Southwest Pacific modulation of abrupt climate change during the Antarctic Cold Reversal–Younger Dryas

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Palaeogeography, Palaeoclimatology, Palaeoecology 260 (2008) 284 – 298 www.elsevier.com/locate/palaeo

Southwest Pacific modulation of abrupt climate change during the Antarctic Cold Reversal–Younger Dryas Lionel Carter a,b,⁎, Barbara Manighetti a , Gerald Ganssen c , Lisa Northcote a a

National Institute of Water and Atmosphere, Private Bag 14 701, Wellington, New Zealand b Antarctic Research Centre, Victoria University, P.O. Box 600, Wellington, New Zealand c Vrije Universiteit, Amsterdam, The Netherlands Received 29 January 2007; accepted 22 August 2007

Abstract The giant piston core, MD97-2121 (2314-m water depth), collected north of the Subtropical Front, New Zealand, provides a well-dated, stable isotopic record of subtropical and sub-Antarctic influences on the surface and deep ocean over the last deglaciation, especially during the Antarctic Cold Reversal (ACR; ~ 14.1–12.4 ka) and Younger Dryas (YD; 13.0–11.5 ka). After the Last Glacial Maximum, benthic foraminiferal δ18Ob changed in phase with Antarctica — the ACR being represented by a pause in the δ18Ob deglacial trend. In contrast, surface waters, as represented by the stable isotopes of 3 planktic foraminifers from 3 different depth zones, and alkenone-based sea surface temperatures, showed no immediate response to the ACR. It was not until the Reversal was at its coldest, starting ~ 13.5 ka that the surface ocean responded. It became less warm, its fertility reduced sharply, and its surface structure changed as shown by the merger of the δ18Opl planktic profiles. Onshore, these changes were accompanied by an expansion of cool climate vegetation, an advancement of alpine glaciers and a likely pause in the post-glacial transgression. These onshore and offshore changes, which continued well into YD time, probably resulted from modification of polar conditions by subtropical influences. The ACR caused cooler temperatures, weakened seasonality and stronger winds. However, its onset and impact were ameliorated by a strengthened inflow of subtropical water, although ACR-driven atmospheric conditions prevailed. By comparison, south of the Subtropical Front, where subtropical influences are weak and water masses have direct links with Antarctica, the surface waters cooled in phase with the ACR. © 2008 Elsevier B.V. All rights reserved. Keywords: Antarctic Cold Reversal; Younger Dryas; Ocean change; SW Pacific

1. Introduction Middle latitudes of the Southern Hemisphere have been the subject of considerable debate regarding the interhemispheric propagation of major abrupt climate ⁎ Corresponding author. Antarctic Research Centre, Victoria University, P.O. Box 600, Wellington, New Zealand. E-mail address: [email protected] (L. Carter). 0031-0182/$ - see front matter © 2008 Elsevier B.V. All rights reserved. doi:10.1016/j.palaeo.2007.08.013

changes strongly recorded in the Northern Hemisphere, particularly the Younger Dryas (YD; ~ 13.0–11.5 ka) (Broecker, 1994; Denton and Hendy, 1994; Singer et al., 1998; Andres et al., 2003; Morigi et al., 2003.). The debate has been fuelled by a paucity of well-dated, palaeoclimatic records from southern oceanic regions (e.g., Steig, 2001), especially records with age models that are independent of Northern Hemisphere reference chronologies. Detailed and robust chronologies are

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needed to resolve the influences of the Antarctic Cold Reversal (ACR; ~ 14.1–12.4 ka) versus those of the YD because these two events overlap (Blunier and Brook, 2001; EPICA, 2004). Recent studies have shown that the Southern Ocean and atmosphere did indeed cool during ACR–YD time, but the timing of this cooling varied between oceanic basins, bearing in mind potential differences in age models. Palaeoceanographic records from the S. Atlantic (Kanfoush et al., 2000; Sachs et al., 2001) exhibit an ocean surface cooling around ACR time. In contrast, the Indian Ocean underwent a cold reversal from ~ 13.2 to 12 ka that fell between the onset of the ACR and end of the YD (Stenni et al., 2001). Similarly, speleothem and palynological records from

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New Zealand also show a cooler atmosphere during the ACR–YD, but this was asynchronous with the Indian Ocean event (Turney et al., 2003; Williams et al., 2004). Here we investigate ocean change over the last deglaciation using new proxy data and an established age model derived from the giant piston core, MD972121, collected from off the eastern North Island of New Zealand (Fig. 1). This record provides an insight into oceanic changes during the abrupt climatic swings of the high-latitude ACR and YD, as well as any modulation by low latitude drivers such as the Subtropical Inflow. In that context a better picture of SW Pacific change is formed to shed further light on inter-ocean variability during this period of marked environmental reorganization.

Fig. 1. Locations of cores MD97-2121 and MD97-2120 located north and south of the Subtropical Front (STF) respectively. Also shown are other elements of the oceanography including the Deep Western Boundary Current (DWBC); Antarctic Circumpolar Current (ACC); Tasman Front (TF); East Auckland Current (EAUC); East Cape Current (ECC); Southland Current (SC); Subtropical (surface) Water (STW); Sub-Antarctic (surface) Water (SAW); Circumpolar Surface Water (CSW), Central Volcanic Region (CVR) and Mernoo Saddle (MS).

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2. Modern environmental setting The SW Pacific Ocean has close atmospheric and oceanic links with Antarctica and the equatorial Pacific. In the south, the interaction of polar easterly and midlatitude westerly weather systems, coupled with climatic shifts such as the Antarctic Oscillation and Antarctic Circumpolar Wave, influence the region on synoptic to decadal time scales (e.g., Kidson, 1975; Gong and Wang, 1999; White and Cherry, 1999). Offshore, cool sub-Antarctic and polar surface waters are introduced via the Antarctic Circumpolar Current, which is forced 5° northwards of its normal path by the New Zealand submarine micro-continent (Carter and Wilkin, 1999; Orsi et al., 1995). At depths N~ 2000 m, the micro-

continent steers the local sector of the global thermohaline circulation into the Pacific Ocean (e.g., Warren, 1973; Carter and McCave, 2002; Rahmstorf, 2003). In the north, synoptic-scale weather of subtropical origin has a strong influence on New Zealand (Sturman and Tapper, 1996), as do El Niño-Southern Oscillation (ENSO) signals out of the tropical Pacific (Basher and Thompson, 1996; Marshall and King, 1998). Typically occurring at 3–7 year cycles, ENSO is modulated by longer cycles such as the Interdecadal Pacific Oscillation (Salinger et al., 2001). Also arriving from the central Pacific is the Subtropical Inflow. It reaches New Zealand as the Tasman Front, which is disrupted by the regional bathymetry into a series of coastal current and eddy systems including the East Cape Current (Fig. 1).

Table 1 AMS 14C dates for planktic foraminifers, and tephra ages for core MD97-2121 Depth (cm)

Tephra/AMS

14

19–20 35–36 40–41 54–55 58–59 106–107 110–111 190–191 196–197 270–271 310–311 342–343 408–409 420–421 480–481 510–511 514–515 514–515 670–671 748–749 748–749 872–873 983–984 984–985 984–985 1041–1042 1126–1127 1154–1155 1234–1235 1309–1310 1353–1354 1439–1440

AMS-P AMS-P AMS-P Taupo AMS-P Waimihia AMS-P Whakatane AMS-P Tuhua AMS-P Rotoma Poronui AMS-P AMS-P Waiohau AMS-P AMS-B Rerewhakaaitu AMS-P AMS-B AMS-P Kawakawa AMS-P AMS-B AMS-P AMS-P AMS-P AMS-P AMS-P AMS-P AMS-P

1032 1139 1748 1800⁎ 2203 3280 3241 4830 5379 6130 8286 8530 9810 10,818 12,265 11,850 12,664 13,448 14,700 17,550 18,060 19,920 22,590 22,450 24,200 24,050 26,250 28,250 31,090 33,590 31,890 38,710

C age (yr)

+/− (yr) 60 60 56 56 20 56 20 57 30 60 10 50 70 96 60 60 65 110 110 150 100 230 130 140 130 150 170 250 550 350 410

Calibrated age intercept (yr BP) 620 680 1290 1718 a 1800 3470 3060 5580 5730 6970 8830 9520 11,190 12,290, 12,230, 12,130 b 13,820 13,830 14,120 15,570 17,710 20,350 20,940 23,180 26,300 25,690 c 27,440 c 27,290 c 29,480 c 31,490 c 34,330 c 36,830 c 35,130 c 41,950 c

2δ range (cal.ages) 686–516 787–609 1393–1202 1921–1675 3215–2888 5875–5622 8927–8617

12,383–11,888 14,057–13,439 15,029–13,864 15,891–15,312 20,799–19,924 21,465–20,437 23,937–22,696

AMS dating was carried out at the Rafter Radiocarbon Laboratory, Institute geological and Nuclear Sciences, Lower Hutt, New Zealand. Calibrated ages are calculated using the program CALIB v.4.2 with a marine calibration dataset and a 30-year moving average to account for the likely age span of foraminifera from a 1 cm-thick sample. a Derived from sequence of 10-year samples from tree rings. b More than one intercept with the calibration curve. c Calibration approximated using constant offset of +3240 years for ages N20,760 14C yr.

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Site MD97-2121 is near the confluence of subtropical (STW) and sub-Antarctic (SAW) surface waters, whose dominance over each other varies with time (Fig. 1). Presently, the surface ocean is occupied by warm STW transported southward by the East Cape Current (Chiswell and Roemmich, 1998). From ~ 600 to 1400 m depth, the water column is occupied by Antarctic Intermediate Water, which is forced periodically to shallower depths by the local bathymetry (Heath, 1985). The remainder of the water column is occupied by upper Circumpolar Deep Water (McCave and Carter, 1997). 3. Methods and age control 3.1. Core details The palaeoceanographic record is from upper 9.15 m of the 34.92 m-long core collected by R.V. Marion Dufresne at 40°22.935′S; 177°59.68′E in a water depth of 2314 m (Fig. 1). This site is subject to a high hemipelagic flux and frequent volcanic eruptions, which together provide a high resolution record constrained by tephra time markers (Carter et al., 2002). The core was visually described and its physical properties determined by a GEOTEK core logger (Nees et al., 1998).

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3.2. Age control The core chronology is based on nine macroscopic tephras and twenty Accelerator Mass Spectrometry (AMS) 14C dates from mixed planktic foraminifers (Table 1; Fig. 2). Tephra and AMS 14C time markers were calibrated to calendar years using the CALIB REV 4, v.4.2 program (Stuiver & Reimer, 1993), and form a highly consistent age-depth trend (Fig. 2). Elements of the age model are presented in Carter et al. (2002) (for the period 0–14 ka), Carter and Manighetti (2006) (0–135 ka) and Pahnke and Sachs (2006) (0–135 ka). However, a key aspect of the model has not been discussed elsewhere, namely the suggestion that 14C reservoir ages may be up to 2000 years for SW Pacific surface waters during the last glacial period (Sikes et al., 2000). Such a reservoir age could produce a discrepancy between “true” ages and the calibrated AMS dates calculated via the marine dataset calibration that incorporates a time-dependent, global ocean reservoir correction of about 400 years (Stuiver et al., 1998). This was assessed qualitatively by examining the calibrated 14C dates against the age-depth trend revealed by the tephrostratigraphy. The regression lines for linear age models, based on tephra and AMS dates, show similar slopes of 0.382 and 0.375 respectively, with

Fig. 2. (a) AMS 14C dates on mixed planktic foraminifers and named tephras in MD97-2121. Regression lines for (b) planktic AMS 14C and (c) tephra data are also shown. Errors are presented in Table 1. Elements of this dataset are presented in Carter et al. (2002), Carter and Manighetti (2006) and Pahnke and Sachs (2006).

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R = 0.999 in both cases (Fig. 2). The slightly steeper slope of the calibrated AMS trend suggests that the regional reservoir correction is only 300 years, on average, over the last 28 kyr. This is identical to the mean (x¯ = 300.63 years, N = 19) of New Zealand regional reservoir corrections calculated from the marine database (Stuiver et al., 1998). When calibrated according to Stuiver et al. (1998), the two AMS dates at 420 and 480 cm core depth, fall outside (older than) the linear trend indicated by the bracketing Waiohau, Poronui and Rotoma tephras (Fig. 2). However, Sikes et al. (2000) calculated a reservoir age for the time of deposition of the Waiohau tephra of 800 years, and for the Rotoma tephra of 390 years. An interpolation between these values suggests reservoir ages of 580 and 726 years for the 14C dates at 420 and 480 cm respectively. Using these instead of the global averages incorporated in the Calib. 4.0 marine dataset produces calibrated ages much closer to the linear trend, and within 2σ error bars for the 480 cm date, thereby adding weight to the Sikes et al. (2000) argument. The 420 cm date can also be forced to fit the trend by applying a reservoir age of 700 or 800 years, suggesting that the period of increased reservoir age in the surface SW Pacific may have lasted at least 2 kyr. Our data show no evidence, however, for an elevated reservoir age during the last glacial as marked by the 26.5 ka Kawakawa tephra (Wilson, 2001). 3.3. Microfauna and stable isotopes To gauge change in surface waters to ~400 m depth, stable isotope profiles were generated for one benthic and three planktic foraminiferal species of which profiles for Globigerina bulloides are presented in Carter and Manighetti (2006). Samples were picked from 1 cm-thick slices positioned at 2–4 cm spacing to give a temporal resolution of 52–104 years. Isotopic analyses of N 150 μm diameter foraminiferal tests were run on a Finnegan MAT 252 mass spectrometer with an automated calcium carbonate preparation line (Type Kiel II) located at Vrije Universiteit, Amsterdam, and NIWA, Wellington. Isotopic results are versus the VPDB scale and are calibrated using the NBS-19 standard. External reproducibility of a routinely run laboratory standard was ±0.08‰ and ±0.05‰ for δ18Opl 13 & b (planktic and benthic) and δ Cpl & b respectively. Replicates were run to check reproducibility, and the isotope curves are an average of the replicate results. The foraminiferal species are as follows: (i) Globigerinoides ruber is one of the shallowest-dwelling foraminifera, preferring warm waters and calcifying chiefly in summer (Ganssen and Kroon, 2000). Its preference for subtropical waters (Thiede et al., 1997) means that G. ruber is scarce or even absent in last glacial age sediments. (ii) G. bulloides is

a surface or near-surface dweller that often resides in the upper 100 m of the water column (Mortyn and Charles, 2003). Its abundance is influenced by the structure of the pycnocline (Mortyn and Charles, 2003) and enhanced nutrient supply (Ganssen & Kroon, 2000). Such conditions are often related to upwelling. (iii) Globoconella inflata is persistent and numerous throughout the core. It has a depth range extending to a sub-thermocline depth of ~400 m (Ganssen and Kroon, 2000; Mortyn and Charles, 2003). Hence, its isotopic signature is likely to be influenced by temperatures colder than co-existing G. ruber and G. bulloides. Comparative δ13Cp records from the planktic species are probably controlled by local ocean fertility (Wefer et al., 1999; Ganssen and Kroon, 2000). Carbon isotopes of G. bulloides tend to increase under fertile conditions, while that of G. ruber and G. inflata become more negative during increased upwelling of subsurface waters enriched in 12C. Isotopic signatures for bottom waters are from Uvigerina peregrina, a shallow infaunal species with intermittent peaks in abundance of up to 12% of the benthic assemblage. It records the isotopic signal of pore waters in the upper layers of sediment (Wefer et al., 1999), which for oxygen is close to that of bottom water (Duplessy et al., 1984). However, its carbon isotopic composition is offset from ambient seawater CO2 by a variable amount, depending upon local conditions, including the accumulation rate of organic matter (Wefer et al., 1999). Profiles of δ18Opl reflect the influence of sea surface temperature (SST), global ice volume and sea surface salinity (SSS). In the case of SSTs, these were determined directly for MD97-2121 and nearby MD97-2120 by Pahnke and Sachs (2006) using alkenone U k ′ 37 palaeothermometry. The ice volume effect was estimated from the record of mean ocean water (δ18Omw) reconstructed by Waelbroeck et al. (2002). Following Pahnke et al. (2003), δ18Omw was subtracted from stable isotope concentrations of planktic foraminifers (δ18Opl) to yield an ice volume-corrected, δ18Oc. Using the alkenone SST data and concomitant δ18Oc (e.g., Doose et al., 1999), the local ocean δ18Ow was determined from the equation of Shackleton (1974) whereby;
T -C ¼ 16:9  4:0 d18 Oc  d18 Ow Following conversion δ18Ow (VPDB) = δ18Ow (VSMOW) −0.27, estimates of salinity were made using the relationship of Savin et al. (1975)
S ¼ d18 Ow þ 23:74 =0:687

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This yielded realistic modern salinity values for site MD97-2121 down to ~400 m water depth (H. L. Neil, personal communication, 2005). Even so, errors relating to these estimations can be significant. The standard error for deriving δ18Ow from a known salinity is 0.2‰, whereas the estimation of salinity from known δ18Ow is 0.4 (Schmidt, 1999). Furthermore, the salinity/δ18Ow relationship may not be valid for glacial conditions due to reduced δ18Ow in melt-water-influenced SAW, and increased fractionation of oxygen isotopes during precipitation (see Pahnke et al., 2003; King and Howard, 2005). Using the results of Schmidt (1999), a minimum error for palaeosalinity derived from alkenone-based temperatures is σs = 0.9 assuming no uncertainty in past freshwater effects, which potentially increase errors to σs = 1.1 to 2.5. In this paper we focus on the SSS profile derived for G. ruber as it is the shallowest dwelling of the three planktic species, and is least likely to be affected by salinity differences associated with upwelling of, for example low salinity Antarctic Intermediate Water (e.g., Heath, 1975). Data are presented as SSS anomalies determined relative to a mean of 34.8‰ calculated for the period 7–4 ka — the youngest part of our palaeosalinity estimates (also see

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Pahnke et al., 2003). The mean was chosen to dampen the variability in SSS that is associated with periodic incursions of SAW (e.g., Heath 1975), as well as strong seasonal and inter-annual variability especially associated with El Niño. 3.4. Foraminiferal assemblages A reconnaissance analysis of planktic foraminiferal assemblages was made to assist with the interpretation of surface ocean conditions. A minimum sample of 300 foraminifers were identified in N 150 μm size fraction obtained from 9 samples with a mean spacing of 90 cm (Fig. 3). From these data, key species were used to define a suite of assemblages characteristic of the following environmental settings (from Imbrie and Kipp, 1971; Weaver et al., 1998): Tropical — G. ruber, Globigerinoides sacculifer Subtropical — Globigerina falconensis, Globorotalia scitula, G. truncatulinoides, G. hirsuta Transitional — G. inflata Subpolar — G. bulloides, Globigerinita glutinata, Neogloboquadrina pachyderma (dextral), N. pachyderma/dutertrei intergrade (P/D) Polar — N. pachyderma (sinistral) Gyre — N. dutertrei, Globorotalia tumida, G. sacculifer, Pulleniatina obliquiloculata. 4. Results 4.1. SSS anomalies The salinity of surface waters is highly variable with excursions of 0.5 units or more (Fig. 4). Nevertheless, there is a broad trend of increasing SSS through the deglaciation to peak between 13 and 12 ka. Thereafter, the ocean freshened slightly to modern salinities, which were maintained, albeit in a fluctuating manner, through the early Holocene. From 5 ka onwards (outside the scope of this study), SSS fluctuations increased although the average remained close to modern salinities.

Fig. 3. Reconnaissance analysis of species assemblages and associations for planktic foraminifera in MD97-2121, based on Imbrie and Kipp (1971) and Weaver et al. (1998). Of note is the increase in transitional (mixed subtropical and sub-Antarctic conditions), subtropical and tropical assemblages at the expense of polar and subpolar forms around ACR time. Profiles are based on irregular sample spacings with a mean interval of ~ 90 cm compared to the 2–4 cm spacings of the isotopic profiles. Note that the key tropical species, G. ruber, occurs intermittently approaching the LGM. ACR = Antarctic Cold Reversal.

4.2. Oxygen isotopes δ 18 Opl G. inflata becomes progressively lighter through the deglaciation until 14 ka when it enters a period of little change after onset of the ACR (Fig. 4). By comparison, the G. bulloides record is more variable. During the glaciation, it is negatively offset from G. inflata, but during the deglaciation δ18Opl G. bulloides hovers around G. inflata. Apart from rare occurrences as

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Fig. 4. (a) δ18Opl profiles of three planktic foraminiferal species in core MD97-2121 from 24 to 4 ka, with profile of G. bulloides also in Carter and Manighetti (2006); (b) alkenone-based sea surface temperatures (SST) from Pahnke and Sachs (2006); (c) estimated sea surface salinity (SSS) anomalies for G. ruber with original and smoothed profiles; (d) δ18Ob profile for the benthic species, U. peregrina, and (e) deuterium (δD) record from the EPICA ice core (EPICA, 2004).

early as 20 ka, G. ruber does not make a permanent appearance until ~16.5 ka, after which it reaches a maximum of 6% of the planktic assemblage. Initially it is negatively offset from the other isotopic profiles, but like the others, δ18Opl G. ruber becomes heavier in the ACR. Thus, the Reversal is a time when the range of planktic δ18Opl profiles is at a minimum of b 0.6‰ (Fig. 4). This lasts through the ACR and YD termination at 11.5 ka when

δ18Opl profiles separate into (from lightest to heaviest) G. ruber, G. bulloides and G. inflata. This hierarchy occupies a band ~1.3‰ wide, and persists through the Holocene. At depth, benthic δ18Ob reaches a maximum of ~ 5.1‰ between 21 and 18 ka. After, it lightens in concert with a warming Antarctic climate (Fig. 4; EPICA 2004). Around 14.1 ka, the rate of change for δ18Ob pauses as the polar climate enters the ACR. This

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reduction was nearly synchronous with a pause in the δ18Opl profile of the deep-dwelling, G. inflata. Near the end of the ACR at ~ 12.4 ka, δ18Ob resumes its trend to lighter values, which reach a minimum of 3.5‰ beginning 9 ka and lasting well into the Holocene. 4.3. Carbon isotopes Emerging from the LGM, δ13Cpl G. inflata exhibits a variable profile with fluctuations of ~0.5‰ that diminish around 16 ka with the main incoming of G. ruber (Fig. 5). Thereafter, the two profiles merge and together gradually

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become heavier in δ13Cpl through the Holocene. In contrast, δ13Cpl G. bulloides shows several prominent positive shifts of up to 1‰ that are broadly centred on 21 ka and 15 ka (Fig. 5). The last shift extended into ACR time when it temporarily joined the other planktic profiles at ~0‰ δ13Cpl. The merger was brief and was followed by a sharp decline of ~−1.2‰ in G. bulloides between 13.7 and 13.2 ka. The level reached was broadly maintained through the Holocene. Compared to the planktic signal, the benthic profile is more muted with an overall positive, up-core excursion at a rate that increases into the Holocene. Small departures from the

Fig. 5. (a) δ13Cpl variation of three planktic species in MD97-2121 from 24 to 4 ka with profile of G. bulloides also in Carter and Manighetti (2006), (b) δ13Cb for benthic U. peregrina, smoothed with a 3-point moving average, and (c) the deuterium and (d) the Na profiles from the EPICA ice core (EPICA, 2004; Röthlisberger et al., 2002).

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trend are a lightening of δ13Cpl around 15 ka and a sharp increment at 11 ka (Fig. 5). 4.4. Foraminiferal assemblages While lacking the resolution of the stable isotopic profiles, the species assemblage data broadly confirm a change in surface waters during ACR–YD time. Subpolar and to a lesser extent, polar assemblages reduced through the ACR whereas the “transitional” assemblage (mix of subpolar and subtropical species), tropical and subtropical groups, increased especially coming out of the YD (Fig. 3). 5. Discussion 5.1. LGM (28–18 ka) to early deglacial times (18–14 ka) Although north of the Subtropical Front, the upper ocean at MD97-2121 was strongly influenced by cold SAW during the LGM as revealed by foraminiferal assemblage and alkenone data (this study; Weaver et al., 1998; Carter et al., 2002; Pahnke and Sachs, 2006). SAW, probably containing upwelled Antarctic Intermediate Water (see Heath, 1975), jetted north across westernmost Chatham Rise under an atmospherically invigorated coastal circulation, reinforced by a branch of the Antarctic Circumpolar Current (Nelson et al., 2000; Neil et al., 2004). SSTs were at their coldest from 23 to 21 ka when temperatures briefly dipped below 12 °C. Thereafter the upper ocean began to warm, leading the EPICA δD record of Antarctic warming (Monnin et al., 2001) and the benthic δ18Ob record by ~3 kyr; a lead that is consistent with the findings of Shackleton (2000) and Mashiotta et al. (1999) (Fig. 4). About this time, winds and polar foraminiferal assemblages began to decline (e.g., Hesse, 1994; Röthlisberger et al., 2002) and, together with a paucity of subtropical species, suggest that warming was encouraged by reduced forcing of the Sub-Antarctic Inflow (Figs. 3, 4). Following the start of the deglaciation at ~18 ka, subtropical G. ruber made its first permanent appearance at 16–15 ka. While G. ruber abundances are only 1–2%, its permanency, together with rising SSS and increased rates of warming from 0.2 °C/kyr to 0.7 °C/kyr, collectively suggest an increasing subtropical influence. Initially δ18Opl G. ruber was negatively offset from the other planktic species; a situation that probably reflects the distinct seasonality of foraminiferal production in the region (King and Howard, 2000; Northcote and Neil, 2005). G. ruber typically calcifies in summer months in the presence of thermally stratified subtropical waters, whereas G. inflata and G. bulloides flourish in early spring when vigorous equinoxial winds encourage deep

mixing of surface waters and incursions of sub-Antarcticderived, upwelled waters at MD97-2121 (e.g., Heath, 1975; Weaver et al., 1998; Nelson et al., 2000). The strengthening Subtropical Inflow coincided with a period, centred on 15 ka, encompassing four positive shifts in δ13Cpl G. bulloides. This is the second period of such peaks; the first occurring around 20 ka (Fig. 5). Marked increments of δ13Cpl G. bulloides can result from increased ocean fertility (Wefer et al., 1999); a point supported by similarly timed peaks in marine phytoplankton productivity measured in nearby core MD97-2120 (Sachs and Anderson, 2005). Presently, phytoplankton flourishes at the Subtropical Front where macronutrientrich SAW meets micronutrient-rich STW, the latter including iron (Boyd et al., 1999; Murphy et al., 2001). Such blooms, together with associated influxes of G. bulloides, peak in spring when light levels and windforced mixing of the surface waters are optimum (Murphy et al., 2001; Northcote and Neil, 2005). Such interaction between SAW and STW also occurred at MD97-2121 in MIS 2 (Weaver et al., 1998; Nelson et al., 2000), and the coincidence of the 15 ka productivity burst with a stronger Subtropical Inflow, as recorded by G. ruber and SSS, may be such a case. However, productive phases also occur prior to a strengthened Subtropical Inflow inferring other contributing factors. Introduction of river-borne iron is a possibility but unlikely in light of the prevailing dry conditions in the LGM (Pillans et al., 1993; McGlone, 2001). Also not favoured is fertilization by airborne iron because the main aeolian fluxes are out of phase with periods of high biogenic carbonate and silica production at MD97-2121 (Carter and Manighetti, 2006). Finally, iron may be introduced via wind-induced upwelling (e.g., Heath, 1972; 1975). Such a mechanism would be encouraged under the vigorous wind regime of LGMearly deglacial times. However, productivity proxies such as the G. bulloides peak at 14 ka, are not in phase with a period of increased wind circulation at 13.5–13.8 ka as suggested by a peak in the Na flux of the EPICA ice core (Röthlisberger et al., 2002) (Fig. 5). The deglacial ocean warming was accompanied onshore by the spread of podocarp trees at the expense of cool climate shrubs and grasslands (Turney et al., 2003). This expansion began around 17 ka in the North Island, at a time when atmospheric methane and carbon dioxide increased as inferred from the EPICA ice core (Monnin et al., 2001). 5.2. Ocean change in ACR time (14.0–12.4 ka) The deglacial trend to lighter benthic δ18Ob slowed at the start of the ACR when Antarctica entered a phase of

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cooling that lasted up to 1.5 kyr (Fig. 4; Blunier et al., 1997). This phase presumably marked an expansion of the cryosphere that may be manifested in Australasia by a still-stand of sea level around 14 ka (Carter et al., 1986; Larcombe et al., 1995). An expansion of ice shelves and sea ice (Shemesh et al., 2002), plus a modest increase in windiness recorded by dust and sea salt proxies in ice cores (Röthlisberger et al., 2002), probably increased deep water production thus strengthening the thermohaline Pacific deep western boundary current (DWBC) off eastern New Zealand (e.g., Hall et al., 2001). Judging by the modern oceanography (e.g., Haine et al., 1998), any Antarctic modification of the deep circulation would be transferred to New Zealand within decades. As the upper reaches of the DWBC probably intrude MD97-2121 (Fig. 1; McCave and Carter, 1997), an invigorated flow may have caused a slight coarsening of sediments similar to that observed in glacial age sediments on Chatham Rise (Hall et al., 2001). Indeed, there is an elevation of silt at MD972121 in ACR time (Carter et al., 2002), but we cannot ignore possible local source effects owing to the site's proximity to a high-input terrigenous margin. Carbon isotopes of bottom waters fluctuated before and during the ACR suggesting temporary influxes of Antarctic-sourced, deep water against a background of increased input of δ13C-rich North Atlantic Deep Water (Raymo et al., 1997). It may be argued that the lowered benthic δ13Cb resulted from an enhanced supply of organic carbon associated with the period of high fertility centred on 15 ka (Fig. 5). However, such fertility spikes are not always synchronous with reduced benthic δ13Cb as attested by the 20 ka event. Change in the upper ocean during the ACR was more complex than that at depth. The shift to lighter δ18Opl G. inflata and G. ruber paused at ~ 14 ka, whereas δ18Opl G. bulloides revealed little change. However, by ~ 13.5 ka all planktic profiles had compressed into a narrow band, equivalent to b0.6‰ δ18Opl, that persisted through the ACR and YD. This change occurred when (i) the rate of ocean warming began to reduce with SSTs hovering around 16 °C — a state that persisted to 11.5 ka when rates increased again to 0.75°/kyr, (ii) SSS was increasing and (iii) ocean fertility declined rapidly as inferred from a sharp negative shift in δ13Cpl G. bulloides that ended at 13.5 ka (Figs. 4 and 5). At first sight, there is a contradiction with reduced rate of ocean warming versus a higher Subtropical Inflow as evinced by G. ruber and increased SSS. However, the data may simply be showing the interaction of two opposing forces. Any increase in SSTs associated with an enhanced Subtropical Inflow may be ameliorated by

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cooler atmospheric temperatures related to the ACR (e.g., Turney et al., 2003) and vice versa. Such an interaction is consistent with the compressed δ18Opl profiles; the three planktic species calcifying in non-warming, mixed surface waters affected by the ACR. Identification of atmospheric conditions over 13.5– 11.5 ka, comes from the following observations: (i) a cooler and drier climate over the lee, eastern side of New Zealand as inferred from an expansion of grassland and shrubland (Newnham and Lowe, 2000; Turney et al., 2003), (ii) cooler air temperatures are also indicated by speleothem records (Williams et al., 2004) and glacial advances along the upwind, western mountainous South Island (e.g., Denton and Hendy, 1994; McGlone, 1995), and (iii) increased windiness as supported by dust and salt proxies in the EPICA ice core (Röthlisberger et al., 2002) and the onshore/offshore deposition of aeolian sediment (Stewart and Neall, 1984; Hesse 1994; Fomento-Triglio et al., 2003) with the caveat that this mode of deposition is also affected by source availability (Shulemeister et al., 2004). Such climatic conditions resemble those of El Niño when (i) winds strengthen and prevail from the W to SW (Salinger and Mullen, 1999), (ii) precipitation increases in the W whereas the E is drier (Sturman and Tapper, 1996), (iii) higher snow accumulation fuels glacier advances (Tyson et al., 1997), (iv) cooler air temperatures prevail, especially in summer (Basher and Thompson, 1996), and (v) generally, but not invariably, the ocean has cooler SSTs and increased depth of mixing (Basher and Thompson, 1996; Hadfield and Sharples, 1996; Sutton, 2001). Yet, despite this strong ENSO character, it was an unlikely driving mechanism for middeglacial change. Palaeoclimatic data from the equatorial Pacific Ocean — the ENSO power house for the region — suggest that ENSO activity was suppressed in ACR–YD time (Tudhope et al., 2001; Koutavas et al., 2002); a contention supported by model simulations (Clement et al., 2000). Significantly, ENSO and the general westerly wind system strengthened into the late Holocene (Tudhope et al., 2001; Gomez et al., 2004; Shulmeister et al., 2004), whereas the upper ocean retained a substantial temperature gradient compared to ACR–YD time (Fig. 4). This association suggests that the thermal structure exhibited by the 3 planktic species is influenced more by weakened seasonality than by physical mixing, although the latter probably played a role. We therefore argue that the cool, windy conditions of 13.5 to 11.5 ka were driven from Antarctica where the ACR cooling presumably forced a northward expansion of cold atmospheric and oceanic temperatures, i.e., a scaled-down version of the LGM (Gersonde et al., 2003). The thermal gradient between pole and equator

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increased, and cold, vigorous westerly weather systems migrated to New Zealand latitudes (Hesse, 1994). Nevertheless, the pronounced oceanic changes (reduced seasonality, inferred deeper mixing and diminished fertility) occurred around 13.5 ka, half way through the ACR. A similarly timed delay is evident in the central New Zealand speleothem record with its well defined atmospheric cooling from 13.5 to 11.2 ka (Williams et al., 2004). More general support comes from local palynological data (e.g., Newnham and Lowe, 2000; Turney et al., 2003). Why the delay especially in light of the close atmospheric/oceanic links between modern Antarctica and New Zealand (e.g., Marshall and King, 1998; White and Cherry, 1999; Rintoul et al., 2001)? One probable cause is amelioration of any local ACR response by STW. This contention can be tested by comparing the timing and/or strength of the ACR-timed SST signals measured north (MD97-2121) and south (MD97-2120; Pahnke et al., 2003; Pahnke and Sachs, 2006) of the Subtropical Front (Fig. 1). In the north, where SAW is mixed with STW, the effect is a reduced rate of deglacial warming starting at ~13.5 ka. To the south, where SAW prevails, the corresponding response to the ACR is variable depending upon the temperature proxy. The alkenone data of Pahnke and Sachs (2006) show a small and brief disruption of the deglacial warming between 13.7 and 13.0 ka, whereas Mg/Ca-based temperatures have a distinct −1.1 °C reversal at ACR time (Pahnke et al., 2003). A better defined northward amelioration of an ACR signal can be inferred from the δ18Opl data from the South Atlantic Ocean (Kanfoush et al., 2000) and SE Pacific (Lamy et al., 2004; Kiefer and Kienast, 2005). Nevertheless, the STW influence was eventually overridden by cold and windy atmospheric conditions around 13.5 ka, which coincided with the onset of peak ACR cooling (Fig. 4; EPICA, 2004). Such conditions persisted through to 11.5 ka, helped by a possible reduction in the Subtropical Inflow as suggested by lower SSS. Such a reduction may reflect a latitudinal shift in the Tasman Front sector of the Inflow. During the LGM, the Tasman Front was about 6° north of its present position along ~32° S (Martinez, 1994). As a result, the Norfolk Ridge system possibly diverted part of the Inflow from the eastern North Island (Carter et al., 2002). We suggest that as the Tasman Front migrated south it intercepted gaps in the Ridge to temporarily enhance the Inflow to the North Island. 5.3. YD (13.0–11.5 ka) into the Holocene Bottom waters resumed a trend to lighter δ18Ob at the ACR termination ~ 12.4 ka, whereas δ13Cb became

irregularly less negative. The latter trend is interpreted as a change in the relative influences of different bottom water sources with southern sources waning and NADW coming to the fore (Hall et al., 2001). It is unlikely that this post-13 ka profile of benthic δ13Cb is markedly affected by surface productivity because it appears to be unrelated to δ13Cpl G. bulloides — a paleaoproductivity indicator. Thus, prominent positive shifts such as that at 11 ka may reflect pulses of NADWinfluenced deep water. At the surface, the compressed δ18Opl signature that began in the ACR, persisted to ~ 12 ka when the G. ruber profile began to separate, followed by G. inflata and G. bulloides in 11.3 ka. By 9.5 ka, the Holocene succession of (from lightest to heaviest) G. ruber, G. inflata and G. bulloides, was established. This trend towards better stratified surface waters reflects warmer SSTs, presumably improved seasonality with warmer summers, and a dominance of STW (Nelson et al., 2000; Carter et al., 2002). Compared to the diffuse signal of the surface ocean, the terrestrial climate warmed in phase with Antarctica. Following the ACR termination, speleothem data reveal a warm shift (Williams et al., 2004) that was accompanied by a rapid expansion of podocarp forest at the expense of cool climate flora (Turney et al., 2003; Vandergoes and Fitzsimons, 2003). Thus, the New Zealand climate was out of phase with the cold YD in the Northern Hemisphere. 6. Concluding comments Deep water at MD97-2121 changed in phase with the ACR (Table 2). A pause in the deglacial lightening of δ18Ob, together with negative shifts in δ13Cb, is consistent with an invigorated deep inflow forced by the expansion of the cryosphere. In contrast, surface waters showed only a muted response to the onset of ACR at ~14.2 ka, reflecting an ameliorating influence of the strengthening Subtropical Inflow (Table 2). Eventually ACR-driven atmospheric influences prevailed, lasting from ~13.5 ka to 11.5 ka. The climate was cool and windy with reduced seasonality and marine fertility. These conditions ended part way through the YD as local climate warmed. This warming was antiphase with the YD suggesting that any high-latitude northern hemispheric forcing of climate was over-ridden by more regional signals; in the case of MD97-2121 by interacting subtropical and southern polar influences. The Indian Ocean also changed out of phase with the ACR– YD, undergoing a 0.8 °C cooling from ~13.2 to 12 ka (Stenni et al., 2001). In contrast, southern hemispheric sites with sub-Antarctic surface waters, cooled in phase with the

Planktic δ18 O [‰]

Upper ocean response

Productivity

Deep ocean response

Climate

Warm SSTs 17.5–18 °C, peaking at 7–9 ka. SSS similar to modern values but fluctuate in late Holocene. Waters well stratified by 9.5 ka reflecting improved seasonality & dominant Subtropical Inflow. Reduced warming 13–11 ka when SSTs ~16 °C; low seasonality 13.5–11.5 ka & more wind; SSS peak 13–12.5 ka as STW competes with SAW Onset of STW ~ 16– 15 ka although SAW still dominant; SSTs rise from 13 to 15 °C through deglaciation. SSTs rise & δ18Opl lightens from ~ 21 ka to lead benthic δ18Ob (& presumably ice volume) by ~3 kyr; coldest SSTs b 12 °C.

Productivity, as reflected by δ13Cpl G. bulloides, stable at ~ −1‰ in early Holocene with minor variability in late Holocene, possibly reflecting slight increase in ENSO.

Deglacial trend to lighter δ18Ob stabilizes at ~9 ka, but δ13C continues to increase suggesting pulses of NADWinfluenced CDW.

Expansion of forests point to moist, mild equable conditions through early Holocene with a strengthening ENSO overprint in late Holocene.

Following phase of high, fluctuating production, rapid decline from 13.8 ka to stable conditions through ACR–YD.

δ18Ob paused in phase with ACR; pulses of light δ13C suggesting influxes of Antarctic source waters.

Cooler windier climate, dry in E; moist in W where alpine glaciers expanded; more polar than ENSO-driven.

High with fluctuating phases of productivity centred on 20 ka & 15 ka; correlate with similar phases south of STF. High, but fluctuating phases of productivity.

δ18Ob begins trend to lighter values at 18 ka, lagging upper ocean by ~ 3 kyr; δ13Cb was light suggesting increased Antarctic source. Heaviest δ18Ob values with LGM-deglacial profile in phase with Antarctic temperatures.

Expansion of forest at expense of grassland & shrubland in N. Island from ~17 ka in concert with ≫ methane & CO2 in Antarctica. Colder (~ 5 °C lower than now) windier, drier with glaciers at maximum extent; grassland & shrubland prevail.

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Table 2 Summary of ocean responses for 24 to 4 ka as interpreted for MD97-2121 with the δ18Opl planktic profiles for reference

STW = Subtropical (surface) Water; SAW = Sub-Antarctic (surface) Water; NADW = North Atlantic Deep Water; CDW = Circumpolar Deep Water; SST = sea surface temperature; SSS = sea surface salinity.

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ACR as indicated for the SW Pacific (Fig. 1; Pahnke et al., 2003), the SE Pacific (Lamy et al., 2004) and the S Atlantic (Kanfoush et al., 2000; Sachs et al., 2001; Shemesh et al., 2002; Gersonde et al., 2003). Thus, the shallow and deep ocean has changed in phase with the ACR where there are direct climatic/oceanic links with Antarctica. However, such changes become modulated at middle latitudes where polar influences may be buffered or over-ridden by subtropical forces. Acknowledgments This study was supported mainly by the New Zealand Foundation for Research Science and Technology grant CO1X0037 with additional funds from the Marsden Grant GNS0401. We are grateful to the officers and crew of R.V. Marion DuFresne, for acquiring the core, and to Helen Neil for helpful discussions. We especially thank Katharina Pahnke for providing the alkenone-based SSTs, and to Will Howard, Dick Kroon, Peter Barrett, Nancy Bertler and Andrew Mackintosh for their constructive reviews of an earlier version of this manuscript. The manuscript also benefited from commentaries by the journal reviewers. References Andres, M.S., Bernasconi, S.M., McKenzie, J.A., Rohl, U., 2003. Southern Ocean deglacial record supports global Younger Dryas. Earth Planet. Sci. Lett. 216, 515–524. Basher, R.E., Thompson, C.S., 1996. Relationship of air temperatures in New Zealand to regional anomalies in sea-surface temperature and atmospheric circulation. Int. J. Climatol. 16, 405–425. Boyd, P., LaRoche, J., Gall, M., Frew, R., McKay, R.M.L., 1999. The role of iron light and silicate in controlling algal biomass in sub-Antarctic water SE of New Zealand. J. Geophys. Res. 104, 13395–13408. Blunier, T., Brook, E.J., 2001. Timing of millennial-scale climate change in Antarctica and Greenland during the last glacial period. Science 291, 109–112. Blunier, T., Schwander, J., Stauffer, B., Stocker, T., Dällenbach, A., Indermühle, A., Tschumi, J., 1997. Timing of the Antarctic Cold reversal and the atmospheric increase with respect to the Younger Dryas event. Geophys. Res. Lett. 24, 2683–2686. Broecker, W.S., 1994. Massive iceberg discharges as triggers for global climate change. Nature 372, 421–424. Carter, L., Wilkin, J., 1999. Abyssal circulation around New Zealand — a comparison between observations and a global circulation model. Mar. Geol. 159, 221–239. Carter, L., McCave, I.N., 2002. Case studies 101–103, Eastern New Zealand Drifts: Miocene to Recent. In: Stow, D.A.V., Pudsey, C.J., Howe, J., Faugeres, J.-C., Viana, A.R. (Eds.), Deep-Water Contourites: Modern Drifts and Ancient Series, Seismic and Sedimentary Characteristics. Geol. Soc. Mem., vol. 22. 472 pp. Carter, L., Manighetti, B., 2006. Glacial/interglacial control of terrigenous and biogenic fluxes in the deep ocean off a high input, collisional margin; a 139 kyr-record from New Zealand. Mar. Geol. 226, 307–322.

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