A possible record of the Younger Dryas event in deep-sea sediments of the Southern Ocean (Pacific sector)

Palaeogeography, Palaeoclimatology, Palaeoecology 198 (2003) 265^278 www.elsevier.com/locate/palaeo

A possible record of the Younger Dryas event in deep-sea sediments of the Southern Ocean (Paci¢c sector) C. Morigi a;b; , L. Capotondi b , F. Giglio b , L. Langone b , M. Brilli c , B. Turi c , M. Ravaioli b a

Dipartimento di Scienze del Mare, Universita' Politecnica delle Marche, Via Brecce Bianche, 60131 Ancona, Italy b Istituto per la Geologia Marina ^ CNR, Via Gobetti 101, Bologna, Italy c Centro di Studio per il Quaternario e l’evoluzione ambientale ^ CNR, P.le Aldo Moro 5, Rome, Italy Received 31 January 2002; accepted 3 March 2003

Abstract The oxygen isotope record combined with radiocarbon dating from two deep-sea cores collected along a transect between New Zealand and the Ross Sea are used to establish a reliable chronostratigraphy for the last 14 kyr. After an integrated geochemical and micropaleontological analysis in this timeframe we detected a cooling interval dated between 12.5 cal kyr BP and 11.4 cal kyr BP. The age control suggests that this event started 1.5 kyr after the onset of the Antarctic Cold Reversal previously observed in several Antarctic ice cores. We infer that the observed cool event corresponds to the Younger Dryas event defined in Northern Europe. This suggests that climate change recorded in this sector of the Southern Hemisphere still shows some synchronicity with Northern Hemisphere variations and that the decoupling of climate change between the two hemispheres likely occurred south of the Polar Front. 9 2003 Elsevier B.V. All rights reserved. Keywords: foraminifera; oxygen isotope; Younger Dryas; Antarctic Cold Reversal; Southern Ocean; Polar Front

1. Introduction The Southern Ocean is a key area for understanding global hydrologic conditions and climatic change through past glacial and interglacial cycles. Although this area has an important role in global ocean circulation and plays a decisive role in climatic change, its paleoceanographic evolution during the last deglacial is still poorly

* Corresponding author. Tel.: +39-071-2204087; Fax: +39-071-2204650. E-mail address: [email protected] (C. Morigi).

understood. On the other hand, the climatic evolution of the Antarctic region during the last deglaciation is relatively well-known because of intensive studies made on ice cores drilled at di¡erent sites on the continent (e.g. Jouzel et al., 1995; Sowers and Bender, 1995; Blunier et al., 1997). The last deglaciation (20^10 kyr BP) contains a sequence of abrupt climatic changes. Warming started around 20^18 kyr BP and was interrupted in the Southern Hemisphere by the Antarctic Cold Reversal at about 14 kyr BP. This cold period lasted for about 1.5 kyr; subsequently, the warming trend continued until the present. In the Northern Hemisphere, the warm-

0031-0182 / 03 / $ ^ see front matter 9 2003 Elsevier B.V. All rights reserved. doi:10.1016/S0031-0182(03)00404-8

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ing trend was frequently interrupted by Dansgaard/Oeschger events, warm^cold oscillations spaced 1.5 kyr, characterized by abrupt terminations (Bond et al., 1993; Alley and Clark, 1999). The most recent of these sharp events was the Younger Dryas that ended approximately 11.5 kyr BP (Alley and Clark, 1999) and lasted for 1 kyr. A comparison of the climatic record from the Greenland ice core (GISP II) with the Vostok and Byrd records (Antarctica) reveals that in the southern continent deglaciation started 3 kyr earlier than in the Northern Hemisphere (Sowers and Bender, 1995). The onset of the Antarctic Cold Reversal, as shown by the isotope data from the Byrd ice core (West Antarctica) as well as in other East Antarctic ice core records (Jouzel et al., 2001), started about 1.5 kyr earlier than the onset of the Younger Dryas event recorded in the Northern Hemisphere (Sowers and Bender, 1995). Results obtained from the studies of ice cores con¢rm that the deglaciation pattern is asynchronous between Greenland and Antarctica (Sowers and Bender, 1995; Blunier and Brook, 2001). The correlation between climatic events recorded in deep-sea cores and those of the ice cores is di⁄cult because of the di¡erent time resolution, which is higher in ice cores than in marine sediments. A further di⁄culty is the application of the appropriate reservoir correction of the 14 C dates on marine samples. The reliability and the correlation potential of 14 C dates are in fact limited by uncertainty regarding the appropriate reservoir correction. In order to compare these deep-sea series with the ice records, the radiocarbon ages must be converted to calendar years. The marine samples are usually several hundred years older than their terrestrial counterparts, and the correction value varies in the di¡erent areas because of the complexity of ocean circulation (Ostlund and Stuiver, 1980; Bard, 1988; Stuiver et al., 1998b). Few studies exist on calcareous microfossils at high latitudes in the Paci¢c sector of the Southern Ocean, mainly because of the very scarce carbonate records from abyssal sites. In this paper, we report the results of the down-core micropaleon-

tological and geochemical analyses of two deepsea cores, collected at the latitude of the Polar Front (PF) during the oceanographic cruise of the R/V Italica in the Antarctic summer 1995^ 1996 (Ravaioli et al., 1995; Langone and Marozzi, 1996). The aim of this paper is to compare the climate record found in the Southern Ocean sediments with ice core records of both hemispheres. In particular, we focus our attention on a cooling event in order to verify if it is equivalent to the Younger Dryas event.

2. Oceanographic setting The Antarctic Circumpolar Current (ACC) is the main oceanographic structure of the Southern Ocean. This current £ows from west to east around Antarctica. In the Paci¢c sector, Antarctic surface waters moving northward sink rapidly below sub-Antarctic waters. Prominent features in the wind-driven ACC are the oceanic fronts, which separate water masses with di¡erent hydrological characters. These are the Sub-Antarctic Front (SAF), the PF, and the Southern Boundary (SB) (Orsi et al., 1995; Russo et al., 1999; Rintoul and Bullister, 1999). The PF shifts seasonally from 59‡50P to 63‡00PS, showing a general tendency to a southward retreat from austral spring to austral summer (Russo et al., 1999). The PF marks an important boundary in terms of air^sea exchange, primary productivity, and downward biogenic £uxes. The surface temperature of the upper layer of this front is lower than 2‡C at depths shallower than 200 m (Orsi et al., 1995; Belkin and Gordon, 1996). The SB has been located between 63‡40PS and 64‡30PS. Its position appears to be less in£uenced by seasons than by geographic location. In fact it has apparently been topographically locked on the northern £ank of the Paci¢c^Antarctic Ridge (Russo et al., 1999) (Fig. 1). The subsurface temperature at its northern limit is Tmin 6 0‡C and T s 1.8‡C at depths s 500 m (Orsi et al., 1995). The SB is the southern end of the Upper Circumpolar Deep Water; an upward £ux of nutrients correlative with this front should increase regional primary production (Tynan, 1998).

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267

180°

A

°E 160

-160

°W

New Zealand

-40 °S

°S -40

-50 °S

°S -50

PF

157

-60

17

°S -60

°S

-70

°S

-70

°S

Ross Sea -80

°S -80

°S

B South

North New Zealand

Antarctica Antarctic Polar Front Zone

Depth (km)

Sub-Antarctic Zone NSAF

0

SSAF

Seasonal Ice Zone

PF

SB

Ross Sea

1 3

Scott Seamount

Anta96-17 Anta95-157

2 Campbell Plateau

4 5 6 46°S

Pacific Antarctic Ridge 48°

50°

52°

54°

56°

58°

60°

62°

64°

66°

68°

70°

72°S

Latitude Fig. 1. (A) Location of cores in the Southern Ocean. (B) N^S section of the study area and core locations. The variation of position of the main hydrological features in the summer season, as documented by Russo et al. (1999), is shown. NSAF, Northern Sub-Antarctic Front; SSAF, Southern Sub-Antarctic Front; PF, Polar Front; SB, Southern Boundary of the ACC. Bathymetric data acquired during the XIII Italian Expedition in Antarctica (1997^1998).

3. Materials and methods The two gravity cores selected for this study are located in the Paci¢c sector of the Southern

Ocean. In particular core 157 was collected in the Antarctic Polar Front Zone (175‡17.34PE, 62‡05.95PS, 4000 m water depth) and core 17 was collected south of the southernmost summer

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Fig. 2. Lithology and CaCO3 wt% content of cores 157 and 17.

position of the PF (177‡14.71PE, 63‡14.68PS, 4260 m water depth) (Fig. 1). The cores were X-radiographed and split into halves. X-ray analysis and visual inspection excluded the presence of hiatuses or postdepositional disturbances. Sediments were mostly composed of biosiliceous clay alternating with calcareous and pelagic clay (Fig. 2). In the core sediment diatom- and foraminifera-bearing sections alternated with sections extremely rich in diatoms and radiolarians and almost barren of foraminifera. We report the results of the study of the upper 40 cm of core 157 and the upper 20 cm of core 17 (foraminifera-bearing section). We limited our analysis to these intervals because the calcium carbonate content in the lower part of the cores was too low to establish a continuous oxygen isotope stratigraphy (Fig. 2).

3.1. Geochemical analyses Carbonate content was obtained by means of a Fisons NA2000 Element Analyzer and calculated as follows : %CaCO3 = (%Ctot 3%Corg )U8.33. Samples for isotope analysis were collected every centimeter from the two cores and oxygen isotope measurements were performed at the Isotope Laboratory, University of Rome by means of a Finnigan MAT 252 mass spectrometer with a KIEL II carbonate preparation device. On average, 40 specimens of Neogloboquadrina pachyderma (sinistral) were picked from the s 150-Wm size fraction. This species was chosen because of its continuous presence along the investigated interval. Specimens were ultrasonically cleaned in distilled water and carefully crushed to release potential sediment in¢lling. The MC200 and NBS-19 carbonate standards were used to calibrate the

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269

Table 1 Core samples Depth (cm)

Sample

Core ANTA95-157 4^6 Mixed planktonic 14^15 Neogloboquadrina pachyderma 22^23 Globorotalia in£ata 29^30 Globigerina bulloides 36^38 Mixed planktonic Core ANTA96-17 0^2 Mixed planktonic 10^11 Neogloboquadrina pachyderma 13^14 Neogloboquadrina pachyderma 16^18 Mixed planktonic

Conventional (yr BP)

14

C age

Calibrated age Maximum and minimum of cal age (yr BP) (yr BP)

1 380 R 310 6 220 R 35 8 880 R 55 10 500 R 160 10 500 R 50

520 6 180 8 920 10 980 10 980

770^260 6 270^6 080 8 980^8 840 11 160^10 600 11 150^10 630

5 970 R 170 11 150 R 60 11 550 R 130 12 350 R 360

5 890 11 820 12 740 13 400

6 110^5 660 12 280^11 500 12 890^12 340 13 820^13 000

Calibrated ages were calculated by applying vR = 462 R 75 yr (Reimer, 2000) and using the software CALIB 4.3 (Stuiver and Reimer, 1993). The calibration ages are rounded to the nearest 10-year interval.

mass spectrometer to the international PDB scale. The reproducibility of these measurements is R 0.10x (1c). The isotope values are reported as delta units per mil on the conventional PDB scale. 14 C accelerator mass spectrometer (AMS) analyses were performed on nine samples from the two cores (Table 1). For each measurement about 1000 specimens of planktonic foraminifera were picked out (where possible a monospeci¢c sample was prepared) in the s 150-Wm size fraction, washed in an ultrasonic bath and analyzed at the AMS laboratory of the Woods Hole Oceanographic Institute. 3.2. Micropaleontological analyses Foraminiferal analyses were conducted on the same samples used for geochemical analysis. All samples were washed through 63- and 150-Wm sieves and each fraction weighed. The foraminiferal species were then identi¢ed and counted from a subsample of the 150-Wm standard aliquot; the 63^150-Wm fraction was stored for future study. A variable number (minimum 93, maximum 624; mean 327 specimens) of planktonic foraminifera tests was counted. The total sample was counted for benthic foraminifera because of the low frequency in the residue. All data were later converted to percentages and the planktonic and benthic foraminifera percentages were then calcu-

lated. The diversity index (Simpson index) was calculated using a ‘BioDiversity’ program (zNatural History Museum, London/Scottish Association of Marine Sciences).

4. Results 4.1. Age assessment An age^depth model for each core was developed based on nine AMS 14 C ages and by correlation between the oxygen isotope records of the two cores (Table 1). The 14 C ages were corrected using a regional mean value of vR = 462 R 75 (Gordon and Harkness, 1992). AMS 14 C ages were calibrated with the CALIB 4.3 software (Stuiver and Reimer, 1993) using the 1998 marine dataset (Stuiver et al., 1998a). According to the calibrated data, the two studied cores span from the last deglaciation to the recent. Ages for the samples were estimated assuming constant sedimentation rate between dated levels. Sample ages for the upper part of core 157 (above the ¢rst 14 C date) were derived assuming an age of zero for the top of the core. Likewise, the sample age for core 17 below 17 cm (13.4 kyr BP) was extrapolated using the sedimentation rate calculated for the core interval between 14 and 17 cm. The core bottom has an age of 13.91 kyr BP (Fig. 3). The age of 5.89 kyr BP at the top

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Depth core (cm) 0 0

10 0

30

20

40

520

Core 157

box core collected at the same site as core 17 indicates that very few benthic organisms survive at this depth, probably due to the shortage of food. The abundance of epifaunal foraminifera, living in an extremely oligotrophic environment, and the total absence of infaunal taxa, usually present with high organic carbon supply (Jorissen, 1999; Morigi et al., 2001), support the hypothesis of low or null bioturbation in recent sediment.

5000 5890

Core 17 Calibrated age (yr BP)

4.2. Isotope record

6180

8920

10000 10980

10980

11820 11820 12740 13400 13910 15000

Fig. 3. Age^depth model of cores 157 and 17. Underlined values indicate extrapolated age.

of core 17 indicates loss of the uppermost part of the core. This loss can either be due to the coring operations or because recent sedimentation is limited by strong dynamic setting of bottom currents (Carter and McCave, 1994). Sediment accumulation rates vary with time and are higher at the northern site than at the PF core (Fig. 3). The mean values during the investigated time intervals are 2.97 and 1.70 cm/kyr for cores 157 and 17 respectively. Compared with cores taken near and south of the PF in the Atlantic and Western Indian Ocean, and Eastern Paci¢c Sector (Howard and Prell, 1992; Labeyrie et al., 1996), our cores have remarkably low sedimentation rates. Given the slow sedimentation rates, bioturbation could degrade resolution in the cores where it would be di⁄cult to clearly de¢ne a climate perturbation with duration of only 1 kyr. However, a study in progress (Sabbatini et al., 2002) on living benthic meiofauna of a

The oxygen isotope record of the cores is shown as a function of depth in Fig. 4. Core 157 shows positive values in the entire investigated interval with values ranging between 3.89x and 2.55x. The most signi¢cant shift (3.89x to 3.23x) occurs at the base of the plot between 39 and 34 cm, showing an abrupt transition from a colder to a warmer phase. The upper part of the oxygen pro¢le has a negative trend interrupted by some short warmer events. Four main warm intervals (low N18 O values) and three cool ones (higher N18 O values) are recognized. In core 17, from 20 to 10 cm, the oxygen isotope record shows a positive trend with three oscillations decreasing in amplitude. At 10 cm the oxygen isotope value reaches the maximum value of 3.85x. An important shift is clearly present from 10 to 9 cm, with a 0.80x decrease. A negative trend characterizes the upper part of the core, with a warm peak recorded between 4 and 0 cm. 4.3. Carbonate dissolution Dissolution reduces the accuracy of the analysis of foraminiferal assemblages through selective removal of species. In fact species-selective dissolution modi¢es the species assemblages below the calcite lysocline; the original diversity of the assemblage is a¡ected below the depth level where carbonate dissolution becomes detectable. In the study area dissolution increases with depth from about 3500 m as the pelagic drape extends into an energetic deep western boundary current (Carter and McCave, 1994). The depth of our cores approaches the carbonate compensation depth,

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Core 157 18 δ O ( ‰ PDB) N. pachyderma 4.00

3.00

Core 17 18 δ O ( ‰ PDB) N. pachyderma

2.00 ‰

4.00

0

5

3.00

0

520

5890

11820

10

12740

6180 Depth (cm)

20

2.00 ‰

5

10

15

271

15 13400 20

8920 25 10980

Depth (cm)

30

35 10980 40

Fig. 4. Oxygen isotopic curves of cores 157 and 17. Arrows indicate

which was recently determined to be ca. 4100 m on the basis of bulk CaCO3 content (Weaver et al., 1997). We cannot exclude a dissolution e¡ect on our cores. In fact in both cores carbonate disappears, below 40 cm in core 157 and below 21 cm in core 17. However, we can argue that dissolution does not strongly in£uence the composition of the planktonic assemblage, in particular in core 17, in the levels corresponding to the cold event. Several authors (e.g. Thunell, 1976; Howard and Prell, 1994; Conan et al., 2002) have shown that dissolution can be estimated in a variety of ways using (1) the ratio of fragmented to whole tests of planktonic foraminifera, (2) the benthic^ planktonic foraminifera ratio, (3) the calcium carbonate content of the sediment sample, (4) the percent of the coarse fraction ( s 150 Wm), and (5) the ratio between dissolution-susceptible and dissolution-resistant planktonic species. We chose some dissolution indices for core 17: (1) % CaCO3 ; (2) the percent coarse fraction = dry weight of s 150-Wm fraction/dry weight of whole

14

C calibrated dates (yr BP).

sample; (3) the percent of benthic foraminifera out of the total foraminiferal assemblage; and (4) the percent of dissolution-susceptible planktonic species (Fig. 5). We do not use the dissolution index obtained by the percentage of fragments out of total planktonic foraminiferal remains, for two main reasons. First, Luz and Shackleton (1975) found that in cores collected below the lysocline the abundance of fragments does not correspond to other dissolution indicators (CaCO3 %; abundance of radiolarians and percent of benthic foraminifera). Second, Volbers and Henrich (2002) stated that in areas of high productivity the most commonly used proxies do not quantify calcium carbonate dissolution within calcareous sediments but instead are primarily related to changes in upwelling. They stated that ‘‘the use of these conventional proxies should therefore be limited to regions with more constant hydrographic conditions’’. Furthermore, we cannot calculate the ratio between dissolution-susceptible and dissolution-resistant planktonic foraminifera, as the diversity of planktonic foraminifera

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Fig. 5. (i) Percentage of CaCO3 weight content, (ii) % coarse fraction ( s 150 Wm), (iii) % benthic foraminifera (%BF) out of the total foraminiferal assemblage and (iv) % susceptible planktonic foraminiferal species out of the total foraminiferal assemblage in core 17. Shaded area corresponds to the cold event.

assemblages south of the PF decreases strongly due to cold surface waters (Niebler and Gersonde, 1998), and they are mainly dominated by Neogloboquadrina pachyderma (sin). We show, however, the percentage of more dissolution-susceptible species, Globigerina bulloides and Turborotalita quinqueloba (Hecht et al., 1975; Conan et al., 2002). All these dissolution indicators show good preservation in the upper 16 cm of the core. This suggests that dissolution does not strongly in£uence the foraminiferal assemblage and the interpretation of the planktonic analysis is reliable. 4.4. Foraminiferal record The planktonic assemblages of the two cores mainly consist of Neogloboquadrina pachyderma (sinistral), Globigerina bulloides and Globorotalia in£ata. Globigerinita glutinata, Globorotalia scitula, Globorotalia truncatulinoides, Neogloboquadrina pachyderma (dextral) and Turborotalita quinqueloba are also present in very small percentages. About 85% of the whole assemblage in

core 157 consists of N. pachyderma (sin) (Fig. 6a). This taxon shows a £uctuating pattern from the bottom to the top of the studied interval, interrupted by three negative peaks at 30, 19, and 7 cm. Neogloboquadrina pachyderma (sin) has a maximum (92%) at 39 cm. Globigerina bulloides shows a decreasing trend towards the top of the core. In particular, this species strongly decreases after 18 cm with percentage less than 10%. Globorotalia in£ata is absent at 39 cm, after this level its percentage is almost uniform from 36 to 22 cm (about 6%); frequency slowly increases up to 14% from 22 to 7 cm. Finally, G. in£ata strongly decreases to frequencies lower than 7% in the uppermost 3 cm of the core. The % benthic foraminifera has a low value in most of the record, except at 39 cm and in the uppermost 5 cm. The Simpson index varies between 0.61 to 0.85, with the highest value at 39 cm. Planktonic foraminiferal assemblage of core 17 consists of the quasi-monospeci¢c Neogloboquadrina pachyderma (sin) (Fig. 6b). This species forms up to 100% of the planktonic assemblages. Only a few specimens of Globorotalia in£ata and

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273

a

b

Fig. 6. Percentage of benthic foraminifera (%BF) and percentages of Neogloboquadrina pachyderma (sin), Globorotalia in£ata, Globigerina bulloides and Simpson index in core 157 (a) and core 17 (b). Shaded area corresponds to the cold event.

Globigerina bulloides are detected at some intervals. The highest values of dominance are recorded at 11 and 7 cm. The % benthic foraminifera of core 17 shows higher values than core 157, due to the lower abundance of benthic foraminifera at the southern site.

5. Discussion The foraminiferal analysis and oxygen isotopes point to the occurrence of a cold event in both

sedimentary records. Although the two sedimentary intervals contain di¡erent planktonic foraminiferal assemblages, we can identify some mutual features probably linked to changes in oceanographic conditions due to the cold event. The investigated sedimentary sequences show a basic di¡erence: (a) a more diversi¢ed microfauna and higher abundance of Globigerina bulloides and Globorotalia in£ata in the northern core (core 157) ; (b) a uniform distributional pattern of planktonic assemblage documented in the southern core (core 17). As testi¢ed by several studies

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C. Morigi et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 198 (2003) 265^278 AMS radiocarbon ages

δ18O (‰ PDB) N. pachyderma

2.0

core 157

core 17

2.5

3.0

3.5

4.0 2

4

6

8

10

12

14

16 yr BP

0

-32

YD

GRIP Greenland ice core 18 δ O (‰)

-36 -40 -44

-32

Byrd Antarctic ice core 18 δ O (‰)

ACR -34 -36 -38

18

18

Fig. 7. N O of cores 17 and 157 compared to the GRIP and Byrd N O records (Blunier and Brook, 2001). Shaded intervals indicate the Antarctic Cold Reversal (ACR) and the Younger Dryas (YD) intervals. The cold events are indicated by gray rectangles. Small black squares show location of AMS radiocarbon ages in each core.

on modern distribution of planktonic microfauna at the highest latitude (Be' and Tolderlund, 1971; Kohnfeld et al., 1996; Bauch et al., 2002), Neogloboquadrina pachyderma (sin) dominates the foraminiferal assemblage. In particular, this species occurs south of the SAF where the diversity of planktonic foraminifera decreases strongly due to cold surface waters (Donner and Wefer, 1994; Boltovskoy et al., 1996; Niebler and Gersonde, 1998). The planktonic foraminiferal assemblage indicates that during the last 14 kyr, site 17 has always been south of the PF. The planktonic microfauna of core 157 is clearly indicative of a subantarctic environment, where G. bulloides and G. in£ata comprise about 10% of the total assem-

blage. During the cold event in both cores (cm 10^11 in core 17; cm 39 in core 157) the planktonic assemblage is characterized by a less diversi¢ed microfauna (highest values of Simpson index), dominated by N. pachyderma (sin). The studied interval of core 17 clearly extends through the Younger Dryas event and the Antarctic Cold Reversal period, whereas core 157 covers the terminations of the Younger Dryas event and the Holocene period (Fig. 3). Both cores, however, have evidence of a cold event in their isotopic record (Fig. 7), although core 157 shows only one heavy value at the bottom of the N18 O record. However, the perfect correspondence of the isotopic data is remarkable. In core 157 the

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C age at 29 and 37 cm yielded the same value implying some sedimentation problem during this interval. In the same interval, our data indicate high frequencies of benthic foraminifera. This likely indicates high bioturbation activity which may indeed be responsible for the thick sediment mixed layer thus resulting in a uniform age for this interval. Nevertheless, the good isotope data ¢t due to the consistency of the N18 O values suggests that the lowest level of core 157 (cm 39 ^ isotopic value 3.89x) corresponds to the 10 cm depth of core 17 (isotopic value 3.85x). In core 17, the calibrated age for this level is 11.82 kyr BP, with a reservoir value of 776 yr. In order to obtain reliable 14 C dates and to establish a trustworthy radiocarbon chronology we have to choose an appropriate reservoir correction. Gordon and Harkness (1992) presented and reviewed radiocarbon data of the Antarctic region and proposed a reservoir correction higher than 750 yr for the polar latitude. For the West Antarctic region Berkman and Forman (1996) suggested a reservoir value of 1115 R 62 yr (vR = 816 R 62), whereas for East Antarctica the mean value of vR is 808 R 102 corresponding to an average reservoir of 1150 R 60 yr (Marine Reservoir Correction Database, Reimer, 2000). For our calibration process we chose the more conservative reservoir value of about 776 R 75 yr (vR = 462 R 75) (Table 1). The reservoir values seem to vary not only in di¡erent areas but also with time. A recent paper (Sikes et al., 2000) suggests older than present surface reservoir ages through the Last Glacial Maximum and deglaciation (including the time of the Antarctic Cold Reversal and the Younger Dryas event) in the Southwest Pacific sector. We therefore re-calibrated the ages using an older reservoir age (reservoir age = 1150 yr); in this way level 10^11 cm of core 17 dates at 11.26 kyr BP. It is clear that in both cases the ages are consistent with the termination of the Younger Dryas event. Although we cannot discriminate the most suitable reservoir age for this area of the Southern Ocean, it is clear that higher reservoir ages would yield younger calibrated ages for the cold event. Several studies report the delay existing in the short climatic £uctuations that occurred in the

275

Northern and Southern Hemispheres (e.g. Charles et al., 1996; Vidal et al., 1999; Mazaud et al., 2000; Blunier and Brook, 2001). In the Northern Hemisphere a cold period between 12.5 and 11.4 kyr BP has been well documented and named the Younger Dryas (e.g. Alley et al., 1993; Alley and Clark, 1999). The Younger Dryas event was also recorded in New Zealand (Denton and Hendy, 1994), the Chilean Andes (Lowell et al., 1995) and the Sulu Sea (Linsley and Thunell, 1990). Furthermore, as summarized by Alley and Clark (1999), some southern Indian Ocean marine records and the Taylor Dome Ice Core in coastal East Antarctica (Steig et al., 1998) show a pattern of deglacial warming more similar to GISP2 (Greenland) than to the Byrd ice core (Antarctica). Although some of these records characterized by a ‘northern’ response (Alley and Clark, 1999) have been questioned (Singer et al., 1998; Mulvaney et al., 2000), these observations suggest the possibility of a more complex regional pattern of millennial-scale temperature variability in high latitudes of the Southern Hemisphere (Blunier and Brook, 2001). Climatic signals recorded in Vostok and Byrd ice cores in Antarctica show a similar oscillation, named the Antarctic Cold Reversal, although the magnitude of the cooling over Antarctica is much smaller than that of the Younger Dryas event (Sowers and Bender, 1995; Charles et al., 1996). The Antarctic Cold Reversal, long thought to be a Younger Dryas equivalent, has been conclusively dated at about 14^12.5 kyr BP and thus preceded the Younger Dryas by about 1.5 kyr (Sowers and Bender, 1995). Furthermore, Stenni et al. (2001) recently inferred that an Oceanic Cold Reversal took place in the Southern Indian Ocean, 800 yr after the Antarctic Cold Reversal. The marked N18 O maximum in the isotopic curve of both cores (Fig. 7) implies a signi¢cant cooling during an overall warming trend. Although our time resolution does not make it possible to discriminate between the Younger Dryas event and the Oceanic Cold Reversal (sensu Stenni et al., 2001), the calibrated ages clearly show that the positive isotopic anomaly occurs between 12.5 and 11.4 kyr BP. This interval occurred at about the same time as the Younger Dryas event (Fig. 7). In any case, the

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lowest 14 C dates of core 17 adequately constrain the timing of the cold event so that the Antarctic Cold Reversal can be excluded.

6. Conclusions Microfaunal and oxygen isotopic investigations performed on two deep-sea cores collected between New Zealand and the Ross Sea yield information on the climatic history of the Paci¢c sector of the Southern Ocean, during the last 14 kyr. The marked N18 O maximum in both isotopic curves implies that a signi¢cant cooling occurred during an overall warming trend. Calibrated ages indicate that the end of this cold event corresponds to the termination of the Younger Dryas event. The break in the synchronicity of climate change has already been thought to occur at the PF instead of at the equator (Broecker, 1996; Bard et al., 1997). The record of a short-term cold spell, synchronous with the Younger Dryas event in sediments collected south of the PF of the Paci¢c Southern Ocean, indicates that the decoupling of the two hemispheres occurs at higher latitude. Collection and analysis of additional Antarctic deep sediment cores of the Paci¢c sector, especially from the PF area, is expected to con¢rm our observations.

Acknowledgements The research was carried out within the framework of the Glaciology and Paleoclimatology Project of the Programma Nazionale di Ricerche in Antartide and was ¢nancially supported by ENEA. The authors wish to thank M. Landuzzi for sample preparation and O. Silvestri for the isotopic sample picking. We also thank A.M. Borsetti, A. Longinelli and A. Negri for useful discussion and comments. We would like to thank L. Burkle and the other anonymous referees for their careful review of the manuscript and their helpful suggestions. We are indebted to the editors for their help and encouragement. This is contribution no. 1325 from the Istituto di Geologia Marina ^ CNR, Bologna.

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