SEARCH

SEARCH BY CITATION

Keywords:

  • paleoceanography;
  • stable isotopes;
  • descriptive regional oceanography

Abstract

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Environment of Campbell Plateau
  5. 3. Methodology
  6. 4. Results
  7. 5. Discussion
  8. 6. Conclusions
  9. Acknowledgments
  10. References

[1] Campbell Plateau occupies a key position in the southwest Pacific sector of the Southern Ocean. The plateau confines and steers the Antarctic Circumpolar Current (ACC) along its flanks, isolating the Subantarctic plateau from cold polar waters. Oxygen and carbon isotope records from Campbell Plateau cores provide new records of water mass stratification for the past 130 kyr. During glacial climes, strengthening of the Subantarctic Front (SAF) caused waters over the plateau flanks to be deeply mixed and ∼3°C cooler. Waters of the plateau interior remained stratified and isolated from the cold southern waters. In the west, waters cooled markedly (∼4°C) owing to reduced entrainment of Tasman Sea water. Marked cooling also occurred north of Campbell Plateau under increased entrainment of polar water by a branch of the SAF. The ACC remained along the flanks of Campbell Plateau during the last interglacial, when interior waters were stratified and warmer by ∼1°C than now.

1. Introduction

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Environment of Campbell Plateau
  5. 3. Methodology
  6. 4. Results
  7. 5. Discussion
  8. 6. Conclusions
  9. Acknowledgments
  10. References

[2] Campbell Plateau, off southern New Zealand, forms one of the major constrictions of the Antarctic Circumpolar Current (ACC). As a result, the flow is intensified, disrupted into large eddies and diverted northward to as far as ∼ 50°S where it resumes its eastward path (Figure 1; Gordon [1972]; Orsi et al. [1995]; Morris et al. [2001]). The role of the ACC in affecting water masses and circulation patterns during paleoclimatic fluctuations has been the subject of much debate. Nelson et al. [1993], for example, suggested that the Subantarctic Front (SAF), the frontal jet defining the northern edge of the ACC, migrated northward in glacial times to ∼45°S where it enhanced the Subtropical Front situated along the crest of Chatham Rise (Figure 1). Certainly, in the open ocean, the Subantarctic Front migrated 3–5° northward in glacial periods [e.g., Howard and Prell, 1992]. Others propose that the ACC remained bathymetrically locked to the edge of Campbell Plateau and that the strong thermal gradients across Chatham Rise resulted from glacial cooling of surface waters [e.g., Weaver et al., 1998]. This second alternative raises a further question: Would a “locked” ACC affect the water masses over the Campbell Plateau interior?

image

Figure 1. Location of Campbell Plateau cores, along with major surface water masses and oceanographic fronts. Circulation after Carter et al. [1998], Davis [1998], and Morris et al. [2001]. Abbreviations are as follows: STW, Subtropical Water; SAW, Subantarctic Surface Water; CSW, Circumpolar Surface Water; STF, Subtropical Front; SAF, Subantarctic Front marking the northern of the Antarctic Circumpolar Current; SF, Southland Front.

Download figure to PowerPoint

[3] In this paper we address issues relating to the SAF/ACC by presenting new data from the Southern Ocean based on stable isotope analysis of 8 cores from Campbell Plateau. Temporal and spatial changes in water mass structure and circulation during the past 130 kyr are defined, providing a new perspective of the paleocirculation southeast of New Zealand under glacial and interglacial conditions.

2. Environment of Campbell Plateau

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Environment of Campbell Plateau
  5. 3. Methodology
  6. 4. Results
  7. 5. Discussion
  8. 6. Conclusions
  9. Acknowledgments
  10. References

2.1. Physiography

[4] Campbell Plateau is a submerged continental platform protruding into the southwest Pacific Ocean for ∼1100 km from the southeastern margin of South Island, New Zealand (Figure 1). For much of its area, plateau depths are 1000–600 m, shoaling to 250 m, and comprising a number of broad rises and depressions. Surface expression occurs in the form of several Subantarctic Island groups, including the volcanic Auckland and Campbell Islands. The Plateau margins are precipitous, plunging from 1000 m to more than 4000 m, often at gradients of 1:4. Campbell Plateau is separated from Chatham Rise in the north by the Bounty Trough [Carter and Carter, 1993], and from the adjacent Bounty Plateau by the ∼1400 m deep Pukaki Saddle which forms a distinct breach in the major topographic boundary steering the SAF of the ACC [Summerhayes, 1969].

2.2. Sediments

[5] The interior of Campbell Plateau is mantled by pelagic carbonate sediments. Similar sediments are also present on the flanks but the cover is patchy due to erosion by the powerful ACC [e.g., Glasby, 1976; Carter, 1989], with mass failure also playing a role. At depths >2000 m, the Deep Western Boundary Current (DWBC), aided by the ACC, have redistributed sediment and formed thin drifts along the plateau base [Carter and McCave, 1997; Carter et al., 1999]. The extensive biopelagic cover is generally fine-grained, white and homogeneous, consisting mainly of foraminifera, coccolith plates, and noncarbonate material (typically ∼10–20%). Quartz-dominant terrigenous sediment and biosiliceous remains of radiolaria and sponge spicules are main constituents of the sand fraction in the noncarbonate component. Holocene carbonate content averages 85% with no distinct variation with depth or geographical location across the shallow plateau [Summerhayes, 1969; Carter et al., 2000]. Sedimentation rates are generally low over Campbell Plateau and dominated by pelagic deposition [Carter et al., 2000].

2.3. Oceanography

[6] Campbell Plateau margin separates Subantarctic from Circumpolar surface waters, as well as constricting the flow of the eastward flowing ACC [Burling, 1961; Gordon, 1972; Orsi et al., 1995]. Constriction causes a pronounced northward deflection of the leading edge of the ACC, marked by the SAF, along the eastern plateau margin [Heath, 1981; Carter and Wilkin, 1999]. As a result, the flow is topographically intensified with mean current speeds of 27–39 cm s−1 recorded within a ∼150 km wide swath centered on the plateau slope [Stanton and Morris, 2004]. Nevertheless, sectors of the ACC resume their eastward flow at ∼55°S [e.g., Orsi et al., 1995] and at ∼50°S [e.g., Bryden and Heath, 1985; Carter et al., 1998]. The western plateau also hosts a northward current, this time associated with the STF. Termed the Southland Current, it has a mean speed of ∼20 cm s−1 [Chiswell, 1996; Sutton, 2003].

[7] At depth the flow is dominated by a large DWBC [Carter and McCave, 1994], which enters the New Zealand region through gaps and around Macquarie Ridge. This thermohaline inflow, extending from 2000 to ∼5000 m deep, finds the western boundary presented by Campbell Plateau and passes north in consort with the ACC to about 50°S, where divergence from the ACC occurs, leaving the DWBC to continue over Bounty Trough, while the ACC veers east across the Pacific Ocean.

[8] Surface oceanographic conditions off southernmost New Zealand are influenced by two distinctly different regimes of circulation: the intense and variable ACC, and weak, gyre-like circulation on Campbell Plateau. The ACC is a composite flow bounded by the strong zonal jets associated with the SAF and Antarctic Polar Front to the north and south respectively [e.g., Peterson and Whitworth, 1989; Paterson and Whitworth, 1990]. The intervening flow is generally weaker. Sharp transitions in water properties and current speed along the fronts extend deep into the water column. Strong northeastward flow in the SAF, and its entrained Subantarctic Mode Water (SAMW), is confined to the shelf edge along Campbell Plateau flanks [Morris et al., 2001] (Figure 2). Trajectories of deep floats reported by Davis [1998] show large eastward displacements, indicating a swift, deep current within the ACC, south of New Zealand. Several deep float tracks head north trapped within the SAF. Other trajectories are more complicated especially in the vicinity of Bounty Plateau and Bounty Trough, although there is a strong tendency to align with the bathymetry and either turn eastward into the southwest Pacific Ocean or pass through Pukaki Saddle, to continue as a cyclonic flow around the head of Bounty Trough [Morris et al., 2001] (Figures 1 and 2).

image

Figure 2. (a). Geostrophic velocities (black and gray vectors) relative to the bottom during hydrographic surveys in May 1998, December 1998, and August 1999. Note the strong geostrophic flow in the SAF along the eastern flank of Campbell Plateau. (b) Time average results for the three surveys for Campbell Plateau section only; note change of scale. The flow vectors on the Campbell Plateau are an order of magnitude smaller than those around the flanks. (c) Repeat upper ocean temperature sections along the southeast track shown in Figure 2a, from South Island, New Zealand, over Campbell Plateau and off the shelf edge. Note the occurrence of seasonal stratification during summer and weak stratification during winter over Campbell Plateau. The contour interval is 0.5°C.

Download figure to PowerPoint

[9] In sharp contrast, current velocities on Campbell Plateau itself are weak, with a mean flow of <10 cm s−1 [Morris et al., 2001] (Figure 2). Although sluggish, these currents are persistent, with weak anticyclonic circulation occurring around Pukaki Rise [Morris et al., 2001], a counter flow inshore of the ACC and cyclonic circulation centered between Campbell Island and Pukaki Rise (Figure 2).

[10] At shallow depths the water over Campbell Plateau is cool, fresh Subantarctic Surface Water (SAW). This water mass is hydrologically and biologically distinct from warmer, saltier Subtropical Surface Water (STW) to the north, and colder, fresher Circumpolar Surface Water (CSW) in the south. These waters overlie Antarctic Intermediate Water (including a shallow SAMW component north of the SAF) and Circumpolar Deep Water (CPDW). Immediately north of the SAF, mixing occurs to ∼400 m; this deep mixing can extend over the Plateau flanks during winter. Additionally, extending from the SAF is a region of low temperature stratification associated with SAMW which extends between ∼200 m and the Campbell Plateau seabed at 600–1000 m depth (Figure 2). Consequently, waters over a large portion of the plateau are weakly stratified during winter, with few periods of active mixing observed. In summer the plateau waters are strongly stratified due to seasonal heating (Figure 2) [Morris et al., 2001].

[11] North of Campbell Plateau lies the Subtropical Front (STF), the northern limit of the Southern Ocean and the boundary with STW. Like the SAF, the STF is deflected around New Zealand. Presently it approximates 45°S in the Tasman Sea, but is deflected around the South Island continental margin to form the Southland Front [Heath, 1985; Chiswell, 1996]. The Southland Front resumes as the STF at Chatham Rise (∼43°S), before shifting south once free of Chatham Rise crest [e.g., Heath, 1985; Chiswell, 1994] (Figure 1).

2.4. Wind Systems

[12] Campbell Plateau currently lies within the influence of the vigorous “Roaring Forties” westerly wind system. Fluctuations in strength and latitude of the westerlies occur seasonally, with the winds expanding northward in winter and receding southward during summer [Markgraf et al., 1992]. Autumn is the windiest season, followed by spring [Reid and Penny, 1982; Reid and Collen, 1983]. Mean wind speeds of 30 km hr−1 with gusts of 96 km hr−1, occur on average for 106 days a year at Campbell Island [New Zealand Meteorological Service, 1981]. During glacial times, atmospheric circulation in the southwest Pacific was enhanced due to an increased pole-equatorial thermal gradient with expanding Antarctic ice shelves and sea ice [Thiede, 1979; Markgraf et al., 1992]. The core of the strengthened westerlies probably shifted significantly northward, and may have lain over New Zealand during such times [e.g., Thiede, 1979; Stewart and Neall, 1984].

3. Methodology

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Environment of Campbell Plateau
  5. 3. Methodology
  6. 4. Results
  7. 5. Discussion
  8. 6. Conclusions
  9. Acknowledgments
  10. References

[13] This study is based on 5 Kasten cores collected from Campbell Plateau during National Institute of Water and Atmosphere (NIWA) research voyages 3034 (1997) and 3046 (1998) on R/V Tangaroa (Table 1). These samples were supplemented by giant piston cores SO136-38 and SO136-55, collected during voyage SO136 TASQWA (1998) on R/V Sonne [Nees et al., 1999], and sediment from Ocean Drilling Program Site 1120 collected during Leg 181 (1998) [Carter et al., 1999]. These define roughly two north-south transects across the eastern margin crossing the northern boundary of the SAF and interior of Campbell Plateau (Figure 1).

Table 1. Summary of Core Locations, Sedimentation, and Accumulation Rates for MISs 1 and 2a
CoreLatitude, SLongitude, EDepth, mCalendar Age, kyrMIS 1MIS 2
Sedimentation Rate, cm/kyrCarbonate MAR, g/cm2/kyrNoncarbonate MAR, g/cm2/kyrSedimentation Rate, cm/kyrCarbonate MAR, g/cm2/kyrNoncarbonate MAR, g/cm2/kyr
V143952° 54.51′168° 07.83′79115.423 ± 0.168 at 48 cm3.32.60.631.91.70.76
    37.630 ± 0.460 at 92 cm      
    > 48.440 at 144 cm      
Y948° 14.21′177° 20.67′12677303 ± 0.112 at 16 cm2.53.41.100.91.931.21
    13.183 ± 0.090 at 52 cm      
    27.270 ± 0.120 at 64 cm      
    48.440 ± 1.800 at 92 cm      
Y1451° 20.11′171° 54.20′523 1.81.60.211.11.40.21
Y1650° 35.43′169° 45.33′60042.640 ± 0.660 at 68 cm2.82.50.6121.60.54
Y1748° 23.31′169° 31.73′691 3.83.70.701.92.30.74
SO136-3850° 13.43′175° 18.73′1359 1.71.80.123.330.30
SO136-5550° 09.61′173° 21.91′563 1.71.60.191.51.50.19
ODP112050° 03.80′173° 22.30′543 2.62.50.2810.90.10

[14] All cores lie beneath SAW (Figure 1). Samples for analysis were collected at 4 cm and 5 cm intervals for NIWA (prefixed Y and V) and TASQWA (SO-136) cores respectively (Table 1). A lower and variable resolution of 5 to 15 cm was used for ODP Site 1120. Magnetic susceptibility was measured using a handheld Bartington Instruments MS2 sensor and probe or, in the case of TASQWA cores, a GEOTEK Multisensor core logger also with Bartington sensor. Samples were disaggregated following drying at 50°C by soaking in buffered solution (pH = 9.4) and wet sieving to separate out the >63 μm size fraction. Sand fractions were subsequently split by dry sieving enabling determination of percent by weight of coarse sand (>125 μm), fine sand (63–125 μm) and mud (<63 μm). These data were used to detect any textural response to the SAF. However, the locations of the cores are insufficient to define the southern extent of the topographically intensified SAF, and we have relied on the local marine geology [Carter and McCave, 1997] and physical oceanography [Carter and Wilkin, 1999; Stanton and Morris, 2004], which show a strong SAF flow, about 150 km wide, and centered over the plateau's steep margin. The carbonate content of dried powdered samples was determined via gasometric quantitative analysis after acidification [Jones and Kaiteris, 1983], with a precision of ±2%.

[15] Stable isotope measurements were performed on the planktonic foraminifera Globigerina (Gg.) bulloides and Globorotalia (Gr.) inflata picked from the >200 μm fraction of washed sediment samples and ultrasonically cleaned in methanol. Specimens were reacted with 2 drops of 100% H3PO4 for 10 min at 75°C in a Finnigan MAT automated individual carbonate (Kiel) reaction device. Oxygen (δ18O) and carbon (δ13C) isotopes were determined from the resultant liberated CO2 analyzed by a Finnigan MAT 252 mass spectrometer. Concurrently run carbonate standards (NBS-19) had an internal precision of ±0.04‰ for δ13C and ±0.08‰ for δ18O, and an external precision (between runs), calculated as a difference between standard means, of ±0.02‰ and ±0.06‰ respectively; results are reported relative to Vienna Peedee belemnite. δ18O results were used to estimate near-surface seawater temperatures from the calibration equation of Epstein et al. [1953] with salinity correction from the relationship of Craig and Gordon [1965] and ice volume effect compensated according to Schrag et al. [2002]. We use the difference in δ18O between Gg. bulloides and Gr. inflata as an index of water stratification. Marine isotope stage classifications are assigned in accord with the description in the later age model section, while AMS dating of a series of monospecific Gr. inflata samples from within Campbell Plateau provides age-depth information that can be combined with the oxygen isotope stratigraphy (Figure 3a). For the purposes of this study, the ages are used to supplement the isotope stratigraphy, with principal emphasis being placed on events defined from the isotope curves.

image

Figure 3. Downcore profiles of (a) δ18O, (b) magnetic susceptibility and calcium carbonate content, and (c) grain size with depth for cores from the plateau interior, western plateau and eastern margin proximal to the Antarctic Circumpolar Current. Shaded intervals correspond to glacial periods. Foraminiferal δ18O records are broadly in phase with the carbonate record while increased coarse fraction percentages often coincide with low carbonate percent and δ18O enrichment. For the plateau interior cores the isotope stratigraphy is significantly more defined than the carbonate stratigraphy, while grain size is often coarser during times of δ18O depletion. Isotope taxonomy is from Prell et al. [1986], and age assignments follow Imbrie et al. [1984], Martinson et al. [1987], and Imbrie et al. [1992] with the MIS 2/3 boundary assigned an age of 27 kyr [Lean and McCave, 1998].

Download figure to PowerPoint

4. Results

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Environment of Campbell Plateau
  5. 3. Methodology
  6. 4. Results
  7. 5. Discussion
  8. 6. Conclusions
  9. Acknowledgments
  10. References

4.1. Age Model

[16] Age models for the Campbell Plateau cores are based upon isotopic profiles of Gr. inflata, supplemented by AMS 14C dates (Figure 3a). AMS dates are calibrated to “calendar years” using the Calib v4.3 program of Stuiver and Reimer [1993]), with the marine calibration data set of Stuiver et al. [1998]. This data set uses a time-dependent global ocean reservoir correction of about 400 years, but we recognize that this age may eventually be revised following the work of Sikes et al. [2000], who suggest that surface waters in the southwest Pacific Ocean may have reservoir ages of up to 2000 years for the Last Glacial Maximum. Additional control for correlation between core sites is provided by calcium carbonate abundance, magnetic susceptibility and grain size data (Figures 3b and 3c). Isotopic events defined by Prell et al. [1986] and age assignments following Imbrie et al. [1984], Martinson et al. [1987]) and Imbrie et al. [1992] are adopted, with the exception that the Marine Isotope Stage (MIS) 3/2 boundary is assigned an age of 27 kyr on the basis of dated local tephra [Lean and McCave, 1998]. Ages between control points are linear interpolations. This study addresses the last 130 kyr using a sample-spacing resolution on the order of 1–2 kyr. The likely core top age has assessed by comparing the thickness of the oxidized layer remaining to that found in box or multicores recovered from the area. This indicates that loss is typically 2–3 cm, which, at an average sedimentation rate of 2–3 cm/kyr, equates to about 1 kyr.

4.2. Isotopes

[17] Oxygen and carbon isotopic profiles from Gg. bulloides and Gr. inflata are presented against age in Figure 4. MISs 1–4 are present in all cores, with five cores extending back as far as MIS 6. MIS 2 planktic δ18O values are characteristically between 3.7 and 2.6‰, while fully postglacial conditions in MIS 1 are around 2 to 1‰. Since sedimentation rates are low, (∼2 cm/kyr), no pause/reversal, or other detail is discernible across the deglaciation. Oxygen isotope values for MIS 5 range from 2.7 to 1.5‰. Unexpectedly, they are in many cases less depleted than MIS 1 values.

image

Figure 4. Downcore profiles of (a) δ18O and (b) δ13C against age for cores of the Campbell Plateau. Shaded intervals correspond to glacial periods. (c) Isotope temperature difference between Gg. bulloides and Gr. inflata during MISs 1, 2, and 5. The isotopic/temperature gradient between Gg. bulloides and Gr. inflata serves as a proxy for stratification, with Plateau margin cores exhibiting small gradients (where deep mixing occurs north of the SAF). Cores from the plateau interior (where today the water masses are stratified) exhibit greater and more variable gradients. δ13C profiles show offsets between the two species are markedly different in time and space. Of note are enriched δ13C values in Gg. bulloides, associated with enriched values of δ18O, especially during MIS 5.

Download figure to PowerPoint

[18] The amplitude of change in δ18O (Table 2) between MISs 2 and 1 varies across the plateau, with generally larger amplitudes recorded in cores from the plateau margins. Likewise the gradient between Gg. bulloides and Gr. inflata, also varies across the plateau. At the plateau margins the absolute values of the two species are similar, while Gr. inflata is usually more enriched than Gg. bulloides across the plateau interior (Figure 4c). Interestingly, however, Gr. inflata in most cores is more depleted than Gg. bulloides during MIS 5 (Figure 4a).

Table 2. δ18O and Inferred SST for Campbell Plateau Region During MIS 1, MIS 2, and MIS 5a
 Observed SST,b °CCoretop SST,c °CMIS 1 δ18OMIS 2 δ18OMIS 5e δ18OΔ1-2 δ18OResidual,d δ18O MIS 2Δ1-5 δ18OΔT1-2, °CΔT1-5, °C
  • a

    Temperature changes (ΔT) are calculated assuming a 1°C temperature change is equivalent to 0.24‰ δ18O [Hecht, 1985].

  • b

    SST values taken from results of three repeat sections across Campbell Plateau reported in the work of Morris et al. [2001].

  • c

    Coretop values determined from the equation of Epstein et al. [1953]; ambient seawater estimated using the relationship of Craig and Gordon [1965].

  • d

    Residual δ18O is δ18O change from MIS 2 to MIS 1, corrected for the global ice volume effect of 1.0‰ [Schrag et al., 2002].

V1439
G. bulloides77.21.833.742.51−1.91−0.91−0.68−3.96−2.96
G. inflata 6.71.913.842.25−1.93−0.93−0.34−4.04−1.48
 
Y14
G. bulloides7.57.11.732.761.60−1.03−0.030.13−0.130.57
G. inflata 6.41.882.841.96−0.960.04−0.080.17−0.35
 
Y16
G. bulloides8.281.723.301.81−1.58−0.58−0.09−2.52−0.39
G. inflata 7.51.733.411.66−1.68−0.350.07−1.520.30
       1.00   
 
Y17
G. bulloides9.59.20.952.99-−2.04−1.04-−4.52-
G. inflata 9.21.213.21-−2.00−1.00-−4.35-
 
SO136-55
G. bulloides8.38.11.101.861.42−0.760.24−0.321.04−1.39
G. inflata 8.81.132.100.95−0.970.030.180.130.78
 
ODP1120
G. bulloides7.57.41.882.581.98−0.700.30−0.101.30−0.43
G. inflata -1.822.491.55−0.670.330.271.431.17
 
SO136-38
G. bulloides8.68.51.523.321.99−1.80−0.80−0.47−3.48−2.04
G. inflata 9.21.423.111.41−1.69−0.690.01−3.000.04
 
Y9
G. bulloides8.48.71.543.362.20−1.82−0.82−0.66−3.57−2.87
G. inflata 6.12.113.031.48−0.920.080.630.352.74

[19] Overall, the foraminiferal δ13C records exhibit less pronounced amplitude than the δ18O records (Figure 4). MIS 2 δ13C values fluctuate between −1.5 and 1‰, depending on core location while MIS 1 values are characteristically between 0.5 and 1.5‰. MIS 5 values are between 0 and 1‰. Throughout most of the records the absolute values of the two species are markedly different with Gg. bulloides more depleted than Gr. inflata. Of note is the relative enrichment of Gg. bulloides during MIS 5 in several of cores (Figure 4).

4.3. Sedimentation

[20] Cores from Campbell Plateau flanks are typically composed of bright white, coccolith-rich foraminiferal ooze belonging to MIS 5, overlain by MISs 2–4 light olive gray foraminiferal mud (dominantly carbonate) grading into light gray foraminiferal ooze at the core top. Lithologies of the plateau interior are similar except that MIS 2 is represented by light gray foraminiferal ooze.

[21] The carbonate records from the plateau flanks are broadly in phase with the foraminiferal δ18O records such that alternating biopelagic and hemipelagic oozes (e.g., Y9, V1439), correspond to periods of relative 18O depletion and enrichment respectively. By contrast, for cores from the plateau interior, the isotope stratigraphy is significantly more defined than the carbonate stratigraphy.

[22] Sedimentation rates and mass accumulation rates are calculated for the Campbell Plateau region (Table 1). Last Glacial (MIS 2) sedimentation rates are in the order of 1 to 2 cm kyr-1, increasing to 2 to 3 cm kyr−1 during the Holocene (MIS 1). Carbonate accumulation rates are generally higher during MIS 1 (≥2 g cm−2 kyr−1) than MIS 2 (∼2 g cm−2 kyr−1), with noncarbonate flux subordinate during both climatic periods (≤1 g cm−2 kyr−1). Carbonate accumulation, although low, varies primarily as a result of biologic production combined with dissolution and changing climate. Noncarbonate accumulation rates are also variable, with the major variations resulting primarily from current transported sediment deposited around the plateau edges, as evinced by the coarser grain sizes of glacial period sediments (Figure 3c). Variations in noncarbonate accumulation may also be partially attributed to increased aeolian input during the Last Glaciation [e.g., McGlone et al., 1994].

5. Discussion

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Environment of Campbell Plateau
  5. 3. Methodology
  6. 4. Results
  7. 5. Discussion
  8. 6. Conclusions
  9. Acknowledgments
  10. References

5.1. Sedimentary Regime

[23] The downcore carbonate profiles (Figure 3b) are consistent with profiles from elsewhere in the New Zealand region with interglacial biopelagites alternating with glacial pelagites−hemipelagites. However, the Campbell Plateau receives little terrigenous material as it is topographically and geographically isolated from New Zealand. The influx of terrigenous sediment from this tectonically active landmass largely bypasses the plateau by way of the Bounty and Solander Channel systems [Carter et al., 1996]. Therefore reduced glacial carbonate percentage (Figure 3b) is mainly attributable to changes in carbonate production as well as to variations in aeolian and biosiliceous input [e.g., Thiede, 1979; Hesse, 1994; Newnham et al., 1999]. Limited carbonate dissolution is reported to occur as shallow as 1370 m [e.g., Weaver et al., 1998; Carter et al., 2000], deeper than the depth of cores in this study. Cores proximal to the plateau edge generally exhibit greater amplitude of carbonate fluctuations as a consequence of a higher terrigenous input due to (1) lateral transport in the ACC, (2) current reworking, and (3) increased incidence of ice-rafted debris [e.g., Carter et al., 2002b].

[24] Downcore fluctuations in bulk sediment texture also express changes in biopelagic input and the hydraulic regime. On the margin, grain size changes from coarser in MIS 2 to finer during MIS 1, whereas in the interior, the effect is reversed. For example, compare V1439 from the margin, to Y14 from the plateau interior (Figure 3c). Cores near the plateau edges are proximal to the vigorous ACC/SAF. Higher coarse fraction percentages exhibited during MIS 2 by cores closest to the plateau flanks (e.g., V1439, SO136-38) (Figure 3c) suggest an intensified ACC/SAF circulation and sediment transport. However, cores of the relatively quiescent plateau interior are more likely to respond to comparative changes in cocclithophorid plates or carbonate fragments as opposed to foraminifera or larger carbonate tests (Figure 3c). During MIS 2, cores of the plateau interior exhibit higher fine fraction percentages consistent with reduced foraminiferal production and a sustained regime of sluggish circulation. Low resolution analysis of the bulk fine fraction (<2 μm, mainly remains of coccoliths) reveal a maxima at this time. Low silicic acid concentrations (<4 μM) occur north of the SAF across the Subantarctic zone [Daly et al., 2001]. Likewise low silica values (<4 μM) were also recorded across Campbell Plateau within the upper 500 m (K. Currie, personal communication, 2002). Low silicic acid concentrations limit opaline productivity in this region without limiting the productivity of the remaining carbonate assemblage. Hence diatoms and radiolaria are rare to absent in the resultant community/sediment composition.

5.2. Sea Surface Temperatures

[25] Gg. bulloides is a transitional to polar foraminiferal species [Hayward, 1983; Hemleben et al., 1989]. It occurs mainly above the thermocline although it is not restricted to the photic zone, and is associated with periods of enhanced nutrient supply [Ganssen and Kroon, 2000]. Plankton tows for foraminifera over Campbell Plateau recovered a high percentage (∼85%) of this species from the upper 100–150 m of the water column (NIWA Cruise Report TAN 9814, 1998). Gr. inflata is a transitional to subpolar species [Hayward, 1983; Hemleben et al., 1989]. In the Campbell Plateau region, it tends to be encrusted by a deposit of smooth calcite in contrast to the crystalline appearance of other New Zealand region cool water specimens. The depth range of this species is between 200 and 600 m [Ganssen and Sarnthein, 1983] and is again supported by foraminiferal tows conducted across the Plateau, which show a high proportion of Gr. inflata at these depths (NIWA Cruise Report TAN 9814, 1998).

[26] Sea surface temperature (SST) estimates from the δ18O of Gg. bulloides and Gr. inflata are presented in Table 2. The SSTs, calculated from coretop δ18O values are close to modern temperatures recorded near the core sites (Figure 2c), with winter temperatures ranging between 6 and 8°C south to north, while summer temperatures are between 8 and 12°C [Morris et al., 2001]. Northern Campbell Plateau exhibits a 5-year mean satellite SST of 10°C, with southern waters close to 8°C [Uddstrom and Oien, 1999]. Values calculated in this study are comparable with temperatures derived from stable isotope and transfer function data for Campbell Plateau core tops recorded by Weaver et al. [1997].

[27] Comparing average δ18O for periods MISs 1, 2, and, where appropriate, 5e (Table 2), it is apparent that MIS 2 values are enriched relative to MIS 1, but that the amplitude of change varies across the plateau. On the ACC-bathed, southern plateau flank, V1439 exhibits a residual (ice-volume compensated) Δδ18O value that is equivalent to a cooling of the surface waters by ∼4°C during MIS 2, assuming no change in salinity. A similar glacial cooling (∼3.5°C) is also recorded on the eastern flank (SO136-38), also beneath the ACC. Core Y17, located in northwest Campbell Plateau, shows the greatest enrichment over and above the global signal during MIS 2, with a reduction in temperature of ∼4.5°C compared to MIS 1. Enrichment in excess of the global signal is also exhibited by Y16, although the cooling is only in order of 2.5°C.

[28] The Δδ18O record from Y9 on the edge of Bounty Plateau is less clear. Gg. bulloides indicates a ∼ 3.5°C cooling of surface waters during MIS 2, in keeping with other plateau edge cores. However, Gr. inflata shows an isotopic change close to the global mean oxygen ice-volume shift of 1‰. Similarly, Δδ18O of both Gg. bulloides and Gr. inflata from the interior cores (SO136-55, ODP1120, Y14) indicate that MISs 2–1 isotopic gradient is close to the mean global shift. During MIS 5, a plateau-wide Gg. bulloides isotopic signature reveals that surface waters were on average ∼1.5°C cooler than those recorded for MIS 1. This is in contrast with Gr. inflata, which shows both enriched and depleted Δδ18O between MISs 5 and 1.

[29] Accordingly the thermal gradient across the flanks of Campbell Plateau has varied over time. During MIS 1 the thermal gradient across the southern flank (V1439 to Y14) was 0.7°C 100 km−1. Similarly, across the eastern flank of Campbell Plateau a thermal gradient of 0.5°C 100 km−1 existed. This thermal gradient increased markedly during MIS 2 with the southern and eastern flanks exhibiting thermal gradients of 2.3°C 100 km−1 and 3.0°C 100 km−1 respectively.

[30] Sea surface salinity during MIS 2 has been estimated to have been ∼0.3‰ higher than the present day, using the equation of Broecker [1989] as applied by Rostek et al. [1993] and Martinez et al. [1997], assuming a linear S:δ18O relationship, and extrapolating the temperature change determined by Weaver et al. [1998] from transfer function analysis. This magnitude of change in salinity could account for up to 0.7°C of the estimated 3–4°C MIS 2 cooling. Rohling and Bigg [1998] consider a linear S:δ18O relationship may not be valid for all sectors of the ocean. However, although the relationship is unknown, Campbell Plateau is a region outside the extent of glacial sea-ice and does not receive direct inflow from any riverine system. While the extent of glacial ice cover on the nearby South Island of New Zealand and the influx of icebergs to this region are likely to have increased during MIS 2 [e.g., Carter et al., 2002b], these changes are less than that for many other sensitive sectors of the global ocean. Consequently, Campbell Plateau could be considered a “stable” region with respect to the linear relationship of S:δ18O, and the error associated with using a linear relationship is probably small.

[31] Irrespective of the explanation for the δ18O residual, higher surface water density is implied by decreased temperature, increased salinity, or a combination of both during MIS 2. As a result conditions are likely to have favored reduced stratification, to some extent.

5.3. Stratification

[32] It is well established that differences in isotopic composition between two species of planktonic foraminifera can serve as a proxy for stratification of the upper water column through time [e.g., Mulitza et al., 1997; Ortiz et al., 1997; Niebler et al., 1999; Ganssen and Kroon, 2000]. In particular, studies from the temperate southern South Atlantic Ocean, as well as in the ACC [Niebler et al., 1999] and the northeast Atlantic [Ganssen and Kroon, 2000] indicate that the shallow-calcifying Gg. bulloides and the deep- to intermediate-calcifying Gr. inflata are most applicable for the reconstruction of past water stratification. The difference between the isotopic signatures of Gg. bulloides and Gr. inflata is independent of ice volume, since the calcite of both species records the global change in ice volume in the same way. Similarly, vital effects or disequilibrium offset due to changing carbonate chemistry of the ocean may be considered constant through time between the two species.

[33] One factor that can introduce uncertainty is that isotopic gradients between the species may, in part, be driven by a seasonal difference in their production peaks rather than vertical stratification or mixing. On Campbell Plateau, sediment trap records (H. L. Neil and L. C. Northcote, manuscript in preparation, 2004) show that 90% of the present-day annual mass flux occurs in a single pulse during spring. The foraminiferal population associated with this event is dominated by Gr. inflata (∼60%) followed by Gg. bulloides (∼20%). Subantarctic sediment traps located north of the Campbell Plateau, on south Chatham Rise, also exhibited a single production peak of foraminifera in spring dominated by Gg. bulloides (∼78%) and Gr. inflata (∼15%) [King and Howard, 2001]. Hence it is probable that seasonality has a negligible effect and the derivation of different isotope values (or a large isotopic gradient) between the two species is a result of stratified waters, whereas similar isotope values will occur in a mixed upper water column [e.g., Niebler et al., 1999; Ganssen and Kroon, 2000]. It would therefore be expected that Gr. inflata would reflect cooler water temperatures than coeval Gg. bulloides within a stratified water mass.

[34] The δ18O isotopic gradient between Gg. bulloides and Gr. inflata from around Campbell Plateau edge (e.g., V1439, SO136-38) is small throughout the cored interval (Figure 4). We attribute this to deep mixing occurring immediately north of the SAF. In contrast, samples from the plateau interior, where today the surface waters are more stratified, exhibit larger, more variable isotopic gradients between the two species. For example, in northwest Campbell Plateau, Y17 exhibits distinct separation between the isotopic profiles of the two species during MIS 1 and MISs 3–4, with Gg. bulloides being more depleted in δ18O than Gr. inflata. However, values of δ18O converge during MIS 2 thereby indicating a weakening of stratification. Conversely, in Y16, from the western interior, behavior is much like that of cores located around the plateau edge, with very little separation between the species isotopic profiles. Isotopic records from Y9, on Bounty Plateau, show Gg. bulloides more depleted that Gr. inflata during MIS 1, with reversed behavior in MIS 2, while the two species signatures converge during MISs 3–4.

[35] An apparent anomaly occurs across much of Campbell Plateau during MIS 5 when isotope records indicate Gr. inflata is consistently more depleted in δ18O than Gg. bulloides. This resembles behavior noted for Y9 during MIS 2 and for large parts of the record from SO136-55 and ODP 1120. We postulate that this incongruity is a consequence of faunal response to the short-term spring bloom, and is discussed further in a later section.

5.4. Source Waters

[36] The possible sources of surface water off eastern New Zealand have been investigated previously using coretop δ13C values [Hendy, 1995; Neil, 1997; Nelson et al., 2000]. A similar approach is adopted here using Gg. bulloides δ13C data from the Chatham Rise, Bounty Trough and Campbell Plateau.

[37] Carbon isotopic variation in foraminiferal tests can provide qualitative information about regional changes in water mass production and distribution, although a range of factors complicates interpretation. Variability in δ13C may be a function of thermodynamic effects at the site of formation [e.g., Oppo and Fairbanks, 1989], “vital effects” causing fractionation during the uptake of carbon [e.g., Ravelo and Fairbanks, 1995; Zeebe et al., 1999], dissolution [e.g., Berger, 1971; Shackleton and Opdyke, 1976] or habitat [e.g., Mulitza et al., 1999]. In addition, global carbon isotopic values are typically 0.3–0.4‰ lower during glacial times compared to the Holocene [Curry et al., 1988; Duplessy et al., 1988] due to terrestrial biomass [Shackleton, 1977; Curry and Crowley, 1987] and/or carbonate ion effects [e.g., Lea et al., 1999].

[38] However, regional changes in δ13C may be assessed, assuming disequilibrium within a species is constant through time, by two main strategies [Mulitza et al., 1999]. The first uses carbon isotope values of different species derived from a single core, thus canceling out global and regional effects, with the aim of highlighting vertical or seasonal differences in water properties [e.g., Curry and Crowley, 1987; Schneider et al., 1994; Ganssen and Kroon, 2000]. The second approach compares records from different sites within the same timeframe, allowing changes in water mass distribution to be distinguished [e.g., Duplessy et al., 1988; Hendy, 1995; Nelson et al., 2000].

[39] Subantarctic waters of the Campbell Plateau, Bounty Trough, and South Chatham Rise region are characterized by planktic foraminiferal δ13C values close to 0‰ during MIS 2, with MIS 1 values of 0.3 to 0.8‰ (Figure 5). By comparison, foraminifera from Chatham Rise, located in subtropical surface waters to the north of the STF, characteristically display δ13C values of 0 to −0.5 ‰ during MISs 2 and 1, respectively [Hendy, 1995; Neil, 1997]. Subtropical surface waters of Tasman Sea and equatorial Pacific origin have variable δ13C during MIS 2, and MIS 1 values close to −1‰ [Nelson et al., 1993, 2000] (Figure 5).

image

Figure 5. Comparison of δ18O and δ13C values in 29 cores from Tasman Sea and off eastern New Zealand, during (a) MIS 1 and (b) MIS 2, revealing water mass affinities. For MIS 2 plotted positions are uncorrected for global δ18O and δ13C shifts. MIS 1 positions are shown in pale symbols. Data are from this study (large symbols) and Nelson et al. [1993], Hendy [1995], Neil [1997], and Nelson et al. [2000] (small symbols). Inset map shows core site positions.

Download figure to PowerPoint

[40] After considering the global shift in δ18O and δ13C for MIS 2, isotopic values from much of Campbell Plateau are consistent with SAW occurring over the core sites during interglacial and glacial times (Figure 5). Exceptions to this are the northern cores, Y17 and Y9 (Figure 5), which lie away from the main trend. During MIS 1, δ13C values at Y17 (∼−1‰) are consistent with waters of Tasman Sea origin, indicating the site is affected by Tasman water entrained around the base of South Island, within the Southland Current. However, during MISs 2–4, δ13C values were strongly enriched, suggesting the site was influenced by an influx of SAW (Figures 4b and 5) resulting from a postulated more northerly position of the STF west of New Zealand [cf. Martinez, 1994] and decreased Tasman water entrainment. Excursions of the δ13C record toward more depleted values during MIS 3 may result from a fluctuating influence of the STF.

[41] On Bounty Plateau, Y9 exhibits fluctuating MIS 1 values (Figure 4b) indicative of both SAW and STW signatures. The more depleted STW signature suggests the influence of sources north of the STF. A depleted δ13C profile is also strongly expressed during the later stages of MIS 5 at Y9 in common with Y16, SO136-55 and ODP1120 (Figure 4b). By contrast, during MIS 2, δ13C values at Y9 are generally consistent with a SAW source (Figure 5).

5.5. Glacial Period Water Mass History

5.5.1. Campbell Plateau Warm Pool

[42] Time series reconstructions of paleotemperatures shed new light on the behavior of the ACC and its interaction with plateau waters (Figure 6). Of note is the maintenance of a relatively warm pool of water over the plateau interior through climatic cycles, compared to the eastern margin and far west. This indicates that the SAF remained constrained along the southern and eastern margin of Campbell Plateau, in contrast to the open ocean situation where the SAF/ACC underwent an equatorward shift of 5–8° in glacial times [e.g., Howard and Prell, 1992]. Despite this constraint, grain size data and the occurrence of a pronounced temperature gradient (up to 3.0°C 100 km−1) across the plateau flanks, signify a more vigorous flow along the SAF. We therefore advocate that the SAF, as for any strong flow with a barotropic component, continued to be steered by the bathymetry during MIS 2 and did not move as a defined front over the flanks of Campbell Plateau during this time.

image

Figure 6. Temperatures (°C) over Campbell Plateau during the past 200 kyr estimated from δ18O in Gg. bulloides. Note the occurrence of warmer water over the plateau during interglacial times, and cool waters at plateau margins during glacial times. (a) Western plateau. (b) Plateau interior. (c) Eastern interior. (d) Eastern margin. Transect locations are shown on the accompanying map.

Download figure to PowerPoint

[43] The southern boundary of the ACC, corresponding to the Antarctic Polar Front, occurs well south of Campbell Plateau and was free to move equatorward as demonstrated for other sections of the Southern Ocean [Hays et al., 1976; Howard and Prell, 1992]. Thus the cold Antarctic waters of the ACC may have been compressed in a manner similar to that in the modern Drake Passage [Peterson and Whitworth, 1989]. This compression of the ACC, combined with increased windiness in glacial times [Thiede, 1979; Klinck and Smith, 1993; Hesse, 1994], probably intensified the paleocirculation along the flanks of Campbell Plateau (Figure 7).

image

Figure 7. Representation of the currents around New Zealand at (a) the present day [from Carter et al., 1998], (b) the Last Glacial [after Carter, 2001], and (c) Last Interglacial. Note the intensified paleocirculation along the flanks of Campbell Plateau, increased inflow through Pukaki Saddle, and migration of the STF during the Last Glacial (MIS 2). The Last Interglacial (MIS 5) is characterized by relaxation of the SAF circulation, reduced Pukaki inflow, and intensification of subtropical incursions.

Download figure to PowerPoint

[44] The magnitude of the last glacial-Holocene shift in oxygen isotopes increases markedly from the plateau interior to its margins (Figure 4a). Signals from the plateau margins exhibit little or no vertical gradient between Gg. bulloides and Gr. inflata, implying deep mixing of surficial waters through most of the past 130 kyr. We further suggest a widening and strengthening of the zone of deep mixing during the glacial, due to intensification of the frontal system and increased westerlies enhancing northward Ekman flow [e.g., Klinck and Smith, 1993].

[45] While Campbell Plateau margin cores show that surface waters within the ACC cooled in MIS 2, δ18O from the eastern plateau interior does not record the expected global cooling; also a weak stratification is implied by the gradient between the foraminiferal species (Figures 4c and 6). At present a relatively warm tongue of surface water appears over the central and eastern plateau [Uddstrom and Oien, 1999] possibly resulting from summer insolation and quiescent flow [Morris et al., 2001]. Isolation of plateau surface waters during MIS 2 may have occurred as a consequence of the intensification of the SAF and ACC along the eastern flank of Campbell Plateau. This is reflected as a greater temperature gradient across the path of flow. Hence, while cooler waters were entrained around the flanks of the plateau, the interior remained as a discrete body of weakly stratified Subantarctic water.

5.5.2. Western Plateau Incursions

[46] Modern observations of the circulation over the western plateau show it is a region of sluggish flow [Morris et al., 2001]. However, pronounced but localized windiness [Reid and Collen, 1983] accounts for the mixed signature of the surface waters indicated by the correspondence of the Gg. bulloides and Gr. inflata isotopic profiles at Y16. By comparison, the waters at Y17 are stratified (Figure 4). However, waters at both sites were substantially cooler during MIS 2 (Table 2), with the cooling at Y17 being the largest recorded on the plateau (∼4.5°C).

[47] Modern warm surface waters over Y17 are, at least partially, of Tasman Sea character as indicated by δ13C (Figure 5). Such water is transported within the Subtropical Front around the base of South Island [Chiswell, 1996] (Figure 1). The pronounced MIS 2 cooling is accompanied by a change in δ13C character to that of mixed Subantarctic waters suggesting a significant change in the circulation (Figure 5). The Subtropical Front, although constrained across Chatham Rise by bathymetry and along-rise current systems [Weaver et al., 1998], was free to migrate within the open Tasman Sea where it moved north by ∼2–5° during MIS 2 [Martinez, 1994; Passlow et al., 1997; Barrows et al., 2000]. Migration of the STF could result in the South Island blocking, or at least decreasing, the volume of warm Tasman Sea surface water entrained around southern New Zealand. In addition, the glacial intensification of the SAF may have allowed cooler Antarctic waters to push further north into Solander Trough [e.g., Gordon, 1975]. These factors, combined with increased windiness probably caused increased mixing and entrainment of cool SAW onto the western plateau during MIS 2 (Figures 6 and 7).

5.5.3. Pukaki Saddle Inflow

[48] Bounty Plateau exhibits a glacial-interglacial shift in oxygen isotopes similar to that on the southern and eastern flanks of Campbell Plateau. At present, an arm of the SAF passes through Pukaki Saddle to feed the strong cyclonic circulation around the head of Bounty Trough [Davis, 1998; Morris et al., 2001] (Figure 2). The proposed intensification of the SAF/ACC during glacial times presumably increased the circulation about Bounty Trough. Indeed, isotopic data from Gg. bulloides at Y9 indicate a cooling of ∼3.5°C during MIS 2, compared to MIS 1, indicating continued or enhanced supply of cool Subantarctic water through Pukaki Saddle. We argue that this provides an explanation for the strong thermal gradients of up to 6°C across the STF over Chatham Rise, previously ascribed to ACC migration [e.g., Nelson et al., 1993; Weaver et al., 1998], and also the strong cool inflow observed through Mernoo Saddle [e.g., Nelson et al., 2000; Carter et al., 2002a].

5.6. Last Interglacial Period Water Mass History

5.6.1. Ecology of the Plateau Interior

[49] Last interglacial (MIS 5) sea surface temperatures of the southwest Pacific Ocean were up to 2°C warmer than at the present day [Nelson et al., 1993; Weaver et al., 1998]. For Campbell Plateau, Gr. inflata data indicate SSTs were variably ∼2°C warmer to ∼1.5°C cooler than now during MIS 5. As for the present day, the SST distribution in MIS 5 (Figure 6) shows the SAF remained the dominant feature, with deep mixing evident immediately north and east of the front. The sluggish paleocirculation inferred for the plateau interior is similar to the modern regime. However, stronger stratification may have resulted from the elevated sea surface temperatures and the relaxation of SAF circumpolar circulation (Figure 7).

[50] While the Gr. inflata data for some sites are consistent with the generally accepted view of MIS 5 being warmer than MIS 1 [e.g., Broecker and van Donk, 1970; Shackleton and Opdyke, 1973, Imbrie et al., 1992], the isotope records of Gg. bulloides imply the MIS 5 surface ocean was at least ∼1.5°C cooler than now (Figures 4a and 6). The enriched δ18O coincides with an enriched δ13C signal in Gg. bulloides. Such increased δ13C values have been observed in upwelling regions [e.g., Ganssen and Sarnthein, 1983; Kroon and Ganssen, 1988; Schneider et al., 1994] and Gg. bulloides is often used as an indicator of upwelling. However, significant upwelling of colder waters onto Campbell Plateau is unlikely. Repeated hydrographic surveys identify weakly stratified waters during winter, with a strengthening of stratification in summer associated with the increase in insolation [Morris et al., 2001]. We infer stronger stratification during MIS 5 on account of increased insolation, reduced windiness and a less intense circulation [e.g., Nelson et al., 1994] (Figure 6).

[51] The recording of colder temperatures and accompanying enrichment of δ13C in Gg. bulloides may be related to pronounced phytoplankton blooms during MIS 5. Sediments of MIS 5 contain four times the abundance of coccoliths than in MIS 1 and at least 10-fold more than MIS 2. Coccoliths have also been recorded in higher numbers during interglacial times at DSDP 594, north of Campbell Plateau [Wells and Okada, 1997]. While recognizing that marked increases in numbers of coccolithophores may be influenced by the mode of productivity and deposition, in general, major increases occur after anomalously warm events in nutrient-rich waters under reduced windiness [Mitchell-Innes and Winter, 1987; Blackburn and Cresswell, 1993]; conditions that likely prevailed during MIS 5. Present-day waters over Campbell Plateau are not considered nutrient limited and Heath and Bradford [1980] suggest that phytoplankton production may be limited by insufficient water column stability. Increased stratification or stability of the upper water column during MIS 5 may account for the increased numbers of phytoplankton recorded.

[52] Lowered or depleted δ13C values in planktonic foraminifera are commonly associated with high nutrient values and enhanced primary productivity, but Gg. bulloides exhibits the opposite response, i.e., enriched δ13C [Hemleben et al., 1989; Ganssen and Sarnthein, 1983; Ganssen and Kroon, 2000]. Enriched δ13C in Gg. bulloides may result from increased primary production and phytoplankton uptake of 12C [e.g., Schneider et al., 1994]. Such an increase is also reflected in the sediment characteristics, which for MIS 5 are coccolith-rich, foraminiferal oozes over the plateau. The accompanying, apparently incongruous, enriched δ18O signature may be due to Gg. bulloides being displaced to a different ecological niche deeper in the water column during times of high marine algae production [Ortiz et al., 1997; Mulitza et al., 1999] (Figures 4a and 4b).

5.6.2. Incursion of Subtropical Waters

[53] In addition to enriched δ13C attributed to primary production, Gg. bulloides in Y9 on the Bounty Plateau, periodically exhibits depleted δ13C during middle to late MIS 5 and the MIS 3/4 boundary, together with a minor excursion across the MIS 1/2 boundary. DSDP 594, located south of the STF, also has depleted δ13C during MIS 5 [Nelson et al., 1993]. While it is possible that depletion may reflect reduced productivity, this option is not favored as depletion appears to be unrelated to the production of biogenic carbonate as recorded by the mass accumulation rates from Y9 and DSDP 594 [Carter et al., 2000]. It is more likely that the excursions are a source effect caused by the southward migration of δ13C-depleted Subtropical Water (Figure 5) from north of the STF along Chatham Rise [Neil, 1997]. Uddstrom and Oien [1999] show the STF to wander and shed eddies as far south as 50°S. This provides a mechanism by which Subtropical Water is introduced to Bounty Plateau and the cyclonic circulation of Bounty Trough. In addition, Subtropical Water is observed to flow south through Mernoo Saddle, separating western Chatham Rise from the South Island (Figure 1; Greig and Gilmour [1992]). These incursions, which may last for several weeks, may account for the reduced δ13C signal in MIS 5 at DSDP 594.

[54] During periods of ameliorated climate, when the intensity of the SAF is inferred to decline, there was presumably a reduction in flow of Subantarctic water through Pukaki Saddle, Mernoo Saddle and around Bounty Trough, thus setting the scene for subtropical incursions (Figure 7). Support for this comes from an increased abundance of the predominantly subtropical Neogloboquadrina dutertrei and tropical coccolithophore Gephyrocapsa recorded during MIS 5 at DSDP 594 [Nelson et al., 1993; Wells and Okada, 1997]. While reporting a predominance of subpolar species during interglacials on the southern flanks of Chatham Rise, Weaver et al. [1998] also note increased abundance of the Globigerinoides ruber, a tropical foraminiferal species.

6. Conclusions

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Environment of Campbell Plateau
  5. 3. Methodology
  6. 4. Results
  7. 5. Discussion
  8. 6. Conclusions
  9. Acknowledgments
  10. References

[55] Over the past 130 kyr, the surface waters over the interior of Campbell Plateau were thermally isolated by the powerful SAF of the ACC which remained constrained and steered by the steep plateau flanks. Surficial waters adjacent to the SAF were deeply mixed over glacial/interglacial cycles with more distant plateau waters stratified on seasonal timescales. However, the lateral extent of mixing varied, and deep mixing increased across Campbell Plateau in glacials, reflecting a more intense SAF under stronger winds and further constriction of the ACC with northward migration of the Polar Front. This resulted in an overall intensification of the paleocirculation around the flanks of Campbell Plateau and consequent isolation of the surficial waters of the central plateau.

[56] Intensification of the SAF during MIS 2 probably reinforced the cyclonic circulation around Bounty Trough via Pukaki Saddle. This influx of cold water increased thermal gradients at the STF as well as feeding cool water flows along the eastern North Island.

[57] Modern waters over the western margin of Campbell Plateau partly originate from the Tasman Sea and are transported to the plateau via the Subtropical Front. Northward migration of the Subtropical Front during glacial times decreased the entrainment of warm Tasman waters, probably in consort with a stronger SAF, resulting in an increased contribution of cold Subantarctic surface waters.

[58] The ameliorated climate of the last interglacial caused warm SSTs, a less vigorous SAF, and altered planktic ecology marked by a change from foraminiferal oozes to mixed nannoforaminiferal oozes. A weakened SAF in MIS 5 encouraged incursions of subtropical-type water via eddies shed from the STF.

Acknowledgments

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Environment of Campbell Plateau
  5. 3. Methodology
  6. 4. Results
  7. 5. Discussion
  8. 6. Conclusions
  9. Acknowledgments
  10. References

[59] The authors thank Lisa Northcote and Rachael Hardwick for assistance in sample processing. The crew and officers of NIWA's Tangaroa and GEOMAR's Sonne are thanked for their professionalism. Stefan Nees and Joern Thiede (GEOMAR, Germany) provided a scientific berth on SO136. Erika MacKay provided assistance with figure generation. Comments from Will Howard and an anonymous reviewer, Barbara Manighetti and Phil Sutton (NIWA) and Penny Cooke (University of Waikato) improved the content of previous manuscripts. This project was run under contract CO1X0037 to the Foundation of Research, Science and Technology.

References

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Environment of Campbell Plateau
  5. 3. Methodology
  6. 4. Results
  7. 5. Discussion
  8. 6. Conclusions
  9. Acknowledgments
  10. References
  • Barrows, T. T., S. Juggins, P. de Deckker, J. Thiede, and J. I. Martinez (2000), Paleoceanography of the southwest Pacific Ocean during the Last Glacial Maximum, Paleoceanography, 15, 95109.
  • Berger, W. H. (1971), Sedimentation of planktonic foraminifera, Mar. Geol., 11, 325358.
  • Blackburn, S. I., and G. Cresswell (1993), A coccolithophorid bloom in Jervis Bay, Australia, Aust. J. Mar. Freshwater Res., 44, 253260.
  • Broecker, W. S. (1989), The salinity contrast between the Atlantic and Pacific Oceans during glacial time, Paleoceanography, 4, 207212.
  • Broecker, W. S., and J. van Donk (1970), Insolation changes, ice volumes and the δ18O record in deep sea cores, Rev. Geophys., 8, 169197.
  • Bryden, H. L., and R. A. Heath (1985), Energetic eddies at the northern edge of the Antarctic Circumpolar Current in the southwest Pacific, Prog. Oceanogr., 14, 6587.
  • Burling, R. W. (1961), Hydrology of circumpolar waters south of New Zealand, N. Z. Oceanogr. Inst. Mem., 10, 66 pp., Wellington, New Zealand.
  • Carter, L. (1989), New occurrences of manganese nodules in the south-western Pacific Basin, east of New Zealand, N. Z. J. Mar. Freshwater Res., 23, 247253.
  • Carter, L. (2001), Currents of change: The ocean flow in a changing world, Water Atmos., 9, 1517.
  • Carter, L., and R. M. Carter (1993), Sedimentary evolution of the Bounty Trough: A Cretaceous rift basin, southwest Pacific Ocean, in South Pacific Sedimentary Basins, edited by P. F. Ballance, pp. 5167, Elsevier Sci., New York.
  • Carter, L., and I. N. McCave (1994), Development of sediment drifts approaching an active plate margin under the SW Pacific Deep Western Boundary Current, Paleoceanography, 9, 10611085.
  • Carter, L., and I. N. McCave (1997), The sedimentary regime beneath the Deep Western Boundary Current inflow to the southwest Pacific Ocean, J. Sediment. Res., 67, 10051017.
  • Carter, L., and J. Wilkin (1999), Abyssal circulation around New Zealand—A comparison between observations and a global circulation model, Mar. Geol., 159, 221239.
  • Carter, L., R. M. Carter, I. N. McCave, and J. Gamble (1996), Regional sediment recycling in the abyssal southwest Pacific Ocean, Geology, 24, 735738.
  • Carter, L., R. G. Garlick, P. Sutton, S. M. Chiswell, N. A. Oien, and B. R. Stanton (1998), Ocean circulation around New Zealand, NIWA Chart Misc. Ser. 76, Natl. Inst. of Land and Atmos. Res., Auckland.
  • Carter, L., H. L. Neil, and I. N. McCave (2000), Glacial to interglacial changes in non-carbonate and carbonate accumulation in the SW Pacific Ocean, New Zealand, Palaeogeogr. Palaeoclimatol. Palaeoecol., 162, 333365.
  • Carter, L., B. Manighetti, M. Elliot, N. Trustrum, and B. Gomez (2002a), Source, sea level and circulation effects on the sediment flux to the deep ocean over the past 15 ka off eastern New Zealand, Global Planet. Change, 33, 339355.
  • Carter, L., H. L. Neil, and L. Northcote (2002b), Late Quaternary incursions of icebergs into the southwest Pacific Ocean off eastern New Zealand, Mar. Geol., 191, 1935.
  • Carter, R. M., et al. (1999), Proceedings of the Ocean Drilling Program, Initial Reports, vol. 181, Ocean Drill. Program, College Station, Tex.
  • Chiswell, S. M. (1994), Acoustic Doppler current profiler measurements over the Chatham Rise, N. Z. J. Mar. Freshwater Res., 28, 167178.
  • Chiswell, S. M. (1996), Variability in the Southland Current, New Zealand, N. Z. J. Mar. Freshwater Res., 30, 117.
  • Craig, H., and L. I. Gordon (1965), Deuterium and oxygen-18 variations in the ocean and the marine atmosphere, in Stable Isotopes in Oceanographic Studies and Paleotemperatures, edited by E. Tongiorgi, pp. 9130, Cons. Naz. delle Rich., Lab. di Geol. Nucl., Pisa.
  • Curry, W. B., and T. J. Crowley (1987), The δ13C of equatorial Atlantic surface waters: Implications for ice age pCO2 levels, Paleoceanography, 2, 489517.
  • Curry, W. B., J. C. Duplessy, L. D. Labeyrie, and N. J. Shackleton (1988), Changes in the distribution of δ13C of deep water ΔCO2 between the last glaciation and the Holocene, Paleoceanography, 3, 317341.
  • Daly, K. L., W. O. Smith Jr., G. C. Johnson, G. R. DiTullio, D. R. Jones, C. W. Mordy, R. A. Freely, D. A. Hansell, and J.-Z. Zhang (2001), Hydrography, nutrients, and carbon pools in the Pacific sector of the Southern Ocean: Implications for carbon flux, J. Geophys. Res., 106, 71077124.
  • Davis, R. E. (1998), Preliminary results from directly measuring middepth circulation in the tropical and South Pacific, J. Geophys. Res., 103, 24,61924,639.
  • Duplessy, J. C., N. J. Shackleton, R. G. Fairbanks, L. Labeyrie, D. Oppo, and N. Kallel (1988), Deepwater source variations during the last climatic cycle and their impact on the global deepwater circulation, Paleoceanography, 3, 343360.
  • Epstein, S., R. Buchsbaum, H. A. Lowenstein, and H. C. Urey (1953), Revised carbonate-water isotopic temperature scale, Geol. Soc. Am. Bull., 64, 13151325.
  • Ganssen, G., and D. Kroon (2000), The isotopic signature of planktonic foraminfera from NE Atlantic surface sediments: Implications for the reconstruction of past oceanic conditions, J. Geol. Soc. London, 157, 693699.
  • Ganssen, G., and M. Sarnthein (1983), Stable isotope composition of foraminifers: The surface and bottom water record of coastal upwelling, in Responses of the Sedimentary Regime to Present Coastal Upwelling: Coastal Upwelling, Its Sediment Record, Part A, edited by E. Suess, and J. Theide, pp. 99121, Plenum, New York.
  • Glasby, G. P. (1976), Surface densities of manganese nodules in the southern sector of the South Pacific, N. Z. J. Geol. Geophys., 19, 771790.
  • Gordon, A. L. (1972), On the interaction of the Antarctic Circumpolar Current and the Macquarie Ridge, in Antarctic Oceanology II—The Australian-New Zealand sector, Antarct. Res. Ser., vol. 19, edited by D. E. Hayes, pp. 7178, AGU, Washington, D. C.
  • Gordon, A. L. (1975), An Antarctic oceanographic section along 170°E, Deep Sea Res., 22, 357377.
  • Greig, M. J., and A. E. Gilmour (1992), Flow through the Mernoo Saddle, New Zealand, N. Z. J. Mar. Freshwater Res., 26, 155165.
  • Hays, J. D., N. J. Lozano, N. J. Shackleton, and G. Irving (1976), Reconstruction of the Atlantic and western Indian Ocean sectors of the 18,000 Antarctic Ocean, B. P. in Investigation of Southern Ocean Paleoceanography and Paleoclimatology, Geol. Soc. Am. Mem., vol. 145, edited by R. M. Cline, and J. D. Hays, pp. 337374, Geol. Soc. of Am., Boulder, Colo.
  • Hayward, B. W. (1983), Planktic foraminifera (protozoa) in New Zealand waters: A taxonomic review, N. Z. J. Zool., 10, 6374.
  • Heath, R. A. (1981), Oceanic fronts around southern New Zealand, Deep Sea Res., 28, 547560.
  • Heath, R. A. (1985), A review of physical oceanography of the seas around New Zealand—1982, N. Z. J. Mar. Freshwater Res., 19, 79124.
  • Heath, R. A., and J. M. Bradford (1980), Factors affecting phytoplankton production over the Campbell Plateau, New Zealand, J. Plankton Res., 2, 169181.
  • Hecht, A. D. (1985), Paleoclimate Analysis and Modelling, 445 pp., John Wiley, Hoboken, N. J.
  • Hemleben, C., M. Spindler, and O. R. Anderson (1989), Modern Planktonic Foraminifera, 363 pp., Springer-Verlag, New York.
  • Hendy, I. (1995), Paleoceanography of the Glacial-Holocene transition in the waters surrounding New Zealand, M.S. thesis, Univ. of Waikato, Hamilton, New Zealand.
  • Hesse, P. P. (1994), The record of continental dust from Australia in Tasman Sea sediments, Quat. Sci. Rev., 13, 257272.
  • Howard, W. R., and W. L. Prell (1992), Late Quaternary surface circulation of the southern Indian ocean and its relationship to orbital variations, Paleoceanography, 7, 79117.
  • Imbrie, J., J. D. Hays, D. G. Martinson, A. McIntrye, A. C. Mix, J. J. Morley, N. G. Pisias, W. L. Prell, and N. J. Shackleton (1984), The orbital theory of Pleistocene climate: Support from a revised chronology of the marine δ18O record, in Milankovitch and Climate, Part I, edited by A. L. Berger et al., pp. 269305, D. Reidel, Norwell, Mass.
  • Imbrie, J., et al. (1992), On the structure and origin of major glaciation cycles: 1. Linear responses to Milankovitch forcing, Paleoceanography, 7, 701783.
  • Jones, G. A., and P. Kaiteris (1983), A vacuum-gasometric technique for rapid and precise analysis of calcium carbonate in sediments and soils, J. Sediment. Petrol., 53, 655660.
  • King, A. L., and W. R. Howard (2001), Seasonality of foraminiferal flux in sediment traps at Chatham Rise, SW Pacific: Implications for paleotemperature estimates, Deep Sea Res. Part I, 48, 16871708.
  • Klink, J. M., and D. A. Smith (1993), Effect of wind changes during the last glacial maximum on the circulation in the Southern Ocean, Paleoceanography, 8, 427433.
  • Kroon, D., and G. Ganssen (1988), Northern Indian Ocean upwelling cells and the stable isotope composition of living planktic foraminifers, in Planktonic Foraminifers As Tracers of Ocean-Climate History, edited by G. J. A. Brummer, and D. Kroon, pp. 299319, Free Univ. Press, Amsterdam.
  • Lea, D. W., J. Bijma, H. J. Spero, and D. Archer (1999), Implications of a carbonate ion effect on shell carbon and oxygen isotopes for glacial ocean conditions, in Use of Proxies in Paleoceanography, Examples From the South Atlantic, edited by G. Fischer, and G. Wefer, pp. 513522, Springer-Verlag, New York.
  • Lean, C. M. B., and I. N. McCave (1998), Glacial to interglacial mineral magnetic and paleoceanographic changes at Chatham Rise, SW Pacific Ocean, Earth Planet. Sci. Lett., 163, 247260.
  • Markgraf, V., J. R. Dodson, A. P. Kershaw, M. S. McGlone, and N. Nicholls (1992), Evolution of late Pleistocene and Holocene climates in the circum-South Pacific land areas, Clim. Dyn., 6, 193211.
  • Martinez, J. I. (1994), Late Pleistocene palaeoceanography of the Tasman Sea: Implications for the dynamics of the warm pool in the western pacific, Palaeogeogr. Palaeoclimatol. Palaeoecol., 112, 1962.
  • Martinez, J. I., P. de Deckker, and A. R. Chivas (1997), New estimates for salinity changes in the Western Pacific Warm Pool during the last glacial maximum: Oxygen isotope evidence, Mar. Micropaleontol., 32, 311340.
  • Martinson, D. G., N. G. Pisias, J. D. Hays, J. Imbrie, T. C. Moore Jr., and N. J. Shackleton (1987), Age dating and the orbital theory of the Ice Ages: Development of a high resolution 0 to 300,000 year chronostratigraphy, Quat. Res., 27, 127.
  • McGlone, M. S., M. J. Salinger, and N. T. Moar (1994), Paleovegetation studies of New Zealand's climate since the Last Glacial Maximum, in Global Climates Since the Last Glacial Maximum, edited by H. E. Wright Jr. et al., pp. 294317, Uni. of Minn. Press, Minneapolis.
  • Mitchell-Innes, B. A., and A. Winter (1987), Coccolithophores: A major phytoplankton component in mature upwelled waters off the Cape Peninsula, South Africa in March, 1983, Mar. Biol., 95, 2530.
  • Morris, M., B. Stanton, and H. L. Neil (2001), Subantarctic oceanography around New Zealand: Preliminary results from an ongoing survey, N. Z. J. Mar. Freshwater Res., 35, 499519.
  • Mulitza, S. H., A. Dürkoop, W. Hale, G. Wefer, and H. S. Niebler (1997), Planktonic foraminfera as recorders of past surface-water stratification, Geology, 25, 335338.
  • Mulitza, S. H., H. Arz, S. Kemle-von Mucke, and C. Moos (1999), The South Atlantic carbon isotope record of planktic foraminifera, in Use of Proxies in Paleoceanography, Examples From the South Atlantic, edited by G. Fischer, and G. Wefer, pp. 427445, Springer-Verlag, New York.
  • Nees, S., et al. (1999), Quaternary variability of water masses in the southern Tasman Sea and Southern Ocean (SW Pacific Sector), F. S. Sonne Cruise Report SO136, Wellington-Hobart, Oct 16–Nov 12, 1998, Geomar Rep., 89, 78 pp., Geomar Res. Cent. for Mar. Geosci., Christian Albrechts Univ., Kiel.
  • Neil, H. L. (1997), Last Glaciation to present paleoceanographic changes, Subtropical Convergence Zone, Chatham Rise, southwest Pacific Ocean, Ph.D. thesis, Univ. of Waikato, Hamilton, New Zealand.
  • Nelson, C. S., P. J. Cooke, C. H. Hendy, and A. M. Cuthbertson (1993), Oceanographic and climatic changes over the past 160,000 years at deep sea drilling project site 594 off southeastern New Zealand, southwest Pacific Ocean, Paleoceanography, 8, 435458.
  • Nelson, C. S., C. H. Hendy, and A. M. Cuthbertson (1994), Oxygen isotope evidence for climatic contrasts between Tasman Sea and southwest Pacific Ocean during the late Quaternary, in Evolution of the Tasman Sea Basin, edited by G. J. van der Lingen et al., pp. 181196, A. A Balkema, Brookfield, Vt.
  • Nelson, C. S., I. L. Hendy, H. L. Neil, C. H. Hendy, and P. P. E. Weaver (2000), Last glacial jetting through the Subtropical Convergence Zone in the southwest Pacific off eastern New Zealand, and some geological implications, Palaeogeogr. Palaeoclimatol. Palaeoecol., 156, 103121.
  • Newnham, R. M., D. J. Lowe, and P. W. Williams (1999), Quaternary environment change in New Zealand: A review, Prog. Phys. Geogr., 23, 567610.
  • New Zealand Meteorological Service (1981), Summaries of climatological observations to 1980, N. Z. Meterol. Serv. Misc. Publ. 177, 172 pp., Wellington, New Zealand.
  • Niebler, H.-S., H.-W. Hubberton, and R. Gersonde (1999), Oxygen isotope values of planktonic foraminifera: A tool for reconstruction of surface water stratification, in Use of Proxies in Paleoceanography, Examples From the South Atlantic, edited by G. Fischer, and G. Wefer, pp. 165189, Springer-Verlag, New York.
  • Oppo, D. W., and R. G. Fairbanks (1989), Carbon isotope composition of tropical surface water during the past 22,000 years, Paleoceanography, 4, 333351.
  • Orsi, A. H., T. Whitworth III, and W. D. Nowlin Jr. (1995), On the meridional extent and fronts of the Antarctic Circumpolar Current, Deep Sea Res. Part I, 42, 641673.
  • Ortiz, J., A. Mix, S. Hosteler, and M. Kashgarian (1997), The California Current of the last glacial maximum: Reconstruction at 42°N based on multiple proxies, Paleoceanography, 12, 191205.
  • Passlow, V., W. Pinxian, and A. R. Chivas (1997), Late Quaternary paleoceanography near Tasmania, southern Australia, Palaeogeogr. Palaeoclimatol. Palaeoecol., 131, 433463.
  • Paterson, S. L., and T. Whitworth III (1990), Physical oceanography of the Pacific sector of the Southern Ocean, in Antarctic Sector of the Pacific, Elsevier Oceanogr. Ser., vol. 51, edited by G. P. Glasby, pp. 5593, Elsevier Sci., New York.
  • Peterson, R. G., and T. Whitworth III (1989), The subantarctic and polar fronts in relation to deep water masses through the southwestern Atlantic, J. Geophys. Res., 94, 10,81710,838.
  • Prell, W. L., J. Imbrie, D. G. Martinson, J. J. Morley, N. G. Pisias, N. J. Shackleton, and H. F. Streeter (1986), Graphic correlation of oxygen isotopes stratigraphy application to the late Quaternary, Paleoceanography, 1, 137162.
  • Ravelo, A. C., and R. G. Fairbanks (1995), Carbon isotopic fractionation in multiple species planktonic foraminifera from core-tops in the tropical Atlantic, J. Foraminiferal Res., 25, 5374.
  • Reid, S. J., and B. Collen (1983), Analyses of Wave and wind reports from ships in the Tasman Sea and New Zealand areas, N. Z. Meteorol. Serv. Misc. Publ., 174, 72 pp., Wellington, New Zealand.
  • Reid, S. J., and A. C. Penney (1982), Upper-level wind frequencies and mean speeds fro new Zealand and pacific island stations, N. Z. Meteorol. Serv. Misc. Publ. 182, 98 pp., Wellington, New Zealand.
  • Rohling, E. J., and G. R. Bigg (1998), Paleosalinity and δ18O: A critical assessment, J. Geophys. Res., 103, 13071318.
  • Rostek, F., G. Ruhland, F. C. Bassinot, P. J. Muller, L. D. Laberyie, Y. Lancelot, and E. Bard (1993), Reconstructing sea surface temperature and salinity using δ18O and alkenone records, Nature, 364, 312319.
  • Schneider, R. R., P. J. Muller, and G. Wefer (1994), Late Quaternary paleoproductivity changes off Congo deduced from stable carbon isotopes of planktonic foraminfera, Palaeogeogr. Palaeoclimatol. Palaeoecol., 110, 255274.
  • Schrag, D. P., J. F. Adkins, K. McIntyre, J. L. Alexander, D. A. Hodell, C. D. Charles, and J. F. McManus (2002), The oxygen isotope composition of seawater during the Last Glacial Maximum, Quat. Sci. Rev., 21, 331342.
  • Shackleton, N. J. (1977), Carbon-13 in Uvigerina: Tropical rainforest history and the equatorial Pacific carbonate dissolution cycles, in The Fate of Fossil Fuel CO2in Oceans, edited by N. R. Anderson, and A. Malahoff, pp. 401427, Plenum, New York.
  • Shackleton, N. J., and N. D. Opdyke (1973), Oxygen isotope and paleomagnetic stratigraphy of equatorial pacific core V28–238: Oxygen isotope temperatures and ice volumes on a 105 year and 106 year scale, Quat. Res., 3, 3955.
  • Shackleton, N. J., and N. D. Opdyke (1976), Oxygen isotope and paleomagnetic stratigraphy of equatorial pacific core V28-239: Late Pliocene to latest Pleistocene, Geol. Soc. Am. Mem., 145, 449464.
  • Sikes, E. L., C. R. Samson, T. P. Guilderson, and W. R. Howard (2000), Old radiocarbon ages in the southwest Pacific Ocean during the last glacial period and deglaciation, Nature, 405, 555559.
  • Stanton, B. R., and M. Y. Morris (2004), Direct velocity measurements in the Subantarctic Front and over Campbell Plateau, southeast of New Zealand, J. Geophys., 109, C01028, doi:10.1029/2002JC001339.
  • Stewart, R. B., and V. E. Neall (1984), Chronology of palaeoclimatic change at the end of the last glaciation, Nature, 311, 4748.
  • Stuiver, M., and P. J. Reimer (1993), Extended 14C database and revised CALIB radiocarbon calibration program, Radiocarbon, 28, 10221030.
  • Stuiver, M., P. J. Reimer, E. Bard, J. W. Beck, G. S. Burr, K. A. Hughen, B. Kromer, F. G. McCormac, J. V. D. Plicht, and M. Spurk (1998), INTCAL98 Radiocarbon age calibration 24,000–0 cal BP. Radiocarbon, 40, 10411083.
  • Summerhayes, C. P. (1969), Marine geology of the New Zealand subantarctic sea floor, N. Z. Oceanogr. Inst. Mem. 50, 92 pp.
  • Sutton, P. J. H. (2003), The Southland Current: A subantarctic current, N. Z. J. Mar. Freshwater Res., 37, 645652.
  • Thiede, J. (1979), Wind regimes over the late Quaternary southwest Pacific Ocean, Geology, 7, 259262.
  • Uddstrom, M. J., and N. A. Oien (1999), On the use of high-resolution satellite data to describe the spatial and temporal variability of sea surface temperatures in the New Zealand region, J. Geophys. Res., 104, 20,72920,751.
  • Weaver, P. P. E., H. L. Neil, and L. Carter (1997), Sea surface temperature estimates from the SW Pacific based on planktonic foraminifera and oxygen isotopes, Palaeogeogr. Palaeoclimatol. Palaeoecol., 131, 7080.
  • Weaver, P. P. E., L. Carter, and H. L. Neil (1998), Response of surface water masses and circulation to late Quaternary climate change, east of New Zealand, Paleoceanography, 13, 7083.
  • Wells, P., and H. Okada (1997), Response of nannoplankton to major changes in sea surface temperature and movements of hydrological fronts over Site DSDP 594 (south Chatham Rise, southeastern New Zealand), during the last 130 kyr, Mar. Micropaleontol., 32, 341363.
  • Zeebe, R. E., J. Bijma, and D. A. Wolf-Gladrow (1999), A diffusion-reaction model of carbon isotope fractionation in foraminifera, Mar. Chem., 64, 199227.