Instabilities in the Labrador Sea water mass structure during the last climatic cycle

Instabilities in the Labrador Sea water massstructure during the last climatic cycle

Claude Hillaire-Marcel and Guy Bilodeau

Abstract: In the modern Labrador Sea, the North Atlantic deep water components are found below the ~2 km deep,intermediate Labrador Sea water (LSW) mass, which is renewed locally through winter convective mixing. This watermass structure remained relatively stable since ~9.5 14C ka BP, as indicated by isotopic studies of foraminifer assem-blages from deep-sea cores. Almost constant differences in δ18O values are observed between major species. These av-erage –0.5‰ between the epipelagic species Globigerina bulloides and the mesopelagic species Neogloboquadrinapachyderma, left coiled, and –1‰ between Neogloboquadrina pachyderma and the benthic species Cibicideswuellerstorfi, after correction for Cibicides wuellerstorfi specific fractionation. These isotopic compositions representthermohaline conditions in surface waters, in the pycnocline with the LSW, and in the deep component of the NorthAtlantic deep water, respectively. A drastically different structure characterized the glacial Labrador Sea. Differences inδ18O values of ~ –2 to –2.5‰ are then observed between Globigerina bulloides and benthic species, indicative of astrong halocline between the corresponding water masses, thus for reduced production of intermediate waters. Duringthe same interval, Neogloboquadrina pachyderma shows 13C and 18O fluctuations of 1 to 1.5‰ amplitude, in phasewith Heinrich-Bond events and higher frequency climate oscillations. The δ18O values in Neogloboquadrinapachyderma vary between those of Globigerina bulloides and of benthic foraminifers, suggesting large amplitudebathymetric fluctuations of the halo–thermocline above and below the bathymetric range occupied byNeogloboquadrina pachyderma. Minimum δ18O values in Neogloboquadrina pachyderma match intervals of maximumice rafting deposition, such as the late Heinrich events, thus intervals with a deeper, more dilute buoyant surface waterlayer.

Résumé : Dans la mer du Labrador moderne, les composantes de la masse d’eau profonde de l’Atlantique Nord(NADW) sont présentes à plus de 2 km de profondeur, sous la masse d’eau intermédiaire de la mer du Labrador(LSW); celle-ci est renouvelée localement par convection hivernale. La structure des masses d’eau est demeurée relati-vement stable depuis -9,5 14C ka Av. Pr., tel qu’indiqué par des études isotopiques d’assemblages de foraminifères decarottes de forage profonds. Les espèces principales présentent des différences presque constantes de teneurs en 18O.Ces différences sont de –0,5‰ entre l’espèce épilagique Globigerina bulloides (Gb) et l’espèces mésopélagique Neo-globoquadrina pachyderma lévogyre (Npl) et de –1‰ entre Npl et l’espèce benthique Cibicides Wuellerstorfi (Cw),après correction du fractionnement spécifique de Cw. Ces compositions isotopiques représentent respectivement lesconditions thermohalines dans les eaux de surface, le long de la pycnocline avec la LSW, et dans la composante pro-fonde de la NADW. Lors du dernier maximum glaciaire, la mer du Labrador était caractérisée par une structure tout àfait différente. Les teneurs en 18O présentaient des différences d’environs –2 à –2,5‰ entre Gb et les espèces benthi-ques, indiquant une forte halocline entre les masses d’eau correspondantes et donc une moins grande production d’eauxintermédiaires. Au cours du même intervalle, Npl montre des fluctuations de teneurs en 13C et en 18O de 1 à 1,5‰d’amplitude, en phase avec les événements Heinrich–Bond et les oscillations climatiques à plus haute fréquence. Lesteneures en 18O de Npl varient entre celles de Gb et celles des foraminifères benthiques, ce qui suggère des fluctua-tions bathymétriques de grande amplitude de la halo-thermocline au-dessus et en-dessous de la zone bathymétrique oc-cupée par Npl. Les teneurs minimales en 18O de Npl s’observent au cours des intervalles de dépôt majeurs par lesicebergs de la fin des événements de Heinrich, soit au cours d’intervalles où la couche d’eau de surface était plusdiluée, plus profonde et plus stratifiée sur la couche sous-jacente.

Hillaire-Marcel and Bilodeau 809

Introduction

Major fluctuations in the production of North Atlanticdeep water (NADW), notably those which occurred since the

last glacial maximum (LGM), are well documented based onvarious sets of information. These notably include sedimen-tological data from the NE American margin (e.g., Johnsonet al. 1988), micropaleontological data (e.g., Streeter andShackelton 1979), stable carbon isotope studies of foramini-feral calcite (e.g., Duplessy et al. 1988), and geochemicalstudies of benthic foraminifer shells (e.g., Boyle andKeigwin 1987; Keigwin et al. 1991). Evidence for a reducedproduction of NADW with spreading of Antarctic BottomWaters (AABW) in the deep North Atlantic has been foundfor the LGM interval (e.g., Duplessy et al. 1988), and also

Can. J. Earth Sci. 37: 795–809 (2000) © 2000 NRC Canada

795

Received December 18, 1998. Accepted November 3, 1999.

C. Hillaire-Marcel1 and G. Bilodeau. GEOTOP-UQAM,B.P. 8888, succ. Centre-Ville, Montreal, QC H3C 3P8,Canada.1Corresponding author (e-mail: [email protected]).

during short climate oscillations, such as Heinrich-Bondevents (e.g., Berger et al. 1987; Keigwin and Lehman 1994).Most data sets indicate rapid switching between the high–low modes of NADW production (e.g., Boyle and Keigwin1987). From this point of view, temporal changes in the ver-tical structure of the water masses occupying the LabradorSea (Fig. 1B) are not as well documented. This is unfortu-nate because this basin has been a major conduit for icemeltwater during the last glacial interval; thus, a major con-duit for the spreading of freshwaters into the North Atlanticleading possibly to a reduction in NADW formation (e.g.,Broecker et al. 1989). Therefore, the reconstruction of pastthermohaline conditions in the Labrador Sea could help toconstrain the mechanisms responsible for rapid changes inthe production of NADW during the last ice age, especiallysince the Labrador Sea was surrounded by some of the ma-jor ice sheets of the last ice age (i.e., the Laurentide,Inuitian, and Greenland ice sheets). It is also located at theoutlet of the Hudson Strait source area for ice-rafted debrisand ice meltwater pulses during Heinrich-Bond events (H-events, Dowdeswell et al. 1995; Alley and MacAyeal 1994;Clarke et al., 1999).The Labrador Sea deserves further attention with respect

to temporal changes in its water mass structure, since it is amajor site for the renewal of intermediate oceanic watersthrough winter mixing and formation of Labrador Sea water(LSW) (Lazier 1988; Clarke and Gascard 1983; see alsoFig. 1). Under present conditions, this intermediate LSWspreads rapidly into the whole North Atlantic (Sy et al.1998), but remains well stratified in the Labrador Sea itself,where it overlies the two components of the NADW (e.g.,Lucotte and Hillaire-Marcel 1994; Fig. 1B). The NADWcomponents include the northeast Atlantic deep water andthe underflowing Denmark Strait overflow water (DSOW),which originates respectively from the Norwegian Sea andthe Greenland Sea. Both components preserve a clearthermohaline signature in the Labrador Sea (Lucotte andHillaire-Marcel 1994). The NADW water masses are carriedaround the deep Labrador Sea by the western boundary un-dercurrent, which sweeps the lower continental slopes. Itshigh velocity core is found at depths between 2500 and3000 m, that is, in the DSOW water layer (Lucotte andHillaire-Marcel 1994; Hillaire-Marcel et al. 1994a).During the LGM, evidence for a reduced outflow of the

western boundary undercurrent in the Labrador Sea has beenestablished (Hillaire-Marcel et al. 1994a; Fagel et al. 1996;Innocent et al. 1997), likely due to a much lesser productionof NADW, notably of its DSOW component. Discrete evi-dence for penetration of AABW into the deepest part of theLabrador Sea basin has also been reported during this inter-val (Bilodeau et al. 1994).In the present study, we examine more closely instabilities

in the structure of the water column of the Labrador Sea dur-

ing the last climatic cycle, with special attention to: (i) shortclimate oscillations; (ii) changes in the bathymetry of thehalo–thermocline between surface and intermediate watermasses; and (iii) the structure of the intermediate versusdeep water mass layers. In this last case, the time intervalexamined spans only the LGM–Holocene transition, mainlybecause of the scarcity of benthic foraminifers in the glacialsedimentary records.Our methodological approach is based on comparative

analyses of oxygen and carbon isotope compositions ofepipelagic versus mesopelagic planktic species (Globigerinabulloides (Gb) versus Neogloboquadrina pachyderma, left-coiling (Npl)), and of the mesopelagic Npl versus benthicspecies (Cibicides wuellerstorfi (Cw) or Uvigerina peregrina(Up)), in deep sea cores. In the past, such approaches havebeen used to document the vertical structure of water massesand (or) seasonality (e.g., Ravelo and Fairbanks 1995). How-ever, the possible mixing of noncontemporaneous popula-tions by bioturbation (e.g., Duplessy et al. 1986) has oftenhampered paleohydrographic interpretations. In the presentcase, the very high sedimentation rates observed at the studysites (Hillaire-Marcel et al. 1994a), which are located on theGreenland and Labrador upper rises (Fig. 1), as well as therestricted development of benthic life during the glacial in-terval resulting from low organic carbon fluxes at the seafloor (Hillaire-Marcel et al. 1994b), both contributed to limitmixing of populations of distinct geological age.

Material and methods

The study sites and sedimentsThe core sites used in this study are located at the en-

trance of the NADW gyre into the deep Labrador Sea, on theGreenland Rise, off Cape Farewell, and at the outlet of thiswater mass into the open North Atlantic Ocean, nearby Or-phan Knoll (Fig. 1A). These sites should thus allow a com-prehensive examination of changes in the inflowing–outflowing water masses of the Labrador Sea.The Greenland Rise site was cored in 1990 by the CSS

Hudson (site 90–013–013; henceforth, core P-013;58°12.59N; 47°22.40W; water depth: 3380 m; Hillaire-Marcel et al. 1990). Core P-013 shows a high Holocene sed-imentation rate (~30 cm@ka–1; Fig. 2) due to winnowingupslope, along the high velocity axis of the western bound-ary undercurrent, with sediment focussing at the coring site.A lower sedimentation rate (~10 cm@ka–1) prevailed duringthe glacial interval, partly due to a weaker western boundaryundercurrent (Hillaire-Marcel et al. 1994a and 1994b). Thesediment consists of hemipelagic mud with abundant bio-genic carbonates during interglacial and interstadial periods(Hillaire-Marcel et al. 1994b). The glacial sediment depictshigh contents in ice rafted lithic fragments with peaks duringHeinrich-events and higher frequency intermediate events

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Fig. 1. (A) Location map of the study sites and deep (dark lines) vs. surface (light lines) current trajectories. The gray surface corre-sponds to the area where winter mixing allows formation of intermediate Labrador Sea water (LSW). (B) Transect across the LabradorSea (see northeast–southwest line in Fig. 1A), showing the major water masses and currents, as well as the bathymetric domains occu-pied by the foraminifer populations used in this study. Abbreviations: NAD, North Atlantic Drift; WGC, West Greenland Current; LC,Labrador Current; WBUC, western boundary under current; NADW, North Atlantic deep water; DSOW, Denmark Strait overflow wa-ter; NEADW, Northeast Atlantic deep water. NAMOC, Northwest Atlantic mid-ocean channel. Npl, Neogloboquadrina pachyderma(left coiling); Gb, Globigerina bulloides; Cw, Cibicides wuellerstorfi; Up, Uvigerina peregrina; CGFZ, Charlie Gibbs fracture zone.

(Fig. 2; Stoner et al. 1995, 1998). Some of the Heinrich lay-ers (H-layers) contain up to 50% of fine, silt size detritalcarbonates originating from the Hudson Strait area (An-drews et al. 1994; Hillaire-Marcel et al. 1994a).

The Orphan Knoll site (Fig. 1A) was cored, in 1991, bythe CSS Hudson (core 91–045–094; henceforth, core P-094;50°12.26′N; 45°41.14′W; water depth: 3448 m; Hillaire-Marcel et al. 1992; Fig. 3), and subsequently in 1995 during

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the first IMAGES (international marine global change study)campaign of the Marion Dufresne (MD-2024; 50°12.40′N;45°41.22′W; water depth: 3539 m; Fig. 4). Due to coring ar-tifacts, the apparent sedimentation rate indicated by coreMD-2024 is about twice that yielded by P-094 (Fig. 5) dueto sediment stretching for MD-2024 and (or) compaction forP-094 (see discussion in Hillaire-Marcel, Turon et al. 1999).Fortunately, these artifacts do not show major impact onpaleoceanographic records. The Orphan Knoll cores yieldedhighly sensitive records of H-events and intermediate higherfrequency oscillations (Stoner et al. 1996; Fig. 4). At thissite, H-layers often depict a composite structure with twopeaks in coarse fraction content encompassing a peak in finedetrital carbonates originating from the Hudson Strait area(Veiga-Pires and Hillaire-Marcel 1999; Clarke et al., 1999).These carbonates spread through debris flow and turbiditeevents triggered at the head of the Northwest Atlantic mid-ocean channel system (Chough and Hesse 1976; Fig. 1).

Subsampling and laboratory techniquesSeveral isotopic data sets are used in this study, which

involve slightly different subsampling procedures and thusmay not be directly comparable. The original data sets forP-013 and P-094 (Hillaire-Marcel et al. 1994a) were basedon 2 cm thick samples collected at 10 cm intervals. The ben-thic foraminifer record of the present study has been ob-tained from this set of samples. The planktic foraminiferrecords were obtained from a new set of samples, collected

continuously and at 1 cm intervals. Based on apparent sedi-mentation rates of ~10 to ~30 cm@ka–1 (Hillaire-Marcel et al.1994a), this corresponds to a theoretical time resolution of~30 to 100 a. During intervals of high productivity and highorganic carbon fluxes at the sea floor, such as the Holocene,benthic life development could be responsible for somesmoothing of the isotopic records with a window of ~10 cm(Wu and Hillaire-Marcel 1994a, 1994b). In the new data setsfrom cores P-094 and P-013, we will essentially refer here tothe interval spanning isotopic stages 2 and 1.The 27 m long core MD-2024 spans approximately the

last 120 ka (Stoner et al. 1998). In this core, the data set cor-responds to 1 cm thick samples collected at 5 cm intervals.Due to the coring artifacts mentioned above, the 5 cm sam-pling interval of MD-2024 corresponds to a 2.5 cm samplinginterval (with 0.5 cm thick samples) in P-094. The theoreti-cal resolution is thus of ~200 years for core MD-2024 data.In order to document the thermohaline structure of the

upper few hundred metres of the water column, isotopicmeasurements were performed when possible on the twoplanktic species Globigerina bulloides (Gb) and Neoglobo-quadrina pachyderma, left coiled, (Npl). Two benthic spe-cies, Cibicides wuellerstorfi (Cw) and Uvigerina peregrina(Up), were used to record conditions in the deepest watermass occupying the basin. The rationale for combining thesetwo benthic species is that Cw is scarcely represented in gla-cial sediments, whereas Up is present, possibly in responseto the northward penetration of the AABW (Bilodeau et al.

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Fig. 2. High-resolution isotopic records in core P-013 (Greenland Rise): The LGM–Holocene interval. A complete data set down toisotopic stage 7, including sedimentological parameters and 14C data not shown here, is in open file (http://www.unites.uqam.ca/geotop/).Dark gray stripes represent H-layers; light gray stripes represent higher frequency oscillations. Cw, Cibicides wuellerstorfi; Npl,Neogloboquadrina pachyderma (left coiling); Gb, Globigerina bulloides. Water mass habitats: WGC, West Greenland Current; LSW,Labrador Sea water; DSOW, Denmark Strait overflow water.

1994; Streeter and Shackleton 1979; Duplessy et al. 1988).Unfortunately, neither Cw nor Up are abundant, thus the res-olution of the benthic isotope records is much lower thanthat of planktic records.For both planktic species, 40 to 50 specimens were hand-

picked from the 150–250 µm fraction in P-094 and P-013.Globigerina bulloides was analyzed only in the Holocenesections of both cores (Figs. 2, 3), where it is relativelyabundant. For the benthic species, 3–15 specimens werehandpicked from the 125–355 µm fraction (Figs. 2, 3). InMD-2024, the planktic specimens were handpicked in the125–250 µm fraction, and most samples yielded enough ma-terial to analyze both Gb and Npl assemblages throughoutthe whole sequence (Fig. 4). The size difference between theNpl samples of P-094 and of MD-2024 results in a 0.39 ±0.16‰ difference in their δ18O-values (Fig. 6), due a largerabundance of small shells depleted in 18O in MD-2024 sam-ples (see also Candon et al. 1999). This difference seems ap-proximately constant throughout the sequence (Fig. 5). Onthe contrary, δ13C values are similar in both size fractions(Fig. 6). Size fractions of Gb depict a lesser and more vari-able difference in their δ18O-values (0.27 ± 0.22‰) thanthose of Npl (Fig. 6).For all samples, CO2 was extracted at 90°C using an

ISOCARB™ device on-line with a VG-PRISM™ massspectrometer. All measurements were made using an in-house standard marble calibrated against the Carrara marbleand other current standard materials of the International

Atomic Energy Agency of Vienna. Results were convertedto the V-PDB scale (Coplen 1996) after usual corrections(Craig 1957). The overall analytical reproducibility, as de-termined from replicate measurements on the in-house stan-dard material, is routinely better than ±0.05‰ (± 1σ) forboth δ13C and δ18O values. δ18O values in Cw were correctedby 0.64‰, following Shackleton and Opdyke (1973) in or-der to allow direct comparison of Cw records with otherbenthic records. Details on other analytical procedures (e.g.,for grain size, CaCO3, 14C measurements, etc.) may befound in Hillaire-Marcel et al. (1994a) and in Stoner et al.(1995). The chronological scale of Fig. 7 is derived fromhigh resolution paleointensity measurements (see Stoner etal. 1998) and peak-to-peak correlation with the compositeSINT-200 paleomagnetic record, the latter already linked tothe SPECMAP (Mapping Species Variability in Global Cli-mate Project) time scale (Martinson et al. 1987) by Guyodoand Valet (1996). The results used for the present study areillustrated in Figs. 2 to 7 and summarized in Table 1. Theoriginal and complete database can be downloaded fromhttp://www.unites.uqam.ca/geotop/.

Results

The Greenland Rise recordThe glacial and Holocene sediments show drastically dis-

tinct isotopic offsets between species. As illustrated by coreP-013, the LGM shows almost identical δ18O values in Npl

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Fig. 3. High-resolution isotopic records in core P-094 (Orphan Knoll): The LGM–Holocene interval. The complete data set down to isoto-pic substage 5d, including sedimentological parameters and 14C data not shown here, is in open file (http://www.unites.uqam.ca/geotop/).Dark gray stripes represent H-layers; light gray stripes represent higer frequency oscillations. Cw, Cibicides wuellerstorfi; Up,Uvigerina peregrina; Npl, Neogloboquadrina pachyderma (left coiling); Gb, Globigerina bulloides. Water mass habitats: LC, LabradorCurrent; LSW, Labrador Sea water; DSOW, Denmark Strait overflow water.

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Fig. 4. High-resolution sedimentological and isotopic records in core MD-2024 (Orphan Knoll). The very late Holocene sediment was not recovered. Note that comparedwith those of Figs. 2 and 3, the δ18O values of planktic foraminifers are shifted by 0.39 ± 0.16‰ (Npl) and 0.27 ± 0.22‰ (Gb) towards lighter values in this core, dueto a size difference in the populations used for isotopic measurements. The δ13C values are directly comparable to those from other data sets (see Fig. 6).

and Cw (Table 1; Fig. 2), but the two species differ by about1‰ throughout the Holocene. During the same interval, Gbshows a slight offset (~ –0.5‰) towards lighter δ18O value,when compared with Npl. Assuming isotopic equilibriumduring calcite precipitation, this offset would indicate a dif-ference of 2°C between their respective growth temperatures(see Duplessy et al. 1991), or slightly less, here, wherestrong salinity gradients in the upper water column couldalso partly account for the offset.Another isotopic feature observed in most Labrador Sea

sequences (e.g., Stoner et al. l996) and clearly depicted hereconcerns the H-events, and to a lesser extent, the intermedi-ate higher frequency oscillations, some correlative toDansgaard-Oeschger oscillations (Bond et al. 1993).Through each of the H-layers, shifts towards lighter δ18Ovalues in Npl are systematically observed, the lightest valuescorresponding to the top of the layer. These isotopic shiftsreach –0.5‰ in H2, –1.5‰ in H1, and –1‰ in H0, whereasslightly reduced isotopic “excursions” are depicted by thehigher frequency oscillations, although some of them mayshow shifts as large as –1‰ (Fig. 2). In most cases, a strong

negative 13C event matches the light 18O peak on top of theH-layer (e.g., Fig. 4). Another feature, also observed in otherLabrador Sea records, (e.g., Hillaire-Marcel et al. 1994a),relates to the slight differences between the early Holoceneisotopic signatures and those of the middle and late Holo-cene (Table 1; Fig. 2). The early Holocene populations aresystematically depleted, in both 13C and 18O, by approxi-mately 0.2–0.4‰, compared with the middle and late Holo-cene assemblages.

The Orphan Knoll recordCore P-094 depicts isotopic features not unlike those of

P-013, notably with respect to the Holocene–LGM differ-ence of isotopic signatures and to the early–late Holocenetrend (Fig. 3; Table 1). At this site, H-layers and high fre-quency oscillations are particularly well recorded (Stoner etal. 1996; Veiga-Pires and Hillaire-Marcel 1999; Clarke et al.,1999). The isotopic “excursions” of H2, H1 and H0 are morepronounced than in P-013, and show shifts in δ18O values of–1.5, –2.5, and –1‰, respectively (Fig. 3). As illustrated bydata from core MD-2024 (Fig. 4), the light Npl-δ18O values

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Fig. 5. Correlation of Npl δ18O-records in P-094 and MD-2024 showing a depth ratio of ~2 between the two cores, likely due to cor-ing artifacts (see text). The almost constant isotopic offset between the two cores is due to a size difference in the populations ana-lyzed (see Fig. 6).

at the top of the H-layers match peaks in the coarse fractioncontent of the sediment, suggesting high ice-rafting deposi-tion and maximum dilution by iceberg melting. At OrphanKnoll, the layers representing higher frequency events depictsedimentological properties resembling those of the H-layers, notably a sharp light δ18O peak matching a light δ13Cpeak in Npl, and a high coarse lithic fragment content(Fig. 4; see also Stoner et al. 1998; Clarke et al., 1999).Core MD-2024 yielded a unique set of isotopic data for

both Npl and Gb from the onset of isotopic substage 5d untilthe Holocene (Fig. 4). In most cases, one may assume thatthe two populations analyzed were contemporaneous, withinthe time slice represented by a 1 cm thick sample (i.e., a few

decades here), for the following arguments: (i) during warmintervals, as today, favorable conditions likely resulted inhigh and simultaneous production of both species (e.g.,Tolderlund and Be 1971); (ii) during cold intervals, due tolow organic carbon fluxes to the sea floor (e.g., Muzuka1996), minimum benthic mixing occurred, as shown by sedi-mentological and geochemical records (e.g., Veiga-Pires andHillaire-Marcel 1999).In order to compare Gb and Npl isotopic records, isotopic

differences (or offsets) between the two species have beenplotted in Fig. 7. Both 18O and 13C offsets show large ampli-tude oscillations matching broadly the major isotopic stages-5 substages and major H-events. The amplitude of the 18Ooffset oscillations averages 1.5‰; that of the 13C offsets,about 1‰, and both are approximately in phase. The onlyexception to this pattern is seen in the isotopic stages-2 sec-tion (between H-events H1 and H3), where isotopic composi-tions of Gb and Npl depict a drastically larger 18O offset(about –2.5‰), but a minimum (close to 0‰) 13C offset.The Orphan Knoll site is located off the Hudson Strait

with respect to meltwater supplies from the NE sector of theLaurentide Ice Sheet (e.g., Andrews et al. 1995). Conse-quently, isotopic records at this site are very sensitive todrainage events. A noticeable one occurred at ca 8.2 ka BP(see Barber et al. 1999) when the glacial lakes Ojibway andAgassiz were drained into the postglacial Tyrrell Sea, theninvading Hudson Bay (Hardy 1977; Hillaire-Marcel et al.1981). During this drainage event, a volume of about1014 km3 of fresh water was evacuated, through Hudson Straitand along the Labrador margin, into the North Atlantic (Bar-ber et al. 1999). As shown in Figs. 4 and 7, this event is re-corded by a drop of almost –0.5‰ in δ18O values of Gb alongthe LC trajectory, whereas δ18O values in Npl remained rela-tively constant, thus suggesting that the freshwater was evacu-ated in the shallow mixed layer occupied by Gb.

Discussion

The contrasted isotopic offsets of the Holocene fora-minifer assemblages, in comparison with the glacial interval,constitute a prominent feature of the Labrador Sea records.The large 18O offset (~1‰) between Npl and benthic fora-minifers, and the relatively low offset between Gb andNpl (~ –0.5‰), which characterize the deglaciation, are 14Cdated at ca. 9.5 ka BP, at both study sites (Figs. 2 and 3).This indicates the early inception of a modern-like situation(Fig. 1B) with an intermediate LSW stratified over theNADW, and moderate thermohaline gradients between thesurface water layer and the LSW, allowing winter convectionto occur down to the Labrador Water–NADW limit(cf. Clarke and Gascard 1983). However, the reconstructionof precise thermohaline conditions in the upper water col-umn, using isotopic data in Gb and Npl, requires further ex-amination of their bathymetric distribution.

The bathymetric calibration of the isotopic“thermohaline probe”In the modern ocean, living assemblages of Gb are found

at depths from zero to 80 m, with a preferential habitat be-tween 50 and 80 m (e.g., Curry and Matthews 1981; Deuserand Ross 1989). In the Labrador Sea, possibly due to tem-

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Fig. 6. A comparison of isotopic compositions of two size frac-tions of Npl and Gb from core MD-2024. The samples were ran-domly selected throughout the sequence illustrated in Fig. 4, andtherefore include glacial sediment samples (high values) andinterglacial ones (light values). The two size fractions yield iden-tical δ13C values, but lighter δ18O values (y-intercepts of theequations).

perature constraints, Gb seems to precipitate its calcite invery shallow waters (i.e., ~50 m; Candon et al. 1999),whereas Npl occupies a much broader and deeper domain.Juvenile specimens are found in shallow waters (~50 to80 m; e.g., Bauch et al. 1997; Kohfeld et al. 1996), butlarger mature specimens (with frequent calcite crystal over-growths) are generally found at much greater depths (asdeep as 500–600 m in the Arctic Ocean; e.g., Aksu andVilks 1988). In the modern Labrador Sea, isotopic studies ofNpl assemblages from box-core tops (Candon 1999) suggestthat the it growths along the thermo–halocline between sur-face and intermediate waters. Therefore, in this basin, com-

parative measurements between Gb and Npl shells shouldhelp to constrain the thermohaline gradients between themixed surface water layer occupied by Gb, notably along thetrajectories of the Western Greenland Current (core P-013)and Labrador Current (cores P-094 and MD-2024; Fig. 1A),and the top of the underlying LSW mass, occupied by Npl(Fig. 1B).A quick examination of late Holocene data confirms this

interpretation. For example in P-094, δ18O values of Nplshells, in the 150–250 µm range, average 2.5‰ (Table 1).This isotopic composition has to be compared with thedepth–dependent isotopic composition of a calcite formed in

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Fig. 7. Isotopic offsets between Gb and Npl at Orphan Knoll throughout the last climatic cycle (from the 5e–5d transition to isotopicstages 1). Large amplitude oscillations of 18O and 13C offsets seem in phase during most of the interval, with the exception of the iso-topic stage 2 (see text).

isotopic equilibrium with ambient water (Fig. 8). We usedthe equation of O’Neil et al. (1969) and Shackleton (1974)to calculate the temperature-dependent fractionation betweencarbonate and water, and the equation of Wu and Hillaire-Marcel (1994a, 1994b), for the regional relationship be-tween the water salinity and the water isotopic composition:

[1] t = 16.9 – 4.38 (δ !A) + 0.10 (δ !A)2

[2] A ~ – 18 + 0.5 S

wheret is temperature (in°C) during precipitationS is salinity during the growth seasonδ is oxygen isotope composition of calcite (vs. PDB)A is oxygen isotope composition of water (vs. SMOW)

corrected by –0.26‰ for conversion into PDB scale.The temperature and salinity are compiled after the WorldOcean Atlas (NODC 1994) and represent the mean valuesfor the months of June, July and August, in the OrphanKnoll area. This period broadly matches that of maximumprimary productivity in the area (NODC 1994), thus likelythat of maximum development of Npl. As shown in Fig. 8,the isotopic composition of the late Holocene Npl assem-blages (Table 1) suggests precipitation on top of thepycnocline, that, at the depth where feeding particulate mat-

ter is likely to be more abundant. A similar observation hasbeen made by Kohfeld et al. (1996) in the Arctic Ocean formodern, live-collected Npl. At Orphan Knoll, the isotopiccomposition of late Holocene Gb shells (Table 1) suggestsprecipitation in shallower water, possibly slightly out ofequilibrium (Fig. 8). Spero and Lea (1996) have observeddeparture from isotopic equilibrium in Gb shells under ex-perimental conditions, but within a much higher temperaturerange (18–22°C) than that prevailing at Orphan Knoll(=13°C; NODC 1994). Therefore, their observation does notnecessarily apply here. Moreover, the isotopic compositionof the late Holocene Gb shells at the study site is in agree-ment with calcite precipitation in isotopic equilibrium withsurface waters in August (i.e., at 13°C). The possibility thatthe growing season of Gb is much shorter than that of Npl,at this site, notably due to temperature constraints, cannot bediscarded.

Instabilities in the shallow–intermediate water massesduring the last climatic cycleThe oscillatory response of the 13C and 18O offsets be-

tween Gb and Npl, during the glacial interval, and their fluc-tuation in phase, constitute striking features of the OrphanKnoll record (Fig. 7B, 7C). They suggest alternating periodsof contrasted hydrographic and climatic regimes (see Fig. 9):

© 2000 NRC Canada


Site & species Age δ18O δ13C N1

Orphan Knoll (91-045-094)Globigerina bulloides Late Holocene 1.74±0.08 –0.24±0.09 56G. bulloides Middle Holocene 1.67±0.07 –0.44± 0.13 22Neogloboquadrina pachyderma (I) Late Holocene 2.49±0.13 0.69± 0.09 8N. pachyderma (I) Early Holocene 2.19±0.11 0.49±0.10 65N. pachyderma (I) Holocene 2.22±0.15 0.51±0.11 73N. pachyderma (I) LGM2 4.55±0.10 0.23±0.11 76Cibicides wuellerstorfi Holocene 3.11±0.13 0.84±0.17 16Uvigerina perigrina LGM2 4.89±0.16 –0.73±0.19 10

Orphan Knoll (MD-2024)4N. pachyderma (I) Middle Holocene 2.10±0.063 0.75±0.09 8N. pachyderma (I) Early Holocene 1.92±0.093 0.52±0.10 25N. pachyderma (I) LGM2 3.85±0.143 0.26±0.07 39G. bulloides Middle Holocene 1.38±0.133 –0.11±0.12 46G. bulloides Early Holocene 0.96±0.213 –0.59±0.12 33G. bulloides LGM2 2.04±0.103 –0.91±0.24 7

Greenland Rise (90-013-013)N. pachyderma (I) Late Holocene 2.53±0.11 0.77±0.09 21N. pachyderma (I) Early Holocene 2.28±0.14 0.49±0.09 145N. pachyderma (I) Holocene 2.38±0.13 0.53±0.13 166N. pachyderma (I) LGM2 4.42±0.12 0.20±0.06 52G. bulloides Early Holocene 1.92±0.14 –0.73±0.14 110C. wuellerstorfi Late Holocene 3.68±0.20 1.10±0.09 12C. wuellerstorfi Early Holocene 3.49±0.18 0.73±0.20 6C. wuellerstorfi Holocene 3.61±0.21 0.97±0.22 23C. wuellerstorfi LGM2 4.78±0.10 0.00±0.20 2Note: N = number of samples representing the corresponding interval and used for the calculation of the

mean values. H-events were excluded for the calculation of mean LGM values. In core MD-2024, Gb and Nplshells were handpicked in the fraction > 125 µm. In this size fraction, Npl shells show a depletion in 18O of0.39 ± 0.16‰, when compared with the size fraction > 150 µm. There is no apparent difference in their δ13Cvalues (see Fig. 6). In core MD-2024, the late Holocene section is missing.

Table 1. Mean isotopic compositions of foraminiferal assemblages (LGM vs. Holocene).

(i) Intervals such as today, with a weak halo–thermoclinebetween the surface and intermediate waters allowing effi-cient renewal of the intermediate LSW through winter con-vection (Fig. 1), and efficient ventilation of this LSW(indicated by minimum 13C offsets); during such intervals,the 18O offset between Gb and Npl averages 0.5‰; whereasthe ~ –1‰ offset between Npl and benthics suggests thepresence of a deep NADW-type water mass (Fig. 9A);(ii) Periods when the (shallow) dilute surface water masswas strongly stratified over a single water mass occupyingthe basin down to the sea floor; maximum 18O offsets be-tween Gb and Npl (i.e., >2‰) and minimum differences be-tween Npl and benthics (i.e., <0.5‰), characterize thisregime, which also corresponds to minimum ventilation ofthe mesopelagic layer occupied by Npl (maximum 13C offsetwith Gb); this situation prevailed notably during the LGM(Fig. 9B); (iii) Periods characterized by the development ofa deeper halo–thermocline than today. Low 18O and 13C off-set values (~ –0.5‰) are then observed between Npl and Gbshells (Fig. 7), suggesting that they grew in a relatively wellmixed upper water layer. The large 18O offsets (~ –1.5 to2‰; e.g., Fig. 3) observed between Npl and benthicforaminifers suggest a strong halo–thermocline between theupper water layer, occupied by Npl and Gb, and an interme-diate-deep water mass represented by the benthics (Fig. 9C).The first regime prevailed during most of the Holocene

and possibly during the warm substages of isotopic stages 5.The third regime characterizes the final stage of Heinrichevents, and some major Dansgaard-Oeschger (DO) oscilla-tions of the isotopic stages 4 and 3 likely correlative ofinterstadial intervals, inland (Walker et al. 1999). The secondregime generally corresponds to cold or cool intervals, start-ing as early as isotopic substage 5d. It shows two modes:one with a –1.5‰ offset in 18O between Gb and Npl(Fig. 7B), which characterizes stadial intervals of isotopicstages 5 to 3; the second with a –2.5‰ offset between thetwo planktic species, which prevailed during cold intervalsof isotopic stages 2. In general, the 18O offsets between Gband Npl respond to climate oscillations (H-events and higherfrequency events) with large amplitude oscillations (of about1‰) between benchmark offset values of approximately−0.5, –1.5 and –2.5‰ (Fig. 7B) representing respectively:(i) interglacial-interstadial type conditions, (ii) stadial typeconditions of the isotopic stages 5 to 3, and (iii) full glacialconditions since H3 and throughout isotopic stage 2.

The “full glacial conditions” of the H3–H1 intervalThe late isotopic stage 3 – stage 2 interval, encompassed

by H-events 1 and 3 (Fig. 7A), shows further specificity thanthose evoked above. It depicts 13C and 18O offsets betweenNpl and Gb not as strongly correlated than those of other in-tervals were (Fig. 7C). This could be due to the fact that itcorresponds to the period with the largest 18O offset value(2.5‰) between Gb and Npl, i.e., to the interval with themost drastically stratified and shallow surface layer of thewhole sequence (Fig. 9B). The low thermal inertia of such ashallow and buoyant surface water layer could account forthe enhanced seasonality, with warm summers and very coldwinters, which is suggested for high latitudes of the NorthAtlantic, during the LGM, by transfer functions based nota-bly on dinocyst assemblages (de Vernal et al. 2000; de Ver-

nal and Hillaire-Marcel 1999). Below this shallow dilutesurface water layer, a more saline water mass was likely oc-cupying the water column down to the sea floor (as sug-gested by minimum 18O offsets between Npl and benthicforaminifers).The isotopic offsets between Gb and Npl cannot be tran-

scribed unequivocally into salinity or temperature gradients,notably because the bathymetric domain occupied by Nplmay have changed, through time, as have the water columntemperature, salinity, and isotopic composition. Moreover,due to glacial–interglacial changes in isotopic compositionsof the oceanic and freshwater end-members, a nonlinear re-lationship links the isotopic composition of paleowaters withthe salinity. Nevertheless, some constraints on salinity andtemperature gradients in the upper water column can be setfor the LGM by combining isotopic data to information fromother proxies. August temperature and salinity of the mixedsurface water layer occupied by Gb can be obtained fromtransfer functions using dinocyst assemblages: values of 2 ±2.5°C and 31.2 ± 0.8 are found for the study site (de Vernalet al. 2000). Mean δ18O values of 2.04 ± 0.10 are found inGb during this interval (Table 1). The eq. [1] above then

© 2000 NRC Canada


Fig. 8. Physical properties of the upper water column and depthdependence of the oxygen isotope composition of a modern calciteformed at equilibrium with ambient water during the months ofJune, July and August at Orphan Knoll. This composition hasbeen calculated using the carbonate-water temperature dependenceequation of O’Neil et al. (1969) and Shackleton (1974), and theequation of Wu and Hillaire-Marcel (1994a, 1994b) for the rela-tionship between the salinity and the isotopic composition of thewater in the area. Temperature and salinity profiles are compiledfrom the World Ocean Atlas (NODC 1994). Mean late Holoceneδ18O value in Npl shells from core P-094 indicate precipitation atmean pycnocline depth, whereas the mean value in Gb shells fromthe same core (arrow) suggests precipitation in shallower waters,either slightly out of equilibrium or during a shorter and warmerinterval (e.g., August; see text).

yields an oxygen isotope composition of –1.7 ± 0.7‰ (vs.SMOW) for the corresponding surface water layer. Based onthese numbers, and assigning a salinity of 36 and an isotopiccomposition of ~ +1.2‰ for the oceanic end-member, dur-ing the LGM (Broecker and Peng 1982; Duplessy et al.1993), it is possible to calculate an isotopic composition forthe regional ice meltwater end-member. However, due tolarge uncertainties in the reconstruction of paleoseasurfaceconditions, this value lacks of precision. It yields a poorlyconstrained eq. [2] above, for the LGM: ALGM ~ –19 ± 9 +0.6 ± 0.2 S. Alternatively, one could constrain the isotopiccomposition for the LGM ice meltwater end-member. Avalue of ~ –25‰ could be suggested, based on the differ-ence of ~6‰ in δ18O values recorded in Greenland Ice coresbetween the modern and LGM ice (e.g., Johnsen et al. 1992;see also Mix and Ruddiman 1984; Duplessy et al. 1993).The LGM eq. [2], above, would then be: ALGM ~ –25 +0.7 S. On this basis, the Npl-Gb 18O offset of ~2.5‰ ob-served during the LGM would indicate a salinity gradient of3.5 between the water layers occupied by Gb and Npl, re-spectively, assuming, of course, the same growing seasontemperature for both species. A lower growing season tem-perature for Npl, as observed today, for example, in the Arc-tic polynias (0°C; Kohfeld et al. 1996), would yield aslightly lower salinity gradient (~3). Nevertheless, the abovevalues yield a reasonable estimate for LGM salinity, at thewater depth then occupied by Npl, in the 34 to 35 and abovesalinity range. A strong halocline was thus present betweenthe surface water layer occupied by Gb and the water massoccupied by Npl. Below this halocline, a single water masswas probably occupying the water column (Fig. 8B), as sug-gested by almost identical δ18O values in Npl and benthics(Fig. 3). The 13C-depleted carbon isotope compositions of

Npl during this interval (Fig. 4) further suggest a low venti-lation rate for this water mass. This could account for alarger reservoir aging effect in North Atlantic waters prior tothe H0–Younger Dryas interval, as documented by Bard etal. (1994) and based on 14C measurements in Npl assem-blages.

Major meltwater and iceberg dispersal eventsDuring late H-events, 18O offsets between Gb and Npl are

strongly reduced (Fig. 7B). This is essentially due to thelarge shift of isotopic compositions in Npl towards lighterδ18O values. Accordingly, this shift resulted in the develop-ment of a large offset between Npl and benthics, since ben-thics show minimum isotopic variations during these events(e.g., H2 and H3 events in Fig. 3). These increasing 18O off-sets, between Npl and benthic foraminifers, matching mini-mum offsets between Npl and Gb, are interpreted as theconsequence of deepenings of the halocline, below the habi-tat of Npl (Fig. 8C). This was probably due to the spreadingand melting of thick icebergs (depicted by the peak in ice-rafted debris at the top of H-layers).

Conclusions

The present study shows that multi-specific isotopic mea-surements in foraminifer assemblages from high sedimenta-tion rate sites may provide indications on high frequencyinstabilities in the overlying water column. Unfortunately, italso shows that unequivocal interpretation of isotopic offsetsbetween epipelagic, mesopelagic and benthic species are un-likely. Other proxies are needed to constrain temperature andsalinity conditions at a given depth in order to transcribe iso-

© 2000 NRC Canada


Fig. 9. Interpretative sketch of the vertical water mass structure in the Labrador Sea during key intervals (not to scale).

topic offsets in the water column into precise thermohalinegradients.At the Orphan Knoll study site, it seems clear that the os-

cillatory nature of the 18O offsets between Gb and Npl, dur-ing the last climatic cycle, is primarily due to the very largeisotopic fluctuations recorded by Npl in comparison to Gband benthics. Benthic foraminifers are generally muchheavier, and Gb, much lighter. The Npl show fluctuationsbetween these two “end members,” likely in response tooscillations of the halo–thermocline in the upper water col-umn, Npl occupying accordingly a bathymetric domain inthe upper or deeper parts of the halo-thermocline.With reference to the role of the Labrador Sea as a place

of renewal of intermediate or deep water, it seems that thisrole might well be restricted to periods such as today, whenthe 18O-offset between the two planktic species does not ex-ceed 0.5‰ (i.e., during the Holocene, the warm intervalsof isotopic stages 5, and possibly during interstadials ofisotopic stages 4 and 3). During other periods, the largeisotopic difference between Gb and Npl suggests salinity-temperature gradients in the upper water column much toohigh to allow for efficient winter mixing.

Acknowledgements

Comments from A. de Vernal and A. Simonetti(GEOTOP-UQAM) helped to clarify several aspects of thepresent paper. Those from J.-C. Duplessy (C.N.R.S., France)and an anonymous reviewer of the Journal also led to majorimprovements of this paper. This study is a contribution tothe Climate System, History and Dynamics project, sup-ported by the National Science and Engineering ResearchCouncil of Canada. Complementary support by the Fondspour la Formation de chercheurs et l’Aide à la Recherche ofthe Quebec Province is also acknowledged. Core MD-2024was retrieved within the framework of the Images program.

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Instabilities in the Labrador Sea water mass structure during the last climatic cycle

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