AK thesis style - Open Research: Home · 5.2. Introduction 89 Figure 5.4: “Sediment Remobilisation Model” for mixed siliciclastic/carbonate sedimentation in the northern GBR

Part II

Palaeoceanography

83

Chapter 5

Siliciclastic/carbonate

sedimentation model for the

Capricorn Channel

5.1 Abstract

Mixed siliciclastic/carbonate sedimentation systems are poorly understood. Several con-ceptual models have been produced to explain the siliciclastic/carbonate sedimentationvariations related to sea level changes. A preliminary model is produced for the hemipelagicsedimentation in Capricorn Channel, using the evidence from grainsize, mineralogy, sta-ble isotopes and radiocarbon ages from a series of seven deep marine sediment corescollected from a depth transect. The model shows that terrigenous material from theFitzroy River influences the Capricorn Channel during the last glacial maximum (LGM)lowstand. Bathymetric evidence shows the Fitzroy River meandered across the conti-nental shelf and entered the Capricorn Channel to the north of Northwest Reef. Highsedimentation rates are found in the cores during the lowstand and early transgressionwith considerable amounts of siliciclastic sediments, the sedimentation rates drop off con-siderably as sea level continues to rise and carbonate, both neritic in the shallower coresand pelagic in the deeper cores, begins to dominate the sediments. There is evidence thatprimarily the clay fraction from the Fitzroy River influences the present sediments in theCapricorn Channel. This model for the southern GBR contrasts with the recently pro-posed models for the northern GBR and the sedimentation to the south of Fraser Island,on the subtropical east coast of Australia. The sedimentation in the Capricorn Channelappears to be a compromise between these two regions to the north and south. We pro-pose that the primary reason for the difference in sedimentation is related to the complexphysiography of the southern GBR region, which exerts a considerable influence on thesediment transport during lowstands and transgressions.

85

86 5. Siliciclastic/carbonate sedimentation model for the Capricorn Channel

Figure 5.1: “Reciprocal Model” for siliciclastic/carbonate sedimentation on a mixed margin.

During a sea level lowstand the terrestrial sediment bypasses the shelf and is deposited offshore on

the continental slope. During the transgression shelfal carbonate begins to be produced as sea level

rises and there is an increase in carbonate deposition offshore. During a sea level highstand the

siliciclastic sediment is deposited near shore, and carbonate is deposited off shore.

5.2 Introduction

Sedimentation on continental margins has traditionally been studied either in a contextof a carbonate or a siliciclastic sedimentary system depending on the dominant source ofsedimentary material. Many continental margins however receive considerable amounts ofboth carbonate and terrigenous material. The generally accepted model for tropical mixedcarbonate clastic systems is the “reciprocal model”. When sea level is low, rivers dischargelarge amounts of siliciclastic material to the continental slope. During a transgressionthe amount of siliciclastic material transported to the slope declines and the majorityof the river sediments are deposited on the shelf. With the rise in sea level, carbonate,from carbonate platforms and pelagic production begins to dominate the outer shelf andcontinental slope (Figure 5.1).

Sea level has varied by approximately 120 m during the Quaternary (Figure 5.2) (Chap-pell et al., 1996). The sedimentation response to these sea level changes on mixed passivecontinental margins varies considerably depending on the sediment supply, composition,physiography and regional climate. The role of sediment reworking and horizontal trans-port by near shore currents is also crucial in understanding sediment redistribution andfocussing. All of these factors must be considered when studying these mixed systems, aseach factor may play a more important role at different stages of a sea level cycle, alteringthe sediment accumulation rates and the carbonate/clastic balance.

In the Great Barrier Reef (GBR), and areas like the Caribbean (Haddad and Droxler ,

5.2. Introduction 87

Figure 5.2: Sealevel fluctuations calculated from corals, coral terraces etc. Data from Chappell

et al. (1996); Pillans et al. (1998); Cabioch and Ayliffe (2001); Lambeck et al. (2002).

1996), it also allows us to understand the evolution and changes in the regional climateand reefs over time. The coral reef of the GBR is thought to have been initiated less than500 ka, related to a change in the frequency of climate cyclicity (Davies and McKenzie,1993; Wilson et al., 2001; Webster and Davies, 2003). Reef growth has been terminatednumerous times due to aerial exposure from fluctuating sea levels throughout the Quater-nary (Figure 5.2) with the latest re-initiation of the modern reef at ∼8 ka (Hopley , 1982;Davies and Hopley , 1983; Peerdeman and Davies, 1993; Larcombe and Carter , 1995). Thisreef exposure results in the weathering and diagenetic alteration of the coral. As a resultthe geochemistry of the aragonite skeletons are altered and cannot be used as a reliablepalaeoclimate record over glacial-interglacial sea level cycles in the GBR. Consequentlythe only continuous palaeoclimate records of the GBR region can be found in deep marinesediment cores, which are influenced, but not interrupted, by sealevel fluctuations.

5.2.1 Previous work on the east coast of Australia

There have been a handful of studies on the temperate narrow shelf of southeast Australia(Marshall , 1979, 1980; Kudrass, 1982; Troedson, 1997; Ferland and Roy , 2000; Troedsonand Davies, 2001; Roberts and Boyd , 2004). South of Fraser Island, the continental shelfis narrow and wave dominated with the occasional subtropical carbonate bank and a steepcontinental slope. Evidence from cores from the narrow (∼30 km) continental shelf offNoosa, 26◦S, on the subtropical, southeast coastal margin of Queensland (GC-25, Figure5.3) display the classic sedimentation of the mixed carbonate/clastic reciprocal model(Troedson and Davies, 2001).

However recent studies on the tropical carbonate/siliciclastic system of the northern


Figure 5.3: Location of the Capricorn Channel at the southern end of the Great Barrier Reef

(GBR). The bathymetric contours at 100 m, 200 m, 500 m, 1000 m, 2000 m and 3000 m are shown

and highlight the relatively shallow gradient of the continental slope down the channel. The locations

of the Fitzroy River and Burdekin River are shown. Also highlighted in the small box is the area in

the northern GBR where the sediment remobilisation models have been developed (Dunbar et al.,

2000; Page et al., 2003; Dunbar and Dickens, 2003a,b). Also highlighted are the cores GC-10 on the

Marion Plateau (Page and Dickens, 2005) and GC-25 just off shore Noosa (Troedson and Davies,

2001).


Figure 5.4: “Sediment Remobilisation Model” for mixed siliciclastic/carbonate sedimentation

in the northern GBR. When sea level is low the rivers aggrade onto the shelf and the siliciclastic

sediment is trapped behind the barrier reefs on the edge of the shelf. During the transgression,

production of carbonate is initiated on the reefs, but when sea level reaches the shelf there is a pulse

of siliciclastic terrestrial sediment which is remobilised and transported offshore onto the continental

slope. During a sea level highstand the siliciclastic sediment is deposited near the coast and the

slope is dominated by carbonate production (after Dunbar and Dickens (2003b)).

Great Barrier Reef and Queensland Trough, between Townsville and Cairns have chal-lenged the reciprocal model (Dunbar et al., 2000; Dunbar and Dickens, 2003a,b; Pageet al., 2003). Their results suggest that maximum siliciclastic sedimentation, primarilysourced by the Burdekin River, do not conform to this conventional model. Instead themaximum siliciclastic deposition is exhibited during the late stage of the transgressions(11-7 ka BP) rather than during the sea level lowstand (18 ka BP). From these studies anew mixed sedimentation model for the northern Great Barrier Reef has been produced,“the sediment remobilisation model” (Dunbar and Dickens, 2003a,b; Page et al., 2003). Inthis model the siliciclastic sedimentation, which dominates during low sea level, is trappedon the continental shelf by the outer barrier fringe reefs. At the last glacial maximum low-stand the Burdekin River appears to meander across the exposed continental shelf andponds when it reaches the exposed reef platforms (Fielding et al., 2003). During the trans-gression the shelf is flooded and the siliciclastic sediment is remobilised and transportedonto the continental slope (Figure 5.4). Thus the maximum in siliciclastic deposition onthe slope is found during the transgression, as a result of the reworking of glacial materialfrom the shelf.

The main problem with this model of sedimentological changes of the GBR resultingfrom fluctuating sea level and climate change is that the majority of the work has beendone on the northern GBR (∼15◦S-18◦S), where the continental shelf is narrow (<100km) and the reefs form a near continuous ribbon along the edge of the continental shelf.


Northern Queensland also experiences a wet tropical climate. This is very different fromthe southern GBR where the shelf is much wider (∼200 km), and the reefs more dispersedacross the shelf (Maxwell , 1968). The southern region also experiences a dry tropicalclimate.

Recent work from the Marion Plateau, an offshore carbonate platform in the southernGBR (∼20.8◦S, GC-10 in Figure 5.3), however has also revealed that the highest fluxes ofsiliciclastic and carbonate material occur at the transgression (Page and Dickens, 2005),although the siliciclastic component shows only a minor increase during the transgression.This is unsurprising given its large distance from the present day sources of terrigenousmaterial.

The sedimentation in the Capricorn Channel, which forms a depression in the southernGBR, does not conform to either the “reciprocal model” or the “sediment remobilisationmodel”. We suggest that this is the result of differing physiography and climate in thesouthern GBR region.

5.2.2 Regional Setting

The Capricorn Channel is situated at the southern end of the Great Barrier Reef, Australia,between 22◦S and 25◦S (Figure 5.3). The Capricorn Channel separates the CapricornGroup and Bunker Group Reefs to the west from the Swain Reefs to the northeast (Figure??). It is situated at the widest point of the GBR (∼200 km). The shallow gradientchannel slopes down from 60 metres below sea level (mbsl) in the northwest to 4000 mbslin the southeast where it merges with the Tasman abyssal plain and the deep waters ofthe Cato Trough (Marshall , 1977). There is no definite shelf break in this region and theaverage gradient of the slope is 2◦, with a slight increase in slope gradient between 600 to1000 mbsl (Fairbridge, 1950).

The geological evolution of the Capricorn Channel began in the Late Cretaceous whenthe Australian and Antarctic plates slowly started separating from the Lord Howe Riseand Pacific Plate. This zone of seafloor spreading in the Tasman Sea continued northwardsinto the Cato Trough and the Coral Sea in the early Palaeocene and continues until theend of the Palaeocene (Mutter and Karner , 1980). The zone of spreading shows an offsetalong the transform fault between the Tasman Abyssal Plain and the Cato Trough and itis likely that the Capricorn Channel formed along a tensional region adjacent to this offset.The Capricorn Channel is made up of a series of north-westerly orientated fault blocksrather than single graben (Ericson, 1976). Drill cores show a terrestrial sequence in thechannel until a major marine transgression in the Late Oligocene. No evidence of coralreefs is found until the Pliocene as the Australian plate drifted north into warmer water(Palmieri , 1974). During the Quaternary the relative subsidence in the Capricorn Channelhad slowed and the shallower shelf (<200 m) has a mixture of terrigenous sediment, reefaland non-reef carbonate. Deeper In the channel the sedimentation has been controlled byglacioeustatic sea level fluctuations (Grimes et al., 1984).

The most substantial previous work on the Quaternary sediments from the Capricorn


Channel was described by Marshall (1977), the results of a 1975 Bureau of Mineral Re-sources (BMR) Cruise. The work looked at the shallow <200 mbsl areas of the channeland involved sea bed sampling, shallow seismic reflection profiling, bathymetric profiles,underwater photographs and some coring was attempted. A couple of interesting featureswere discovered during this survey. These included; symmetrical and asymmetrical quart-zose sand waves at a water depth of 60-80 mbsl; a small submerged shallow trough parallelto the isobaths, possibly an ancient shoreline; a series of east-west orientated reefal shoalsand banks from 55-60 mbsl up to 8 mbsl, some of these banks may be drowned reefs; highmagnesium (Mg) calcite ooids from depths of 100-120 mbsl (Marshall and Davies, 1975;Marshall , 1977). Radiocarbon dates of the ooids give a calibrated age of 16,800 ka BP(Yokoyama et al., in prep).

Marshall (1977)s work also provided a good survey of surface sediment distributionon the continental shelf, highlighting the extent and range of terrestrial river sedimentsfrom the rivers in the region. The surface sediments display a plume of high feldspar,quartz and rock fragments coming out of the Fitzroy River into Keppel Bay and onto thecontinental shelf. These terrestrial river sediments are found up to ∼120 km away from thepresent Fitzroy River mouth. However recent sedimentological work in Keppel Bay hassuggested that these sediments may be reworked relict sediments from a period of lowersea level (Ryan et al., 2005). The shallow seismic profiles and acoustic data highlightseveral channels cutting across the outer shelf of Hervey Bay and Maxwell (1968) andKrause (1967) revealed clearly defined drainage patterns around Hervey Bay and northof Fraser Island with channels cut to a base level of 64 m, corresponding to a Pleistocenelow sea-level. These are probably the palaeo-channels of the Mary, Burrum and ElliottRivers (Marshall , 1977). Marshall (1977) also suggested that during the glacial lowstand,the Fitzroy River meandered northeast across the shelf before being diverted down theCapricorn Channel.

5.2.3 Regional Climatology and Oceanography

The present climate of the Capricorn Region is dominated by the subtropical high-pressurezone with its axis about 25◦S during winter and 35◦S during summer, the result of asoutherly shift in the low pressure intertropical convergence zone (ITCZ) in the tropics.The combination of the tropical low pressure system and subtropical high pressure systemcreates the gradient that drives the southeast trade winds. These trade winds form astable system from April to December and vary little in intensity and direction. Duringthe summer the southeast trade winds are displaces to the south and the winds becomeweaker, except during tropical cyclones (Wolanski , 1994). The cyclone season lasts be-tween November and May and bring the majority of the precipitation to the region with60-70% falling between January and March. However there is considerable annual vari-ability in precipitation linked to the erratic nature of cyclones and the El Nino SouthernOscillation. The magnitude of river run off can vary significantly between years and usuallyoccurs as short lived floods. The main contributor of freshwater and sediment discharge


to the southern GBR is the Fitzroy River with a pre-industrial discharge of ∼2.5 Mt/yr(Neil et al., 2002).

Sea surface temperatures (SSTs) are an important control on the abundance and di-versity of marine biota in the region. Of special importance for the GBR and its tropicalcoral reefs is the southern most extent of the 18◦C isotherm, the so-called “Darwin Point”(Grigg , 1997), which is the limiting lower temperature at which tropical hermatypic coralcan survive. The Capricorn Channel is the present day limit of tropical coral reef growthin the GBR. The sea surface temperatures in the southern GBR range from 20.5-27.5◦CPickard1977, although temperatures as high as 29◦C have been recorded during coralbleaching events at Heron Island in the Capricorn Bunker Group (Hughes et al., 2003).The relatively high annual SSTs are the result of the main surface current in the region,the East Australian Current (EAC). This current originates as the Southern EquatorialCurrent (SEC) flowing east west along the equator as a result of the tropical trade winds.On collision with the Queensland Plateau the SEC bifurcates and the southern arm formsthe EAC. Using NOAA-9 AVHRR satellite images spanning a period of 2 years, Kleypasand Burrage (1994) showed that the EAC follows the 200 m contour outside the reefs of theGBR until it reaches the Capricorn Channel. Annual variation in regional oceanographicconditions results in the EAC either continuing to follow the slope contour westward alongthe shelf, or flowing directly south until it hits the shelf break near Fraser Island. Duringconditions of southward current flow, the current tends to bifurcate, producing a southerncurrent that continues along the coast and a northern component that becomes a cycloniceddy within the Capricorn Channel. The sea surface salinity ranges from 34-36 over theGBR, however it may be much lower close to a river mouth during the summer monsoon.

5.3 Methods and Samples

A series of gravity cores were taken down the Capricorn Channel from water depthsof 166 to 2892 mbsl (Table 5.1 and Figure 5.5) during the RV Franklin 1/97 Cruise.All these depths are above the present calcite lysocline of >3000 mbsl, although GC-13,GC-14 and GC-15 are below the present aragonite lysocline at 1100 mbsl (Figure 5.6).This is similar to the aragonite saturation depth of ∼1000 mbsl that was suggested byHaddad et al. (1993) for the Coral Sea. The lysocline (where saturation state =1) iscalculated from the alkalinity and total inorganic carbon (Appendix A for calculationdetails). These calculated saturation states suggest that the cores should be relativelyunaffected by carbonate dissolution.

Stable isotopes have been used to determine the chronostratigraphy of the cores usingthe δ18O of planktonic foraminifera Globigerinoides ruber compared with Martinson et al.(1987) SPECMAP and this was combined with AMS radiocarbon ages in several of thecores (Table 5.2). Methods for these analyses are outlined in Chapter 2. The chronos-tratigraphy was used to calculate sedimentation rates in the cores. There is no evidenceof hiatuses or turbidite sequences in these cores and therefore sedimentation between

5.3. Methods and Samples 93

Station # Core # Longitude Latitude Water depth (mbsl) Core length (cm)

Stn23 GC-09 23◦53’70S 152◦38’10E 166 123

Stn23b GC-10 22◦59’76S 152◦48’00E 335 436

Stn24 GC-11 23◦23’07S 153◦22’00E 502 578

Stn25 GC-12 23◦34’37S 153◦46’94E 990.5 564

Stn26 GC-13 23◦47’30S 154◦13’56E 1482 350

Stn27 GC-14 23◦49’52S 154◦21’67E 2004 260

Stn28 GC-15 23◦49’15S 154◦37’34E 2892 no recovery

Table 5.1: Details of cores collected from the Capricorn Channel during cruise R.V.FR1/97.

Figure 5.5: The complex bathymetry of the southern GBR and Capricorn Channel with the

location of the cores taken during RV FR1/97. The modern Fitzroy River meets the coast near

Rockhampton.


Figure 5.6: The calcite and aragonite lysocline calculated from the alkalinity and total inorganic

carbon data from the World Ocean Circulation Experiment (WOCE) cruise P21W. Saturation of

calcite (grey crosses) and aragonite (black squares) are plotted against depth in metres below sea

level (mbsl). The saturation state = 1 is marked on the graph (dashed black line) and is the position

of the lysocline. Below this depth the saturation sate is <1 and dissolution will start o occur. The

calculated aragonite saturation gives a lysocline (dotted black line) of ∼1100 mbsl, whilst the calcite

lysocline (dotted grey line) is at >3000 mbsl.

chronostratigraphic tie points is considered to be constant.

Carbonate content, sieved grainsize and x-ray diffraction (XRD) techniques were alsoundertaken on sub-samples from the cores (Chapter 2). The main minerals that wereanalysed for were quartz, feldspar, clays, calcite, high Mg calcite and aragonite (Figure5.7. Quartz, feldspar and clays are primarily terrigenous in origin and transported intothe marine environment by dust or riverine fluxes. Calcite is primarily pelagic in origin,whilst Mg calcite is neritic, utilised by benthic foraminifera and red algae living on andaround the reefs. Aragonite has two primary sources: pelagic pteropods, which live inthe upper 500 m of the ocean (Be and Hutson, 1977) and green algae and scleractiniancorals which live on the reef flats. Strontium concentrations in the sediment can be usedto determine between these two sources of aragonite as pteropods have low concentrations∼1500 µg/g, while coralline algae and corals contain >5000 µg/g (Dunbar and Dickens,2003a). Unfortunately strontium concentrations were not measured in this study.

Using the XRD results from the cores, variations in terrigenous, reef platform andpelagic sediments were determined. XRD was also used to analyse the clay mineral as-semblages from several sub-samples. Early low resolution maps of clay mineral assemblagesfound a latitudinal pattern with warm, humid, low latitude regions dominated by kaoli-nite, while physically weathered regions display a high percentage of micas (illite). Cold,high latitudes produced chlorite (Biscaye, 1965; Griffin et al., 1968). This data suggestedthat the distribution of different clay minerals were climate sensitive (Chamley , 1989).


Figure 5.7: Results of XRD analyses from samples within the cores. Samples from interglacial

periods are on the left hand side, whilst those from glacial periods are on the right. The shallowest

core is on the left in each group across to the deepest on the right. A) bulk sample, B) carbonate

minerals and C) siliciclastic components.


Lab Code Sample # Conventional

Radiocarbon

Age (years)

Error in Radio-

carbon Age (±years)

Calibrated

age (Calendar

years BP)

Error in Calibrated

Age (± Calendar

years)

ANUA 20305 GC-10 20cm 5550 210 6325 330

ANUA 22220 GC-10 70cm 11500 220 13550 350

ANUA 20306 GC-10 130cm 15875 330 18950 500

ANUA 20316 GC-12 20cm 4380 190 4537 250

ANUA 20313 GC-12 40cm 7930 210 8380 230

ANUA 20319 GC-12 80cm 10590 220 11720 400

ANUA 20320 GC-12 90cm 11910 350 13425 300

ANUA 22216 GC-12 190cm 16665 250 19270 410

ANUA 22218 GC-12 250cm 21880 260 25200 300

ANUA 22217 GC-12 300cm 23930 300 27580 350

ANUA 22219 GC-12 350cm 30290 480 34870 550

Table 5.2: Radiocarbon AMS dates and their calibrated ages for samples from GC-10 and GC-12.

Calibration up to 20 kyr BP was calculated using Calib 4.3, Marine 98 database (Stuiver et al., 1998).

Dates beyond 20 kyr BP are calibrated using the Bard et al. (1998)) polynomial. The reservoir age

for Heron Island in the southern GBR is 344±14 years (Druffel and Griffin, 1999).

Recent work on the Australian continent (Gingele and DeDeckker , 2004) which comparesriver clay assemblages with marine core clay assemblages has suggested a “reservoir con-cept”, where the distribution of clay minerals is determined by the rock types in thecatchment as well as climate. Therefore if rivers have distinct clay signatures they canbe used as a unique fingerprint of a particular clay provenance. However studies fromthe northern GBR highlight changing distributions of clay mineral between mud flats andmangroves (Aliano, 1978). Aliano (1978) work suggests that clay minerals, transportedto the coast by river suspended loads undergo significant diagenesis within environmentslike mangroves. Therefore interpretation of clay data must be treated with caution.

5.4 Results

5.4.1 GC-09

GC-09 was collected from a water depth of 166 mbsl, and 123 cm of core was recovered.It is difficult to determine an unambiguous age model from the relatively few isotopicdata points measured in this core (Figure 5.8). It is possible that the core reaches thelast glacial maximum (LGM) at ∼60 cm depth, as determined by the isotopically heavierδ18O. This would give a relatively low average sedimentation rate of 3 cm/kyr, comparedto 50 cm/kyr for similar depth cores in the west Queensland Trough (Dunbar et al., 2000).At this shallow depth within a broad channel like the Capricorn Channel it is likely thatthere is limited supply of sediment at this location in the channel or significant scouringby bottom currents through the channel. From Figure 5.8 there is a clear correlation be-tween grainsize and carbonate content, which suggests that the larger grains are primarilybioclastic carbonate fragments. During the LGM there is an increase in percentage of siltand mud which appears to be related to a reduction in bioclastic carbonate production in

5.4. Results 97

Figure 5.8: FR1/97 GC-09, grainsize, carbonate % and δ18O G. ruber plotted against depth in

the core. Grainsize is displayed as % >400 µm and the % <100 µm. The depths of XRD samples are

indicated by arrows on the right hand side, and the estimated depth (60 cm) of the LGM is shown

by a dashed black line.

and around the adjacent reefs when sea level was ∼120 m lower than today (Lambeck andChappell , 2001). XRD results from two samples in the core at 5 cm and 60 cm (Figure 5.7)exhibit high Mg calcite and aragonite compared with calcite, suggesting that the primarycarbonate influx to the core is neritic, rather than pelagic. At 60 cm depth there is also a10% increase in siliciclastic material, primarily made up of feldspars and quartz, indicatinga relative increase in terrestrial sediment during the LGM. XRD clay analyses of samplesfrom GC-09, highlights that the clay assemblages are dominated by kaolinite (>50%), butalso made up of illite (∼20%) and smectite (25-30%) (Table 5.3). During the glacial thekaolinite increased by ∼5%, the illite remained uniform, whilst the smectite decreasedby 5%. Data from the Fitzroy River (Gingele and DeDeckker , 2004) gives similar clayassemblages of 44% kaolinite, 18% illite and 37% smectite. The smectite content of theFitzroy is especially high, compared with other rivers discharging into the GBR lagoon,as a result of the basalt outcrops in the catchment. This suggests that the clay fractiontransported by the Fitzroy River is influencing the sediments of the Capricorn Channel.However, areas where mangroves are present also exhibit high montmorillonite (smectite)concentrations of up to 40-60% (Aliano, 1978). It is plausible that the high smectite foundat present is related to the coastal mangroves in this region.

5.4.2 GC-10

GC-10 was taken from a water depth of 335 mbsl, and 436 cm of core was recovered. Thiswater depth is presently dominated by the thermocline waters of the western south Pacificcentral water (WSPCW) and below the salinity maximum in the water column. Although


Sample # Smectite % Illite % Kaolinite %

GC-09 5cm 29.04 19.93 51.03

GC-09 60cm 23.67 21.57 54.76

GC-11 160cm 22.41 14.24 63.36

GC-11 260cm 26.03 16.77 57.2

Fitzroy River 36.93 17.57 45.1

Burdekin River 16.36 23.04 59.6

Table 5.3: % of different clay minerals within the clay fraction for 4 samples, compared with two

river samples (Gingele ,pers comm.).

at a lowered sea level it is likely to have been influenced by surface water circulation. Thislonger core (Figure 5.9A and B) contains adequate isotopic variability for correlation withSPECMAP (Martinson et al., 1987), and combined with three radiocarbon dates (Table5.2) provides an age model and calculation of approximate sedimentation rates.

There appears to be a high sedimentation rate during the last termination, which givesthe δ18O curve an unusual appearance compared to the normal asymmetric marine isotopiccurve. There is also a peak in sedimentation rates during the penultimate terminationalthough this is only brief and hence this transition gives a more typical asymmetricalcurve. The sedimentation rates vary considerably from a high of 12 cm/kyr down to <1cm/kyr (Figure 5.9B). During the glacials the grainsize increases and there is a higherpercentage of quartz and feldspar, but reduced clay content in the siliciclastic fraction.This suggests that the terrigenous fraction has increased and the core site is affected bya relatively high energy environment, which is too high for clays to settle and they aretransported further offshore. The carbonate content appears to correlate with the δ18O

signal with the minimums during and just lagging the glacial maximums, suggesting anincrease in siliciclastic material is not matched by an increase in carbonate at this time.The carbonate rapidly increases during the deglaciation to reach a peak of ∼75% duringthe mid interglacial. XRD results show that the greater abundance of CaCO3% duringinterglacials is the result of increase aragonite relative to calcite (Figure 5.7, presumablythe result of increased neritic material being transported down the channel.

5.4.3 GC-11

GC-11 was taken from a water depth of 502 mbsl, and 578 cm of core was recovered.This core sits in the thermocline waters of the WSPCW, although it maybe influenced bymixing with the intermediate waters of the AAIW below. From the δ18O of G. ruber, fourglacial cycles have been determined from correlation to SPECMAP (Martinson et al., 1987)(Figure 5.10A and B). No radiocarbon dates were analysed as it would only have beenpossible to date the top 50 cm of the core due to the low sedimentation rates of <1 cm/kyrto ∼4 cm/kyr (Figure 5.10B). These sedimentation rates are within the normal range ofabyssal cores of 1-2 cm/kyr, but this is much lower than the shallower and deeper cores ofGC-10 and GC-12 (discussed below). Higher sedimentation rates are evident during the

5.4. Results 99

Figure 5.9: A) FR1/97 GC-10, grainsize, carbonate % and δ18O G. ruber plotted against depth

in the core. Grainsize is displayed as % >400 µm and the % <100 µm. The depths of XRD samples

are indicated by arrows on the right hand side, and the depths of radiocarbon AMS dates (Table 5.2)

are displayed as grey triangles with the calibrated ages in text adjacent. The estimated depths of the

glacial maximums are displayed by a dashed black line, ∼130 cm for the LGM and ∼270 cm for MIS

6.2. B) Age model (black line) determined from radiocarbon ages and correlation with SPECMAP

(Martinson et al., 1987) with calculated sedimentation rates (dashed grey line).


terminations which drop off rapidly in the interglacials. The silt and mud fraction andcarbonate content display an inverse relationship throughout the core. Carbonate increasesrapidly at the terminations with highs during the interglacials of 85%, whilst during theglacial periods carbonate is as low as 50%. The carbonate highs of coincide with the verylow sedimentation rates of <1 cm/kyr and are probably dominated by pelagic productivity.This is evident in the XRD results which show that the majority of the carbonate in theinterglacial sample at 260 cm is made up of calcite (Figure 5.7). The glacial sample from160 cm is only 50% carbonate with the rest of the sample predominantly quartz andclays. Clay XRD displays a similar percentage of constituents to the Fitzroy River clayassemblages, with high smectite values during both the glacial and interglacial samples(Table 5.3).

5.4.4 GC-12

GC-12 was taken from a water depth of ∼990 mbsl, and 564 cm of core was recovered.The core sits within the AAIW. This is a similar length core to GC-11, however it hasa significantly higher accumulation rate as determined by nine radiocarbon dates and acorrelation to the SPECMAP curve (Martinson et al., 1987) (Figure 5.11). From the agemodel there is evidence for marine isotopic stages (MIS) 1 through to 5a, with sedimen-tation rates varying from 2 cm/kyr to 20 cm/kyr. This is a high sedimentation rate fora sedimentary basin, however, GC12 is situated on the increased slope of the CapricornChannel, therefore sediments scoured off the shallow gradient slope above may feasiblybe focussed at this depth on the slope. The sedimentation rate exhibits an unorthodoxpattern, with a very high rate (25 cm/kyr) during the middle of MIS 3. This anomalouslyhigh sedimentation rate highlights problems with calibrating the radiocarbon ages beyond20 kyr BP, especially within the age range of 20-40 kyr BP where there are large varia-tions in the flux of 14C in the atmosphere (Edwards et al., 1993; Bard , 1998; Hughen et al.,2004). The carbonate content highlights some interesting results with a higher percentage(>60%) during the middle of MIS 3, levels similar to those present at the top of the corein MIS 1. This was also found by Correge and DeDeckker (1997), who suggested thatthe temperatures in the AAIW during MIS 3 were similar to those during MIS 1 and5. Throughout MIS 2 and the termination the CaCO3% remains relatively constant at∼60%. Carbonate drops to 50% prior to a possible Antarctic cold reversal (between 14ka BP and 12 ka BP) at 80 cm in the core, before recovering to 65-70% in the Holocene.Carbonate content shows no correlation to sedimentation rates in the core. This suggeststhat the carbonate sources have kept pace with any changes in siliciclastic sedimentationfrom terrestrial sources. The XRD results suggest that the majority of the carbonate iscalcite during the interglacials, primarily from pelagic productivity. During the glacialsand transitions more quartz, feldspar, aragonite and Mg calcite are present (Figure 5.7).As well as fluctuations in sea level it has also been suggested that the intermediate watersof the AAIW have varied throughout the last glacial cycle (Chapter 6)and (Bostock et al.,2004).

5.4. Results 101

Figure 5.10: A) FR1/97 GC-11, grainsize, carbonate % and δ18O G. ruber plotted against depth

in the core. Grainsize is displayed as % >400 µm and the % <100 µm. The depths of XRD samples

are indicated by arrows on the right hand side. The estimated depths of the glacial maximums are

displayed by a dashed black line, ∼30 cm for the LGM and ∼150 cm for MIS 6.2, ∼330 cm for MIS

8 and ∼460 cm for MIS 10. B) Age model (black line) determined from correlation with SPECMAP

(Martinson et al., 1987) with calculated sedimentation rates (dashed grey line).


Figure 5.11: A) FR1/97 GC-12, grainsize, carbonate % and δ18O G. ruber plotted against

depth in the core. Grainsize is displayed as % >400 µm and the % <100 µm. The depths of XRD

samples are indicated by arrows on the right hand side, and the depths of radiocarbon AMS dates

(Table 5.2) are displayed as grey triangles with the calibrated ages in text adjacent. The estimated

depth of the glacial maximum is displayed by a dashed black line, ∼175 cm for the LGM and ∼450

cm for MIS 4. B) Age model (black line) determined from radiocarbon ages and correlation with

SPECMAP (Martinson et al., 1987) with calculated sedimentation rates (dashed grey line).

5.4. Results 103

5.4.5 GC-13

GC-13 was taken from a water depth of 1482 mbsl, and 350 cm of core was recovered(Figure 5.12A). The core was taken from the circumpolar deep waters (CPDW) of theTasman Sea. The age model was produced from correlation with SPECMAP (Martinsonet al., 1987) and plausibly extends back to MIS 6, with a sedimentation rate calculatedaccording to this age model (Figure 5.12B). Sedimentation rates fluctuation from 5 cm/kyr during the last transition and interglacial to 1 cm/kyr during the glacial maximum.Variations in sedimentation rates are difficult to calculate when few data points can becorrelated with SPECMAP. However, it is not dissimilar to the sedimentation rate curvesof the core GC-10 and GC-12 with increased sedimentation at the glacial/interglacialtransition. Peaks in the larger grainsize fractions correlate with pteropod accumulationsin the sediment, especially in the lower half of the core at 260 cm and 290 cm. Theabsence of complete pteropods at the top of the core may be related to the fact that GC-13 is presently situated below the aragonite lysocline. Therefore the aragonite lysoclinemay have been deeper during the glacials. The XRD results show a slight increase incarbonate percentage during the interglacial and increased clay and quartz during theglacial maximum (δ18O minimum) at 90 cm (Figure 5.7). During the interglacial thereare minor amounts of Mg calcite present, which suggest that there is little influence fromthe reef platform on the sediments during the present sea level or that the core is presentlybelow the Mg calcite compensation depth.

5.4.6 GC-14

GC-14 was taken from a depth of 2004 mbsl, and a 260 cm core was recovered (Figure5.13A). This core is also taken from the CPDW. An age model was estimated from acorrelation with SPECMAP (Martinson et al., 1987), from which it appears the coreextends back to MIS 3 (Figure 5.13B). From limited tie points with SPECMAP it is difficultto determine a sedimentation rate, however, given this age model a relatively uniform rateof ∼4 cm/kyr is found throughout the core. This is a relatively high sedimentation rate fora distal abyssal core, however without radiocarbon ages it is our best estimate. The XRDresults for the core top and glacial are similar to those from GC-13 (Figure 5.12). Againthere is minimal Mg calcite in the sediments today. There is a slight difference betweenGC-14 and GC-13 during the glacial with more quartz, and some feldspar present, alsoan increase in Mg calcite in GC-14. It is possibly that the position of the GC-14 coresite is receiving more sediment transported from the shelf than GC-13 due to submarinecanyons or gulleys transporting the sediments down the shelf bypassing large areas andfocussing sediment in other areas (Boyd pers. comm..) At water depths of >1500 mbsl,there is little variation in the quartz content between the glacial and interglacial periods.In the shallower cores the quartz:clay ratio changes considerably between the glacial andinterglacial with increased ratios during the glacial (Figure 5.13) This is not the case forthese two deeper cores, which display little change in this ratio. Therefore it appears that


Figure 5.12: A) FR1/97 GC-13, grainsize, and δ18O G. ruber plotted against depth in the

core. Grainsize is displayed as % >400 µm and the % <100 µm. The depths of XRD samples are

indicated by arrows on the right hand side. The estimated depth of the glacial maximum is displayed

by a dashed black line, ∼90 cm for the LGM and ∼340 cm for MIS 6?. B) Age model (black line)

determined from correlation with SPECMAP (Martinson et al., 1987) with calculated sedimentation

rates (dashed grey line).

5.5. Discussion 105

these cores are distal enough and deep enough that they are not experiencing any dramaticchanges.

5.4.7 GC-15

GC-15 was attempted at 2892 mbsl, unfortunately there was no recovery as the corebarrel hit a relatively solid surface and was bent. This suggests that there is minimalsedimentation at this location.

5.5 Discussion

The study of hemipelagic sedimentation is notoriously difficult to interpret as it is in-fluenced by many different factors including; transport, deposition, reworking, diagenesisand changing sediment sources over time. In the Capricorn Channel there is evidenceof riverine influences and pelagic productivity as well as transport off the proximal reefplatforms. It is also affected by a range of oceanic processes including wave, tides and cur-rents, as well as processes related sea level variations. These seven cores provide a previewof an interesting and complex sedimentary story within the Capricorn Channel during thelate Quaternary. A preliminary conceptual model has been produced to compare withthe mixed siliciclastic/carbonate models developed for the northern GBR and subtropicaleastern Australian margin.

5.5.1 Sedimentary Model

Glacial sedimentation

During the LGM the sea level was ∼120 m below present (Chappell et al., 1996; Yokoyamaet al., in prep), therefore the palaeo-shoreline was situated around the outside of the GBR,resulting in a subaerially exposed reef platform (Figure 5.14). It has been suggested thatthe Fitzroy River may have meandered across the shelf, initially diverted north around theCapricorn and Bunker Group reefs before retroflecting south down the Capricorn Channel(Marshall , 1977). New evidence from bathymetry data displays a bathymetric featurecorresponding to a major palaeo-channel, or relict incised valley of the Fitzroy River,present offshore from Keppel Bay (Figure 5.15) (Ryan et al., 2005; Webster and Petkovic,2005). The channel appears to run in a north-easterly direction (between North West Reefand Douglas Shoal), and is of comparable width and sinuosity to the modern day FitzroyRiver. The sinuosity ratio of the channel is 1.56, consistent with a stable meandering riversystem, which typically occur on slopes of <2◦ (Rosgen, 1994). Cut-off loops (or billabonglakes) are also apparent (Ryan et al., 2005). There is also evidence from echosoundingbathymetry data that the palaeo-Kolan, Burnett and Burnun Rivers reached the shorelineduring the LGM offshore Bundaberg, just north of Hervey Bay, where they incised the shelfand entered the Capricorn Channel from the southwest (Maxwell , 1968). Similar evidencefrom a survey using shallow seismic also indicates that the Burdekin River, which presently


Figure 5.13: A) FR1/97 GC-14, grainsize, and ?18O G. ruber plotted against depth in the core.

Grainsize is displayed as % >400 µm and the % <100 µm. The depths of XRD samples are indicated

by arrows on the right hand side. The estimated depth of the glacial maximum is displayed by a

dashed black line, ∼80 cm for the LGM. B) Age model (black line) determined from correlation with

SPECMAP (Martinson et al., 1987) with calculated sedimentation rates (dashed grey line).

5.5. Discussion 107

enters the northern GBR lagoon near Ayr, north Queensland, also incised the inner shelfand meandered across the shelf during the glacial lowstand (Fielding et al., 2003). Howeverthe palaeo-Burdekin River channel becomes progressively smaller and underfilled as itpasses near karstified reefs on the outer 10 km of the shelf. It is presently hard to tracethe palaeo-Fitzroy River channel deeper than 60 mbsl and although shallow seismic wasacquired in the Capricorn Channel (Marshall , 1977) it is of poor quality and not targetedin the right locations. There is evidence for minor depressions and infilled structures,which could indicate possible palaeo-Fitzroy River channels. The LGM evidence fromthe cores also supports the hypothesis that the Fitzroy River continued to influence thesediments within the Capricorn Channel. GC-09 displays a low carbonate content andhigh percentage of fine mud and silt sediments. The XRD clay analyses are very similarto the modern Fitzroy River clay proportions. This type of sedimentation is seen in thepresent day Fitzroy River estuary channels, just off shore from the estuary mouth), withhigh concentrations of quartz and feldspar (Ryan et al., 2005; Radke et al., 2005). GC-10 displays a very different pattern of sedimentation, and it would have probably beensituated below the storm wave base during the LGM. In this core high sedimentation ratesare evident during the LGM. GC-10 displays a similar low carbonate content and highquartz and feldspar content to GC-09, however it also displays an increase in grainsize.This can also be related to the modern Fitzroy River estuary where the distal end of theestuary channel and the outer Keppel Bay shelf displays coarser sediment than closer tothe mouth (Ryan et al., 2005). These two shallow cores suggest that the Fitzroy Riversediment plume was having a major influence on this area of the Capricorn Channelduring the sea level lowstand. This is further highlighted by the indicative 10% feldsparcontent which is characteristic of the modern Fitzroy River sediment plume influence. TheXRD results for the siliciclastic component of the sediments in the deeper cores displaya decreasing trend in quartz and feldspar with depth during the glacial, suggesting thatthere is a reduction in the terrestrial influence down the channel and with distance fromthe coast, as would be expected. The amount of quartz is higher in the LGM samples thanthe present, with the exception of cores GC-13 and GC-14. These deeper cores display anincrease in the clay content, however further analyses on the clays is required to determinewhether they also show a Fitzroy River signal. Therefore the terrestrial river influx wasprobably influencing the channel at least as deep as GC-12, and probably had a minimalinfluence on the deeper cores. GC-13 and GC-14 display abundant pteropods duringthe glacial and the presence of Mg calcite. This suggests that the aragonite lysoclineand the Mg calcite lysocline were potentially deeper during the glacial period. A deeperglacial carbonate lysocline was initially suggested for the central equatorial Pacific byFarrell and Prell (1989), however the evidence is controversial. In the northern GBR, anincreased depth in the aragonite lysocline is disputed by Haddad et al. (1993). Dissolutionand preservation cycles vary between the ocean basins and are related to atmosphericpCO2 as well as variations in the water mass carbonate saturations. All these factorsare unconstrained for the glacial ocean where dissolution proxies must be used to assess


the carbonate saturation. From the bathymetry evidence and the sediments in the coreswe propose that the Fitzroy River had a much greater influence on the sediments of theCapricorn Channel during the glacial than it does at present. Sedimentation rates in thecores are generally higher during the LGM, suggesting there was more terrestrial materialtransported to the channel. Palaeoclimate evidence from pollen and dune fields suggeststhat the LGM was much drier than present (Kershaw and Nanson, 1993). The riverswould probably not have transported high sediment loads during this period unless therewas a large reduction of vegetation in the catchment at this time. Evidence from modernsystems suggests that high precipitation after a period of drought will cause an increasein fluvial sediment discharge. It is possible that even though precipitation was lower theepisodic flooding may have caused large sediment discharge at the estuary and onto themarine shelf.

Transgression sedimentation

During the glacial/interglacial transition the sea level rose, with several hiatuses or periodsof sea level stasis during the transgression (Larcombe and Carter , 1995). Sedimentationrates which increased during the glacial and early transgression, rapidly drop off as thesea level rises. Evidence for these periods of stasis is seen as a series of ridges in both theseismic surveys from the Burdekin River (Fielding et al., 2003) and the Capricorn Channelregion (Marshall , 1977). It is suggested that the ooids found at 100-120 mbsl were formedwhen sea level was 95-115 m lower than present during one of these periods of sea levelstasis, prior to rapid sea level rise during meltwater pulse 1a (Yokoyama et al., in prep).The presence of ooids suggests that either at this particular point in time there was minorterrestrial influence allowing their formation, or that the river discharge into the channelwas relatively restricted in its spatial distribution. Unfortunately the sedimentation andthe lack of resolution on the majority of the cores from the Capricorn Channel, discussedin this study does not allow a detailed history of the changes in sedimentation duringthe transgression. Evidence from GC-12, which has the most well constrained chronology,suggests that the high sedimentation rates in these Capricorn Channel Cores were reachedprior to 11 ka BP and that with only a minor decrease in the carbonate content at this timeit appears that the increase in siliciclastic sediments was matched by a rise in carbonateproduction. The high sedimentation rate in the early transgression is considerably reducedduring the later stages of sea level rise. This is also evident from GC-10, but harder todetermine on the other cores as the chronology is poorly resolved.

Present day sedimentation

From the core tops it is apparent that GC-09 contains a large amount of quartz andconsiderably more Mg calcite than the deeper cores. It is also evident from the XRD clayresults that it is probably still influenced by the Fitzroy River clays. Therefore at 160 mbslthe present sediments of the Capricorn Channel are dominated by neritic carbonate and

5.5. Discussion 109

Figure 5.14: Sedimentation model for the Capricorn Channel. A) Last Glacial Maximum (∼18

ka BP), B) Transgression (∼12 ka BP), C) Mid-Holocene to Present day (∼6 to 0 ka BP). The

shoreline is shown by a thick black line. The influence of the feldspar and quartz plume is shown by

dashed black lines, whilst the estimated influence of the river clay sediments is shown in grey.


Figure 5.15: Bathymetric data of the continental shelf west of the Capricorn Channel. A) A

palaeo-channel leaving Keppel Bay and heading north of Northwest Reef is evident. B) Cross sections

through the palaeo-channel showing that as the channel heads further offshore it incises deeper into

the shelf, or that there has been less infilling in the deeper sections of the channel (Ryan et al.,

2005).

5.5. Discussion 111

some terrigenous river sediment from the Fitzroy River. GC-10 displays a slight increase inthe low Mg calcite content and a reduction in quartz and feldspar. This deeper and moredistal core therefore, shows an increase in pelagic carbonate influx. GC-11 continues thistrend but at a reduced sedimentation rate. GC-12 shows an increase in aragonite content,but very similar carbonate content to GC-11. It also has relatively high sedimentationrates. With the increase in the slope gradient of the channel between 600 and 1000 mbsl itis possible that the sediment on the shallow gradient shelf are being scoured and focusseddown slope in the region of GC-12. GC-13 and GC14 both show minimal Mg calcite andthere are two explanations for this; that they are distal enough that they are not influencedby the neritic carbonate or they are below the present Mg calcite compensation depth.An excess of 12% Mg in the Mg calcite produces a more soluble mineral than aragonite(Walter and Morse, 1984; Haddad and Droxler , 1996), and therefore it is plausible hat thedepth of 1500 mbsl is below the present compensation depth for Mg calcite. However inthe northern GBR studies have suggested that the Mg calcite concentration correlates withdistance from the reef rather than depth (Dunbar and Dickens, 2003a), although it maybe a combination of the two factors as well as accumulation rate. There is also an absenceof pteropod layers in the Holocene sediments of the cores which fits with the calculatedpresent aragonite lysocline of ∼1100 mbsl (Figure 5.6). This is slightly shallower thanthe present aragonite lysocline depth (∼1700 mbsl) suggested by Correge (1993) frompteropod abundances and fragmentation in the Queensland Trough. These deeper coresare primarily influenced by pelagic calcite productivity with only minor abundances ofquartz in the deeper cores, suggesting a minimal terrigenous influx, potentially by aeoliandeposition due to the westerly air currents coming off the Australian continent (Hesse,1994). No sediment was recovered from the deepest core at 2892 mbsl. It is unlikelythat the carbonate compensation depth (CCD) has been breached at this depth as thecalculated calcite lysocline depth is >3000 mbsl (Figure 5.6). The location of GC-15 wasright at the end of the Capricorn Channel near the entrance to the Cato Trough (Figure??) and sitting in the Antarctic Bottom Waters (AABW). The Cato Trough is the mainconduit of deep water from the Tasman Sea north to the Coral Sea. The narrowness of thispassage north and the exchange of water through suggest that there may be some relativelystrong currents at this location. An increase in current strength has been suggested inmodels by Hughes (1993). Hughes (1993) calculated an average current speed of up to 3cm/s using a Smagorinsky mixing run model. This current is not great enough to causecomplete scouring of the sediment, however it may be stronger episodically. Anotherexplanation is the submarine canyons down the Capricorn Trough (Marshall , 1977) havediverted sediment away from this location. Therefore along with possible winnowingthis location is receiving little sediment from shallower regions of the Capricorn Channel.Within the present Capricorn Channel there is evidence for a transition from a hemipelagicsedimentation, with a large influence of terrestrial sediments from the Fitzroy River andcarbonate shelfal material at shallow depths in the northwest of the channel to morepelagic dominated sediments at >1500 mbsl.


5.5.2 Comparisons with other siliciclastic/carbonate regions

Comparing this mixed sedimentation model with the subtropical southeast Australiancoast (Troedson and Davies, 2001; Roberts and Boyd , 2004) and the northern GBR (Dun-bar et al., 2000; Page et al., 2003; Dunbar and Dickens, 2003a,b) highlights that duringa sea level highstand similar siliciclastic/carbonate sedimentation is occurring in all threeregions. During a highstand the mid shelf siliciclastic sediments from the rivers are de-posited and distributed along the coast by proximal longshore currents. At the same timethe outer shelf is starved of sediment, which allows the formation of carbonate banks andreefs. These carbonate factories provide neritic carbonate sediments to the continentalslope.

With a fall in sea level, however all these regions display differing sedimentationregimes. The southeast coast transports the terrestrial sediment across the narrow shelfdown the continental slope, similar to the classic “reciprocal model” (Figure 5.1). Whilstthe temperate carbonate banks are exposed and very little carbonate is produced. Inthe northern GBR the “sediment remobilisation model” (Figure ??) shows that the riversmeander across the shelf, but the ribbon reefs at the edge of the shelf act as a barrier andthe sediment aggrades on the shelf behind the reefs, with very little siliciclastic sedimenttransported to the continental slope. The Capricorn Channel however displays continu-ous terrestrial sedimentation throughout the sea level lowstand. In the shallower coresan increase in quartz and feldspar is evident during the glacials suggesting an increasedinfluence of river sediment in the channel.

During the transgression the subtropical southeast coast displays the classic decreasein siliciclastic sedimentation and increase in carbonate content. The northern GBR showsa massive increase in sedimentation dominated by the siliciclastic material between 12 and8 ka BP. This is suggested to be the result of the higher sea levels remobilising the sedimentfrom the continental shelf that aggraded behind the reef during the lowstand. The Capri-corn Channel shows an initial increase in sedimentation during the early transgressionbut then decreasing sedimentation rates from 12-15 ka BP in cores GC-10 and GC-12. Inboth of these cores the carbonate content remained relatively constant, between 50-70%,throughout the transgression, suggesting that any changes in siliciclastic sedimentationwere matched by carbonate production. The changes in siliciclastic/carbonate sedimen-tation in these three regions are summarised in Figure 5.16.

The differences evident in these mixed siliciclastic/carbonate regions are related toseveral factors. The primary factors for the east coast of Australia are the physiographyof the regions and the spatial and temporal variations in climate and possibly ocean cur-rents. Northern Queensland presently experiences a wet tropical climate, whilst southernQueensland is still tropical but much drier. Both regions experience summer monsoonalrains. The present sediment yield from the Burdekin River is very similar to that of theFitzroy River (Neil et al., 2002). During the glacial the palaeo-climate evidence from veg-etation changes (Moss and Kershaw , 2000; Kershaw et al., 1993) and dust flux (Hesse,1994) suggests that precipitation was considerably reduced. There is also evidence that

5.5. Discussion 113

Figure 5.16: Carbonate v siliciclastic sedimentation variations with sea level in the three regions;

A) subtropical east coast, B) northern GBR and C) Capricorn Channel. Carbonate is shown in grey,

siliciclastic sedimentation is in black. (Adapted from Dunbar and Dickens (2003b); Page et al.

(2003))

the Australasian monsoonal system was altered (Gingele et al., 2002). The majority ofstudies have suggested that glacial periods are cool and dry, whilst interglacials are warmand wet. However better resolved records suggest that precipitation lags temperaturechanges and that the transgression was the driest period during the last glacial cycle, withthe wettest phases during the interstadials (Kershaw and Nanson, 1993). As a result ofthese climatic changes, it is highly likely that variations in terrestrial sediment yields oc-curred. The evidence for a dry climate during the transgression may explain why there isonly a small increase in terrestrial sedimentation in the Capricorn Channel. There is alsoevidence for palaeoceanographic changes in the south Pacific region during the last glacialcycle. Changes in the strength and pathway of the East Australian Current may haveaffected carbonate production in the southern GBR during the transgression (Chapter7). However the most likely explanation for the differences in the sedimentation betweenthese regions is the physiography. The antecendant topography has a significant affect onthe sedimentation during lowstands. The southern GBR has very complex physiographycompared to the northern GBR and the subtropical southeast coast. In the southern GBRthe coral reefs of the Capricorn and Bunker Group do not form a barrier at the edge ofthe shelf, but are a series of widely spaced patch reefs. This allows the rivers to flow be-tween the karstified reef structures at low sea level and continue to contribute terrestrialsiliciclastic sediments to the slope, although some siliciclastic sediment may get trappedon the shelf. Compared to the wave dominated subtropical southeast coast, however, thesouthern GBR has a very wide continental shelf which is dominated by tides and longshorecurrents and therefore any terrestrial sediment that is deposited on the continental slopehas to be transported considerable distances and is probably reworked and remobilised byseveral processes. From this data it appears that it the southern GBR region is a mix-ture of the classic “reciprocal model” and the “sediment remobilisation model”. This is


a preliminary model for the mixed siliciclastic/carbonate sedimentation of the CapricornChannel and the southern GBR region, more work will be required to improve the model.It is a spatially complex region with many processes affecting the shallow shelf and thecontinental slope.

5.6 Conclusions

Mixed siliciclastic/carbonate systems are highly complex with many factors affecting thesedimentation during sea level rise and fall. Using the sedimentary data from a seriesof cores in the Capricorn Channel in the southern GBR we have produced a preliminarysedimentation model for the region. During the glacial the Fitzroy River meandered acrossthe shelf and enters the Capricorn Channel in the northwest. As a result the sediments ofthe Capricorn Channel are influenced by riverine terrestrial sediments during the LGM.During the early transgression there is a high sedimentation rate in the channel withincreases in both carbonate and siliciclastic material (probably reworked from the shelf).At higher sea levels the sedimentation rate drops off considerably and by 11 ka BP theaccumulation rates are low, dominated by neritic and pelagic carbonate. At present theFitzroy River deposits the majority of is sediment proximal to the coast and only the finefraction clays influence the Capricorn Channel. Comparing this model with the classic“reciprocal model” for mixed sedimentation systems and the recently developed “sedimentremobilisation model” for the northern GBR, highlights that the southern GBR is a crossover of these two models. We believe the main reason for this difference is the complexphysiography of the southern GBR region. The physiography and antecedant topographyplay a important role in the sedimentation during sea level lowstands and transgressions.We thus propose that the morphology of the continental shelf is as important as theclimate, rivers, currents and the presence of tropical, subtropical and temperate carbonatebanks, in understanding the mixed sedimentation of hemipelagic regions.

Also evident in the data is the changing carbonate lysoclines and compensation depthsfor different carbonate minerals during the last glacial cycle. The data suggests that thearagonite and Mg calcite lysoclines were deeper during the glacial.

Chapter 6

Changes in the AAIW Circulation

Accepted as a paper entitled “Carbon isotope evidence for changes in Antarctic Inter-mediate Water circulation and ocean ventilation in the southwest Pacific during the lastdeglaciation” to Paleoceanography in July 2004, co-authored by Opdyke, B. N., Gagan,M. K. and Fifield, L. K.

6.1 Abstract

Deep-sea sediment core FR1/97 GC-12 is located 990 metres below sea level (mbsl) inthe northern Tasman Sea, southwest Pacific, where AAIW presently impinges the conti-nental slope of the southern Great Barrier Reef (GBR). Analysis of carbon (δ13C) andoxygen (δ18O) isotope ratios on a suite of planktonic and benthic foraminifera revealsrapid changes in surface and intermediate water circulation over the last 30 ka BP. Duringthe LGM, there was a large δ13C offset (1.1 ) between the surface-dwelling planktonicforaminifera and benthic species living within the AAIW. In contrast, during the lastdeglaciation (Termination 1) the ∆δ13Cplanktonic-benthic offset reduced to 0.4 prior toan intermediate offset (0.7 ) during the Holocene. These variations may be related to thedominance and direction of AAIW circulation in the Tasman Sea, and increased oceanicventilation, can account for the rapid change in the water column ∆δ13Cplanktonic-benthicoffset during the glacial-interglacial transition. The results support the hypothesis thatintermediate water plays an important role in propagating climatic changes from the polarregions to the tropics. In this case, climate variations in the Southern Hemisphere mayhave led to the rapid ventilation of deep water and AAIW during Termination 1, whichcontributed to the postglacial rise in atmospheric CO2.

115

116 6. Changes in the AAIW Circulation

6.2 Introduction

The exact sequence of climatic events that occurred between the LGM and the Holoceneis still controversial. Evidence from trace gases and isotopes trapped in ice cores fromGreenland (GRIP and GISP II) and Antarctica (Vostok, Dome C, Taylor, Byrd) hasrevealed some of the leads and lags in propagating global climate change between thehemispheres (Sowers and Bender , 1995; Blunier et al., 1998; Alley , 2000; Steig and Alley ,2002). An important challenge in paleoceanography is to link the atmospheric changesobserved in polar ice core records with variations in the tropics, as recorded in marinesediment cores. The ultimate aim is to determine the progression of events, leads andlags, and thereby understand the causes of feedback loops within the climate system.

One popular theory has argued that the last deglaciation was initiated by the for-mation and strengthening of the North Atlantic Deep Water (NADW) and the returnto interglacial thermohaline circulation (Oppo and Fairbanks, 1987; Howard and Prell ,1994; Flower et al., 2000). However, there is a growing body of work suggesting that theNorthern Hemisphere lags the changes in the Southern Hemisphere (Labeyrie et al., 1996;Charles et al., 1996; Ninneman and Charles, 1997; Blunier et al., 1998; Petit et al., 1999;Brathauer and Abelmann, 1999; Henderson and Slowey , 2000; Loubere, 2000; Lassen et al.,2002; Spero et al., 2003). A recent review paper by Alley et al. (2002), which re-analysedthe existing data sets, reaffirms that the glacial/interglacial cycles appear to be controlledby northern insolation and the Northern Hemisphere leads the deglaciation. The NorthernHemisphere records are complicated, though, by millennial events, thereby often makingthe timing of the deglaciation difficult to identify. (Alley et al., 2002) also propose thatthere was considerable inertia caused by the presence of the large Northern Hemisphereice caps, which helped maintained the glacial conditions in the north, thus delaying theonset of deglacial changes compared with the relatively rapid sequence of events in thesouth.

This supports the original theory of Milankovitch (1941) and Imbrie et al. (1992). In-solation primarily controls the commencement of the transition from glacial to interglacialconditions, however there are many factors that accelerate and enhance the weak insola-tion signal and its associated warming. It has been suggested that rising air temperaturesand the release of CO2 through rapid ventilation of deep and intermediate waters in theSouthern Hemisphere may have augmented the deglaciation (Francois et al., 1997; Speroand Lea, 2002). The temperature increase in the high latitudes of the Southern Hemi-sphere had a direct effect on the surface conditions of the Southern Ocean, which alteredthe character of the source waters of the AAIW (Lynch-Stieglitz et al., 1994). The AAIWis considered the most plausible vehicle for transporting these changes from the SouthernHemisphere polar front to the tropics during deglaciation (Ninneman and Charles, 1997;Loubere, 2000; Spero and Lea, 2002).


The majority of studies of the south Pacific have concentrated on reconstructing thecomplex dynamic oceanic processes in the high productivity region of the eastern equato-rial Pacific (EEP; (Mix et al., 1991; Loubere, 2001, 2000; Koutavas et al., 2002; Feldbergand Mix , 2003; Koutavas and Lynch-Stieglitz , 2003; Spero et al., 2003)). However there isa paucity of studies that investigate reconstructing circulation changes in the southwestPacific (Martinez , 1994, 1997; Passlow et al., 1997; Correge and DeDeckker , 1997; Kawa-gata, 2001). Most of these studies, however, have concentrated on the surface waters,whilst the intermediate waters in this region are poorly understood.

This Chapter aims to highlight changes in ocean circulation in the northern TasmanSea for the last 30 kyr using foraminiferal carbon and oxygen isotope ratios in deep-seasediment core FR1/97 GC-12. Core GC-12 is strategically located at 990.5 mbsl wherethe core of the AAIW impinges on the continental slope off the southern end of the GBR,(Figure 6.1).

6.2.1 Oceanography

The present oceanography of the south Pacific and the Coral Sea and Tasman Sea, thesouthwest Pacific, has been discussed in Section 1.3. The present distribution of the AAIWhas also been discussed in detail in Chapter 4. Therefore this is just a short summary ofthe water masses in the water column of the southwest Pacific, highlighting the variationsin δ13C with depth in this region (Figure 6.1 and 6.2).

The surface waters of the Coral Sea and Tasman Sea are dominated by the EastAustralian Current (EAC), sourced from the southern arm of the South Equatorial Current(SEC). Below this are the Subtropical Mode Waters (STMW), which form at the edge ofthe WBC by wintertime mixing and cooling prior to summer re-stratification. The AAIWdominates the depth range of 600 to 1100 mbsl. This is evident in the water column profileby the salinity minimum across this depth range. In the southwest Pacific there are twosources of AAIW (Chapter 4). These are evident from the variations in the salinity andδ13C in the north and south of the region (Figure 6.1). The major source of AAIW entersfrom the northeast through the Coral Sea part of the main subtropical gyre circulation(following the surface currents). This northeast AAIW source has lower salinity (Figure6.2) as it is ultimately sourced from the southeast Pacific, although it has experiencedsignificant mixing during its transport around the gyre. It also exhibits lower δ13C, as ithas accumulated organic matter during its transport around the south Pacific gyre. Thesouthern, more minor, source of AAIW, comes directly from the Southern Ocean (Rintouland Bullister , 1999). This is more saline as it is a combination of the highly saline watersfrom the central waters of the Indian Ocean and the Southern Ocean. It has slightly higherδ13C, than the northeast AAIW, primarily as it has not had the chance to accumulateorganic matter.

Below the intermediate waters sit the deep waters. These are considerably different inthe Coral Sea and Tasman Sea below 1500 mbsl as a result of the topographic restrictionsbetween the two basins (Figure 6.1, which allow only a minor transfer of water below this


Figure 6.1: Map of southwest Pacific with the location of GC-12 in the north Tasman Sea.

The location of the two water column data sets in Figure 6.2, are also highlighted. Grey square for

the waters entering the Coral Sea and black diamond for waters in the Tasman Sea. The general

circulation of the EAC (black line), AAIWN and AAIWS (grey dashed lines), and CPDW/AABW

(light grey dotted line), are also shown.

depth. In the Tasman Sea these deep waters are made up of the Circumpolar Deep Waters(CPDW) down to 2500 mbsl and below this the Antarctic Bottom Waters (AABW). Boththese deep water masses have low δ13C (Figure 6.2).

6.2.2 Stable Isotopes

The analysis of stable isotopes in multiple species of planktonic foraminifera has providedinsight into changes in the gradients of physical and biological properties in the watercolumn (Mulitza et al., 1997; Faul et al., 2000; Elderfield et al., 2002; Spero et al., 2003).Minor fluctuations of oxygen isotope ratios in foraminiferal calcite can be used to trackchanges in seawater temperature and salinity that can provide clues about ocean circula-tion and vertical mixing. Larger variations in oxygen isotope ratios (δ18O), however, arebrought about by changes in sea level and ice volume on glacial-interglacial timescales,


Figure 6.2: Temperature, salinity and δ13C data from two sites one in the Tasman Sea and the

other representing water entering the Coral Sea. See Figure 6.1, for location of sites. The water

masses, delineated according to the water properties, are shown on the right hand side. Surface

EAC, STMW, TW (thermocline waters), AAIW, CPDW and AABW.

and provide a approximate chronology for deep-sea cores when correlated to the globalSPECMAP compilations of benthic δ18O (Martinson et al., 1987).

Carbon isotope ratios are used to track changes in biological productivity and havebeen the traditional tool with which palaeoceanographers trace water masses (e.g. NADW,Duplessy et al. (1988)). The carbon isotopic composition of the dissolved total inorganiccarbon (TIC) in the ocean is influenced by three factors 1) air-sea exchange, 2) biologicalproductivity and 3) ocean circulation (discussed in Section 1.2). The TIC is used byforaminifera to secrete their tests, and the δ13C is incorporated into the calcite.

There are also other factors that alter the δ13C signal analysed from the calcite testsof foraminifera. Well-known biological vital effects (Urey , 1947) cause each foraminiferalspecies to fractionate carbon and oxygen isotopes differently as they form their tests.Numerous planktonic foraminiferal species also show significant differences in δ13C fordifferent size fractions (Berger et al., 1978; Oppo and Fairbanks, 1989). These isotopicoffsets probably reflect the different physiologies and feeding habits of foraminifera duringtheir life cycles (Hemleben et al., 1985). These biologically mediated isotopic offsets areassumed to be constant for a particular size fraction of foraminiferal species. Thereforecomparisons, which are restricted to one species and size-fraction, should reflect changes inthe δ13C of seawater through time. It is therefore important to use a similar size fractionfor each foraminifera species; preferably the largest adult tests which give a representativesignal of the complete life cycle. The only problem with this is that integrates all thevariations with depth, which can be considerable for deeper dwelling foraminfera.


Species δ18O Normalization to G. ruber (white) at 25◦C δ13C Normalization to TIC

G. ruber (white) [250-359 µm] 0 + 0.94

G. sacculifer [250-350 µm] - 0.11 + 0.12

Gr. menardii [600-850 µm] -0.13 0

Table 6.1: Stable isotope corrections for different species of planktonic foraminifera. Adapted

from Spero et al. (2003). No correction is available for Gr. truncatulinoides G.ruber - correction

from Bemis et al. (1998) O. universa high light equation, G.sacculifer - correction from Spero

et al. (2003) laboratory derived equations, Gr. menardii - correction from Mielke (2001) laboratory

derived equations.

δ13C and δ18O values in foraminiferal tests are also influenced by the carbonate ioneffect, whereby variations in the carbonate ion concentration [CO2−

3 ] alters isotope frac-tionation between seawater and carbonate (Spero et al., 1997)(refer to Appendix E). Theeffect of changes in [CO2−

3 ] on isotope fractionation is minor compared to the role of vitaleffects and, again, the effect on fractionation is species-specific in foraminifera. Numerousworkers have published empirical relationships to correct biologically-mediated offsets inisotope ratios between different planktonic species of foraminifera and the δ13C of TICin seawater (Bouvier-Soumagnac and Duplessy , 1985; Deuser , 1987; Niebler et al., 1999;Bemis et al., 1998; Spero et al., 2003). The isotope ratios for the planktonic species usedin this study are corrected according to the empirical values in Table 6.1.

6.2.3 Foraminifera

Previous workers have concentrated on the abundance and distribution of planktonic andbenthic foraminifera within the Tasman Sea. These studies have highlighted variations inprimary productivity (Thiede et al., 1997; Kawagata, 2001), SST changes temporally andspatially (CLIMAP , 1981; Anderson et al., 1989; Barrows et al., 2000) including shiftsin the various oceanic fronts in the region, such as the shift in the Tasman Front fromits present latitude of 33◦S to as far north as 26◦S during the LGM (Martinez , 1994;Kawagata, 2001)(this is discussed in Chapter 7). Suggestions have also been made thatintermediate water circulation (AAIW) may have varied on a glacial/interglacial cycle(Martinez , 1997; Correge and DeDeckker , 1997; Haddad et al., 1993).

In this study, a suite of foraminiferal species have been analysed for their stable iso-topes. These species are from various life cycles and ecology, and inhabit different watermasses. The stable isotope results are used to provide data to highlight changes in oceancirculation at different water depths throughout the ∼30 ka BP period of core FR1/97GC-12.

A brief description of the foraminifera used in this Chapter is highlighted in the listbelow. For more details about the foraminifera analysed refer to Appendix E.

Globigerinoides ruber is restricted to the upper 50 m of the water column.

Globigerinoides sacculifer lives predominantly in the mixed layer, within the upper25-75 m.


Globorotalia menardii appears to increase in abundance in upwelling tropical watersand is common in shallow thermocline waters where primary productivity is high.

Globorotalia truncatulinoides is the deepest dwelling planktonic foraminifera usedand although it appears in the upper few hundred metres it spends the majority ofits life cycle below 100 metres adding its final calcite crust when it reaches the 10◦Cisotherm corresponding approximately to 1000 mbsl.

Cibicidoides spp. are predominantly epifaunal benthic genera. δ13C of the genusis considered to reflect the δ13C of bottom water with relatively little fractionationrelative to equilibrium calcite. However the δ18O for Cibicidoides spp. is fractionatedby + 0.64 compared to δ18O seawater.


Marine core RV Franklin 1/97 GC-12 (23◦34.37S, 153◦46.94E) is from a water depth of990.5 mbsl, the core of the AAIW mass. Samples were extracted every 5 cm at the top ofthe core and every 10 cm below 1.5 m. The samples were sieved and picked for the fourspecies of planktonic foraminifera and benthic genera. These samples were run for stableisotopes (see Chapter 2 for more details).

6.4 Results

The δ18O values of the different foraminiferal species down the depth of the core showrelatively consistent inter-species offsets throughout the glacial and interglacial periods(Figure 6.3). This provides us with confidence that each individual species presumablyremained within the same water mass. Comparison of the δ18O values with SPECMAPMartinson et al. (1987) and the calibrated radiocarbon dates (Figure 6.3) provide us witha chronology for core GC-12 (refer to Chapter 5).

The results for Gr. menardii and Gr. truncatulinoides are sporadic due to the absenceof these species within some intervals of the core (Figure 6.4). Gr. menardii is foundthroughout the glacial but disappears during the early Holocene. Gr. truncatulinoides isabundant during the glacial, absent during the deglaciation, and then present again in theHolocene, although rare individuals are present in the Holocene (Figure 6.4).

The δ13C curves corrected according to Table 6.1, plotted against age, vary consider-ably throughout the last 30 kyr (Figure 6.5A). The δ13C values of the benthic Cibicidoidesspp. are the lowest, whilst those for the planktonic species are comparable, once correctedfor their vital effects (except Gr. truncatulinoides because no correction for vital effects isavailable). It is apparent from the δ13C curves that there are three main δ13C phases dur-ing the glacial/interglacial transition (Figures 6.5A and B). The three phases associatedwith major changes in the carbon isotope ratios are described below:


Figure 6.3: δ18O for a suite of foraminifera from core GC-12 against the depth in the core. The

foraminifera are labelled on the right hand side, with the δ18O for each foraminifera matching the

species depth range. Radiocarbon AMS dates are highlighted by red triangles. See Chapter 5 Table

5.2.

Figure 6.4: Abundance of Gr. menardii (dark grey) and Gr. truncatulinoides (light grey) in

core GC-12. Where Rare = 1-5 individual foraminifera seen per sample, Common = 5-20 individuals,

Abundant = >20 individuals.

6.4. Results 123

Figure 6.5: A) δ13C for a suite of foraminifera in core GC-12. The foraminifera have the

same symbols and colour as for Figure 6.3. B) ∆δ13Cplanktonic-benthic offset. ∆δ13CG.ruber-

Cibicidoidesspp. is shown by dark blue diamonds, ∆δ13CG.sacculifer-Cibicidoides spp. is shown in

pink squares, and ∆δ13CGr.menardii-Cibicidoides spp. is shown in blue triangles. Dashed black

line represents the average offset of G. ruber-Cibicidoides spp. and G. sacculifer-Cibicidoides. The

three phases of δ13C changes are highlighted.


1. Glacial: [∼30-18 ka BP] This phase shows large fluctuations of up to 0.7in the δ13C of each planktonic and benthic foraminifera species (Figure 6.5A). The∆δ13Cplanktonic-benthic offset averages ∼1.1 (ranging from 0.75 - 1.5 ) (Figure6.5B).

2. Deglaciation: [∼18-12 kyr)] This phase was initiated by a rapid ∼1 decreasein δ13C of the planktonic foraminifera and is associated with a synchronous in-crease in the δ13C of the benthic foraminifera. This resulted in a decrease in the∆δ13Cplanktonic-benthic offset to a minimum of 0.4 (Figure 6.5B). The δ13C ofthe foraminifera species, both planktonic and benthic, appears to be less variableduring this period (Figure 6.5A).

3. Present : [∼12 kyr to present] This phase is marked by an increase in bothplanktonic and benthic foraminiferal δ13C and a minor increase in ∆δ13Cplanktonic-benthic offset to 0.7 (Figure 6.5B).

6.5 Model

For each of the three phases above a possible ocean circulation scenario has been producedfor the Tasman Sea (N-S transect including the equator, Coral Sea, and Southern Ocean)to try to explain the changes in δ13C of the different foraminifera, and their associatedwater mass (Figure 6.6).

1. Phase 1 - Glacial (∼30-18 ka BP): The large ∆δ13Cplanktonic-benthic offsetof the foraminifera suggests that ocean circulation was slower, allowing nutrientstratification and a carbon isotope gradient to develop in the Tasman Sea. Theincreased abundance of Gr. inflata and G. bulloides during the LGM in the EasternEquatorial Pacific (EEP) has been suggested to represent relatively limited upwellingand than advection was the main oceanographic mechanism in the region (Martinezet al., 2003) and increased density stratification in the eastern Pacific (Loubere,1981, 2000). As a result the EEP would have been reduced in nutrients, whichare controlled by the Equatorial Undercurrent (EUC), but ultimately sourced fromthe Subantarctic Mode Water (SAMW) (Toggweiler et al., 1991). This would havelead to a reduction in the nutrient content of the SEC. Foraminiferal abundances anddistributions also indicate that the Tasman Front shifted northward to approximately26◦S during MIS 2/3 (Martinez , 1994). This suggests that the EAC surface layerwas considerably altered during the LGM (discussed in Chapter 7).

There are abundant Gr. menardii and Gr. truncatulinoides during the LGM, signi-fying that the thermocline was shallower (Andreasen and Ravelo, 1997) and that theintermediate waters were thicker and more dominant. Reduced δ13C of the benthicforaminifera may have been the result of a change in the intermediate water, possiblyrelated to a change in the main source of the AAIW in this region with a greater

6.5. Model 125

Figure 6.6: 2D model of changes in the AAIW circulation, through a transect from the Coral

Sea on the left hand side to the Southern Ocean on the right hand side. 1) Glacial [MIS 2/3],

2) Deglaciation [18 to 12 ka BP], and 3) Holocene [12 ka BP to present]. The star highlights the

approximate location of core GC-12. Water masses: EAC, AAIW, CPDW and AABW are all shown.

TF - Tasman Front, SAF - Subantarctic Front. Approximate δ13C for each of the water masses are

shown.


contribution from the south. Evidence from intermediate water benthic foraminiferaδ13C in the EEP exhibit an increase during the glacial (Mix et al., 1991). If thisrecord from core V19-27 in the intermediate depths of the EEP (Mix et al., 1991)is also subjected to the AAIW, then the FR97/1 GC-12 δ13C benthic results fromthis study must be controlled by an alternative AAIW source.

There are several potential explanations the decrease in δ13C in the AAIW from thesouth. There could be a change in the δ13C of the source waters, this could reflecta decrease in productivity, an increase in nutrient supply to the surface waters, ora reduction in air-sea exchange in the AASW. A more plausible explanation is thatthere was great mixing between the underlying CPDW and the AAIW, providingcool, nutrient rich waters at intermediate depths. Recent evidence from benthicforaminifera from DSDP 594 in the Bounty Trough to the east of New Zealand, sup-ports this upwelling scenario (Hayward et al., 2004). The benthic faunal abundancesdisplay an increase in species which require a high carbon flux and cold, low oxygenwaters during the glacials. This suggest that the oxygen minimum zone of the upperCPDW was considerably shallower during the LGM (Hayward et al., 2004).

The presence of Gr. menardii, which inhabits thermocline water, suggests a shallow-ing of the thermocline during the LGM, as is often associated with greater exportproduction caused by upwelling of nutrient rich, cold intermediate waters. Unfor-tunately there is no independent proxy to constrain the nutrient content in thethermocline waters at this time. However, a water column regime, where the major-ity of the productivity is within the stratified layers below the surface mixed layer,has been suggested to account for the high production of diatoms in the SouthernOcean and Mediterranean Sea (Kemp et al., 1999, 2000).

The observed glacial fluctuations in δ13C of 0.7 , which appear to be relativelycontemporaneous for different species of foraminifera, may be related to millennialchanges in deep water circulation during MIS 2/3 (e.g. Kanfoush et al. (2000);Pahnke et al. (2003)). However there may be problems with small sample size andhigher resolution sampling will be required to confirm this.

2. Phase 2 - Deglaciation (∼18-12 ka BP): Temperature increases in Antarctica(Petit et al., 1999) during deglaciation would have initiated the melting of permanentsea ice and the return of wind mixing in the surface waters of the Southern Oceanat ∼19 ka BP (Shemesh et al., 2002). This resulted in the upwelling of deep waters,Pacific Deep Waters and CPDW into the Antarctic surface waters and increased air-sea exchange, leading to the rapid release into the atmosphere of 13C-depleted CO2

stored within the deep waters (Jesse-Smith et al., 1999). Seasonal sea ice formationand brine rejection also increased AABW formation and the return to thermohalinecirculation regimes.

With the return to normal circulation within the south Pacific subtropical gyre, theAAIW switched its main source from the south of the Tasman Sea to the northeast

6.5. Model 127

in the Coral Sea. The shift in the δ18O of Cibicidoides spp. slightly leads the plank-tonic foraminiferal δ18O at the deglaciation. This change in the δ18O of Cibicidoidesspp. in the intermediate waters my indicate an initial rise in temperature, a switchin the source of AAIW and the reduction in cool upwelling from CPDW. The thick-ness and dominance of intermediate waters in the Tasman Sea were probably alsoreduced, which resulted in a decline in the abundance of Gr. truncatulinoides. TheAAIW, entering the Coral Sea in the northeast, forming in the southeast Pacific, isinfluenced by the upwelling of 13C-depleted deep waters (Haddad et al., 1993; Sikeset al., 2000). The 13C-depleted northeast AAIW, however, had higher δ13C thanthe glacial intermediate water in the Tasman Sea, thus increasing the intermediatewater benthic δ13C.

The sudden reduction in the ∆δ13Cplanktonic-benthic offset between the surfaceand intermediate waters at the initiation of the deglaciation is most likely causedby the rapid ventilation of intermediate waters affecting the tropics (Loubere, 2000;Spero and Lea, 2002). The upwelling of the AAIW into the equatorial regions influ-ences the δ13C signature of the EUC, which, in turn, feeds the SEC and ultimatelythe EAC. The surface currents would also have been influenced by air-sea exchangeof the newly released 13C-depleted CO2, which would have mixed rapidly into theatmosphere, thus allowing rapid transfer of the signal from the poles to the equator(Spero and Lea, 2002)). This ventilation event essentially homogenises the upperocean surface with intermediate water. The minor ∆δ13Cplanktonic-benthic off-set between the surface and intermediate waters resulted from the slight differencebetween productivity in the surface waters (higher δ13C), and respiration and rem-ineralisation of organic carbon in the intermediate waters (lower δ13C) during theircirculation around the subtropical gyre.

Another reason for the decrease in intermediate waters in the Tasman Sea duringdeglaciation is that the EAC probably strengthened and began to dominate the sur-face waters. This is caused by the development of SST gradients between the east-west equatorial Pacific (Lee and Slowey , 1999) and the initiation of a more modernconfiguration of the subtropical gyre circulation. The Tasman Front is eventuallydriven south from its glacial position ∼26◦S to 33◦S at the start of the Holocene(Chapter 7). The result of this increased dominance of the EAC is an increase in thedepth of the thermocline causing a reduction in the abundance of Gr. menardii (An-dreasen and Ravelo, 1997). No attempt was made to calculate depth of thermoclinevalues in this study because the only quantification of species was into categories ofabsent, rare, common and abundant.

The ongoing decline in Gr. menardii through the deglaciation suggests that changesin the surface waters were more gradual than changes in the intermediate waters,or that Gr. menardii may have changed its depth range to accommodate the watercolumn changes.


3. Phase 3 - Present (∼12 ka BP to 0): The initiation of this final phase was theincrease of both surface and intermediate δ13C signatures after 12 ka BP. There is aslight increase in the ∆δ13Cplanktonic-benthic offset between surface and interme-diate water depths from its minimum of ∼0.4 during the initial deglaciation to∼0.7 in the Holocene, similar to that seen in the modern water profiles (Figure6.2). This change heralds the onset of modern oceanic circulation in the region.

Similar to the deglaciation phase, the small offset in ∆δ13Cplanktonic-benthic be-tween the surface and intermediate waters is because the AAIW contributed to up-welling in the tropics, and therefore influences the nutrient and chemical signaturesof the surface waters as well as the intermediate waters. The rise in both surface andintermediate water δ13C after 12-13 ka BP is more difficult to interpret, but maybe related to the strengthening of the NADW at around 14 to 15 Ka radiocarbonyears (Charles and Fairbanks, 1992; Spero and Lea, 2002). Therefore, it may be thefinal piece in the global modern thermohaline circulation puzzle, with the initiationof present-day deep water circulation and NADW contributing 13C-enriched waterto the upwelling CPDW (Duplessy et al., 1988).

An alternative interpretation of the rise in surface-ocean δ13C is that it is relatedto the uptake of isotopically light atmospheric carbon by the expanding terrestrialbiosphere during the deglaciation (Maslin et al., 1995). Jesse-Smith et al. (1999)discuss the post-glacial rise in atmospheric pCO2 and δ13CCO2 (Figure 6.7. Theysuggest that because the increase in δ13CCO2 outpaced the rise in SSTs the δ13C

of the additional flux of CO2 must have been higher than that of the pre-existingatmosphere, thus eliminated the terrestrial biosphere as a primary source of CO2.

At present, the EAC is a deep mixed, oligotrophic surface layer in the Tasman Seadominated by G. ruber and G. sacculifer. The absence of Gr. menardii indicatesthat, as well as having low nutrients, the thermocline was significantly deeper duringthe Holocene compared to the glacial and deglacial periods (Andreasen and Ravelo,1997). AAIW is the main intermediate water presently found between 700 and 1100mbsl, and is primarily sourced from the northeast through the Coral Sea at the endof the south Pacific subtropical gyre. The intermediate waters are also relativelyreduced in thickness compared to the glacial period, as indicated by the absenceor rarity of Gr. truncatulinoides in the southwest Pacific during the Holocene (asnoted by Martinez (1994, 1997). Gr. truncatulinoides is highest in the east Pacificwhere the AAIW is 600 m thick, compared to 300-400 m in the north Tasman Sea,and therefore it has been proposed to correlate with the thickness of the SAMW andAAIW (Lohmann and Schweitzer , 1990; Martinez , 1997).

6.5. Model 129

Figure 6.7: Summary of the correlation and timing of the model phases 1-3 in foraminiferal δ13C

with; atmospheric pCO2; and δ13CCO2 measured in ice cores. Planktonic and benthic foraminiferal

δ13C (bottom) is represented by G. sacculifer and Cibicidoides spp., respectively. The atmospheric

pCO2 (ppmv) (middle) is shown for the Taylor Dome, Byrd and Vostok ice cores from Antarctica.

The δ13CCO2 (top) was measured in the Taylor Dome core (Jesse-Smith et al., 1999).


6.6 Discussion

Changes in ventilation rates of the deep and intermediate oceans have been invoked asa driver of the rapid fluctuations in atmospheric CO2 that are evident in ice cores overglacial/interglacial cycles (Barnola et al., 1987). The oceans are the largest exchangeablecarbon reservoir on Earth (50 times greater than the atmosphere, Broecker et al. (1980))and changing gradients in δ13C between the surface, intermediate and deep waters reflectchanges in the total carbon reservoir, carbon storage in different water masses, and changesin ocean circulation and ventilation (Boyle, 1986).

Present day deep-water ∆14C ventilation ages were originally collected as part of theGEOSECS program (Stuiver and Ostlund , 1980; Ostlund and Stuiver , 1980) providing theevidence for the “thermohaline circulation” (Broecker and Peng , 1982). In the Pacific sev-eral studies from deep-sea cores using 14C data from planktonic and benthic foraminiferahave shown that ventilation rates of the deep oceans were much slower during the glacial(Sikes et al., 2000), whilst during the deglacial ventilation rates were faster (<500 yrs)than at present (Shackleton et al., 1988). This pattern of change in global deep-waterventilation and variations in the thermohaline circulation during the last deglaciation isgenerally accepted, although there may be local variations (e.g. Hall et al. (2001)).

Intermediate waters, however, also play a major role in the global thermohaline cir-culation (Schmitz , 1995; Saenko et al., 2003), and in the case of the AAIW these watersare exported northwards and ventilate the equatorial thermocline of the Pacific (Oppo andFairbanks, 1989; Toggweiler et al., 1991). The importance of the intermediate waters isbeginning to be understood and explored in a range of recent paleoceanographic studies.

The δ13C minimum evident in planktonic foraminifera during the last deglaciationis coeval (within the errors of radiocarbon chronology) over an extensive region, includ-ing the EEP (Shackleton et al., 1983; Spero and Lea, 2002; Spero et al., 2003), Coral Sea(Haddad et al., 1993), Tasman Sea (this study), the Southern Ocean (Lynch-Stieglitz et al.,1994; Ninneman and Charles, 1997), Caribbean (Oppo and Fairbanks, 1989; Haddad andDroxler , 1996), and western equatorial Atlantic (Curry and Crowley , 1987; Mulitza et al.,1998). Each of the records shows a similar 0.5-1 decrease in δ13C spanning the Sub-antarctic to the tropics for a range of foraminifera species, which show different seasonalabundances. The result has prompted several workers to suggest that the decrease in theδ13C is a signal spread by the AAIW (Ninneman and Charles, 1997; Spero and Lea, 2002;Spero et al., 2003). It has also been proposed that it is a preformed signal in the surfacewaters of the Southern Ocean, as there is no major increase in productivity accompany-ing the δ13C minimum (Oppo and Fairbanks, 1989; Spero and Lea, 2002). How the δ13C

signal is formed is still a matter of considerable debate, but it may be caused by changesin air-sea gas-exchange at the circumpolar surface oceans with the return of wind mixingduring the deglaciation (Ninneman and Charles, 1997).

The planktonic foraminiferal δ13C decrease also appears to have occurred concurrentlywith the rapid rise of atmospheric CO2 concentrations and decrease in the δ13C of CO2

6.7. Summary and Conclusions 131

observed in ice core records from Antarctica (Jesse-Smith et al., 1999) (Figure 6.7). Speroand Lea (2002) suggested that the rapid ventilation of Southern Ocean surface watersduring the initiation of deglaciation released a large volume of 13C-depleted CO2, whichhad built up in the deep ocean throughout the glacial. The gradual δ13C increase in theatmosphere and surface oceans after the δ13C minimum ∼12 kyr ago is probably relatedto the strengthening of the NADW at this time (Spero and Lea, 2002). The feedback ofwarming caused by the release of CO2 into the atmosphere could have contributed to themelting of the Northern Hemisphere ice-sheets and helped the initiation of NADW andthe return to interglacial conditions.

Despite the synchronous decrease of planktonic foraminiferal δ13C and atmosphericδ13C during the deglaciation, there seems to be little agreement between intermediatewater benthic foraminiferal δ13C curves across the same region. In the EEP, intermediatewater benthic foraminifera exhibit a δ13C minimum spike during the deglaciation (Mixet al., 1991). In the southwest Pacific and Coral Sea, the δ13C of benthic foraminifera inintermediate water sites 817A and 818B (ODP leg 133) is higher than that for planktonicforaminifera, but both are lowest during glacial/interglacial transitions (Haddad et al.,1993). In this study, the δ13C of benthic foraminifera increases during the deglacial.More studies of benthic records from intermediate depth marine cores are required tounderstand the changes in the glacial/interglacial AAIW circulation and its distributionin the Southern Hemisphere, especially within the south Pacific where a series of basinsand the variable bathymetry create a more complex distribution and circulation patternas suggested in Chapter 4. The distribution and influence of AAIW circulation may wellhave been considerably different during the LGM as suggested by studies in the Caribbeanand Atlantic (Haddad and Droxler , 1996).

6.7 Summary and Conclusions

Changes in δ13C in benthic foraminifera and a range of planktonic foraminifera in theTasman Sea indicate that variations in the dominance and direction of flow of the AAIWand surface waters occurred during the last deglaciation.

The results of this study suggest that, during the glacial, the main flow of the AAIWmay have been from the south, through the Tasman Sea. The AAIW were thicker andprobably subject to considerable mixing with nutrient rich deep water. As a result, thethermocline was shallower due to upwelling caused by the dominance of this intermediatewater mass. Overlying the intermediate water was a relatively thin, low nutrient capdue to a weakening of the EAC. This difference between the oligotrophic surface waterand the high nutrient intermediate waters formed a large surface to intermediate water∆δ13Cplanktonic-benthic offset of ∼1 .

With the initiation of the deglaciation, there was a rapid decrease in δ13C in theplanktonic foraminifera, accompanied by an increase in benthic foraminiferal δ13C, thusreducing the ∆δ13Cplanktonic-benthic surface to intermediate water offset to a minimum


of 0.4 . This supports the hypothesis that this is the result of rapid ventilation of inter-mediate waters in the south Pacific, which evidence suggests caused a switch in the AAIWsource from the south to the northeast in the Tasman Sea.

Due to the wide geographical extent of the δ13C minimum in planktonic foraminifera,which spans the subantarctic to the Southern Hemisphere tropics, it has been suggestedthat the AAIW was the main conduit for the transport of the preformed δ13C signalto the surface waters of this region (Oppo and Fairbanks, 1989; Ninneman and Charles,1997; Spero et al., 2003; Spero and Lea, 2002). Exactly how this preformed δ13C signal isproduced without an equivalent increase in productivity is still a quandary; one plausibleexplanation is a change in air-sea exchange in the surface waters of the Southern Ocean(Ninneman and Charles, 1997).

The δ13C of AAIW starts to increase again after 12-13 kyr, possibly as a result of thestrengthening of the NADW production, and its renewed influence on the global thermo-haline circulation (Spero and Lea, 2002). This supports proposals that intermediate waterplays an important role in the propagation of oceanic changes from the polar regions tothe surface waters of the tropics.

Such a connection may well have served to transfer climate variations in the SouthernHemisphere to trigger the initiation of the Northern Hemisphere ice sheet collapse andthe commencement of increased NADW formation during the deglaciation. This dataalso supports the notion of a more stagnant deep ocean during glacials, which providesa vast 13C-depleted carbon reservoir and source of atmospheric CO2. Given this oceanicsetting, rapid ocean overturning and ventilation events at glacial terminations could releasesignificant quantities 13C-depleted CO2 to the atmosphere.

Chapter 4 has highlighted that the modern EEP is not influenced by the same inter-mediate waters as the Tasman Sea and Coral Sea, and therefore a range of cores fromaround the south Pacific gyre is required to determine the distribution and circulationof the AAIW during the LGM. Combining the δ13C data with other water mass tracerse.g. Cd/Ca in foraminifera, may help extricate the ocean circulation information from theother factors that affect δ13C (Section 1.2).

Chapter 7

Changes in the East Australian

Current

Submitted as a paper entitled “Glacial/Interglacial changes in the East Australian Cur-rent” to Climate Dynamics in October 2004, co-authored by Opdyke, B. N., Gagan, M.K., Kiss, A. E. and Fifield, L. K.

7.1 Abstract

The East Australian Current (EAC) is the western boundary current of the south Pacificgyre transporting warm tropical waters to higher southern latitudes. Recent modellingshows that the partial separation of the EAC (∼32◦S) and the coupled formation of theTasman Front (∼34◦S) are caused by a steep gradient in the zonally integrated windstress curl. Analysis of oxygen isotope ratios (δ18O) in the planktonic foraminifer, Glo-bigerinoides ruber, from sediment cores from the Coral Sea and Tasman Sea indicatesthat the EAC separation shifted northward to between 23◦S and 26◦S during the lastglacial. These results suggest a significant change in the Pacific wind stress curl duringthe glacial. Given recent evidence for El Nino-like conditions in the Pacific during thelast glacial, with a reduction in the east-west sea surface temperature (SST) gradient, wesuggest that weaker trade winds combined with more northerly, stronger westerlies wereassociated with a change to the wind stress curl, which repositioned the EAC separationand Tasman Front. In contrast, by ∼11 ka BP the EAC separation was forced south of26◦S. This southward shift was synchronous with a rapid warming of tropical SSTs, andthe onset of a La Nina-like SST configuration across the tropical Pacific. It appears thatthe south Pacific trade winds strengthened accordingly, causing the EAC to readjust itsflow. This readjustment of the EAC marks the onset of modern surface-ocean circulationin the southwest Pacific, but the present EAC transport was only achieved in the lateHolocene, after 5 ka BP.

133

134 7. Changes in the East Australian Current

7.2 Introduction

Western boundary currents are essential components of ocean circulation. In subtropicalgyres, these fast, narrow, deep currents transport warm equatorial waters to higher lati-tudes and are important for poleward heat transfer associated with large ocean-atmosphereheat flux. Warm western boundary currents are also vital to the modulation of coastalclimates on adjacent landmasses. The classic example of this is the north Atlantic GulfStream, which allows Britain and coastal northern Europe to maintain a relatively mild,warm and wet climate compared to other areas at similar latitudes. The east coast of Aus-tralia and islands of New Zealand experience a similar moderation of climate as a result ofthe East Australian Current (EAC) (Sprintall et al., 1995). The EAC flows south along theeast coast of Australia and separates at ∼32◦S to form the Tasman Front (33-35◦S) whichheads across the Tasman Sea towards the north of New Zealand. Surface currents in theocean are primarily driven by sea-surface temperature (SST) gradients, ocean-atmospherefeedbacks and basin-scale wind systems. Today, these feedbacks between the ocean andatmosphere are clearly observed during El Nino events and the seasonal transition of themonsoons. During the glacial the atmospheric circulation was considerably different asa result of the effects of large ice sheets over the Northern Hemisphere landmasses (e.g.Marshall and Clarke (1999)) and increased sea ice around Antarctica (Armand , 2000).Changes in SST gradients may have occurred in the tropical Pacific Ocean during theglacial (Lea et al., 2000; Koutavas et al., 2002; Stott et al., 2004; Koutavas and Lynch-Stieglitz , 2003). Understanding past changes in ocean circulation and their impact onatmospheric circulation in the tropics provides clues about ocean-climate feedback sys-tems (Matsumoto and Lynch-Stieglitz , 2003). Such information provides insight into theprocesses involved in future changes in the ocean circulation and climate.

7.2.1 The EAC and Western Boundary Current Dynamics

The EAC is the western boundary current of the south Pacific subtropical gyre and is theresult of the bifurcation of the southern arm of the South Equatorial Current between15◦S and 20◦S when it collides with the Queensland Plateau (Figure 7.1) (Wyrtki , 1962;Church, 1987; Ridgway and Dunn, 2003). A large proportion of the flow is trapped againstthe Queensland and Papuan boundaries and directed northward to form the New GuineaCoastal Undercurrent (Andrews and Clegg , 1989), whilst the rest flows south to form theEAC observed along the east coast of Australia from 18 to 35◦S (Boland and Church,1981).

After formation, the EAC intensifies, accelerates, and deepens, following the coastlineuntil ∼32◦S, where the main flow of the current (primarily the upper layer) separatesand flows east across the Tasman Sea to form the Tasman Front (Andrews et al., 1980).The Tasman Front is the main outflow of the EAC western boundary current. The frontoccurs as a series of jets, which follow different paths across the Tasman basin, with themain component of the eastward flowing Tasman Front occuring between 33◦S and 35◦S


(Ridgway and Dunn, 2003). A more minor flow and the deeper layers of the EAC continuesouth along the coast (Figure 7) (Wyrtki , 1962; Ridgway and Godfrey , 1994).

The exact reasons for western boundary current separation are still not well understoodand at least four theories have been put forward (see review by Dengg et al. (1996)):response to a change in wind stress curl (Munk , 1950), collision with another westernboundary current (Cessi , 1991; Agra and Nof , 1993), “crises” in the angular momentumbudget (Kiss, 2002) and interaction with bottom topography or boundary shape (Pedlosky ,1965). Unlike other western boundary currents, such as the Gulf Stream and the KuroshioCurrent, the EAC is significantly smaller and more variable. The EAC separation pointwas originally suggested to be 32◦16S, at Sugarloaf Point, where there is a change in theorientation of the east Australian coastline (Godfrey et al., 1980). Recent work has nowshown that EAC separation can occur over a range of 150 km (Ridgway and Dunn, 2003).

Model studies indicate that the EAC separation point and formation of the eastwardflowing Tasman Front can be captured by fairly simple, linear, wind-driven models usingthe dynamics of Sverdrup (1947); Stommel (1948); Munk (1950) with globally appliedhorizontal friction (De Szoeke, 1987; Godfrey , 1989; Ridgway and Godfrey , 1994; Tilburget al., 2001). Tilburg et al. (2001) employed Munk’s theory to explain the latitude atwhich the EAC separates from the east Australian margin to form the Tasman Front.Tilburg et al. (2001) hypothesise that the latitude of the Tasman Front is determined by amaximum in the second meridional derivative of the zonal wind stress (see the appendixfor theoretical background).

Western boundary currents must separate completely from the coast at the latitude ofzero wind stress curl because there is no meridional transport at these latitudes. Figure 7.2(adapted from Tilburg et al. (2001)), however, shows that wind stress curl is not zero at thelatitude of EAC separation. Thus, the EAC does not separate completely from the coastand some southward EAC transport is needed to compensate the net northward Sverdrupflow. Most of the EAC does leave the coast at this latitude because the meridional gradientof the wind stress curl is large, yielding a large eastward flow velocity. This EAC outflowforms the Tasman Front and reduces the remaining unseparated EAC transport to matchthe much smaller Sverdup transport further south in the region of weak wind stress curl.It has previously been suggested that the presence of New Zealand affects the separation ofthe EAC and the latitude of the Tasman Front due to its blocking of westward propagatingRossby waves south of 34◦S (Warren, 1970). Tilburg et al. (2001) repeated their models,removing the New Zealand landmass and observed no change in the EAC separation point,or the location of the Tasman Front.

The real EAC separation, however, is much more complex than these simple linearwind-driven models suggest. The EAC continues south even after detachment from thecoast, most probably as a result of the coastline sloping away from the separation point (Ouand de Ruijter , 1986), and then retroflects northward before forming the Tasman Front.The Tasman Front exhibits several complex meanders (Stanton, 1981; Mulhearn, 1987;Hamilton, 1992; Ridgway and Dunn, 2003) these are reproduced using models with real-


Figure 7.1: Map of the East Australian Current (EAC) circulation (black line) in the Coral Sea

and Tasman Sea, showing the complex bathymetry of the southwest Pacific Ocean. Bathymetric

contours are drawn at 100 m, 1500 m and 3000 m below sea level. The EAC is formed by the

bifurcation of the South Equatorial Current (SEC) as it collides with the Queensland Plateau. The

EAC separates from the east coast of Australia ∼32◦S and flows eastward as the Tasman Front

in a series of meanders associated with the north/south bathymetry of the Lord Howe Rise, New

Caledonia Trough and the Norfolk Ridge. A more minor flow and the deeper layers of the EAC

continue south along with several eddies, shown by a narrower arrow. Bold arrows mark the average

positions of the southeast trade winds and the mid-latitude westerlies. Circles mark the locations

of the sediment cores: PC-27a, GC-12, GC-25 and GC-9, with the dotted boxes highlighting the

location of the suite of cores from Table 7.3 and Figure 7.4.


Figure 7.2: Simplified diagram illustrating the effect of wind stress curl τ on the surface currents

of the Coral Sea and Tasman Sea. Where the latitude of curl τ = 0 is the boundary of the gyre

because the flow is purely E-W e.g. the SEC. Because the curl τ does not reach 0 at any other

point there is always some flow south along the western boundary, hence, the EAC exhibits only

partial separation and the deeper layers continue south. At the minima and maxima of curl τ , the

curve has no gradient, therefore there is no E-W zonal flow and currents flow only N-S (north in the

Tasman Sea due to the positive curl). Flow velocity is greatest when curl τ is at a maximum point,

and least at the minimum. The maximum gradient in the integrated curl τ is seen at ∼34◦S in the

Tasman Sea, this coincides with the separation point of the EAC and location of the Tasman Front.

This maximum in the derivative of curl τ causes the easterly flow across the Tasman, with only a

minimal northwards flow because the curl τ itself is small. The real situation is, however, affected

by numerous other factors including complex bathymetry.


istic bottom topography and non-linear dynamics, indicating the importance of couplingbetween the upper ocean and bottom topography (Tilburg et al., 2001). The bathymetryof the southwest Pacific Ocean is complex and includes a prominent series of roughlynorth/south-orientated ridges that radiate northward from New Zealand (Figure 7.1).The Tasman Front flows east between 33◦S and 35◦S and the meanders veer northwardas they hit the relatively shallow bathymetry of these ridges, then southward as the slopedescends into the adjacent basin (Ridgway and Dunn, 2003). Therefore, the present-daycirculation of the EAC and location of the Tasman Front is primarily controlled by theresponse of the ocean to the wind stress curl, but the meanders and eddies are stronglyinfluenced by the complexities of the basin bathymetry.

Based on recent palaeoceanographic studies in this region, we suggest that the windstress curl may have differed considerably during the last glacial and deglaciation, thusaltering the flow of the EAC.

7.2.2 Previous Research

Previous researchers have also suggested that the EAC was considerably different dur-ing the glacial. CLIMAP (1981) reconstructed a warm SST anomaly off the east coastof Australia and suggested that the EAC was even stronger during the glacial than atpresent. However, subsequent work based on the abundances of both planktonic and ben-thic species of foraminifera, from cores in the Tasman Sea and Coral Sea, indicates coolerSSTs (Anderson et al., 1989; Barrows et al., 2000) and several researchers have suggesteda dramatic northward shift of the Tasman Front from its present latitude of 33◦S to ∼25◦Sduring the glacial (Martinez , 1994; Kawagata, 2001; Martinez et al., 2002). The majorityof the cores analysed in these previous studies were from the Lord Howe Rise and had slowsedimentation rates (<2 cm/ka), therefore, the detailed history of EAC behaviour sincethe last glacial has been difficult to elucidate.


Marine core RV Franklin 1/97 GC-12 (23◦34.37S, 153◦46.94E) is a 546 cm core froma water depth of 990.5 mbsl in the Capricorn Channel off the southern Great BarrierReef. Core GC-12 is compared with 3 other recently published/unpublished cores fromthe east Australian margin (Table 7.1; Figure 7.1). These cores were chosen for theirlatitudinal range and their well dated chronologies from radiocarbon dating (Table 7.2).These hemipelagic cores also have relatively high sedimentation rates (greater than 10cm/kyr, except for GC-9 which averages ∼2.5 cm/kyr). Further comparisons were alsomade with a suite of cores from the Australian margin, which are not so well dated andhave wide range of sedimentation rates (Table 7.3).

The average δ18O for G. ruber in cores GC-25 and GC-9 were previously published inTroedson (1997) and the data are used here with the permission of Troedson (pers comm.).


Core Depth

(mbsl)

Latitude Longitude Reference

RV Franklin 5/90 PC-27a 2163 15◦17’S 145◦ 57’E Dunbar and Dickens (2003b)

RV Franklin 1/97 GC-12 990.5 23◦34’S 153◦13’E this thesis and Bostock et al.

(2004)

RV Rig Seismic 105 GC-25 1022 26◦35’S 153◦51’E Tsuji et al. (1997); Troedson

(1997); Troedson and Davies

(2001)

RV Rig Seismic 112 GC-9 1467 33◦57’S 151◦55’E Troedson (1997); Troedson and

Davies (2001)

Table 7.1: List of cores used in this Chapter.

Lab Code Core Depth

down

core

(cm)

Conventional

radiocarbon

age (yr)

Error in radio-

carbon age (±yr)

Calibrated cal-

endar age (yr

BP)

Error in calen-

dar age (± yr

BP)c

OZ G450 PC27a 7 1960 40 1510a 50

OZ G451 52 4760 50 4970a 90

OZ D814 82 7160 90 7610a 70

OZ D815 156 9320 90 9860a 230

OZ D816 200 11250 160 12870a 170

OZ D817 224 12900 150 14340a 580

OZ G449 318 22150 150

ANUA 20316 GC-12 20 4380 190 4540a 250

ANUA 20313 40 7930 210 8380a 230

ANUA 20319 80 10590 220 11720a 400

ANUA 20320 90 11910 350 13425a 400

ANUA 20321 150 14110 480 16330a 640

ANUA 22216 190 16665 250 19270a 410

ANUA 22218 250 21880 260 25200b 300

ANUA 22217 300 23930 300 27580b 350

ANUA 22219 350 30290 480 34870b 550

NZA 3574 GC-25 30 2070 60 1630a 70

NZA 3575 120 5850 70 6245a 80

NZA 3576 180 9200 75 9830a 200

NZA 3577 270 11900 80 13470a 300

NZA 3578 300 13460 120 15600a 240

NZA 3588 350 18680 180 21590a 390

OZ 272 GC-9 34 13450 90 15570a 220

Oz273 50 22150 150

Table 7.2: Radiocarbon ages and calibrated ages for cores PC-27a, GC-12, GC-25 and GC-9.a Calibrated age using CALIB 4.3 and (Druffel and Griffin, 1999) reservoir age of 35010 years b

Calibrated age using the polynomial fit for the data set from Bard (1998) c Error in calibrated age

is the maximum one sigma range for the calibrated radiocarbon date, determined from CALIB 4.3.


Core Latitude◦S

Depth

(mbsl)

LGM Holocene ∆δ18OLGM−Hol

FR90/5 PC-22a 15.02 1907 -0.9 -2.5 1.6

FR90/5 PC-27aa 15.17 2163 -0.8 -2.4 1.6

FR90/5 PC-29a 15.42 1834 -0.8 -2.4 1.7

FR92/4 PC-36a 15.82 1824 -0.8 -2.3 1.5

FR92/4 PC-35a 16.00 1857 -0.5 -1.9 1.4

FR92/4 PC-39a 16.37 1334 -0.6 -2.2 1.6

FR92/4 PC-42a 16.65 1389 -0.7 -2.2 1.5

FR92/4 PC-14a 16.72 1574 -0.4 -2.3 1.9

FR92/4 PC-16a 17.43 1043 -0.5 -2.3 1.8

FR92/4 PC-6a 17.53 1193 -0.4 -2.2 1.8

RS105 GC-36b 23.05 251 -0.2 -1.8 1.6

JK PC-2b 23.53 206 -0.4 -2.0 1.6

JK PC-4b 23.55 271 -0.7 -1.9 1.2

FR97/1 GC-09d 23.88 166 -1.0 -2.0 1.0

FR97/1 GC-10d 22.98 335 -1.0 -2.0 1.0

FR97/1 GC-11d 23.38 502 -0.9 -2.2 1.3

FR97/1 GC-12d 23.57 991 -0.5 -2.0 1.5

FR97/1 GC-13d 23.78 1482 -0.4 -1.9 1.5

FR97/1 GC-14d 23.82 2004 -0.3 -2.2 1.9

RS105 GC-22b 26.48 621 -0.1 -1.9 1.8

RS105 GC-23c 26.05 842 -0.2 -2.2 2.0

RS105 GC-25c 26.58 1022 0.3 -1.9 2.2

RS112 GC-09c 33.95 1467 0.4 -1.3 1.7

RS112 GC-10c 33.98 2007 0.6 -1.3 1.9

FR91/5 14PC-1b 34.12 2373 0.8 -0.9 1.7

Table 7.3: Summary of Holocene and glacial δ18O and ∆δ18Oglacial−Holocene for G. ruber in

cores from the Coral Sea and Tasman Sea. a - Dunbar et al. (2000), b - Troedson (1997), c - Troedson

and Davies (2001), d this thesis.

7.4. Results 141

Figure 7.3: δ18O G. ruber from cores PC-27a, GC-12, GC-25 and GC-9 plotted against calendar

age. PC-27a (Dunbar and Dickens, 2003b; Page et al., 2003) is shown by black dotted line with

diamonds, GC-9 (Troedson, 1997; Troedson and Davies, 2001) is shown by a grey dot-dash line

with crosses. Thick lines are used for GC-12 (grey) and GC-25 (Tsuji et al., 1997; Troedson, 1997;

Troedson and Davies, 2001) (black) are 3-point running averages of all the data (thin lines) for each

core. There is a large difference between the δ18O of G. ruber in GC-12 and GC-25 during the glacial

(highlighted in grey) compared to present, the records converge between 12-11 ka BP (highlighted

in grey), with a final convergence to the present day offset at 5ka.

Data for several of the cores displayed in Table 7.3 are also unpublished, as summarisedin Troedson (1997).

7.4 Results

The core tops from all of the cores are missing, however the Holocene values from allof the cores (Figure 7.3) show a 1.5 range which corresponds to a temperature offsetof 7.5◦C between 33◦S offshore Sydney (GC-9) and 15◦S offshore Cooktown (PC-27a).This compares with a 6-7◦C offset of mean modern SSTs from the two locations. Theδ18O G. ruber in cores GC-12 and GC-25 display a very small offset (0.2 ) during thelatest Holocene. This is to be expected, as average SSTs differ by ∼1◦C between the twolocations throughout the year. Average sea surface-salinity (SSS) is also similar for theGC-12 and GC-25 sites, averaging 35.4 ± 0.2 .

The δ18O of G. ruber in core GC-12 decreases by ∼1.5 between the last glacial (themaximum in δ18O) and the Holocene (Figure 7.3). The last glacial to Holocene change ofglobal ocean seawater δ18O is approximately 1 ± 0.1 (Schrag et al., 2002). Assumingno other significant salinity changes in the southwest Pacific, the additional 0.5 shiftin the δ18O of G. ruber must reflect a 2-3◦C cooling of SSTs during the glacial at 23◦S(using -0.2 ◦C for G. ruber, Duplessy et al. (1981)). This is in good agreement withthe foraminiferal Modern Analog Technique SST estimates from the southwest Pacific


(Anderson et al., 1989; Barrows et al., 2000). PC-27a and GC-12 are also offset by 0.2-0.3 during the glacial, similar to the Holocene offset. This suggests that these two coresare experiencing the same water mass throughout the last glacial to Holocene period.

In contrast, the δ18O of G. ruber in GC-25 decreases by ∼2.2 between the glacialand the Holocene, which is ∼0.7 greater than the shift observed in GC-12 (Figure 3).Even after accounting for the ∼1.0 shift in seawater δ18O resulting from ice volume,assuming there are no variations in salinity, the remaining ∼1.2 shift equates to ∼6◦Ccooling. During the last glacial GC-25 displays more similar values to GC-9 in the southand therefore we suggest experiencing similar water mass conditions to this more southerlycore.

The GC-12 and GC-25 δ18O G. ruber records rapidly converge between 12 and 11 kaBP (Figure 7.3). The δ18O values for the two records are offset by 0.8-1.0 during theearly deglaciation, but converge to an averge offset of ∼0.4 at ∼11 ka BP. The averageδ18O curves remain offset by ∼0.4 from 11 ka BP to 5 ka BP when they converge to anoffset of ∼0.2 This minor offset continues from 5 ka BP to the present, with some of thedata overlapping between the two core sites. During the deglaciation GC-25 diverges fromGC-9 around 10 ka BP, although the resolution on GC-9 is low and therefore difficult toascertain the exact timing of this event. GC-12 also diverges slightly from PC-27a around12 ka BP, again limited data make this difficult to interpret.

Table 7.3 lists the last glacial and Holocene δ18O values of G. ruber for a suite ofrecently published and unpublished records for 25 cores collected along the continentalmargin of east Australia. The cores are grouped according to location: ∼15-17◦S inthe Queensland Trough (Dunbar et al., 2000; Dunbar and Dickens, 2003b; Page et al.,2003); ∼23◦S in the Capricorn Channel, southern Great Barrier Reef (this study); ∼26◦Soffshore southern Queensland (Troedson, 1997); and ∼33◦S offshore Sydney, New SouthWales (Troedson, 1997).

Given the uncertainties associated with the lack of age constraint on many of thesecores and the potential for hiatuses along with wide variations in sedimentation rates,the maximum δ18O value in each record was chosen for the last glacial value, whilst theminimum or core top δ18O value was used for the Holocene. The last glacial and Holoceneδ18O values are plotted as a function of latitude in Figure 7.4. An average for eachlatitudinal cluster of cores is plotted and joined by a line to give an estimate of the G.ruber δ18O gradient with latitude. The spread of the data in each of the clusters couldbe the result of several factors, including the different size fractions used, seasonality anddiagenesis. Despite the scatter in the data from this suite of cores we believe it is importantto highlight the wide range of results for the Australian margin and that our conclusionswere reached through the analyses of δ18O in G. ruber data from numerous cores.

7.4. Results 143

Figure 7.4: Summary of δ18O G. ruber as a function of latitude for a suite of cores in the Coral

Sea and Tasman Sea (See Table 7.3). A) δ18O G. ruber for the Holocene (black dots), and last

glacial (grey crosses). The thick lines join the average δ18O G. ruber value for each cluster of cores.

B) The glacial-Holocene difference in δ18O (∆δ18Oglacial-Holocene) of G. ruber (black triangles).

Large black squares highlight the location of GC-12 and GC-25 in each of the plots.


7.5 Discussion

The difference in the reconstructed SSTs from the ◦ G. ruber data (Figure 7.3) for coresGC-12 and GC-25 indicate ∼3-4◦C temperature change across ∼3◦ of latitude during thelast glacial. Such a steep north-south SST gradient is evident in satellite images of SSTacross the Tasman Front today, and supports foraminiferal evidence for the repositioningof the EAC separation and the Tasman Front north of 26◦S during the glacial (Martinez ,1994; Kawagata, 2001; Martinez et al., 2002).

The δ18O G. ruber, however, is not only affected by ice volume and temperature.Variations in the size fraction used to measure δ18O of G. ruber may have had someaffect on the values for each core (e.g. for PC-27a δ18O G. ruber was measured on thesize fraction 250-355µm). The seasonality of the G. ruber may also influence the stableisotopes. In the Coral Sea and northern Tasman Sea G. ruber are present all year round.The presence of G. ruber in GC-9, however, is likely to be seasonal as the present EAConly reaches this far south during the Austral summer and autumn. Although G. ruberdoes have a large temperature and salinity tolerance it is surprising that they are found inGC-9 during the last glacial. This implies a small EAC current continued to reach 33◦Sthroughout the glacial period. It has been suggested that G. ruber may be transportedfrom their tropical/subtropical habitat south in the discrete eddies which separate from theEAC at the Tasman Front and continue south (Nelson et al., 1993; Tomczak and Godfrey ,1994; Sokolov and Rintoul , 2000). There may have been some diagenetic alteration ofthe δ18O signature with water depth, although this should be relatively minor as all thecores are above the calcite lysocline estimated as 3600 mbsl in the Tasman Sea (Martinez ,1997).

As a result of choosing hemipelagic marginal cores, which exhibit higher sedimentationrates, there are also problems as a result of the proximity to shore and salinity variationscaused by coastal river discharge of freshwater and sediments. This is especially importantduring the glacial with the change in palaeo-shoreline as a result of the 120 m fall insea level (Chappell et al., 1996). Terrestrial evidence from the continent suggests thatAustralia was more arid during this period (Nanson et al., 1995), thus reducing riverdischarge. Without independent estimates of palaeo-SSTs (e.g. foraminiferal Mg/Ca andabundance, or alkenones) for the cores, variations in δ18O caused by changes in salinitycannot be ruled out.

One possibility to explain the δ18O G. ruber difference between the GC-12 and GC-25during the glacial is their respective proximity to the palaeo-shoreline. GC-25 is locatedon a narrower section of the continental shelf and would have been relatively proximalto the glacial coastline (∼20 km, compared with the present ∼70 km). Therefore it isplausible that GC-25 was preferentially influenced by the upwelling of cold, high salinitywater or a cooler coastal counter current. In comparison, GC-12 would have remainedmore than 100 km away from the shoreline during the glacial, as a result of the gentlebathymetric gradient of the Capricorn Channel, although the Fitzroy River may have

7.5. Discussion 145

discharged directly into the channel at this time (Chapter 7).However, we believe that the primary reason for the δ18O G. ruber offset between

GC-12 and GC-25 is related to a more northerly position of the EAC separation andthe Tasman Front. This is supported by the data from the suite of cores from the eastAustralian margin (Table 7.3, Figure 7.4) where any local influences on δ18O in G. rubershould be conspicuous from the regional trends. It is evident from Figure 4A that a changein the G. ruber δ18O gradient occurs south of GC-25 (∼26◦S) during the Holocene, and thatthe inflexion in the gradient occurs between GC-12 (∼23◦S) and GC-25 (∼26◦S) duringthe last glacial. The largest glacial-Holocene difference (∆δ18O G. ruberGlacialHolocene)(Figure 7.4B) occurs offshore at ∼26◦S.

If the modelled physical oceanography models of Tilburg et al. (2001) are adopted,the average latitudinal positions of the EAC separation and Tasman Front are most likelyrelated to an altered wind stress curl during the glacial. Such a change in the Pacific windstress curl may be achieved through changes in the position or strength of the equatorialtrade winds and westerlies.

7.5.1 Glacial Conditions

Recent research has provided ample evidence for changes in the equatorial Pacific SSTgradients and trade winds during the glacial. Palaeo-temperature estimates for the tropicalwestern Pacific region derived from Mg/Ca in G. ruber show cooling of 2-4◦C during thelast glacial maximum (Lea et al., 2000; Stott et al., 2002; Rosenthal et al., 2003; Visseret al., 2003). In contrast, the east Pacific cold tongue exhibits a much smaller cooling of∼1◦C (Koutavas et al., 2002). This reduction in the east-west Pacific SST gradient duringthe glacial suggests El Nino-like conditions predominated and the southeast trade windswere relatively weak (Koutavas et al., 2002; Koutavas and Lynch-Stieglitz , 2003).

Evidence from studying the surface ocean of the modern tropical Pacific during ElNino events has highlighted several factors that have a significant influence on the EAC.During the 1989-1990 El Nino the width and strength of the westward South EquatorialCurrent decreased and the flow direction reversed during westerly wind bursts (McPhadenet al., 1992). A reduced South Equatorial Current would considerably decrease the flowof the EAC, which it sources.

Evidence for an El Nino-like state during the glacial is also provided by pollen anddust accumulation records from northern Australia and Indonesia, which indicate drierconditions as a result of weaker monsoons over the relatively cool western Pacific (Hesseand McTainsh, 2003).

There is also a growing body of evidence for a more northerly position of the subtrop-ical westerlies during the glacial. In the southwest Pacific region, measurements of dustconcentrations in deep-sea cores from the Tasman Sea show a dramatic 3-7 fold increase inthe dust flux from Australia during the glacial (Hesse, 1994; Hesse and McTainsh, 1999).Deep-sea sediment records suggest that the belt of dust deposition during the last twoglacials shifted 3◦ to the north as a result of a northward shift of the westerlies during


summer (Thiede, 1979; Hesse, 1994). The distribution of quartz in deep-sea sedimentcores also appears to extend further east into the Tasman Sea during the glacial (Thiede,1979), however there is no increase in particle size, which has been interpreted as evidencefor no increase in westerly wind velocity (Hesse and McTainsh, 1999). Instead Hesse andMcTainsh (1999) propose that the increase in dust flux during this time is related to agreater sediment supply as a result of a drier Australian continent.

This northward shift in the subtropical westerlies coincides with the northward ex-pansion of dunes in central Australia (Nanson et al., 1995) and evidence for a northwardshift of the subtropical front in the south Tasman Sea (Passlow et al., 1997). Furthersouth, a marine core off southeast Tasmania (FR94-GC3) displays an increase in aeoliangrain-size during glacials (MacPhail and De Deckker, unpublished data) coeval with anincrease in terrestrial plant material (Calvo et al., 2004), also suggesting a strengtheningof the westerlies.

Palaeo-evidence from more distal areas around the Southern Hemisphere also implicatechanges in the position and strength of the westerlies during the glacial. Larger dustparticles have been found in the glacial interval of the Vostok ice core, along with an overallincrease in dust concentrations (Petit et al., 1999). The decrease in Antarctic temperatureevident in the δD and δ18O of last glacial maximum ice (Petit et al., 1999) would haveincreased the pole-equator thermal gradient and increased the subtropical westerly winds,as seen in the modern winter scenario (Shulmeister et al., 2004). Analysis of marinecores from the continental margin of western South Africa (Walvis Ridge) (Stuut et al.,2002), and from western South America (Lamy et al., 1998, 1999, 2001), also suggest anequatorward movement of up to 5◦, along with a strengthening of the subtropical westerliesduring the last glacial.

Several general circulation models, however, suggest a southerly migration of the west-erlies during the last glacial based on storm track proxies and a strengthening of themeridional temperature gradient (Wyroll et al., 2000; Wardle, 2003). Wyroll et al. (2000)suggest that the winter double jet structure over Australia was expressed as a weakenedsubtropical jet and an intensified polar jet. The evidence from the last glacial palaeo-datais still ambiguous. The general consensus from the marine and terrestrial records avail-able for the last glacial suggest that the Southern Hemisphere westerlies expanded to thenorth and increased in strength, especially in the south. At the same time, the equatorialtrade winds were probably weaker in response to a reduced east-west SST gradient acrossthe tropical Pacific (Figure 7.5), so there was a westerly wind anomaly at both low andmid-latitudes. However the wind stress curl of this anomaly cannot be determined fromthe available data, therefore we cannot say whether the total curl became more or lessanticyclonic. The location of the maximum curl gradient (which should determine thelatitude of the palaeo-Tasman Front) is even harder to determine. Given the significantchanges in the wind, it is plausible that there was a sufficient curl anomaly to shift thelatitude of the maximum curl gradient, thereby explaining the shift in the Tasman Front.

Evidence from the modern ocean “Tasman Box” World Ocean Circulation Experiment

7.5. Discussion 147

Figure 7.5: Simplified diagram of one potential scenario explaining oceanic and atmospheric

patterns in the southwestern Pacific during the last glacial. The coastline of Australia is the approx-

imate coast given a change in sea level of ∼120 m (Chappell et al., 1996). The bathymetric contours

of 1500 m and 3000 m are also shown. Changes in the strength and position of the trade winds

and subtropical westerlies are shown by the size and location of arrows. The probable circulation

of the EAC is shown by a black line. This line is thinner than that for the present-day situation

(see Figure 7.1) because a weaker east-west SST gradient in the tropical Pacific during the glacial

causing relatively weak southeast tradewinds and El Nino-like conditions. The subtropical wester-

lies expanded to the north and increased in strength to the south, as shown by the larger arrow.

Separation of the EAC from the east coast of Australia is shown at ∼25◦S. It is most likely that the

Tasman Front meandered across the Tasman Sea at this latitude in response to the bathymetry, as

does the present-day front. The locations of cores PC-27a, GC-12, GC-25 and GC-9 are shown and

the dotted boxes highlight the areas covered by the suite of cores from Table 7.3 and Figure 7.4.


(WOCE P14C) was compared with high-resolution expendable bathy-thermograph (XBT)temperature measurements from the region during the prolonged El Nino of 1991-1993.This study highlighted a thinning and reduction of the warm EAC waters around northernNew Zealand, which is thought to have been responsible for the exceptionally cold 1992winter in northern New Zealand (Sprintall et al., 1995). However, Sprintall et al. (1995)also noted unusually strong southwesterly winds in records from northern New Zealandcoastal sites (Thompson and Basher , 1993) and Auckland airport during this period. Adirect link between the Southern Oscillation Index (SIO) and the subtropical westerlieswas postulated by Gordon (1986), who showed a correlation between a low SIO andanomalous southwesterly air flow in New Zealand. Gordon (1986) argued that this wasthe result of changing air pressure fields in the subtropical gyre. Therefore the reductionin the tropical trade winds may have a direct link to an increase in the strength of thesubtropical westerlies.

7.5.2 Onset of the EAC at 11 ka BP

The G. ruber δ18O records for cores GC-12 and GC-25 indicate that the EAC separationwas rapidly forced south of 26◦S between 12-11 ka BP. The timing of this event coincidedwith substantial climate change in the tropical Pacific and Southern Hemisphere. Recordsof Mg/Ca in planktonic foraminifera from the tropical western Pacific show that rapidpost-glacial warming of SSTs led deglaciation by ∼3000 years and reached modern SSTsby ∼11 ka BP (Lea et al., 2000; Stott et al., 2002; Visser et al., 2003; Rosenthal et al., 2003).Comparison of foraminiferal Mg/Ca and coral Sr/Ca records show that modern east-westSST gradients across the Pacific were established by ∼11 ka BP (Koutavas et al., 2002;Stott et al., 2002; Gagan et al., 2004). Given that the Indo-Pacific Warm Pool, east Pacificcold tongue and deep overturning atmospheric Hadley (meridional) and Walker (zonal)circulations form a tightly coupled system (Liu and Huang , 1997; Liu, 1998), the rapidpost-glacial increase in the east-west SST gradient would serve to enhance the easterlytrade winds, alter the Pacific wind-stress curl, and modify surface water circulation in thesouth Pacific subtropical gyre.

Other evidence from around the south Pacific region also supports major changes inoceanic and atmospheric circulation between around 11 ka BP. Changes in clay concentra-tions in deep-sea sediment cores from the eastern Indian Ocean indicate the presence ofdistinct wet and dry seasons and the onset of an active monsoon system around 12 ka BP(Gingele et al., 2002). Pollen sequences from southeastern Australia exhibit an expansionof forest and woodland communities throughout the region from 11.5 ka BP, suggestingboth increases in moisture and temperature in this region (Kershaw et al., 1993). There isalso evidence for increased air temperature at Taylor Dome in coastal eastern Antarctica(Steig et al., 1998) and a poleward movement and weakening of the westerlies at this time(Lamy et al., 1998; Shulmeister et al., 2004). Taken together, these climatic shifts at ∼11ka BP are likely to represent the initiation of modern ocean circulation and climate in thesouthwest Pacific. Between 11 ka BP and 5 ka BP an offset of 0.4 in δ18O still exists

7.5. Discussion 149

between the G. ruber records for cores GC-12 and GC-25. Coupled ocean-atmospheremodel results show a La Nia-like state in tropical Pacific SSTs between ∼8 ka BP and ∼5ka BP (Bush, 1999; Liu et al., 2003), with cooling in the east and subtle warming in thewest, in good agreement with SST reconstructions (Gagan et al., 1998; Koutavas et al.,2002; Stott et al., 2004). There is also evidence for a northerly shift in the Inter TropicalConvergence Zone (ITCZ), with higher rainfall in northern South America evident in theCariaco Basin sediments during the early Holocene and increasingly arid conditions de-veloping since ∼5 ka BP (Haug et al., 2001). A maximum in Asian monsoon intensity isalso evident at ∼9 ka BP, with a reduction after 6 ka BP (Overpeck et al., 1996; Morrillet al., 2003).

The GC-12 and GC-25 records of G. ruber δ18O converge to within ∼0.2 at ∼5ka BP (Figure 7.3), and remain relatively constant from 5 ka BP to present. Modernconditions are only exhibited in the tropical Pacific after about 5 ka BP (Gagan et al.,2004). This period includes the onset of modern ENSO periodicities of 2-8 years between7 ka BP and 4 ka BP shown in laminated lake sediments from Ecuador (Rodbell et al.,1999; Moy et al., 2002).

In summary, the strength of the EAC circulation and separation point are primarilycontrolled by changes in SST gradients, which affect the position and strength of theequatorial trade winds and the subtropical westerlies. The palaeo-evidence and models(e.g.(Lyle et al., 1992; Beaufort et al., 1997; Clement et al., 1999; Liu et al., 2003; Rodgerset al., 2003) indicate that small changes in insolation seasonality and tropical Pacific SSTgradients may be brought about by precession of the earths equinoxes. Variations inthe tropical SSTs can result in changes in the higher latitude subtropics, through thecirculation of western boundary currents, and potentially drive global climate change.

7.5.3 Other western boundary currents

The shift in the separation of the EAC western boundary current resulting from variationsin the wind field during the last glacial cycle is not unique. Foraminiferal abundancesin marine sediments near Japan also suggest that the eastward flowing Kuroshio Currentshifted equatorward during the last glacial in response to a southward shift of the NorthernHemisphere westerlies evident in terrestrial records from Asia (Kawahata and Ohshima,2002; Ono and Irino, 2004). However, records of benthic and planktonic foraminiferal?18O within the Gulf Stream show a weaker flow through the Florida Straits during theglacial (Lynch-Stieglitz et al., 1999), and that the separation point did not change from itspresent position at Cape Hatteras (Matsumoto and Lynch-Stieglitz , 2003). Matsumoto andLynch-Stieglitz (2003) argue that the wind-induced theory for western boundary currentseparation does not apply to the Gulf Stream separation because it is inconsistent withtheir palaeoceanographic reconstruction. The different results for the EAC, Kuroshioand Gulf Stream reconstructions highlight the complex dynamics of western boundarycurrents and the range of influences that affect their separation points and the locationof the meandering jet or front that forms in the open ocean. As a consequence, western


boundary currents need to be treated independently.

7.6 Conclusions

Analysis of δ18O in the planktonic foraminifera, Globigerinoides ruber, in deep-sea coresfrom the east Australian continental shelf indicates that the EAC separation and TasmanFront shifted north of 26◦S during the last glaciation. From palaeo-evidence we suggestthere was a reduction in the east-west SST gradient across the equatorial Pacific and adecrease in the strength of the south Pacific trade winds during the last glaciation. At thesame time, the Southern Hemisphere subtropical westerlies appear to have shifted up to5◦ equatorward and increased in strength, especially in the southern sector of the westerlystorm belt. The combination of these changes would have had a significant impact on thepalaeo-wind stress curl in the Pacific during the glacial, and potentially served to forcethe EAC to separate from east Australia at a lower latitude than today.

In contrast, a threshold was reached by ∼11 ka BP, forcing the EAC separation andthe Tasman Front south of 26◦S. This southward shift was synchronous with the end ofthe Younger Dryas, and the onset of a La Nina-like SST configuration across the tropicalPacific. In terms of the coupled ocean-atmosphere system, the shift to a La Nina-like statein Pacific SSTs after 11 ka BP implies that the Hadley-Walker circulation system wasinvigorated at that time. These changes in the Pacific ocean-atmosphere system wouldhave enhanced the easterly trade winds and altered the Pacific wind-stress curl, thusmodifying the surface water circulation in the south Pacific subtropical gyre. Therefore,it appears that the EAC separation and the location of the Tasman Front readjustedaccordingly. This readjustment of the EAC ∼11 ka BP marks the onset of near-modernsurface-ocean circulation in the southwest Pacific. The present EAC transport, however,was only achieved after 5 ka BP with the onset of modern conditions, including presentday El Nino Southern Oscillation periodicities. The link between EAC flow and El Ninoevents allows us to speculate that with increasing frequency of these events the EAC maypotentially weaken in the future with dramatic effects on the climate of eastern Australiaand New Zealand.

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