glacial/interglacial control of terrigenous and biogenic fluxes in the deep ocean off a high input,...

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Marine Geology 226

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

Lionel Carter *, Barbara Manighetti

National Institute of Water and Atmosphere, Private Bag 14 901, Kilbirnie, Wellington, New Zealand

Received 23 May 2005; received in revised form 19 October 2005; accepted 11 November 2005

Abstract

Sediment input to the deep ocean, off the NewZealand collisional plate boundary, is large and variable as revealed by a 139 kyr-old

record from the 34.92 m-long core, MD97 2121. Mass accumulation rates (MARs) were derived for terrigenous and biogenic

(carbonate and silica) sedimentary components, with temporal control from a stable isotope age model verified by tephra and

radiocarbon ages. Terrigenous MARs changed in phase with glacio-eustatic fluctuations of sea level. Highest rates (z30 g/cm2/kyr)

coincided with the late regressive-lowstand-early transgressive phase of marine isotope stage (MIS) 2 when rivers discharged at or

near the shelf edge. Lesser, but still high terrigenous rates (20–30 g/cm2/kyr) characterized peak warm phases of early MIS 5 and 1,

when a strengthened Subtropical Inflow probably introduced sediment from fluvial and seabed sources to the north of MD97 2121.

However, during prolonged highstands the flux declined as more sediment was retained on a tectonically subsiding shelf with a shore-

parallel current. Paradoxically, the lowest MARs (10–5 g/cm2/kyr) occurred in MIS 4; either sea level was too high to allow much

sediment to escape off-shelf, or the fluvial input was modest, or both these factors.

Despite the terrigenous dominance, biogenic MARs are high, and are 2–4 times larger than SW Pacific Ocean rates. MARs

were lowest during MIS 4 and 3, possibly due to reduced marine production under a lowered input of fluvial micronutrients. In

contrast, MARs increased irregularly though glacial maxima to peak at 5–6 g/cm2/kyr carbonate and 2–2.5 g/cm2/kyr silica in

early-mid MIS 5 and 1. Such rates reflected the interaction of macronutrient-rich subantarctic waters, sourced from the south, with

local, micronutrient-rich subtropical waters. Production also responded to warmer temperatures, elevated nutrient runoff and a more

stratified surface ocean. Later in MIS 5 and 1, fluxes reduced as less productive subtropical waters prevailed. That biogenic and

aeolian MAR profiles are out of phase, suggests Fe-fertilization by aerosolic dust was not a major influence.

D 2005 Elsevier B.V. All rights reserved.

Keywords: mass accumulation rates; terrigenous; biogenic; glacial/interglacial; New Zealand

1. Introduction

In a landmark paper, Milliman and Syvitski (1992)

drew attention to the significance of mountainous rivers

on oceanic islands as major suppliers of sediment to the

0025-3227/$ - see front matter D 2005 Elsevier B.V. All rights reserved.

doi:10.1016/j.margeo.2005.11.004

* Corresponding author. Fax: +64 4 3862 153.

E-mail address: [email protected] (L. Carter).

ocean. In the western Pacific, islands from the Philip-

pines to New Zealand currently contribute more than a

quarter of the estimated 20�109 t of sediment annually

entering the world’s ocean. Papua New Guinea, for

example, provides 1.9�109 t/yr (Milliman, 1995),

whereas New Zealand delivers a more modest, but

nonetheless substantial 0.21�109 t/yr (Hicks and Shan-

kar, 2003).

(2006) 307–322

L. Carter, B. Manighetti / Marine Geology 226 (2006) 307–322308

A common thread linking these islands is their lo-

cation at active margins where pronounced tectonism,

coupled with erosion-prone source rocks and high rain-

fall, ensure a copious sediment supply to the coast. In

the case of New Zealand, its position astride the bound-

ary between the colliding Pacific and Australian plates

results in the rapid uplift of mountain ranges composed

of easily eroded rocks. Sediment supply is favoured

further by frequent earthquakes, numerous volcanic

eruptions, and a vigorous climate (e.g., Berryman et

al., 1989; Froggatt and Lowe, 1990; Hicks et al., 2000;

Lewis and Pettinga, 1993). Finally, there are the effects

of human activities, which have increased modern river

discharges several times above pre-human rates (Milli-

man and Syvitski, 1992). On the eastern North Island

shelf, for example, sedimentation is estimated to have

increased fourfold following Polynesian and European

colonization (Gomez et al., 2004).

In general, the deep ocean off high input margins

receives much terrigenous sediment, the delivery of

which is regulated by the bathymetric and sedimentary

framework of the intervening continental margin. This

framework has many forms reflecting the wide range of

tectonic, hydraulic, climatic and geologic processes that

shape the onshore and offshore realms (e.g., Walsh and

Nittrouer, 2003; Orpin et al., 2002). Additionally, the

effects of glacio-eustatic fluctuations of sea level, while

global in extent, have a marked but regionally variable

influence (e.g., Carter et al., 1986). Despite the terrige-

nous influx, the New Zealand margin still has a signifi-

cant biogenic component, e.g., Stewart and Neall (1984).

This paper is based on a giant piston core, MD97

2121, acquired as part of the International Marine

Global Change Study (IMAGES). The core site (40822.93VS, 1778 59.68VE; 2314 m water depth) is at the

southern and down-current end of the continental mar-

gin, between East Cape and Hawke Bay (Fig. 1). With

an average fluvial discharge of 80�106 t/yr of sus-

pended load, this is a prime example of a high terrig-

enous input, oceanic island in a temperate climate (see

Milliman and Syvitski, 1992). Site MD97 2121 was

chosen to reflect this input. Because the core record

encompasses a full glacial/interglacial cycle back to

early MIS 6, it allows an evaluation of eustatic, palaeo-

climatic and palaeoceanographic controls on the terrig-

enous and biogenic (carbonate, silica) fluxes to the deep

ocean. The biogenic component is intriguing because

although b20% of the total flux, it is still high by local

standards (Carter et al., 2000). To identify any plate

boundary-related influences on the biogenic flux, com-

parisons are made with other flux records from cores

located near to and far from the plate boundary.

1.1. Regional sedimentary framework

The east coast margin sits within the Hikurangi

subduction system (Fig. 1). There, a partly prograded

and partly off-scraped sediment thrust wedge is de-

formed where the southwest moving Pacific plate des-

cends below the Australian plate (Lewis and Pettinga,

1993). The resultant imbricate thrust wedge forms the

coastal hills and submergent continental margin, sea-

ward of a frontal ridge represented by the actively rising

mountain ranges of the North Island.

Fluvial input to the east coast is dominated by

the Waiapu and Waipaoa rivers, which discharge

35�106 t/yr and 15�106 t/yr of suspended load, re-

spectively (Hicks and Shankar, 2003). Although they are

New Zealand’s largest contributors of sediment, the

rivers have modest catchments (V2200 km2). This situ-

ation reflects rapid uplift of up to 3 mm/yr in nearby

ranges (Reyners and McGinty, 1999), which exposes

soft and sometimes sheared sedimentary rocks (Trustrum

et al., 1999; De Rose et al., 1998). Numerous earth-

quakes, (Berryman et al., 1989; Davey et al., 1997),

frequent eruptions from the nearby Central Volcanic

Region (Froggatt and Lowe, 1990), and a vigorous

climate that is presently dominated by the ENSO-mod-

ulated, Roaring Forties westerly weather system, all

enhance the sediment supply (Gomez et al., 2004).

The Waipaoa and Waiapu, together with other high

discharge rivers, empty onto a 14–37 km-wide conti-

nental shelf (16 km wide near Site MD97 2121; Fig. 1).

Part of the fluvial load is trapped in subsiding basins on

the inner to middle shelf. Off the Waipaoa River, for

example, the middle shelf is subsiding at 1–2 mm/yr.

Subsidence has encouraged deposition as attested by

substantial thicknesses of last post-glacial sediment on

the North Island eastern margin, e.g., 110 m on the East

Cape shelf (Lewis et al., 2004) and N35 m on the

Poverty Bay shelf (Foster and Carter, 1997). Locally,

shelf capture is encouraged by growing ridges that

deform the outer shelf (Lewis, 1980; Barnes et al.,

2002). Sediment that escapes to the continental slope

via hyperpycnal flows, turbidity currents or hemipela-

gic processes, accumulates within slope basins formed

by imbricate thrusting of the frontal wedge. These

sediments are prone to mass wasting that encompasses

small slumps of b1 km3 to giant debris avalanches and

flows of up to 3750 km3 (Barnes and Lewis, 1991;

Collot et al., 2001).

MD97 2121 is from such a slope basin. Yet despite

widespread redeposition along the margin, the site

appears to be undisturbed as shown by; (i) a coherent

isotopic stratigraphy (see Results), (ii) an intact hemi-

Fig. 1. Locality of MD97 2121 with the main elements of the surface circulation and water masses: EAC = East Australian Current, ECC = East

Cape Current, WE = Wairarapa Eddy, WCC = Wairarapa Coastal Current, DC = D’Urville Current, STF = Subtropical Front, SC = Southland

Current, STW = Subtropical Water, SAW = Subantarctic Water. Wp = Waiapu River, Wa = Waipaoa River, HB = Hawke Bay, CVR = Central

Volcanic Region, CS = Cook Strait, MS = Mernoo Saddle. Inset outlines the plate boundary, accretionary thrust wedge (hatched) and core locations.

L. Carter, B. Manighetti / Marine Geology 226 (2006) 307–322 309

pelagic drape outlined in 3.5 kHz profiles (Fig. 2), (iii)

an absence of proximal turbidity current channels and

landslide scars on the multibeam chart of Lewis et al.

(1999), and (iv) a lack of mass transport deposits in the

core itself. This analysis is supported by data from the

nearby piston core P69 (Stewart and Neall, 1984).

1.2. Oceanography

The surface oceanography has three elements.

1. The modern continental shelf is bathed by the north-

ward flowing Wairarapa Coastal Current (WCC)

(Brodie, 1960; Chiswell, 2000). This flow transports

cool, low salinity water comprised of (i) Subantarctic

Surface Water (SAW) and upwelled Subantarctic

Mode Water transported from the south by the

Southland Current (Heath, 1975; Sutton, 2003),

and (ii) surface water transported by the D’Urville

Current from Cook Strait (Fig. 1). The WCC extends

at least to northern Hawke Bay and possibly further

north when forced by winds. The volume transport is

~1.6�106 m3/s, but diminishes northwards.

2. The East Cape Current (ECC) is the local component

of the Subtropical Inflow to New Zealand (Fig. 1).

Located seaward of the WCC, it carries warm, high

salinity Subtropical Surface Water (STW) along the

outer continental shelf and slope to Chatham Rise,

Fig. 2. 3.5 kHz profile from MD97 2121 highlighting the hemipelagic character of the substrate. Outline of core stratigraphy for Marine Isotope

Stages 1–6 is also presented.


before veering east along the Rise’s northern flank.

The current is up to 110 km wide and transports 10–

20�106 m3/s (Chiswell and Roemmich, 1998).

3. Near the latitude of MD97 2121, part of the ECC

recirculates to help form the Wairarapa Eddy (Roem-

mich and Sutton, 1998). This 200 km diameter,

anticyclonic gyre is centred on 418 S, 1788 30V E

(Fig. 1). Previously regarded as a semi-permanent

gyre, a new analysis by Chiswell (2005) reveals a

more dynamic situation. Eddies pass southwest

along the continental margin at a rate of 2 to 3 per

year. During their passage, eddies may stall or merge

with older features to alter eddy size and strength.

Five years of satellite-based, sea surface temperature

data (Uddstrom and Oien, 1999) outline a marked

variability in the surface circulation. The WCC com-

monly reaches 408–418 S, but periodically extends to

388 S. Sometimes this flow is accompanied by an

offshore displacement of the ECC and adjacent Wair-

arapa Eddy. Whether such displacements result from an

increased coastal current or a weakened ECC is uncer-

tain. At other times, expansion of the Wairarapa Eddy

brings warm, productive waters to MD97 2121 (Roem-

mich and Sutton, 1998).

The deep oceanography is less well known. A geo-

strophic section along 388 S, shows that the ECC

extends to N1000 m depth (Stanton et al., 1997).

Even at it deepest, the current is still substantial with

a velocity of ~10 cm/s to the south. Currents are

anticipated to diminish to the regional level of no

motion, which has been arbitrarily taken at ~2000 m

depth (e.g., Warren, 1973), but may locally extend

deeper (P. Sutton, NIWA, pers. comm.). Below

~2000 m, the circulation may be affected by shallow

reaches of the Southwest Pacific deep western boundary

current (DWBC) (McCave and Carter, 1997). Although

the DWBC velocity has not been measured at the Hikur-

angi accretionary prism, geostrophic sections nearby

suggest velocities of 1–3 cm/s to the north (McCave

and Carter, 1997).

The relative roles of the shallow and deep circulation

on suspended sediment transport may be gauged by

recent hydrographic and nephelometric transects off

Poverty Bay (Carter, 2005). Dense nepheloid plumes

spread over the shelf and the upper slope; the deeper

plumes forming beneath the fast flowing core of the

ECC. In contrast, deep-ocean nepheloid layers are ei-

ther weak on absent at the depth of MD97 2121, and

only become significant at ~3000 m where the DWBC

generates a widespread benthic plume (McCave and

Carter, 1997).

2. Methods and age model

Core MD97 2121 was taken with a 45 m-long

Calypso corer deployed from the RV Marion Dufresne

on 30th May, 1997. Once onboard, time constraints

permitted only the upper 12 m of the 34.92 m-long

core to be analysed for physical properties on a multi-

sensor track (MST) equipped to measure magnetic

susceptibility, gamma rays (density) and P-wave velo-


city (Nees et al., 1998; Carter et al., 2002). Light reflec-

tance readings were also taken; in this instance with a

hand-held, Minolta CM-508I spectral photometer.

Magnetic susceptibility and density measurements

were rerun at 2 and 10 cm intervals respectively for

the whole core at the NIWA laboratory in Wellington. A

Bartington MS2 sensor, similar to that of the shipborne

MST system, measured magnetic susceptibility. Dry

bulk density (DBD) was determined from a known

volume of sediment, which was weighed wet, dried at

50 8C and reweighed to obtain the dry weight. DBD

was taken as the ratio of dry weight/wet volume, in-

cluding a correction for a salt content of 3.47%. The

results correlated well with those collected at sea. Con-

centrations of calcium carbonate were determined at

10 cm intervals for the whole core using a vacuum

gasometric procedure (Jones and Kaiteris, 1983). The

grain size of subsamples from 10 cm intervals in the

upper 6 m of the core was measured with a Sedigraph

5100 particle size analyzer. Biogenic silica or opal con-

tent was resolved at 50 cm spacings by X-ray diffraction

following a 24 h, 1000 8C heat conversion of opal to

cristobalite (Ellis and Moore, 1973). Although the

reported accuracy for this method is F5%, it was pre-

ferred to alkaline extraction techniques because of a risk

of leaching additional silica from volcanic glass present

throughout the core. Furthermore, the X-ray diffraction

data permitted comparison with core P69 of Stewart and

Neall (1984). Their concentrations of biogenic silica for

MIS 1–2 averaged 19.2% of total sediment in contrast to

amean of 4.2% for the same time interval inMD97 2121.

Other biogenic silica concentrations from cores off cen-

tral New Zealand tend to be b5% for MIS 1–2 (e.g.,

Carter et al., 2000). Differences between our data and

that of Stewart and Neall (1984) may relate to the re-

spective analytical techniques. Our analyses were made

on carbonate-free sediments whereas Stewart and Neall

(1984) added calcite to their samples as an internal

standard. It is possible that the carbonate acted as a

flux to liberate additional silica from the ubiquitous

volcanic glass.

The age model is based on stable isotope profiles

that were correlated with the SPECMAP profile of

Martinson et al. (1987) with some changes based on

more recent revisions of boundary ages such as for MIS

1/2 (e.g., Lean and McCave, 1998). d18O and d13C data

came from the planktic foraminifer, Globigerina bul-

loides. Samples were collected at 50 cm intervals,

equivalent to an average time interval of 1.77 kyr.

More detailed data, at decadal to century resolution,

were generated for last 15 kyr for the past deglaciation

and Holocene period (Carter et al., 2002). Isotopic

analyses were carried out at the Institute of Earth and

Life Sciences at the Vrije Universiteit, The Netherlands,

and at the National Institute of Water and Atmospheric

Research, New Zealand. Both institutes operate Finne-

gan MAT 252 mass spectrometers with automated cal-

cium carbonate preparation lines (Type Kiel II). Results

are versus the V-PDB scale and calibrated using the

NBS-19 standard. External reproducibility of a routine-

ly run laboratory standard was F0.10x and F0.05xfor d18O and d13C respectively, in both laboratories.

The resultant age model was refined by several

tephra layers that were correlated with dated tephras

onshore using the major element geochemistry of glass

shards and heavy mineral compositions (Carter et al.,

2002; Carter et al., submitted for publication). Addi-

tional time control is from 23 AMS dates on single

species foraminifers. Some of these determinations

post-dated publication of the last post-glacial MAR

record in Carter et al. (2002). Therefore, the profiles

presented here supercede the earlier curves. All ages in

this paper are expressed in calendar years following

correction based on the CALIB v.4.2 programme of

Stuiver et al. (1998).

The sediment burial flux or mass accumulation rate

(MAR) was derived according to the relationship;

MAR g=cm2=kyr� �

¼ linear sedimentation rate cm=kyrð Þ

� dry bulk density g=cm3� �

� g of component=g dry bulk sedimentð Þ

Linear sedimentation rates were calculated from

time intervals defined by the marine isotopic stage

boundaries, dated tephra layers and AMS dates. Dry

bulk density values came from the direct measurements

coupled with MST data. Finally, MARs were calculated

for three sediment components; biogenic carbonate,

biogenic silica and terrigenous sediment. The last com-

ponent includes aeolian and airfall volcanic detritus as

well as river-borne material.

Apart from a few macroscopic, pure tephras, which

can be confidently described as airfall, other volcanic

detritus is intimately mixed with terrigenous sediment

and is included with the terrigenous flux. Volcanic

detritus in some of these mixed units probably arrived

as primary airfall to be later mixed into the sediment by

physical processes or bioturbation (e.g., Kennett, 1981).

In other units the ash component was delivered to the

margin by rivers draining ash-covered landscapes and

redistributed to the slope (Lewis and Kohn, 1973;


Orpin, 2004). Additional ash was probably reworked

from terrestrial deposits by winds, as suggested for

loess derived from the Kawakawa Tephra of 26.5 ka,

(Pillans et al., 1993).

3. Sediment properties

3.1. Lithology

Cored sediments are mainly massive to faintly la-

yered, olive grey mud. Textural data, determined for

101 samples from MIS 1–2 (Carter et al., 2002), reveal

an average of 5% sand, 34% silt and 61% clay, with sand

rising to 25% in volcanic ash-rich layers. Magnetic

susceptibility and density measurements suggest the

presence of at least 88 ash-rich layers. The prevalence

ofmudwith terrigenous and biogenic components, a lack

of mass transport deposits, and the distinctive 3.5 kHz

signature of thin, parallel continuous sub-bottom reflec-

tors (Fig. 2), collectively support a mainly hemipelagic

mode of deposition.

3.2. Stable isotopes

The down-core profile of planktic d18O (Fig. 3)

correlates well with the MIS 1–6 sequence of the

SPECMAP record of Martinson et al. (1987) and

other profiles from off eastern New Zealand (Lean

and McCave, 1998; Nelson et al., 1985; Pahnke et

al., 2003). For interglacial maxima MIS 1 and 5e,

d18O values are 0–0.2x, whereas d18O reached

3.2x. in the last glacial maximum.

Using revised SPECMAP ages for stage and sub-

stage boundaries, the MIS 5–6 boundary of 129 ka is

positioned at 32.5m core depth (Fig. 3). The core base

is estimated to be ~139 ka through correlation with the

SPECMAP curve. The average linear sedimentation

rate is 28.3 cm/kyr, but the variance among stages is

high with averages for individual stages varying from

13.4 cm/kyr in MIS 4 to 40.0 cm/kyr in MIS 2.

d13C values for G. bulloides range between 0.4xand �2.05x with most positive concentrations tending

to mark cool periods (MIS 2, 5d) as well as MIS 3 (Fig.

3). Most negative d13C values mainly identify warm

periods, in particular substages 5a, c and e.

3.3. Magnetic susceptibility

Volume magnetic susceptibility, j (dimensionless

units), exhibits a fluctuating signal in MIS 1, but rap-

idly becomes subdued below the MIS 1–2 boundary

(Fig. 3). Initially, this change was regarded as an aber-

ration of the instrument’s sensitivity, but it was later

verified with another Bartington magnetic susceptibility

meter. The subdued signal is unlikely to reflect any

increase in diamagnetic calcite as the signal exhibits a

roughly inverse relationship with the calcium carbonate

profile, especially from MIS 3 to 5. More likely the

signal change reflects diagenetic dissolution of magne-

tic minerals (Greenwood, 2002).

Strong magnetic spikes in MIS 1 and more subdued

spikes in older sediments coincide with layers of tephra

and dispersed ash (Fig. 3). However, not all tephras

yield a strong magnetic response. The prominent Kawa-

kawa Tephra occurs at 9.84 m, but is not evident in the

j profile because of an abundance of glass shards and a

corresponding dearth of magnetic minerals. Instead, the

layer was accompanied by elevated dry bulk density

and depressed CaCO3 profiles.

3.4. Dry bulk density (DBD)

The general down-core increase in DBD is punctu-

ated by subtle elevations during MIS 2, 4 and 5/6 (Fig.

3). Superimposed on this broad trend is a series of

pronounced spikes associated with ash-rich layers. In

fact, DBD is more sensitive to tephra layers than mag-

netic susceptibility.

Oceanic sediments frequently have DBDs that di-

rectly correlate with the weight percent carbonate con-

tent (e.g., Snoeckx and Rea, 1994). Off eastern New

Zealand, the regional correlation between DBD and

carbonate is only moderate (r =0.71; Carter et al.,

2000). By comparison, DBDs from MD97 2121 are

only weakly correlated with carbonate content, and in

MIS 2 and 4, the correlation is negative (Figs. 3 and 4).

This weak relationship with carbonate presumably

reflects the influence of volcanic ash as confirmed by

the coincidence of elevated densities with ash-rich

deposits (Fig. 3). Volcanic material is dispersed

throughout the core, and may overwhelm any density

effect of carbonate, especially as carbonate contents are

low.

3.5. Biogenic carbonate

Carbonate contents fall within 4.2–24% of the total

sediment. The interglacial average of 12.8% is mar-

ginally higher than the glacial average of 10.2%.

However, sufficient stage differences exist to positive-

ly correlate the carbonate and stable isotope profiles

(Fig. 3). In MIS 1, CaCO3 increases down-core to a

maximum of 24% at 11.5 ka before declining to 5% at

the MIS 2–3 boundary. The MIS 3 profile increases

Fig. 3. Profiles of depth versus; d18O and d13C from Globigerina bulloides, magnetic susceptibility, dry bulk density, biogenic CaCO3 and biogenic silica. Marine isotope stages (shaded) are based

mainly on SPECMAP (Martinson et al., 1987).

L.Carter,

B.Manighetti

/Marin

eGeology226(2006)307–322

313

Fig. 4. Core profiles of mass accumulation rates for terrigenous (including volcanic and aeolian components), biogenic carbonate and silica, together with the eustatic sea level curve of Pillans et al.

(1998), and a profile of aeolian quartz accumulation in core E26.1 from the Tasman Sea (Fig. 1 inset; Hesse, 1994).

L.Carter,

B.Manighetti

/Marin

eGeology226(2006)307–322

314


down-core to Holocene levels, followed by an abrupt

decline in MIS 4 when carbonate fell to 4.2%. Last

interglacial values vary in consort with substages 5a–

e. Late MIS 6 has slightly reduced CaCO3 (8–15%)

similar to that of other cold periods. This basic profile

follows other published carbonate data from offshore

New Zealand (Carter et al., 2003; Lean and McCave,

1998; Nelson et al., 1985).

Superimposed on the stage-related trends are higher

frequency changes associated with tephras, which

markedly reduce carbonate. For example, the carbonate

minimum of early MIS 2 is affected by the Kawakawa

Tephra (see Lean and McCave, 1998). Similarly, the

low values in MIS 4 may relate to a suite of tephras.

3.6. Biogenic silica

Concentrations of biogenic silica sit between 1.8–

8.4% weight of total sediment with a mean of 4.3%

(Fig. 3). Silica is more abundant in interglacials (MIS

1=5%; MIS 5=5.4%) than glacials/stadials (MIS

2=3.5%; MIS 4=3.9%), with lowest contents reserved

for MIS 3 (2.9%).

4. Mass accumulation rates

4.1. Terrigenous flux

Given the overall high linear sedimentation rate at

MD97 2121, it is no surprise that total mass accumu-

lation is high with a mean rate of 27.0 g/cm2/kyr

(r =8.0; n=347). Accumulation is dominated by the

terrigenous flux with a mean of 22.7 g/cm2/kyr (r =7.1;

n =344). These MARs fluctuate (Fig. 4) the main pulse

occurring in MIS 2 when rates typically exceed

30 g/cm2/kyr (mean of 31.5 g/cm2/kyr; r =5.9, n =59).

Apart from an influx ~6 ka, interglacial MARs are lower,

but still substantial with a mean of 21.7 g/cm2/kyr

(r =5.9, n =174) for MIS 1 and 5 together. MIS 4 has

lowest rates of 12 g/cm2/kyr, with marginally higher

values in MIS 5b/c and 3.

4.2. Biogenic carbonate flux

Compared to terrigenous MARs, the mean carbonate

flux is lower and less variable with a mean of

3.1 g/cm2/kyr (r =1.1; n =347). Down-core profiles

show a distinct temporal variability that transcends

isotopic stage boundaries (Fig. 4). High rates of

N4 g/cm2/kyr characterize the glacial to interglacial

transitions, and the optimum warm phases in MIS 1 and

5e. MIS 2 and 5d witnessed a moderate accumulation of

biogenic carbonate of ~3 g/cm2/kyr, that locally peaked

at 5 g/cm2/kyr at the last glacial maximum. The lowest

biogenic fluxes of b2.5 g/cm2/kyr, are found in MIS 3, 4

and 5b/c.

4.3. Biogenic silica flux

Despite the lower resolution of the biogenic silica

profile, it resembles the carbonate record (Fig. 4). The

mean rate is 1.2 g/cm2/kyr (r =0.54, n =70). MARs

N2 g/cm2/kyr coincide with the glacial to interglacial

transitions and peak warm periods in MIS 1 and 5e,

whereas lowest rates are restricted to MIS 3–4.

5. Palaeoenvironmental controls of terrigenous

MARs

The terrigenous flux at MD97 2121 is influenced by

supply and transport/deposition, which are regulated by

palaeoenvironmental and geologic factors associated

with glacial/interglacial cycles as well as tectonic/vol-

canic events. We also recognize that these major cli-

matic cycles are modulated by more frequent changes

especially, in New Zealand’s case, by the El Nino-

Southern Oscillation (ENSO) (Gomez et al., 2004).

However, the resolution of the MD97 2121 record is

too coarse to identify individual phases of ENSO or

other activity, even at the sub-millennial scale of the last

15 kyr (Carter et al., 2002). We therefore concentrate on

the major climatic oscillations.

5.1. Supply

5.1.1. Fluvial supply

A lack of quantitative information on glacial/inter-

glacial discharges from east coast rivers allows for only

speculation based on circumstantial evidence. Palyno-

logical records for the North Island (McGlone et al.,

1994; Turney et al., 2003) reveal a glacial period flora

dominated by grasses and shrubs. This ephemeral cover

was frequently disturbed by climatic extremes. Thus,

the landscape was exposed to fluvial and wind erosion;

a contention supported by an abundance of Tertiary

pollen from erosion of the regolith in MIS 2 sediments

(McGlone, 2001). Nevertheless, fluvial erosion was

potentially offset by a drier glacial climate. In contrast,

interglacial erosion and river inputs probably reduced

under the stabilizing effect of an expanded forest cover.

An example of that effect is the 3–4 fold increase in

shelf sedimentation following the loss of forest that

accompanied European colonization of the region

(Gomez et al., 2004).


5.1.2. Volcanic supply

Eruptives from the rhyolitic centres of the Central

Volcanic Region were delivered to the ocean as fluvial

load and airfall. Large events, such as the Taupo erup-

tion of 1.718 ka, produced a flush of sediment to the

coast in response to; (i) denudation of the landscape

(Wilmshurst and McGlone, 1996), (ii) emplacement of

readily eroded ash and non-welded pyroclastic deposits

in river catchments (Pillans et al., 1993), and (iii) direct

offshore dispersal of airfall (Carter et al., 1995a,b).

Indeed, eruptive material is a significant component

of MD97 2121, which has over 88 tephras and ash-

rich layers, the latter formed through bioturbation of

thin airfall deposits (Kennett, 1981), redepositional

processes (Lewis and Kohn, 1973) or river discharge.

5.1.3. Aeolian supply

New Zealand contributes aeolian sediment to the

deep ocean, reflecting its position within the Roaring

Forties westerly wind belt. (Fenner et al., 1992; Stew-

art and Neall, 1984). Aeolian quartz was not measured

in MD97 2121. However, grain size and quartz anal-

yses from nearby core, P69, were used by Stewart and

Neall (1984) to infer an aeolian flux of ~3.4 g/cm2/kyr

(~14% of total flux) for MIS 2, and b0.5 g/cm2/kyr

(b5% of total flux) for MIS 1. These values are

considered estimates because of potential contamina-

tion by hemipelagic silt of similar grain size (e.g.,

Hesse, 1994). Perhaps the argument could be made

that the MIS 1 baeolian fluxQ is mainly hemipelagic

silt in light of diminished aeolian sources under the

expanded vegetational cover of warm periods. If cor-

rect, the MIS 2 aeolian silt would be closer to 10% of

the total budget. Regardless of quantities, the broad

trend of the Stewart and Neall data is supported by the

MIS 1–6 record in core E26.1 from the Tasman Sea

(Fig. 4; Hesse, 1994). Located upwind of New Zeal-

and (Fig. 1 inset) and sufficiently far from terrestrial

sources to be uncontaminated by hemipelagic detritus,

the Hesse record confirms the increased aeolian flux in

glaciations. It also reveals no significant aeolian in-

crease during the cold periods of MIS 4, 5b and 5d

(Fig. 4). We therefore conclude that aeolian contribu-

tion to MD97 2121 was small, possibly ~10% of the

total flux, in glacial times and reducing to b5% in

interglacial periods.

5.2. Changes of sea level and palaeocirculation

Correlation of the terrigenous MAR profile with the

last 140 kyr of sea level change (Fig. 4) confirms the

late regression-lowstand-early transgression phase as

the main period of high terrigenous input to the

ocean. For MIS 2, sea level was 113–120 m below

present (e.g., Carter et al., 1986). Thus, the Hawke Bay

shelf was exposed to about the modern ~140 m isobath

(assuming a �120 m palaeoshoreline that subsequently

subsided at 1 m/kyr for the past 20 kyr). Rivers extend-

ed over the emergent shelf to discharge their loads at

the palaeoshoreline. A similar scene is envisaged for

MIS 6 (Fig. 4). Although the core captured only the

uppermost part of MIS 6, it exhibits increased accumu-

lation over the MIS 6/5 transition (Fig. 4). In contrast,

the deep-ocean flux responded little to the lowstands of

cold periods MIS 4, 5b and 5d (Fig. 4). Then, sea level

was generally shallower than �75 m, and we contend

that an along-shelf sediment transport regime helped

divert sediment from moving off-shelf in a manner

similar to now. Presently, the active alongshore trans-

port zone occupies the 0–30 m depth (e.g., Gibb and

Adams, 1982). Whether these cold periods had reduced

fluvial inputs is uncertain, but their dry climates suggest

that this was the case (Mildenhall, 1994; Shane and

Sandiford, 2003).

During lowstands, MD97 2121 probably captured

sediment mainly from local rivers, with reduced input

from northern Waipaoa and Waiapu sources under a

waning ECC as proposed by Carter et al. (2002). There

was also an unquantified contribution from southern

sources suggested by incursions of SAW and associated

planktic foraminifers (Hendy, 1995; Nelson et al., 2000;

Weaver et al., 1998). At the time, transport in the upper

ocean was probably forced by vigorous glacial winds

(see Palaeoclimatic forcing of biogenic MARs).

The highstands of MIS 5a, c, e and 1 were also

accompanied by increased terrigenous fluxes (Fig. 4).

Using d13C to identify surface water masses, Nelson et

al. (2000) concluded that MIS 1 was dominated by

STW transported by a strengthened ECC. Thus, warm

periods favoured the introduction of suspended load

from prominent fluvial sources to the north of Hawke

Bay; this sediment shifting into path of the ECC as

hypopycnal or less common hyperpycnal plumes

(Hicks et al., 2004; Carter, 2005). Furthermore, an

invigorated ECC picked up additional sediment by

seabed erosion, as attested by current-scoured moats

along Lachlan Ridge (Barnes et al., 2002). Despite

this input, the flux profile for MIS 1 (our best-dated

period) exhibits a general decline over the last 4 kyr; a

trend that Carter et al. (2002) attributed to reinforce-

ment of along-shelf transport by the WCC. However,

over the last century, the decline in the flux stabilized

(but was not reversed) as human-induced erosion in-

flated river loads (e.g., Gomez et al., 2004).


6. Palaeoclimatic forcing of biogenic MARs

6.1. Glacial times MIS 4 to MIS 2

Despite the cold conditions throughout much of MIS

4–2 (e.g., Weaver et al., 1998; Mildenhall, 1994;

Pahnke and Zahn, 2005), the biogenic carbonate flux

was highly variable. Rates were lowest in MIS 4, and

then increased slightly in early MIS 3 before acceler-

ating through the MIS 3/2 transition to peak at the Last

Glacial Maximum (LGM) (Fig. 4). While cold-period

MARs may be affected by carbonate dissolution

(Weaver et al., 1998), those authors note that dissolu-

tion at nearby core P69 was minor in MIS 2. Thus,

assuming little post-depositional alteration at MD97

2121, MAR variability may be interpreted in terms of

changing marine production, which is controlled by

nutrient and light availability, temperature, and physical

oceanographic processes such as upwelling and ocean

surface mixing (e.g., Longhurst, 1995).

To identify possible causes of MAR variability, we

turn to MIS 2 which has the best palaeoenvironmental

control. In that stage, the surface waters at MD97 2121

comprised STW with a prominent SAW component

(Weaver et al., 1998; Nelson et al., 2000). This situation

resulted from (i) northward forcing of SAW through

Mernoo Saddle as the Southland Current intensified

under strong winds and increased density gradients

(Fig. 1; Barnes, 1992; Chiswell, 1996; Carter et al.,

2004), and (ii) the lowstand closure of Cook Strait

(Proctor and Carter, 1989), which inhibited any west-

erly diversion of the northward intrusion (see Heath,

1975). The resultant mixing of macronutrient-rich SAW

with the micronutrient-rich STW favoured production,

just as occurs at the modern Subtropical Front along

Chatham Rise (Boyd et al., 1999; Murphy et al., 2001).

Upwelling under strengthened glacial winds potentially

enhanced the nutrient supply. However, conditions

were not entirely favourable. Strong winds probably

increased the depth of surface mixing to carry phyto-

plankton below the photic zone (Sachs and Anderson,

2005). Furthermore, the supply of river-borne nutrients

may have reduced under the dry glacial climate (Turney

et al., 2003; Pillans et al., 1993).

Compared to the LGM, MIS 3–4 had very low

biogenic MARs (Fig. 4). MIS 4 is particularly puzzling

because its cold sea surface and terrestrial temperatures,

windiness and dry climate were similar to that of MIS 2

(Mildenhall, 1994; Pahnke and Zahn, 2005). Further-

more, it is likely that these cold periods also had mixed

SAW–STW waters, but we suggest that micronutrients

in the STW were depleted due to reduced fluvial dis-

charge. Compared to MIS 2, MIS 4 had a lower a

terrigenous MAR due to higher sediment capture on

the shelf that was partly submerged to �75 m (Fig. 4).

The reduced terrigenous and dissolved nutrient input to

the deep ocean helped drive down productivity. That

productivity is influenced by land runoff is inferred

from the association of high biogenic fluxes with the

high input margin of central New Zealand (see Region-

al perspective). Biogenic MARs climbed slightly in

early MIS 3, possibly because of increased river input

under a wetter climate (Mildenhall, 1994; Shane and

Sandiford, 2003), concomitant with short periods of

elevated productivity, which Sachs and Anderson

(2005) attribute to upwelling and/or a stratified oceanic

conditions. These productive phases declined into late

MIS 3 as climate became harsher into the LGM. The

subsequent drop in sea level and flush of fluvial sedi-

ment and nutrients directly to the deep ocean triggered

off the next flush in biogenic productivity.

Another consideration is productivity enhancement

through biron fertilizationQ by airborne dust, as mooted

for the glacial Southern Ocean by Kumar et al. (1995)

and Ikehara et al. (2000). Certainly, there was an east-

ward influx of aeolian detritus to the ocean in glacial

times (Thiede, 1979; Hesse, 1994), and this would be

the main opportunity for iron fertilization because

MD97 2121 received a strong if not dominating influx

of iron-limited SAW. In interglacial periods, the site

was dominated by STW, which is not iron limited.

Comparison of the aeolian MAR record of Hesse

(1994) with the carbonate and silica MARs (Fig. 4),

shows the profiles are out of phase. High aeolian fluxes

are restricted to glacial periods, whereas the highest

biogenic fluxes are an interglacial feature. There is

overlap between the LGM aeolian and carbonate fluxes,

but the peaks are offset by ~5 kyr (Fig. 4).

6.2. Interglacial periods—MIS 5 and 1

During glacial–interglacial transitions, biogenic

MARs increased markedly to reach maxima in early

MIS 5e and 1 (MIS 1 has two MAR peaks that probably

reflect its better age control compared to MIS 5).

Flushes of productivity also marked warm sub-stages

5c and 5a. (Fig. 4). Conditions became even more

favourable for plankton growth than in MIS 2. In

addition to the mixed micronutrient-rich STW and mac-

ronutrient-rich SAW, MIS 5e and 1 had warmer sea

surface temperatures (data of Nelson et al., 2000; Weav-

er et al., 1998) and weakened winds (Newnham et al.,

2003; Hesse, 1994). Thus, the surface ocean became

less mixed and better stratified. Rising temperatures


were accompanied by more precipitation (McGlone,

2001; Newnham et al., 2003), which presumably trans-

lated to higher fluvial discharges.

Following interglacial maxima, biogenic MARs re-

duced as the inflow of SAW waned and STW increased

(Carter et al., 2002; Nelson et al., 2000). Presently, the

prominent oceanographic feature at MD97 2121 is the

Wairarapa Eddy (e.g., Chiswell and Roemmich, 1998).

Significantly, this warm core eddy is a site of elevated

productivity (Bradford et al., 1982; Bradford and Chap-

man, 1988), which is concentrated at the eddy’s periph-

ery as outlined by remotely sensed distributions of

chlorophyll a (Murphy et al., 2001). Feeding such pro-

duction are nutrients from local rivers, upwelled water

from East Cape via the ECC (Bradford and Chapman,

1988), and occasional incursions of SAW (Heath, 1985).

However, these incursions were much less than glacial

counterparts, judging by the foraminiferal assemblages

(Weaver et al., 1998; Nelson et al., 2000).

6.3. Regional perspective

To appreciate the significance of the MD97 2121

biogenic fluxes, we compare the carbonate component

Fig. 5. Comparison of biogenic carbonate fluxes near the central New Zeala

and south of the Subtropical Front (STF). Sites are located on Fig. 1.

with the equivalent from other sites, including open-

ocean sites far from terrigenous supply. In that manner,

near-margin affects on biogenic production can be bet-

ter evaluated. Accordingly, comparisons are made with

Y9 on the pelagic carbonate-draped Campbell Plateau

in subantarctic waters, S924 on eastern Chatham Rise

in subtropical waters and, to confirm margin effects,

DSDP 594 located south of the Subtropical Front (Fig.

1, inset; 5).

Interglacial MARs, recorded along the continental

margin to the north (MD97 2121) and south (DSDP

594) of the Subtropical Front, are high with values

exceeding 4 g/cm2/kyr. MARs reduce in glacials and

stadials, with larger losses recorded at DSDP 594. By

comparison, carbonate MARs from more distant sites

are distinctly lower overall (Fig. 5). At subtropical site

S924, interglacial MARs average 1.0 g/cm2/kyr, while

average glacial rates are slightly less at 0.9 g/cm2/kyr.

Subantarctic site Y9 has a mean interglacial rate of

2.2 g/cm2/kyr, about half that recorded at margin sites,

whereas its glacial flux is 0.8 g/cm2/kyr.

Clearly, the eastern New Zealand margin has eleva-

ted productivity, but controls of this process appear to

vary between the two sites as inferred from the modern

nd continental margin and open Southwest Pacific Ocean, to the north


physical and biological oceanography. In today’s inter-

glacial conditions, DSDP 594 resides in a zone of high

productivity mainly caused by the interchange and

mixing of macronutrient-rich SAW with micronutri-

ent-rich water (Boyd et al., 1999; Murphy et al.,

2001). The latter water has two potential sources; (i)

the Southland Current carrying micronutrients from

local South Island rivers, and (ii) STW that periodically

passes south through Mernoo Saddle (Fig. 1; Greig and

Gilmour, 1992). Productivity at DSDP 594 may be

boosted further by upwelling related to (i) flow diver-

gence within the Southland Front (Fig. 1), and (ii) the

upward forcing of Subantarctic Mode Water in the

Saddle itself (Heath, 1985). In contrast, warm climate

productivity at MD97 2121 appears to be associated

more with the warm core Wairarapa Eddy, but with

possible enhancements from occasional incursions of

SAW, and upwelled water from the north (e.g., Chis-

well, 2000; Sutton, 2003; Chiswell and Sutton, 1998).

Cold periods were accompanied by changes in pro-

ductivity instigated by the increased northward intru-

sion of SAW to Hawke Bay and beyond. This intrusion,

together with lowered sea level, meant there was little

or no interchange with STW, south of Chatham Rise.

Fig. 6. Summary of terrigenous and carbonate MARs together with likely pal

to large rhyolitic eruptions from the Central Volcanic Region, the size/frequen

1–3. ECC = East Cape Current, WCC = Wairarapa Coastal Current, STW

temperatures. References are; 1 = Nelson et al. (2000), 2 = Carter et al. (2002

(2001), 6 = Mildenhall (1994, 1995), 7 = Shane and Sandiford (2003), 8 = Fr

al. (1989).

Indeed, subtropical planktic foraminifers are absent at

DSDP 594 during glacials (Weaver et al., 1998; Nelson

et al., 1993). We surmise that glacial productivity at

DSDP 594 resulted from increased upwelling under the

strong ocean circulation, coupled with the direct injec-

tion of iron and other micronutrients by local South

Island rivers. To the north, at MD97 2121, productivity

was influenced more by the direct mixing of SAW and

STW, the latter carrying micronutrients from fluvial

sources.

The lower biogenic fluxes at the distal ocean sites

(Fig. 5; Carter et al., 2000) probably resulted from a

reduced nutrient supply from land and possibly unfa-

vourable conditions for topographically influenced up-

welling. Certainly, remotely sensed distributions of

chlorophyll a confirm an offshore reduction in nutrient

availability and uptake (Murphy et al., 2001).

7. Conclusions

High terrigenous MARs off the collisional Austra-

lian/Pacific plate boundary are modulated by changes in

eustatic sea level and palaeoceanography (Fig. 6). Max-

ima of z30 g/cm2/kyr occurred during late regression-

aeoenvironmental controls. In the case of volcanism, this refers mainly

cy of which may be gauged by the amount of airfall estimated for MIS

= Subtropical Water, SAW = Subantarctic Water, SST = sea surface

), 3 = Weaver et al. (1998), 4 = Pahnke and Zahn (2005), 5 = McGlone

oggatt and Lowe (1990), 9 = Carter et al. (1995a,b), 10 = Berryman et


lowstand-early transgression phases when rivers dis-

charged at or near the shelf edge. This lowstand flux

was supplemented by aeolian detritus of uncertain

quantities. Intermediate terrigenous rates of 20–30 g/

cm2/kyr characterized warm sub-stages MIS 5a, c, e

and mid-MIS 1, when declining fluvial loads under an

expanding forest cover, development of an along-shelf

transport regime, and continued shelf subsidence,

favoured sediment capture. However, these warm per-

iods were also marked by a strengthened ECC, which

introduced sediment from sources north of MD97 2121.

Lowest fluxes of b20 g/cm2/kyr coincided with MIS 4–

3 when possible dry conditions and sea level helped

reduce input to the deep ocean.

Despite an overwhelming terrigenous signal, near-

margin sediments have biogenic carbonate MARs 2–4

times larger than those recorded away from the plate

boundary. Fluxes were lowest in MIS 4–3, probably

due to cold temperatures, reduced nutrient input, and

deep mixing under strong winds. Similar environmen-

tal conditions persisted into MIS 2 with the exception

that the accompanying lowstand allowed fluvial dis-

charge directly to the deep ocean thus fuelling a

revival of biogenic fluxes. Fertilization by aerosolic

iron was not a major influence as glacial biogenic

MARs are out of phase with coeval aeolian MARs.

Biogenic fluxes rose during deglaciations and peaked

in interglacial periods. Direct and circumstantial evi-

dence suggests that increased productivity resulted

from; (i) the mixing of macronutrient-rich SAW with

micronutrient-rich STW, (ii) reduced windiness and

surface layer mixing, (iii) increased precipitation and

nutrient run-off, and (iv) localised upwelling. Follow-

ing such flushes, MARs decreased as full subtropical

conditions became established.

Following identification of a glacial/interglacial sig-

nal for MD97 2121, the next research phase is to refine

key sections of the core record to decadal scales. In this

manner, flux responses to high frequency climatic dri-

vers such as ENSO, may be identified along with the

impacts of individual earthquakes and volcanic erup-

tions. Of special interest is the effect of interglacial

climatic optima on marine productivity as gauged by

the biogenic fluxes. If such optima are analogues of a

warmer ocean in the next century, then a well dated,

high resolution palaeo-record should improve predic-

tions of ocean change.

Acknowledgements

We are indebted to the officers and crew of the R.V.

Marion Dufresne for the acquisition of the giant piston

core, MD97 2121, which was funded by the Foundation

for Research Science and Technology (FRST). Thanks

also go to Gerald Ganssen of Vrije Universiteit,

Amsterdam, for providing stable isotope data, Brent

Alloway, IGNS, for tephra identifications and Lisa

Northcote for sediment analyses. The paper benefited

from the diligent critiques of Bob Carter of James Cook

University, Penny Cooke of Waikato University, and

the two journal referees.

References

Barnes, P.M., 1992. Mid-bathyal current scours and sediment drifts

adjacent to the Hikurangi deep-sea turbidite channel, eastern New

Zealand: evidence from echo-character mapping. Mar. Geol. 106,

169–187.

Barnes, P.M., Lewis, K.B., 1991. Sheet slides and rotational failures

on a convergent margin: the kidnappers slide, New Zealand.

Sedimentology 38, 205–221.

Barnes, P.M., Nicol, A., Harrison, T., 2002. Late Cenozoic evolution

and earthquake potential of an active listric thrust complex above

the Hikurangi subduction zone, New Zealand. Geol. Soc. Amer.

Bull. 114, 1379–1405.

Berryman, K.R., Ota, Y., Hull, A.G., 1989. Holocene palaeoseismi-

city in the fold belt and thrust belt of the Hikurangi subduction

zone, eastern North Island, New Zealand. Tectonophysics 163,

185–195.

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.

Bradford, J.M., Chapman, B.E., 1988. Epipelagic zooplankton

assemblages and a warm-core eddy off East Cape, New Zealand.

J. Plankton Res. 10, 601–619.

Bradford, J.M., Heath, R.A., Chang, F.H., Hay, C.H., 1982. The effect

of warm-core eddies on oceanic productivity off northeastern New

Zealand. Deep-Sea Res. 29, 1501–1516.

Brodie, J.W., 1960. Coastal surface currents around New Zealand.

N.Z. J. Geol. Geophys. 3, 235–252.

Carter, L., 2005. Suspended sediment dispersal over the high dis-

charge Waipaoa margin, New Zealand. N.Z. Mar. Sci. Conf.

(Abstracts). ISBN: 0-473-10299-4, p. 29.

Carter, R.M., Carter, L., Johnson, D.P., 1986. Submergent shorelines

in the SW Pacific: evidence for an episodic post-glacial transgres-

sion. Sedimentology 33, 629–649.

Carter, L., Nelson, C.S., Neil, H., Froggatt, P., 1995a. Offshore

occurrences of the Kawakawa and other late Quaternary tephra

in the Southwest Pacific Ocean. N.Z. J. Geol. Geophys. 38,

29–46.

Carter, L., Nelson, C.S., Neil, H.L., Froggatt, P.C., 1995b. Correla-

tion, dispersal, and preservation of the Kawakawa Tephra and

other late Quaternary tephra layers in the Southwest Pacific

Ocean. N.Z. J. Geol. Geophys. 38, 29–46.

Carter, L., Neil, H.L., McCave, I.N., 2000. Glacial to interglacial

changes in non-carbonate and carbonate accumulation in the SW

Pacific Ocean New Zealand. Palaeogeogr. Palaeoclimatol.

Palaeoecol. 162, 333–356.

Carter, L., Manighetti, B., Elliot, M., Trustrum, N., Gomez, B., 2002.

Source, sea level and circulation effects on the sediment flux to


the deep ocean over the past 15 ka off eastern New Zealand. Glob.

Planet. Change 33, 339–355.

Carter, R.M., Gammon, P.R., Millwood, L., 2004. Glacial–interglacial

(MIS 1–10) migrations of the Subtropical Front across ODP Site

1119, Canterbury Bight, Southwest Pacific Ocean. Mar. Geol.

205, 29–58.

Carter, L., Manighetti, B.M., Ganssen, G., submitted for publica-

tion. Surface-and deep-ocean change in the SW Pacific during

Antarctic Cold Reversal-Younger Dryas time. Earth Plan. Sci.

Letters.

Chiswell, S.M., 1996. Variability in the Southland Current, New

Zealand. N.Z. J. Mar. Freshw. Res. 30, 1–17.

Chiswell, S.M., 2000. The Wairarapa coastal current. N.Z. J. Mar.

Freshw. Res. 34, 303–315.

Chiswell, S.M., 2005. Mean and variability in the Wairarapa and

Hikurangi eddies. N.Z. J. Mar. Freshw. Res. 39, 121–134.

Chiswell, S.M., Roemmich, D., 1998. The East Cape Current and two

eddies: a mechanism for larval retention? N.Z. J. Mar. Freshw.

Res. 32, 385–397.

Chiswell, S.M., Sutton, P.J.H., 1998. A deep eddy in the Antarctic

Intermediate Water north of the Chatham Rise. J. Phys. Oceanogr.

28, 535–540.

Collot, J.-Y., Lewis, K., Lamarche, G., Lallemand, S., 2001. The giant

Ruatoria debris avalanche on the northern Hikurangi margin, New

Zealand: result of oblique seamount subduction. J. Geophys. Res.

106, 19271–19297.

Davey, F.J., Henrys, S., Lodolo, E., 1997. A seismic crustal section

across the East Cape convergent margin, New Zealand. Tectono-

physics 269, 199–215.

De Rose, R.C., Gomez, B., Marden, M., Trustrum, N., 1998.

Gully erosion in Mangatu Forest, New Zealand, estimated

from digital elevation models. Earth Surf. Process. Landf.

23, 1045–1053.

Ellis, D.D., Moore, T.C., 1973. Calcium carbonate, opal and quartz in

Holocene pelagic sediments and calcite compensation level in the

South Atlantic Ocean. J. Mar. Res. 31, 210–227.

Fenner, J., Carter, L., Stewart, R., 1992. Late Quaternary paleocli-

matic and paleoceanographic change over northern Chatham Rise,

New Zealand. Mar. Geol. 109, 383–404.

Foster, G., Carter, L., 1997. Mud sedimentation on the continental

shelf at an accretionary margin—Poverty Bay, New Zealand. N.Z.

J. Geol. Geophys. 40, 157–173.

Froggatt, P.C., Lowe, D.J., 1990. A review of late Quaternary silicic

and some other tephra formations from New Zealand: their stra-

tigraphy, nomenclature, distribution, volume and age. N.Z. J.

Geol. Geophys. 33, 89–109.

Gibb, J.G., Adams, J., 1982. A sediment budget for the east coast

between Oamaru and Banks Peninsula, South Island, New Zeal-

and. N.Z. J. Geol. Geophys. 25, 335–352.

Gomez, B., Carter, L., Trustrum, N.A., Palmer, A.S., Roberts, A.P.,

2004. ENSO signal associated with mid-Holocene climate change

in inter-correlated terrestrial and marine sediment cores. Geology

32, 653–656.

Greig, M.J., Gilmour, A.E., 1992. Flow through the Mernoo Saddle,

New Zealand. N.Z. J. Mar. Freshw. Res. 26, 155–165.

Greenwood, S., 2002. Assessing the effects of climate and diagenesis

on the magnetic properties of a sediment core from Hawke’s Bay,

New Zealand. Unpublished B.Sc (hons) thesis GS399, University

of Southampton. 30 pp.

Heath, R.A., 1975. Oceanic circulation and hydrology off the south-

ern half of South Island, New Zealand. N.Z. Oceanogr. Inst. Mem.

72, 1–36.

Heath, R.A., 1985. A review of the physical oceanography of the

seas around New Zealand—1982. N.Z. J. Mar. Freshw. Res. 19,

79–124.

Hendy, I.L., 1995. Paleoceanography of the Glacial–Holocene transi-

tion in the waters surrounding New Zealand. Unpub. M.Sc. thesis,

University of Waikato, Hamilton, New Zealand, pp. 1–194.

Hesse, P.P., 1994. The record of continental dust from Australia in

Tasman Sea sediments. Quat. Sci. Rev. 13, 257–272.

Hicks, D.M., Shankar, U., 2003. Sediment from New Zealand Rivers.

National Institute of Water and Atmosphere Chart. Misc. Ser., 79.

Hicks, D.M., Gomez, B., Trustrum, N.A., 2000. Erosion thresholds

and suspended sediment yields, Waipaoa River basin, New Zeal-

and. Water Resour. Res. 36, 1129–1142.

Hicks, D.M., Gomez, B., Trustrum, N.A., 2004. Event suspended

sediment characteristics and the generation of hyperpycnal plumes

at river mouths: East Coast continental margin, North Island New

Zealand. J. Geol. 112, 471–485.

Ikehara, M., Kawamura, K., Ohkouchi, N., Murayama, M., Naka-

mura, T., Taira, A., 2000. Variations of terrestrial input and marine

productivity in the Southern Ocean 480S during the last two

deglaciations. Paleoceanography 15, 170–180.

Jones, G.A., Kaiteris, P., 1983. A vacuum-gasometric technique for

rapid and precise analysis of calcium carbonate in sediments and

soils. J. Sediment. Petrol. 53, 655–660.

Kennett, J.P., 1981. Marine Tephrochronology. In: Emiliani, C. (Ed.),

The Sea, vol. 7. John Wiley, New York, pp. 1373–1436.

Kumar, N., Anderson, R.F., Mortlock, R.A., Froelich, P.N., Kubik, P.,

Dittrich-Hannen, B., Suter, M., 1995. Increased biological pro-

ductivity and export production in the glacial Southern Ocean.

Nature 378, 675–680.

Lean, C.M.B., McCave, I.N., 1998. Glacial to interglacial mineral

magnetic and palaeoceanographic changes at Chatham Rise, SW

Pacific Ocean. Earth Planet. Sci. Lett. 163, 247–260.

Lewis, K.B., 1980. Quaternary sedimentation on the Hikurangi

oblique subduction and transform margin, New Zealand. Spec.

Pub. Inst. Prof. Sediment. 4, 171–189.

Lewis, K.B., Kohn, B.P., 1973. Ashes, turbidites, and rates of sedi-

mentation on the continental slope off Hawkes Bay. N.Z. J. Geol.

Geophys. 16, 439–454.

Lewis, K.B., Pettinga, J.R., 1993. The emerging, imbricate frontal

wedge of the Hikurangi Margin. In: Ballance, P.F. (Ed.), South

Pacific Sedimentary Basins: Sedimentary Basins of the World.

Elsevier Science Publishers B.V., Amsterdam, pp. 225–250.

Lewis, K.B., Barnes, P.M., Collot, J.-Y., Mercier de Lapinay, B.,

Geodynz Team, 1999. Central Hikurangi Geodynz swath maps.

National institute of water and atmosphere chart. Misc. Ser. 77.

Lewis, K.B., Lallemand, S.E., Carter, L., 2004. Collapse in a Qua-

ternary shelf basin off east Cape, New Zealand; evidence for

passage of a subducted seamount inboard of the Ruatoria giant

avalanche. N.Z. J. Geol. Geophys. 47, 415–429.

Longhurst, A., 1995. Seasonal cycles of pelagic production and

consumption. Prog. Oceanogr. 36, 77–167.

Martinson, D.G., Pisias, N.G., Hays, J.D., Imbrie, J., Moore, T.C.,

Shackleton, N.J., 1987. Age dating and the orbital theory of the

ice ages: development of a high resolution 0–300,000-year chron-

ostratigraphy. Quat. Res. 27, 1–29.

McCave, I.N., Carter, L., 1997. Recent sedimentation beneath the

Deep Western Boundary Current off northern New Zealand.

Deep-Sea Res. 44, 1203–1237.

McGlone, M.S., 2001. A late Quaternary pollen record from marine

core P69, southeastern North Island New Zealand. N.Z. J. Geol.

Geophys. 44, 69–77.


McGlone, M.S., Salinger, M.J., Moar, N.T., 1994. Paleovegetation

studies of New Zealand’s climate since the Last Glacial Maxi-

mum. In: Wright, H.E., et al. (Eds.), Global Climates Since the

Last Glacial Maximum. University Minnesota Press, Minneapolis,

pp. 294–317.

Mildenhall, D.C., 1994. Palynostratigraphy and paleoenvironments of

Wellington, New Zealand, during the last 80 ka, based on paly-

nology of drill holes. N.Z. J. Geol. Geophys. 37, 421–436.

Mildenhall, D.C., 1995. Pleistocene palynology of the Petone and

Seaview drillholes, Petone, Lower Hutt valley, North Island, New

Zealand. J. R. Soc. N.Z. 25, 207–262.

Milliman, J.D., 1995. Sediment discharge to the ocean: the New

Guinea example. Geo Mar. Lett. 15, 127–133.

Milliman, J.D., Syvitski, J.P.M., 1992. Geomorphic tectonic control

of sediment discharge to the ocean: the importance of small,

mountainous rivers. J. Geol. 100, 525–544.

Murphy, R.J., Pinkerton, M.H., Richardson, K.M., Bradford-Grieve,

J.M., Boyd, P.W., 2001. Phytoplankton distributions around New

Zealand derived from SeaWiFs remotely sensed ocean colour

data. N.Z. J. Mar. Freshw. Res. 35, 343–362.

Nees, S., Jellinek, T., Suhonen, J., Winkler, A., Helmke, J., Shipboard

Scientific Party, 1998. Images III- IPHIS. Indian and Pacific

Ocean Pleistocene and Holocene History: an IMAGES Study.

Cruise MD106-Leg 1.2; RV Marion Dufresne II, Hobart (Aus-

tralia)- Christchurch (New Zealand) May 6–21, 1997.

Nelson, C.S., Hendy, C.H., Cuthbertson, A.M., Jarrett, G.R., 1985.

Late Quaternary carbonate and isotope stratigraphy, subantarctic

Site 594, Southwest Pacific. In: Kennett, J.P., von der Borch, C.C.,

et al. (Eds.), Initial Reports of the Deep Sea Drilling Program Leg,

vol. 90, pp. 1425–1436.

Nelson, C.S., Cooke, P.J., Hendy, C.H., Cuthbertson, A.M., 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,

435–458.

Nelson, C.S., Hendy, I.L., Neil, H.L., Hendy, C.H., Weaver, P.P.E.,

2000. Late glacial jetting of cold waters through the Subtropical

Convergence zone in the Southwest Pacific off eastern New

Zealand, and some geological implications. Palaeogeogr. Palaeo-

climatol. Palaeoecol. 156, 103–121.

Newnham, R.N., Eden, D.N., Lowe, D.J., Hendy, C.H., 2003. Rere-

whakaaitu Tephra, a land-sea marker for the Last Termination in

New Zealand, with implications for global climate change. Quat.

Sci. Rev. 22, 289–308.

Orpin, A.R., 2004. Holocene sediment deposition on the Poverty-

slope margin by the muddy Waipaoa River, East Coast, New

Zealand. Mar. Geol. 209, 69–90.

Orpin, A.R., Carter, L., Kuehl, S.A., Trustrum, N.A., Lewis, K.B.,

Alexander, C.R., Gomez, B., 2002. Deposition from very high

sediment yield New Zealand rivers is captured in upper margin

basins. Margins Newsl. 9, 1–4.

Pahnke, K., Zahn, R., 2005. Southern Hemisphere water mass con-

version linked with North Atlantic climate variability. Science

307, 1741–1746.

Pahnke, K., Zahn, R., Elderfield, H., Schulz, M., 2003. 340,000-year

centennial-scale marine record of Southern Hemisphere climatic

oscillation. Science 301, 948–952.

Pillans, B., McGlone, M., Palmer, A., Mildenhall, D., Alloway, B.,

Berger, G., 1993. The last glacial maximum in central and

southern North Island, New Zealand: a paleoenvironmental re-

construction using the Kawakawa Tephra Formation as a chron-

ostratigraphic marker. Palaeogeogr. Palaeoclimatol. Palaeoecol.

101, 283–304.

Pillans, B., Chappell, J., Naish, T., 1998. A review of the Milankovich

climatic beat: template for Plio–Pleistocene sea-level changes and

sequence stratigraphy. Sediment. Geol. 122, 5–21.

Proctor, R., Carter, L., 1989. Tidal and sedimentary response to the

late Quaternary closure and opening of Cook Strait, New Zeal-

and: results from numerical modelling. Paleoceanography 4,

167–180.

Reyners, M., McGinty, P., 1999. Shallow subduction tectonics in the

Raukumara Peninsula, New Zealand, as illuminated by earthquake

focal mechanisms. J. Geophys. Res. 104, 3025–3034.

Roemmich, D., Sutton, P.J.H., 1998. The mean and variability of

ocean circulation past northern New Zealand: determining the

representativeness of hydrographic climatologies. J. Geophys.

Res. 103, 13041–13054.

Sachs, J.P., Anderson, R.F., 2005. Increased productivity in

the subantarctic ocean during Heinrich events. Nature 432,

1118–1121.

Shane, P.A., Sandiford, A., 2003. Paleovegetation of marine isotope

stages 4 and 3 in Northern New Zealand and the age of the

widespread Rotoehu tephra. Quat. Res. 59, 420–429.

Snoeckx, H., Rea, D.K., 1994. Dry bulk density and CaCO3 relation-

ships in upper Quaternary sediments of the eastern equatorial

Pacific. Mar. Geol. 120, 327–333.

Stanton, B.R., Sutton, P.J.H., Chiswell, S.M., 1997. The East

Auckland Current, 1994–95. N.Z. J. Mar. Freshw. Res. 31,

537–549.

Stewart, R.B., Neall, V.E., 1984. Chronology of palaeoclimatic

change at the end of the last glaciation. Nature 311, 47–48.

Stuiver, M., Reimer, P.J., Bard, E., Beck, J.W., Burr, G.S., Hughen,

K.A., Kromer, B., McCormack, G., van der Plicht, J., Spurk, M.,

1998. INTCAL 98 radiocarbon age calibration, 24,000–0 cal BP.

Radiocarbon 40, 1041–1084.

Sutton, P.J.H., 2003. The Southland Current: a subantarctic current.

N.Z. J. Mar. Freshw. Res. 37, 645–652.

Thiede, J., 1979. Wind regimes over the late Quaternary Southwest

Pacific Ocean. Geology 7, 259–262.

Trustrum, N.A., Gomez, B., Page, M.A., Reid, L.M., Hicks, D.M.,

1999. Sediment production, storage and output: the relative role of

large magnitude events in steepland catchments. Z. fur Geomorp.

Suppl. Band 115, 71–86.

Turney, C.S.M., McGlone, M.S., Wilmshurst, J.M., 2003. Asynchro-

nous climate change between New Zealand and the North Atlantic

during the last deglaciation. Geology 31, 223–226.

Uddstrom, M.J., Oien, N.A., 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, 20729–20751.

Walsh, J.P., Nittrouer, C.A., 2003. Contrasting styles of off-shelf

sediment accumulation in New Guinea. Mar. Geol. 196,

105–125.

Warren, B.A., 1973. TransPacific hydrographic sections at latitudes

438S and 288S; the SCORPIO Expedition—deep water. Deep-Sea

Res. 20, 9–38.

Weaver, P.P.E., Carter, L., Neil, H., 1998. Response of surface water

masses and circulation to late Quaternary climate change, east of

New Zealand. Paleoceanography 13, 70–83.

Wilmshurst, J.M., McGlone, M.S., 1996. Forest disturbance in the

central North Island, New Zealand, following the 1850 BP Taupo

eruption. Holocene 6, 399–411.

glacial/interglacial control of terrigenous and biogenic fluxes in the deep ocean off a high input,...

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