Variation in surface and deep water circulation in the Denmark Strait, North Atlantic, during marine isotope stages 3 and 2

Transcription

1 PALEOCEANOGRAPHY, VOL. 17, NO. 4, 1061, doi: /2001pa000632, 2002 Variation in surface and deep water circulation in the Denmark Strait, North Atlantic, during marine isotope stages 3 and 2 Sveinung Hagen 1 Department of Geology, University of Tromsø, Tromsø, Norway Morten Hald Department of Geology, University of Tromsø, Tromsø, Norway Received 12 February 2001; revised 24 April 2002; accepted 3 July 2002; published 23 October [1] The paleoceanography in the Denmark Strait, North Atlantic, is reconstructed based on down core analysis of planktonic foraminifera, carbon and oxygen isotopes, carbonate, and ice rafted debris (IRD). The stratigraphy spans 57 to 6 calibrated (cal) ka years and reveals sea surface temperatures that covary with Greenland ice core temperatures throughout marine isotope stages (MIS) 3 and 2. Planktonic d 18 O minima, interpreted as meltwater events, follow sea surface warming. IRD from both the Greenland and Iceland Ice Sheets was delayed by 2 ka relative to the North Atlantic Heinrich events. Ventilation of bottom water was mostly sustained through open ocean convection associated with North Atlantic Deep Water production or by convection of glacial North Atlantic Intermediate Water, except for two episodes at 54 cal ka (MIS 3.31) and cal ka (termination I). During these two periods we suggest that deep water ventilation was accomplished mainly by brine formation. INDEX TERMS: 3022 Marine Geology and Geophysics: Marine sediments processes and transport; 3344 Meteorology and Atmospheric Dynamics: Paleoclimatology; 3030 Marine Geology and Geophysics: Micropaleontology; KEYWORDS: marine isotope stages 3 and 2, Denmark Strait, sea surface temperature, deep water characteristics, ice rafting debris Citation: Hagen, S., and M. Hald, Variation in surface and deep water circulation in the Denmark Strait, North Atlantic, during marine isotope stages 3 and 2, Paleoceanography, 17(4), 1061, doi: /2001pa000632, Introduction 1.1. North Atlantic Paleoceanography [2] The last glacial period was characterized by several climatic oscillations on a millennial scale, first documented in pollen and ice core records [Dansgaard et al., 1982; Woillard and Mook, 1982]. These fluctuations, known as Dansgaard-Oeschger cycles in the ice core records, were superimposed on an orbitally controlled cooling trend [Hays et al., 1976; Imbrie et al., 1984]. One explanation for these climatic fluctuations is related to ocean circulation including northward surface ocean heat flux, and deepwater production [Boyle and Keigwin, 1982, 1987; Broecker, 1991; Keigwin et al., 1994; Maslin, 1995; Oppo and Lehman, 1995; Rasmussen et al., 1996; Broecker, 1997]. [3] The northern North Atlantic is a locus for thermohaline convection including northward transport of warm Atlantic Water by the Norwegian Current and Irminger Current, southward surface transport of cold polar water in the East Greenland Current (EGC) and formation of deep water. Deep water formed in the northern North Atlantic leaks over the Denmark Strait and Iceland Faeroe-Shetland- Scotland ridges and contributes to the North Atlantic Deep 1 Now at Statoil ASA, Bergen, Norway. Copyright 2002 by the American Geophysical Union /02/2001PA Water (NADW) [Worthington and Volkmann, 1965; Swift, 1980; Lonsdale and Hollister, 1979; Malmberg, 1984; McCave and Tucholke, 1986; Dickson et al., 1990]. [4] A reconstruction of the North Atlantic during the last glacial shows a very variable oceanographic regime. During the cold phases the oceanic polar front was located west of Portugal, the surface circulation was inert, the northward heat flow was strongly reduced and deep water formation in the Nordic Seas (Greenland, Icelandic and Norwegian seas) was suppressed [CLIMAP Project Members, 1981]. However, during warmer phases Atlantic Water advected into the polar North Atlantic and production of bottom water was sustained through thermohaline convection and/or brine formation [Veum et al., 1992; Dokken and Hald, 1996; Rasmussen et al., 1997; Dokken and Jansen, 1999]. In the North Atlantic south of the Nordic Seas, ventilated intermediate water termed Glacial North Atlantic Intermediate Water (GNAIW) was formed during periods of the last glacial cycle. GNAIW is interpreted to be derived through advection of surface water during periods when the polar front was suppressed south of Iceland [Duplessy et al., 1988]. According to Oppo and Lehman [1993], GNAIW occupied the intermediate water layer above 2000 m depth during the LGM in the North Atlantic, while d 13 C-depleted and nutrient enriched Southern Ocean water (SOW) replaced North Atlantic Deep Water (NADW) below 2000m. [5] Based on a multiproxy paleoceanographic record from a sediment core retrieved at 1683 m water depth 13-1

2 13-2 HAGEN AND HALD: SURFACE AND DEEP WATER IN THE DENMARK STRAIT south of the Denmark Strait in the North Atlantic, the purpose of this study is to elucidate the paleoceanography and climate history during marine isotope stages (MIS) 3 and 2. Situated in the northern Irminger Basin just south of the Denmark Strait sill (present sill-depth approximately 600 m), the core is located in an area that is sensitive to changes in surface ocean heat flux of North Atlantic Water through the Irminger Current. Also, various modes of deep water formation in the North Atlantic and Nordic Seas may be detected. Today, the core site is bathed by a combination of Denmark Strait Overflow Water (DSOW) and returned Norwegian Sea Bottom Water (NSBW). The site should also be favorably situated for detecting marine evidence for Greenland and Iceland glacial history, and is close to marginal moraines from the last glacial period on SE Greenland and SW Iceland continental shelves Glacial History [6] The glacial history of Greenland and Iceland during MIS 3 and 2 has been addressed in several studies both on land and the adjoining continental shelves [Dowdeswell et al., 1995; Andrews et al., 1996; Eiriksson et al., 1997; Ingólfsson et al., 1997; Funder et al., 1998; Andrews et al., 1998, 2000]. During the LGM, the Greenland Ice Sheet covered an area of approx. 4 million km 2 (slightly less than the Scandinavian and Barents Sea Ice Sheets combined), and more than 50% of its volume survived into the present interglacial conditions [Funder et al., 1998]. Today, the Greenland Ice Sheet is the second largest continental ice cap after the Antarctic Ice Sheet complex. Although evidence of large climate fluctuations is found in the ice cores during MIS 3 and 2, associated ice sheet fluctuations appear to have been minor [Funder et al., 1998]. The Flakkerhuk Stade (MIS 2 and 3) lasted for ca 50 ka (the ka unit is used for 1000 years) culminating between about 14 and C ka, as indicated by ice rafted debris deposits on the continental slope [Stein et al., 1996; Funder et al., 1998]. During the Flakkerhuk stade, the ice margin advance was restricted to cold based outlet glaciers that filled the fjord troughs from Scoresby Sound (68 N) and northward, while southwards of 68 N most of the 300 km wide continental shelf was glaciated [Stein et al., 1996]. Evidence from sediment cores in the outer Kangerluggsuaq Trough suggests that the Greenland Ice Sheet margin reached the shelf break during the LGM Oceanography [7] On Iceland, the glacial extent during the LGM glaciation is poorly known because ice margins were off the present shore. Some evidence, however, suggests that the ice drained through troughs on the shelf, and was grounded at 200 m below present sea level [Ingólfsson et al., 1997]. The Látra end moraine complex (LM in Figure 1), having a seaward maximum age of C ka, may mark the Late Weichselian maximum extent of the western Icelandic Ice Sheet [Syvitski et al., 1999]. Based on seismic evidence there are indications for at least three separate episodes of glaciation on the SW Iceland shelf, although it is not clear which ages these episodes represent [Syvitski et al., 1999]. Radiocarbon dates on marine cores and terrestrial deposits reveal an extremely rapid deglaciation of the SW Icelandic shelf during the Bølling interstadial C ka, placing the Younger Dryas end moraine complex above present sea level [Geirsdottir et al., 1997; Ingólfsson et al., 1997; Syvitski et al., 1999]. 2. Material and Methods [8] The twin gravity cores JM /1 and JM / 2 (henceforth 1225/1 and 1225/2) (64 54,3 0 N, 29 17,4 0 W) were retrieved from the lower continental slope southwest of Iceland in the northern Irminger Basin at 1683 m water depth [Hald et al., 1996] (Figure 1). We combined the two because magnetic susceptibility (MS) analyses showed that 1225/1 has a slightly higher resolution in its younger part, while 1225/2 penetrated deeper. [9] MS was measured every 1 cm on the longitudinally split core, using a Bartington instrument with a MS2E1 sensor. Bulk sediment samples (1 2 cm intervals) were washed and sieved at >63 mm, >100 mm, >150 mm, >500 mm and >1000 mm, and foraminifera census and stable isotope measurements were performed on the > mm size fraction. Between 100 and 200 individuals were counted for planktonic foraminiferal censuses and benthonic species were registered in order to compute benthonic foraminifera/gram. [10] Stable oxygen and carbon isotope measurements were carried out on the planktonic foraminifera Neogloboquadrina pachyderma sinistral (sin, 6 8 specimens) and the benthonic foraminifera Cibicides wuellerstorfi (Cibicides kullenbergi were used in 10 samples). It is assumed that both C. wuellerstorfi and C. kullenbergi closely record d 13 C values of ambient water and contain an equal disequilibrium (0.64%) ford 18 O measurements [Belanger et al., 1981]. Neither the benthonic nor planktonic foraminifera that were picked for isotope measurements showed any sign of mechanical abrasion that could indicate reworking. The isotope measurements were performed on a Finnigan MAT 251 at the GMS laboratory, University of Bergen, which operates with a reproducibility of 0.07% for d 18 O and 0.06% for d 13 C (reported at 1s. They are calibrated to Pee Dee Belemnite using NBS-19. Grain counts were performed on the >500 mm and >1000 mm size fractions, and were divided into lithic grains and basaltic glass. [11] The chronology is based on 12 AMS radiocarbon dates performed on the planktonic foraminifera N. pachyderma sin, and the position of ash layer I (Vedde ash) and ash layer II in the core (Table 1). We identified the tephra layers by quantitative counts of rhyolitic and basaltic glass shards from the mm size fraction. Geochemical analyses of the volcanic glass shards were carried out by Dr. Haflidi Haflidason, using a standard wavelength dispersal technique on an ARL-SEMQ electronic microprobe at the University of Bergen. Radiocarbon dates were corrected for a marine reservoir effect of 400 years assuming that the dated foraminifera calcified in waters dominated by an Atlantic Water component [Sveinbjörns-

3 HAGEN AND HALD: SURFACE AND DEEP WATER IN THE DENMARK STRAIT 13-3 Figure 1. Map of the studied area, and locations of the analyzed core JM com and others used in this study (007 = HU , 609 = DSDP609, 1726 = PS1726). The surface ocean circulation is generalized [Hopkins, 1991], showing warmer current from the south (bold, filled arrows) and cooler currents from the north (bold, open arrows); NAC = North Atlantic Current, IC = Irminger Current, EGC = East Greenland Current. Deep water masses are indicated by thin arrows (DSOW = Denmark Strait Overflow Water, NSBW = Norwegian Sea Bottom Water, NADW = North Atlantic Deep Water). Locations referred to in the text are indicated; RPL = Rockall Plateau, FSR = Faeroe-Scotland Ridge, IFR = Iceland- Faeroe Ridge, RR = Reykjanes Ridge, CGF = Charles Gibbs fracture zone, LM = Látra Moraine, DS = Denmark Strait, KT = Kangerluggsuaq Trough, SS = Scoresby Sound. The map was based on Online Map Creation, available at dóttir et al., 1993]. Conversion to calibrated years was made using the INTCAL-98 calibration data set for radiocarbon dates younger than 15 cal ka [Stuiver et al., 1998], and using an equation provided by Bard et al. [1992] for the older dates. 3. Results 3.1. Chronology, Core Correlation and Sedimentation Rates [12] Two ash layers are identified (Table 1; Figure 2). The upper one (7 24 cm) has an identical geochemical composition to ash zone I, also termed the Vedde ash layer. This layer has been described previously from marine cores in the North Atlantic, Greenland ice cores, and on land surrounding the Nordic Seas [Mangerud et al., 1984; Jansen, 1987; Kvamme et al., 1989; Norddahl and Haflidason, 1992; Bard et al., 1994; Grönvold et al., 1995; Lacasse et al., 1996; Haflidason et al., 2000]. The Vedde ash layer is dated to Cka[Birks et al., 1996] and cal ka years in the GRIP Greenland ice core [Grönvold et al., 1995] (Figure 2). The lower layer ( cm) corresponds to ash zone II (Figure 2) described from marine cores [e.g., Ruddiman and Glover, 1972; Kvamme et al., 1989] and dated to 52 cal ka cal years in the central Greenland ice core (GRIP) record [Grönvold et al., 1995]. [13] Among the twelve 14 C dates obtained, three were excluded. One of these is the core top date of core 1225/1 which yield an age of 2.38 cal ka (Table 1). This date was omitted since we suspect the upper 5 cm was disturbed during coring, causing a mixed age. The two other dates omitted are from the lower portion of core 1225/2. The dates bracket ash zone II layer at Cka and C ka, respectively. Interpolating between these dates yields an age of ash zone II of C ka, which is about 7 ka too young relative to the ice core record [Grönvold et al., 1995]. A possible explanation is that the two dated samples were contaminated by young carbonate. Contamination by relatively little modern (high 14 C

4 13-4 HAGEN AND HALD: SURFACE AND DEEP WATER IN THE DENMARK STRAIT Table 1. Radiocarbon Dates From the Two Records Used in This Study to Construct an Age Model a Laboratory Reference Core Interval, cm Material Dated 14 C Age, Uncorrected Error, 1s 14 C-400 Calendar Ages Used in Age Model KIA /1 0 1 N. pachyderma sin b 1225/ Vedde ash KIA / N. pachyderma sin Tua / N. pachyderma sin AA / N. pachyderma sin Tua / N. pachyderma sin Tua / N. pachyderma sin Tua / N. pachyderma sin Tua / N. pachyderma sin AA / N. pachyderma sin AA / N. pachyderma sin AA / N. pachyderma sin b 1225/ ash II Tua / N. pachyderma sin b a Vedde ash and ash zone II was additionally used as correlation points and in the age model that were based on extrapolation between dated points. Radiocarbon ages were corrected for a marine reservoir effect of 400 years. b Omitted from the age model. activity) carbonate of dates older than C ka would reduce the reported age by several thousand years [Bowman, 1990]. Alternatively ash zone II could be reworked, e.g., by ice berg rafting. Infinite 14 C ages, e.g., from the last interglacial were not obtained, as the core did not penetrate below MIS 3. We are also aware of the ongoing discussion regarding possible larger reservoir (ventilation) ages of surface water masses north of the Denmark Strait [Voelker et al., 1998], that may also have affected the study area. [14] We interpolated between the dates, assuming constant sedimentation rates. The record captures the period 57 to 6 cal ka, from early MIS 3 to middle MIS 1. Average sedimentation rates vary between >10 cm/ka in MIS 2, to 4 6 cm/ka in MIS 1 and 3 (Figure 3d). [15] The two cores 1225/1 and 1225/2 were spliced to produce one record from the site, using the age model, the MS record, and the planktonic oxygen isotope record (Figure 3). The correlation between the two cores was based on characteristic peaks in the MS records (labeled A K) and the planktonic d 18 O record, (Figures 3a and 3b). Based on this we decided to use the upper 118 cm of core JM /1 and the lower 262 cm (from 370 to 108 cm in core) from core JM /2 interval between. The chronology of each of these cores is based on the AMS dates and ash zones as discussed above. In the following, the data are presented as one record, termed 1225com, either versus a spliced corrected core depth (Figures 4 and 5), or versus calibrated years (Figures 6 8) Lithology [16] Visual examination of the core and X-ray photographs shows that the sediment mainly is composed of grayish olive and gray mud, with various amounts of sand grains and gravel. Grain-size analysis indicates a mud (<63 mm) content of between 40 and 98% (Figure 4a). The CaCO 3 record shows fluctuations of 5% around an average content of about 10% during most of the record, except from a marked increase to 70% in the upper 15 cm, which corresponds to the Pleistocene-Holocene (glacialinterglacial) transition (Figure 4b). Percentage total organic carbon is mainly between 0.1 and 0.3%. [17] The grain count records for the >500 mm and >1000 mm grain-size fractions show an overall parallel trend, so we present only the >500 mm size fraction for simplicity. The sediments contain between 0 50 lithic grains per gram and 0 40 basaltic glass grains/gram (Figure 4d). Background Figure 2. Geochemical analyses of samples from tephra layers identified in cores 1225/1 and 1225/2 plotted as weight percent of FeO*:CaO on a variation diagram. Total iron is expressed as FeO*. The circled areas mark the geochemical characterization of the two main pre-holocene rhyolitic components identified in the study area. The geochemical data for the Vedde ash are based on Mangerud et al. [1984], Kvamme et al. [1989], and for ash zone II, it is based on Sigurdsson [1982]. The depth location of the analyses carried out on the respective cores, shown by sample numbers in parentheses, is shown in the legend.

5 HAGEN AND HALD: SURFACE AND DEEP WATER IN THE DENMARK STRAIT 13-5 values are lower for basaltic glass than lithics (5 and 25, respectively), and both records reveal large fluctuations that are mainly synchronous. We identified 6 distinct peaks in the basaltic grain record (labeled B1 B6 in Figure 4d) between 360 and 20 cm that all are associated with elevated content of lithic grains, except for B3 at 160 cm. A general parallel trend between IRD and magnetic susceptibility is indicated (Figures 4c 4d) Isotope Records [18] Planktonic d 18 O values range from 4.7% during upper MIS 2 to a mean value of 2.0% in early Holocene (Figure 4e). This corresponds to a total glacial-interglacial isotope shift of 2.7%. The shift is characterized by two d 18 O reductions, at cm and 15 cm, respectively. There are several light d 18 O excursions of % amplitude during MIS 3 and 2 (numbered 1 9; Figure 6f ), superimposed on a general increasing trend of 1% between 375 and 60 cm (Figure 4e). [19] Benthonic d 18 O varies from 4.7% during upper MIS 2to2.8% in the early Holocene (Figure 4f ). Most of this 2.0% shift parallels the planktonic d 18 O decrease between 60 and 50 cm. During MIS 3 and 2, the benthonic record shows an overall d 18 O increase of approximately 1.2%, but it is also interrupted by some light isotope oscillations of as much as 0.5% (Figure 4f).The benthonic and planktonic d 18 O are not always exactly in phase. Part of the discrepancy may be due to SST changes. [20] Benthonic d 13 C values vary mainly between values of % and 1.1% (Figure 4g). Two pronounced depletions having d 13 C values of 0.6% are recorded at cm and cm. Both the benthonic d 13 C minima correspond to d 18 O minima in both the planktonic and benthonic records Planktonic Foraminifera [21] The planktonic foraminiferal fauna is dominated by Neogloboquadrina pachyderma sin (>80%) during both MIS 3 and 2 (Figure 5). The percentage of N. pachyderma sin is rapidly reduced to 10% in the Holocene portion of the record (Figure 5a). Other important species include Neogloboquadrina pachyderma dextral, Globigerina quinqueloba, Globigerina bulloides, and Globigerinita glutinata. Although N. pachyderma sin dominates during MIS 3 and 2, species that favor warmer water, such as G. bulloides and G. glutinata [Bé and Tolderlund, 1971], are found in most of the record except between 120 and 160 cm, around 200 cm and around 310 cm (Figures 5b and 5c). [22] Planktonic and benthonic foraminifera/g dry, bulk sediment show an order of magnitude increase during the glacial-interglacial transition around 15 cm (Figures 5d and 5e). During late MIS 3 and MIS 2, both records document low foraminiferal abundance (<2000 planktonics/gram and <30 benthonics/gram). In contrast, middle and early MIS 3 show spikes of planktonic foraminifera that reach >10,000 species/gram, at which levels the benthonic record also, shows elevated values. Although the benthonic foraminiferal assemblage was not counted, we noticed a large increase in the benthonic species Bulimina marginata around cm (Figure 5f ). A similar B. marginata peak at comparable depth and during the same time period is also found east of the Reykjanes Ridge (Figure 1) (H. Ebbesen, personal communication, 2001). 4. Discussion 4.1. Ice Rafting Debris [23] The frequency of minerogenic grains >500 mm has proven to be a valuable tool for elucidating changes in iceberg rafting [Baumann et al., 1995; Elverhøi et al., 1995]. We interpret the ice rafted debris (IRD) peaks (Figures 6a and 6b) to represent increased iceberg rafting during periods when the Greenland Ice Sheet (GIS) and/or the Icelandic Ice Sheet (IIS) may have reached the nearby continental shelves. In addition, increased IRD may reflect glacial retreats, or large surging episodes. [24] We distinguished between basaltic/rhyolitic glass and lithic grains, assuming basaltic glass originates mainly from Iceland, whereas lithic grains mainly have their source in Greenland [Conolly and Ewing, 1965; Jakobsson, 1979]. Parallel maxima in basaltic glass and lithic grains in the core suggest synchronous ice shelf dynamics of the Greenland and Icelandic Ice Sheets around cal ka, cal ka, cal ka and cal ka (Figures 6a and 6b). In addition, there are lithic IRD-peaks at cal ka, cal ka and around 13 cal ka suggesting a more persistent IRD flux from the Greenland Ice Sheet to the site. A basaltic peak, B3 at cal ka, has no counterpart in the lithic record, and thus indicates iceberg rafting mainly from the Icelandic Ice Sheet (Figure 6a) Sea Surface Temperature and Meltwater [25] Based on the planktonic foraminiferal assemblage records (Figure 5), we calculated sea surface paleotemperature (SST) using the SIMMAX program [Pflaumann et al., 1996], applying the modern analogue technique approach [e.g., Overpeck et al., 1985] (Figure 6c). Each sample was compared to 738 modern samples from the Atlantic Ocean, and five analogues were identified having a similarity index of >0.9. SST was calculated for all four seasons at 0 m water depth. The results show parallel trends for all seasons. We present only the summer SST, defined as August October [Pf laumann et al., 1996] since this season represents the main growth period for planktonic foraminifera [Ostermann et al., 1998] in this area. Spero and Lea [1996] showed that G. bulloides, included in the present SST reconstructions, requires different paleotemperature equations depending on its size. This factor has not been considered here. However, the frequency of this species is so low during MIS 2 and 3 that it would have little effect on the calculations. The standard error of the reconstructed temperatures is of the same order as the magnitude of the SST changes. However, the good correlation between percent CaCO 3 and SST (Figure 6) further support the correlations. [26] SST varies from 3.5 C around 26 cal ka, to a Holocene mean of 10 C (Figure 6c). The Holocene SST estimates correspond well both to instrumental measurements (10 12 C) [Malmberg, 1984] and planktonic d 18 O. During the Holocene, planktonic d 18 Oof2.0% indicates SST of about 9 C during foraminiferal calcification, assuming a d 18 O water of 0.15% [Craig and Gordon, 1965b; Shackleton, 1974].

6 13-6 HAGEN AND HALD: SURFACE AND DEEP WATER IN THE DENMARK STRAIT

7 HAGEN AND HALD: SURFACE AND DEEP WATER IN THE DENMARK STRAIT 13-7 Figure 4. Paleoceanographic proxy data versus core depth in core 1225com. The age model (left horizontal axis) is discussed in section 3.1 and is indicated by the solid diagonal curve. (a) Percent silt and clay (<63 mm) in bulk sediment, (b) percent calcium carbonate in bulk sediment, (c) magnetic susceptibility, (d) number of lithic grains basaltic glass (shaded) per gram dry sediment, (e and f ) d 18 Oon N. pachyderma sin Cibicides wuellerstorfi, and (g) d 13 CinC. wuellerstorfi. Marine isotope stages (MIS) are shown. [27] SST fluctuations of C during MIS 2 and 3 can be translated to d 18 O shifts of around % [Shackleton, 1974]. However, the planktonic isotope record contains d 18 O shifts of as much as 1%, and we interpret the light d 18 O excursions to mostly reflect meltwater discharges. Most of these d 18 O excursions, numbered 1 9 (Figure 6f ) can be correlated to well documented meltwater events in North Atlantic marine cores during the last glacial cycle (Figure 6e) [Elliot et al., 1998, and references therein]. [28] In core 1225com, SST warming was either slightly leading, or synchronous to the onset of each of the meltwater events. Approximately 2 ka after each meltwater Figure 3. (opposite) (a and b) Magnetic susceptibility (MS), d 18 O N. pachyderma sin and radiocarbon dates (in kiloyears; cf. Table 1) in core JM /1 (Figure 3a) and JM /2 (Figure 3b). Thick vertical line indicates where the two records were spliced, and the labels a k mark synchronous MS peaks. (c) Core interval used from each core. (d) Composite core depth versus calibrated (cal) ka converted 14 C dates versus depth.

8 13-8 HAGEN AND HALD: SURFACE AND DEEP WATER IN THE DENMARK STRAIT Figure 5. Paleoceanographic proxy data versus core depth in core 1225com. The age model (left horizontal axis) is discussed in section 3.1 and is indicated by the solid diagonal curve. (a c) Down core distribution of the principal planktonic foraminiferal species in percent relative to total planktonic fauna, (d) number of planktonic foraminifera/gram bulk sediment, (e) benthonic foraminifera/gram bulk sediment, (f ) B. marginata/ram bulk sediment. peak, the IRD records a maximum that partly occurred during, and partly after, the termination of each meltwater event (Figures 6a, 6b, and 6f). This succession seems fairly robust between 57 and 25 cal ka (e.g., in MIS 3). Apparently, the glacial ice sheet responded to these circulation changes, as indicated by the IRD maxima Bottom Water Characteristics [29] The epibenthonic species C. wuellerstorfi is found to be a reliable indicator of bottom water d 13 C[Mackensen et al., 1993]. Based on down core variations in both d 13 C and d 18 O measured on this species, we identify three different bottom water types, related to (1) North Atlantic Deep Water (NADW), (2) glacial North Atlantic Intermediate Water (GNAIW), and (3) Brine formation and/or Southern Ocean Water (SOW), respectively. The interpretations of these water masses as given below should read with some caution, as we have not taken into account global isotope changes. However, such changes are very difficult to evaluate on the short timescales discussed here.

9 HAGEN AND HALD: SURFACE AND DEEP WATER IN THE DENMARK STRAIT 13-9 Figure 6. Selected proxy data versus calibrated (cal) years from core 1225com and GISP2 ice core. To suppress the scatter in the curves, we smoothed the 1225com records by using a three-point running average. (a) Basaltic glass grains/gram (peaks discussed in section 3.2 are denoted B1 B6), (b) Lithic grains/gram, (c) summer, sea surface temperature (SST) based on the SIMMAX modern analogue technique [Pflaumann et al., 1996], (d) percent CaCO 3 in bulk sediment, (e) oxygen isotope record from GISP2 ice core, (f ) d 18 O N. pachyderma sin, Bm indicate position of the B. marginata spike from Figure 6f (g) d 18 O and (h) d 13 C C. wuellerstorfi NADW [30] The most frequent bottom water type is characterized by d 13 C between 0.9% and 1.3% and d 18 O between 3.4% and 4.8%. These values are typical for NADW [Oppo and Lehman, 1993] and appear to have dominated during the Holocene, and in several shorter periods during MIS 2 and 3. The formation of NADW, known from instrumental monitoring, is mainly related to open ocean thermohaline convection both in the Norwegian-Greenland seas and in the Labrador Sea.

10 13-10 HAGEN AND HALD: SURFACE AND DEEP WATER IN THE DENMARK STRAIT Figure 7. Correlation between our record and the DSDP609 core from the North Atlantic (Figure 1) for various paleoceanographic proxies [Bond et al., 1992]. The 1225com records are smoothed using the running average (three adjacent samples) method GNAIW [31] The second important bottom water type is characterized by enriched benthonic isotope values within % d 18 O and % d 13 C. We associate this water mass with GNAIW previously described from the North Atlantic by, e.g., Kroopnick [1980] and Oppo and Lehman [1993]. GNAIW is associated with the glacial periods, and should therefore contain an ice-volume-affected d 18 O

11 HAGEN AND HALD: SURFACE AND DEEP WATER IN THE DENMARK STRAIT (+1.2%) signal. This water mass is probably additionally enriched in d 18 O due to having a convection site farther south and thus of higher salinity compared to NADW. GNAIW is further interpreted to contain a heavy d 13 C signal of around 1.5% due to its high productive source from the subtropical Atlantic [Oppo and Lehman, 1993]. GNAIW and NADW alternately bath the core site throughout most of MIS 2 and MIS 3. [32] During periods of the last glacial, it has been suggested that GNAIW occupied the water column at intermediate depths above 2000 m [Boyle and Keigwin, 1987; Oppo and Lehman, 1993]. This was explained by a shift in the convection from deep to intermediate mode during the colder phases of the last glaciation Brine/SOW [33] The third bottom water mass is marked by a strong depletion in benthonic d 13 C (values are typically between 0.8 and 0.4%) and a depletion in benthonic d 18 O( %). This water mass dominated only during two short intervals during the last deglaciation (16 14 cal ka), and around 54 cal ka (Figure 6h). The formation of this bottom water can be explained either by (1) a halt in thermohaline convection in the North Atlantic that results in an increased influence of Southern Ocean water (SOW), or (2) by brine formation. 1. SOW is known to be depleted in benthonic d 13 C due to its increased nutrient content and longer residence time in the ocean [e.g., Kroopnick, 1980; Duplessy et al., 1988]. Glacial d 13 C/d 18 O values during the last glacial in the South Atlantic core RC11-8B are 0,8%/5.0% [Charles and Fairbanks, 1992]. Oppo and Lehman [1993] argued that SOW in the North Atlantic was entrained by NADW to a 50/50 proportion, leaving SOW in the northern Irminger basin with a d 13 C signature of about %, that is somewhat lighter than observed in the present data. The depletion in benthonic d 18 O associated with this water mass in the Nordic Seas is suggested to reflect increased bottom water temperatures [Labeyrie et al., 1995; Rasmussen et al., 1997]. This is supported by the distribution of benthonic foraminifera during the last glacial in the Faeroe-Shetland channel [Rasmussen et al., 1997]. 2. The alternative explanation for depleted benthonic isotope values (both d 18 O and d 13 C) is bottom water produced through brine formation. Brine formation occurs where surface water, through salt rejection, becomes more saline (and thereby more dense), and is associated with insignificant isotopic fractionation leaving the water masses almost unchanged in isotopic composition [Craig and Gordon, 1965a]. Hence during periods of depleted d 13 C and d 18 O values at the surface, for example due to meltwater discharges, brine may carry an equally light signal to depth. This process is well known from the Antarctic [Jacobs and Fairbanks, 1985] and the Barents Sea [Midttun, 1985] today. It also contributed to the composite water mass NADW observed in paleorecords in the Nordic Sea and North Atlantic [Veum et al., 1992; Dokken and Jansen, 1999; Vidal et al., 1997; Kreveld et al., 2000]. [34] We suggest this bottom water to be due to brine formation. Such a process may explain the synchronous depletion seen in both d 13 C and d 18 O between and 54 cal ka. Further, brine formation may explain the parallel d 18 O depletion observed at the surface and in the bottom water. This depletion is much too rapid to be explained by a global d 18 O shift. If the depletion was mainly forced by temperature, it would require an increase of more than 5 C [Shackleton, 1974]. There is no evidence to support such a large warming at cal ka and around 54 cal ka (Figure 6). This suggests the bottom water d 18 O depletion most likely image the surface water freshening, and therefore reflects a salinity signal. A transfer of d 18 O depleted surface water to bottom is possible through brine formation. 5. Correlations 5.1. Atmospheric Temperatures and Sea Surface Conditions [35] When comparing reconstructed SST to atmospheric temperatures as reflected by the Greenland Ice Sheet Project 2 (GISP2) d 18 O record [Grootes et al., 1993], a correlation is evident (Figures 6c and 6e). Both larger and smaller interstadials from the ice core record can be correlated to periods of increased SST in the marine record of core 1225com. However, there is an offset of 2 3 ka in the older portion (>35 ka) of the core. This is probably due to the less accurate dating of sediments beyond the radiocarbon method. [36] The H events reflect synchronous iceberg rafting, a depletion in planktonic d 18 O and lowered SST [Heinrich, 1988; Bond et al., 1992], and is correlated to various climate proxy data indicative of cooling throughout the North Atlantic region [Broecker et al., 1992; Grimm et al., 1993; Andrews et al., 1994, 1998; Keigwin and Lehman, 1994; Dowdeswell et al., 1995; Auffret et al., 1996; Zahn et al., 1997]. We correlated core 1225com to H events by assuming that the onset of each H event occurred during major SST coolings (Figure 6, vertical lines). From the ice core record, the onset for each H event is seen to occur during peak stadials (cooling) [Bond et al., 1993]. This infers onset of H events in our record at approx. 18 cal ka (H1), 24 cal ka (H2), 30 cal ka (H3), 37.5 cal ka (H4), and 44.5 cal ka (H5). These ages compare well to reported ages of the H events [e.g., Vidal et al., 1997, and references therein]. We also place H event 6 before a SST increase at 56 cal ka, arguing that this reflects the warming into interstadial 17 which is where H event 6 took place [Bond et al., 1993]. A SST increase and a salinity decrease in the 1225com record (Figure 6) follow each H event. [37] We compared in detail H events 3 and 4 in core Deep Sea Drilling Project (DSDP) 609, [Bond et al., 1992] to core 1225com (Figure 7). DSDP 609 lies below the main route of the North Atlantic Current, and should therefore record changes in northward heat transport. In addition, the DSDP 609 core is located in the middle of the Ruddiman Belt [Ruddiman, 1977; Heinrich, 1988; Bond et al., 1992]; hence correlation between DSDP 609 and core 1225com has the potential to reveal possible temporal and spatial differences between climatic successions related to the H events. [38] The correlation clearly demonstrates the delay of the IRD events in the Denmark Strait, relative to the North

12 13-12 HAGEN AND HALD: SURFACE AND DEEP WATER IN THE DENMARK STRAIT

13 HAGEN AND HALD: SURFACE AND DEEP WATER IN THE DENMARK STRAIT Atlantic Heinrich events. Another important difference between the two records is that SST increase and SSS reduction are more or less synchronous in 1225com, while in DSDP 609, warming lags behind freshening. Further, in DSPD 609 the Heinrich events are more or less synchronous to freshening indicated by planktonic d 18 O (Figure 7). In contrast, in core 1225com IRD lags behind the warming and freshening by 2 cal ka. Thus the Iceland and SE Greenland Ice Sheets apparently responded to increased SST by ice berg rafting, while the H events from the Laurentide Ice Sheet are linked to cooling. This difference is not fully understood and should be further explored. [39] The correlation also demonstrates that the IRD flux associated with the H event 3 is relatively small in DSDP 609, while in 1225com H event 3 is followed by IRD event B3, dominated by basaltic glass assumed to be of an Icelandic origin. Core HU close to the 1225com core site, and PS1726 north of the Denmark Strait (Figure 1), both show a relatively small IRD event of assumed Greenland origin [Stein et al., 1996; Andrews et al., 1998] that can be correlated to B3. Since only the elevated content of basaltic glass is associated with H event 3 in core 1225com, we suggest the Iceland Ice Sheet was the main IRD contributor in the Denmark Strait some 2000 years after H event 3 (Figure 7). It has been suggested that H event 3 was a larger event in the eastern Nordic Seas indicating a larger discharge from the Fennoscandian Ice Sheet [Dokken and Jansen, 1999]. [40] The discrepancies documented between cores 1225com and DSPD 609 indicate a dynamic system linked to the H events in different regions of the North Atlantic. Studies of planktonic foraminifera and coccoliths in the Nordic Seas indicate that Atlantic Water reached as far north as 78 N several times during the last glacial-interglacial, during the so-called High Productive (HP) zones [Dokken and Hald, 1996] Bottom Water, SST, and Heinrich Events [41] A more direct expression of the various bottom water masses in the 1225com record can be made by a detailed comparison of benthonic d 18 O and d 13 C. In Figure 8a, we plotted d 18 O versus d 13 C and we labeled each d 13 C/ d 18 O data point with a quadrant numbered (hereafter termed q number) from 1 to 88 (after model from T. Dokken et al., manuscript in preparation, 2002). Thus the q number combine the benthonic d 18 O and d 13 C at any given stratigraphic level and the numbers are further grouped into the different main bottom water masses (GNAIW, NADW, Brine) on the basis of their isotopic composition as discussed above. The q number are plotted versus age (Figure 8c) and compared to SST in the 1225com record and the Heinrich events (Figure 8b). H events are placed during peak SST minima, as discussed earlier (Figure 8b). In general, the deep water was characterized of NADW alternating with GNAIW. Most H events are associated with a synchronous transition from GNAIW to NADW. The time resolution in core 1225com does not allow determining whether the shift to NADW occurred just prior to, synchronously with, or immediately after the H event. The shift from GNAIW to NADW may have been caused by the lowered surface ocean salinity that took place during the H events in the North Atlantic. If H events led to a temporary halt of GNAIW production, how was NADW production possible? We have no good answer, although there may possibly have been a corridor of northward transport of Atlantic Water in the eastern North Atlantic, as indicated by Lassen et al. [1999], that conveyed saline water to the northern Nordic Seas [Dokken and Hald, 1996] where open ocean thermohaline circulation could occur. [42] Brine formation is found during two intervals, around 54 cal ka and cal ka, respectively. However, transitional values like samples in quadrants 28 to 35, may belong either to GNAIW or NADW. Likewise, values plotted in quadrants may equally reflect either brine or NADW. However, regardless of what deep water mode the respective samples belong to, the succession of deep water mode indicated in Figure 8c, would not change. [43] We also hypothesize on the possibility of mixed GNAIW and brine. A 50/50 mixing would actually yield d 13 C/d 18 O values equal to NADW. Hence we suggest there may be a potential pitfall in interpreting a mixed GNAIW/ brine signal as NADW. [44] A relative long period with NADW-dominated deep water occurred between 37 and 33 cal ka following the H event 4. This indicates a long period of warming during the middle Weichselian period, indicated also by the ice core data (interstadial 8). This correlates with the Ålesund interstadial in Fennoscandian [Larsen et al., 1987]. 6. Conclusions [45] Based on multiproxy analysis of core 1225com south of the Denmark Strait sill, we offer new information on the climate history of the region spanning 57 to 6 cal ka, which can be summarized as follows: 1. Existence of an oscillating heat conveyor around Iceland (e.g., strength of Irminger Current) that covaried with atmospheric temperatures measured in Greenland ice core during MIS 2 and Three modes of deep water producing ventilated deep water in the north Atlantic during MIS 3 and 2 are recorded south of the Denmark Strait. Between 57 and 6 cal ka, ventilation was either sustained by glacial North Atlantic Intermediate Water or North Atlantic Deep Water, except from two periods (55 53 cal ka and cal ka) when deepwater production was accomplished by brine formation. Figure 8. (opposite) Benthonic d 18 O versus d 13 C measurements in core 1225com. Each data point is given a quadrant number (q number) and is further grouped into main bottom water masses GNAIW, NADW, and brine on the basis of their isotopic composition as discussed in section 5.2. (a) Small panel in lower right indicates relative position of major deep water masses. (b) Summer sea surface temperature (SST), plotted versus cal years, adopted from Figure 6c. (c) The q numbers defined in Figure 8a plotted versus cal years. Vertical bars (H1 H6) indicate the position of Heinrich events.

14 13-14 HAGEN AND HALD: SURFACE AND DEEP WATER IN THE DENMARK STRAIT 3. Repetitive episodes of iceberg rafting took place in the Denmark Strait, but were delayed relative to the North Atlantic Heinrich events by 2 ka. 4. Marine evidence suggests the Greenland Ice Sheet reached the shelf break between cal ka and cal ka. Our data indicate a possible advanced position of the Icelandic Ice Sheet prior to the Basaltic glass events (e.g., >55 cal ka, >44 cal ka, >37 cal ka, >28 cal ka, >16 cal ka and >12 cal ka). 5. Most of the Denmark Strait rafting events contains both an Icelandic and Greenland component, except the B3 event (28 cal ka), following after H-3 that seems to contain only iceberg-rafted debris from Iceland. [46] Acknowledgments. Funding was provided from the University of Tromsø, and the Research Council of Norway (grant /720) and the Roald Amundsen Research Centre. Mary Raste, Marit Berntsen, Rune Sørås and Odd Hansen provided laboratory services. Haflidi Haflidason carried out geochemical analyses of the volcanic glass shards. Jan P. Holm helped with computer drawings. Anne Jennings, John T. Andrews, Mikie Smith, Donald Barber, Jan Sverre Laberg, Torbjørn Dahlgren and Trond Dokken discussed earlier drafts of the manuscript, and David Anderson helped with the sea surface temperature calculations. To all these persons and institutions we offer our sincere thanks. References Andrews, J. T., H. Erlenkeuser, K. Tedesco, A. E. Aksu, and A. J. T. Jull, Late Quaternary (stage 2 and 3) meltwater and Heinrich events, northwest Labrador Sea, Quat. Res., 41, 26 34, Andrews, J. T., A. E. Jennings, T. Cooper, K. M. Williams, and J. Mienert, Late Quaternary paleoceanography of the North Atlantic margins, in Late Quaternary Paleoceanography of the North Atlantic Margins, GeoSoc. Spec. Publ., edited by J. T. Andrews et al., pp , Geol. Soc. Of Am., Boulder, Colo., Andrews, J. T., T. A. Cooper, A. E. Jennings, A. B. Stein, and H. Erlenkeuser, Late Quaternary IRD events on the Denmark Strait/southeast Greenland continental slope (65 N): Related to North Atlantic Heinrich events?, Mar. Geol., 149, , Andrews, J. T., et al., The N and W Iceland Shelf: Insights into last glacial maximum ice extent and deglaciation based on acoustic stratigraphy and basal radiocarbon AMS dates, Quat. Sci. Rev., 19, , Auffret, G. A., A. Boelaert, C. Vergnaud-Grazzini, C. Müller, and R. Kerbrat, Identification of Heinrich layers in core KS 01 northeastern Atlantic (46 N, 17 W): Implication for their origin, Mar. Geol., 131, 5 20, Bard, E., M. Arnold, and B. Hamelin, Present status of the radiocarbon calibration for the last Pleistocene, GEOMAR Rep. 15, pp , Kiel, Germany, Bard, E., M. Arnold, J. Mangerud, M. Paterne, L. Labeyrie, J. Duprat, M.-A. Mélières, E. Sønstegaard, and J.-C. Duplessy, The North Atlantic atmosphere-sea surface 14 C gradient during the younger Dryas climatic event, Earth Planet. Sci. Lett., 126, , Baumann, K.-H., K. S. Lackschewitz, J. Mangerud, R. F. Spielhagen, T. C. W. Wolf-Welling, R. Henrich, and H. Kassens, Reflection of Scandinavian ice sheet fluctuations in Norwegian Sea sediments during the past 150,000 years, Quat. Res., 43, , Bé, A. W. H., and D. S. Tolderlund, Distribution and ecology of living planktonic foraminifera in surface waters of the Atlantic and Indian Oceans, in The Micropaleontology of Oceans, edited by B. M. Funnel and W. R. Riedel, pp , Cambridge Univ. Press, New York,1971. Belanger, P. E., W. B. Curry, and R. K. Mathews, Core top evaluation of benthonic foraminiferal isotopic ratios for paleoceanographic interpretations, Paloegeogr. Paleoclim. Paleoecol., 33, , Birks, H. H., S. Gulliksen, H. Haflidason, J. Mangerud, and G. Possnert, New radiocarbon dates for the Vedde ash and the Saksunarvatn ash from western Norway, Quat. Res., 45, , Bond, G., et al., Evidence for massive discharge of icebergs into the North Atlantic Ocean during the last glacial period, Nature, 360, , Bond, G., W. Broeker, S. Johnsen, J. McManus, L. Labeyrie, J. Jouzel, and G. Bonani, Correlations between climate records from North Atlantic sediments and Greenland ice, Nature, 365, , Bowman, S., Radiocarbon Dating (Interpreting the Past), British Museum Publ. Ltd., London, Boyle, E. A., and L. D. Keigwin, Deep circulation of the North Atlantic over the last 200,000 years: Geochemical evidence, Science, 218, , Boyle, E. A., and L. Keigwin, North Atlantic thermohaline circulation during the past 20,000 years linked to high-latitude surface temperature, Nature, 330, 35 40, Broecker, W., The great ocean conveyor, Oceanography, 4(2), 79 89, Broecker, W., Thermohaline circulation, the Achilles heel of our climate system: Will man-made CO 2 upset the current balance?, Science, 278, , Broecker, W., G. Bond, M. Klas, E. Clark, and J. McManus, Origin of the northern Atlantic s Heinrich events, Clim. Dyn., 6, , Charles, C. D., and R. G. Fairbanks, Evidence from Southern Ocean sediments for the effect of the North Atlantic deep-water flux on climate, Nature, 355, , CLIMAP Project Members, Seasonal reconstructiions of the Earth s surface at the last glacial maximum, Geol. Soc. of Am. Map and Chart Ser. MC-36, Boulder, Colo., Conolly, J. R., and M. Ewing, Pleistocene glacial marine zones in North Atlantic deep-sea sediments, Nature, 208, , Craig, H., and L. I. Gordon, Stable isotope in oceanographic studies and paleotemperatures, in Stable Isotopes in Oceanographic Studies and Paleotemperatures, edited by T. Tongiorgi, pp , Cons. Naz. delle Ric. Lab. de Geol. Nucl., Pisa, 1965a. Craig, H., and L. I. Gordon, Deuterium and oxygen 18 variation in the ocean and the marine atmosphere, in Stable Isotopes in Oceanographic Studies and Paleotemperatures, edited by T. Tongiorgi, pp , Cons. Naz. delle Ric. Lab. de Geol. Nucl., Pisa, 1965b. Dansgaard, W., et al., A new Greenland deep ice core, Science, 218, , Dickson, R. R., E. M. Gmitrowicz, and A. J. Watson, Deep-water renewal in the northern North Atlantic, Nature, 344, , Dokken, T. M., and M. Hald, Rapid climatic shifts during isotope stages 2 4 in the polar North Atlantic, Geology, 24(7), , Dokken, T. M., and E. Jansen, Rapid changes in the mechanism of ocean convection during the last glacial cycle, Nature, 401, , Dowdeswell, J. A., M. A. Maslin, J. T. Andrews, and I. N. McCave, Iceberg production, debris rafting, and the extent and thickness of Heinrich layers (H-1, H-2) in North Atlantic sediments, Geology, 23(4), , Duplessy, J. C., N. J. Shackleton, R. G. Fairbanks, L. Labeyrie, D. Oppo, and N. Kallel, Deepwater source variations during the last climatic cycle and their impact on the global deepwater circulation, Paleoceanography, 3, , Eirikson, J., L. A. Simonarson, K. L. Knudsen, and P. Kristensen, Fluctuations of the Weichselian ice sheet in SW Iceland: A glaciomarine sequence from Sudurnes, Seltjarnarnes, Quat. Sci. Rev., 16, , Elliot, M., L. Labeyrie, G. Bond, E. Cortijo, J.-L. Turon, N. Tisnerat, and J.-C. Duplessy, Millennium timescale iceberg discharges in the Irminger Basin during the last glacial period: Relationship with the Heinrich events and environmental settings, Paleoceanography, 13, , Elverhøi, A., E. S. Andersen, T. Dokken, D. Hebbeln, R. Spielhagen, J. I. Svendsen, M. Sørflaten, A. Rørnes, M. Hald, and C. F. Forsberg, The growth and decay of the Late Weichselian ice sheet in western Svalbard and adjacent areas based on provenance studies of marine sediments, Quat. Res., 44, , Funder, S., C. Hjort, J. Y. Landvik, S.-I. Nam, N. Reeh, and R. Stein, History of a stable ice margin-east Greenland during the middle and upper Pleistocene, Quat. Sci. Rev., 17, , Geirsdottir, A., J. Hardardottir, and J. Eirikson, The depositional history of the younger Dryas; Preboreal Budi moraines in south-central Iceland, Arct. Alp. Res., 29, 13 23, Grimm, E. C., G. L. Jacobson, W. A. Watts, B. C. S. Hansen, and K. A. Maasch, A 50,000-year record of climate oscillations from Florida and its temporal correlation with the Heinrich events, Science, 261, , 1993.