Quat. Int. 73/74 (2000) 111–125.

[73] M. Bar-Matthews, A. Ayalon, A. Kaufman, G.J. Wasserburg, Theeastern mediterranean paleoclimate as a reflection of regionalevents: Soreq cave, Israel, Earth Planet. Sci. Lett. 166 (1999)85–95.

[74] N. Combourieu-Nebout, J.L. Turon, R. Zahn, L. Capotondi, L.Londeix, K. Pahnke, Enhanced aridity and atmospheric high-pressure stability over the western mediterranean during theNorth Atlantic cold events of the past 50 k.y. Geology 30 (2002)863–866.

[75] A. Moreno, I. Cacho, M. Canals, J.O. Grimalt, M.F. Sanchez-Goni, N. Shackleton, F.J. Sierro, Links between marine andatmospheric processes oscillating on a millennial time-scale. Amulti-proxy study of the last 50,000yr from the Alboran Sea(Western Mediterranean Sea), Quat. Sci. Rev. 24 (2005)1623–1636.

[76] G.R. Bigg, T.D. Jickells, P.S. Liss, T.J. Osborn, The role of theoceans in climate, Int. J. Climatol. 23 (2003) 1127–1159.

[77] M.J. Vautravers, N.J. Shackleton, C. Lopez-Martinez, J.O.Grimalt, Gulf stream variability during marine isotope stage 3,Paleoceanography 19 (2004 (04-30)) 1–12.

[78] M. Rogerson, E.J. Rohling, P.P.E. Weaver, J.W. Murray, Theazores front since the last glacial maximum, Earth Planet. Sci.Lett. 222 (2004) 779–789.

[79] L. de Abreu, N.J. Shackleton, J. Schoenfeld, M. Hall, M.Chapman, Millennial-scale oceanic climate variability off theWestern Iberian margin during the last two glacial periods, Mar.Geol. 196 (2003) 1–20.

[80] NGRIP members, High-resolution record of northern hemisphereclimate extending into the last interglacial period, Nature 431(2004) 147–151.