Late Quaternary (marine isotope stages 6-1) seismic sequence stratigraphic … › ~eosterberg › Osterberg et al. 2006b... · 2008-12-11 · Late Quaternary (marine isotope stages

(2006) 159–178www.elsevier.com/locate/margeo

Marine Geology 229

Late Quaternary (marine isotope stages 6-1) seismicsequence stratigraphic evolution of the Otago

continental shelf, New Zealand

Erich C. Osterberg ⁎

Department of Geology, University of Otago, P.O. Box 56, Dunedin, New Zealand

Received 7 October 2005; received in revised form 26 February 2006; accepted 17 March 2006

Abstract

A proposed chronostratigraphic model for the Late Quaternary evolution of the Otago continental shelf, New Zealand, includesfluvially incised sequence boundaries eroded during marine isotope stages 2, 4 and 6, bounding three sequences subdivided intolowstand, transgressive and highstand/regressive systems tracts. Datable material is limited to the uppermost 1 m of the sequence,and consequently the model is based on correlation of new high-resolution seismic reflection data (Geopulse™) to a compilation ofindependent global sea-level curves. A second chronostratigraphic model that lacks a stage 4 sequence boundary and distinct stage3 highstand/forced regressive deposit is discounted because its subsidence-corrected paleoshoreline depths do not match lowstandsin the eustatic curves or temporally constrained sequences at other far-field locations. In the preferred chronostratigraphic model,sea-level lowstands during stages 2, 4 and 6 reached maximum depths of 126–131 m, 101–111 m and 116–135 m, respectively,producing sequence boundaries, landward-pinching deltaic wedges on the outer Otago shelf, and submarine canyon incision.Lowstand paleoshoreline depths have been corrected for subsidence from thermal cooling, sediment compaction and loading, butmay be erroneously deep owing to erosion of paleoshoreline indicators. Clastic, shallow marine wedges and back-barrier valley-filldeposits accumulated during the stage 6–5, 4–3 and 2–1 marine transgressions. Contrary to the traditional sequence stratigraphicmodel, deltaic and strandline units deposited during highstand through falling sea level (stages 5–4, 3–2, 1) volumetricallydominate Otago Late Quaternary sequences. Discrete units and minor marine erosion surfaces within the transgressive andregressive systems tract deposits are interpreted as strong evidence for seventh- (∼ 20 ka) or higher-order sea-level fluctuationsinfluencing sedimentation on the Otago shelf.© 2006 Elsevier B.V. All rights reserved.

Keywords: sequence stratigraphy; paleoshoreline; subsidence; regressive systems tract; sea level; New Zealand

1. Introduction

The Otago continental margin, South Island, NewZealand, is located in the heart of a convergent margin

⁎ Present address: Department of Earth Sciences, University of Maine,Orono, ME 04469, USA. Tel.: +1 207 581 2112; fax: +1 207 581 1203.

E-mail address: [email protected].

0025-3227/$ - see front matter © 2006 Elsevier B.V. All rights reserved.doi:10.1016/j.margeo.2006.03.005

sedimentary system that stretches∼ 1000 km from activesediment source regions to abyssal fans in the Pacif-ic Ocean (Fig. 1a). Terrigenous sediment is shed fromthe rapidly uplifting Southern Alps, transported to thecoast via large rivers, and deposited either on the con-tinental shelf or in the submarine canyons feeding theBounty Channel and Fan Complex (Carter et al., 1985).The sedimentological evolution of the margin since its

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Fig. 1. a) New Zealand South Island showing the location of the Otago continental shelf with respect to the Pacific–Australian plate boundary,sediment source areas in the Southern Alps, and the Bounty Channel/Fan complex. The path of sediment is shown by open arrows. The location ofODP Site 1119 and the Galleon-1 and Endeavor-1 petroleum exploration boreholes is also shown. b) Study area bathymetric map showing the grid ofhigh-resolution seismic reflection profiles and the location of figures in this paper (bold lines).

160 E.C. Osterberg / Marine Geology 229 (2006) 159–178

inception ∼ 80 Ma (Molnar et al., 1975) has conse-quently depended on the transient interplay of the upliftand erosion rates of the Southern Alps, the competenceand location of rivers draining the Alps, marine currentsassociated with the Southland Front, and relative sealevel, all of which depend on global and local climateconditions. Like most passive margins, Milankovitch-order (20–120 ka) sea-level fluctuations induced by thewaxing and waning of continental ice sheets have beenthe primary control on the Late Quaternary (last∼ 500 ka) evolution of the Otago margin. In this paper,the sedimentary response to these eustatic changes isinvestigated from a sequence stratigraphic perspectivebased on newly acquired high-resolution seismicreflection data from the Otago shelf.

Late Quaternary shelf sequences have increased ourunderstanding of the relationship between sedimenta-tion, erosion and sea level change because they weredeposited during a period for which sea level has beenwell constrained by independent proxy records (e.g.Chappell and Shackleton, 1986; Shackleton et al., 1990).High-resolution seismic reflection studies have foundthat contrary to the traditional sequence stratigraphicmodel, units deposited during falling relative sea leveloften volumetrically dominate Late Quaternary shelfsequences because of the longer duration of fifth-(∼ 120 ka) and sixth-order (∼ 40 ka) glacio-eustaticregressions compared to rapid transgressions (e.g.Posamentier et al., 1992; Naish and Kamp, 1997;Haywick, 2000; Browne and Naish, 2003). Such studies

161E.C. Osterberg / Marine Geology 229 (2006) 159–178

have also described discrete sediment units and minorerosion surfaces attributed to rapid (1–10 ka), climate-driven sea-level cycles superimposed on the dominantMilankovitch-order glacio-eustatic signal (Hernandez-Molina et al., 2000; Hamberg and Nielsen, 2000; Plintand Nummedal, 2000). The shelf sequences describedhere provide examples of both patterns in a region thathas not been studied extensively.

While sea-level proxies can be valuable for under-standing sedimentary responses to base level changes,they can also provide chronological constraints on LateQuaternary sequences in lieu of datable material fromcores or boreholes. Using sequence stratigraphic princi-ples, highstand and lowstand shorelines are identifiableon seismic profiles and can be compared to eustaticcurves after taking account of subsidence. This methodhas been used to chronologically constrain shelfsequences in New Zealand (Browne and Naish, 2003)and Spain (Hernandez-Molina et al., 2000), and is usedhere to define two chronostratigraphic models for Otagoshelf evolution over the past 250 ka, incorporating adetailed analysis of subsidence in the study area. The twomodels differ primarily in their interpretation of sea levelduring oxygen isotope stages 3 and 4, which is not astightly constrained as during other periods of the LateQuaternary, and has been the subject of some debate (e.g.Wellner et al., 1993; Cann et al., 2000; Rodriguez et al.,2000). The preferred Otago chronostratigraphic modelrequires that the stage 4 lowstand and stage 3 highstandwere significant sea-level fluctuations capable ofproducing a complete sequence on the Otago shelf.

2. Regional setting

2.1. Otago shelf description and Tertiary evolution

The modern Otago continental shelf stretches fromShag Point in the north to Nugget Point in the south,ranging in width from 12 to 45 km (Fig. 1). The 220 km2

study area discussed here covers a high-gradient (0.006–0.007), 15–30 km-wide portion of the shelf directly eastof the Otago Peninsula (a remnant Miocene volcaniccomplex; Fig. 1b). South of the peninsula, most of theOtago coastline is cliffed as a result of uplift alongnortheast-trending (roughly shore-parallel), eastward-dipping reverse faults within the Otago reverse faultprovince (Litchfield and Norris, 2000). North of thepeninsula, the northwest-trending, reverse Waihemofault system has uplifted Shag Point and sedimentsfurther eastward on the continental shelf (Allan, 1990).The study area discussed here lies northeast of the Otagoreverse fault province (Johnstone, 1990) and south of the

Waihemo fault system (Allan, 1990), and has beenrelatively passive during the Late Quaternary. The outerOtago shelf is incised by seven major submarinecanyons, including Papanui Canyon within the studyarea (Fig. 1b), which have periodically channeled terri-genous sediment to the abyssal Bounty Channel and Fancomplex since the Early Miocene (Fig. 1a; Carter andCarter, 1987).

Geophysical surveys and petroleum exploration bore-holes on the Otago and Canterbury (located north of thestudy area) shelves have revealed a 2–5 km-thick trans-gressive–regressive megasequence spanning ∼ 80 Ma,the deposition of which was controlled primarily by theevolution of the Pacific–Australian plate boundary(Molnar et al., 1975; Carter, 1988; Fulthorpe and Carter,1991; Fulthorpe et al., 1996). Jurassic schist basementunderlies a Late Cretaceous to Oligocene marinetransgressive sequence consisting of basal fluvial depositsand coal measures, overlain by a deepening succession ofmarine sandstone, limestone and mudstone (Wilson,1985; Carter, 1988). Maximum transgression in theOligocene (∼ 30 Ma) is represented by a limestone unit,followed by an eastward thickening, Miocene-to-presentregressive marine sediment wedge consisting predomi-nantly of sandstone and mudstone (Wilson, 1985; Carter,1988; Fulthorpe and Carter, 1991; Fulthorpe et al., 1996;Lu and Fulthorpe, 2004; Lu et al., 2005). The uppermost,Late Quaternary portion of the megasequence is the focusof this paper.

2.2. Hydraulic regime

The dominant hydrologic feature on the Otago conti-nental margin is the Southland Front, which is anextension of the global Subtropical Front separatingrelatively warm and salty Subtropical Water on thecontinental shelf from relatively cold and fresh Subant-arctic Water offshore (Heath, 1972). Associated with theSouthland Front is the Southland Current, a northwardflow (∼ 8.3 Sv; ∼ 0.15–0.30 m/s) of dominantlySubantarctic water (Chiswell, 1996; Sutton, 2003) thathas been responsible for the formation of the CanterburyDrifts north of the study area since at least the Pliocene(Fulthorpe and Carter, 1991; Carter et al., 2004b; Lu andFulthorpe, 2004). The local semidiurnal tide averages1.5 m in height with an average velocity over the Otagoshelf of 0.1 to 0.2 m/s (Andrews, 1973; Carter et al.,1985). Southerly gales mobilize sand on the middle andouter shelf up to several times a year by accentuatingthe Southland Current and producing storm swell capa-ble of stirring sand at N100 m depth (Andrews, 1973;Carter and Herzer, 1979; Carter et al., 1985). Bed-load


mobilization during calm weather is believed to be re-stricted to the surf zone where active longshore process-es transport sediment to the northeast (Andrews, 1973;Carter and Herzer, 1979).

2.3. Sediment supply and surface distribution

Terrigenous sediment on the Otago shelf is sourcedprimarily from the Clutha (pre-dam sediment flux of∼ 3.2 million tons per annum (Mt a−1)), Taieri (∼ 0.6 Mta−1) and Tokamairiro (∼ 0.1 Mt a−1) Rivers (Carter,1986). The Clutha River drains Lakes Wanaka, Haweaand Wakatipu nested within the high-relief outboard(eastern) portion of the Southern Alps (Fig. 1a). All threerivers drain catchments enriched inMesozoic, low-gradeprehnite–pumpellyite and greenschist facies, and con-sequently have a predominantly quartzofeldspathic bed-load with a high mica content (Andrews, 1973; Carter etal., 1985).

Four, roughly shore-parallel surface sediment beltshave been identified in previous Otago shelf studies: afine, well-sorted, modern quartz sand facies from 0 to∼ 55m depth; a rounded, iron-stained, relict quartz gravelfacies from∼ 55 to∼ 70m depth; a coarse, poorly-sorted,iron-stained, relict/palimpsest quartz sand facies from∼ 70 to ∼ 100 m depth; and a poorly-sorted, patchy,biogenic sand and gravel facies from ∼ 100 to ∼ 150 mdepth (Andrews, 1973; Carter et al., 1985; Orpin et al.,1998). The modern sand and biogenic gravel facies havebeen interpreted as Holocene highstand deposits, whilethe relict terrigenous sand and gravel facies have beeninterpreted as post-glacial transgressive units deposited inresponse to episodic changes in the rate of post-glacialsea-level rise (possibly including pauses and/or reversals)(Carter et al., 1985; Orpin et al., 1998).

3. Methods

3.1. Data collection and processing

Two hundred linear km of single-channel, high-resolution seismic reflection data were collected in a gridpattern over a 220 km2 area of the Otago continentalshelf and Papanui Canyon during 14 individual day-longcruises on the University of Otago research vesselMunida between March 9, 2000 and February 1, 2001(Fig. 1b). Position fixes were acquired at a 5-s intervalusing a Fugro OmniSTAR™ differential global posi-tioning system. The seismic data were collected using aFerranti-ORE Geopulse Sub-bottom Profiling Sys-tem™, and recorded digitally in raw form for reproces-sing. In general, power levels of 175 and 350 J were used

over the continental shelf and Papanui Canyon, respec-tively, with a 0.5–2 pulse/s firing rate. Vessel surveyspeed ranged from approximately 3–5 knots, resulting ina ∼ 1.5–3 m seismic shot spacing interval. A Ferranti-ORE 5210A Receiver™ was used to process the rawseismic data with flat gain, time-varied gain, bandpassfilter (500–2000 Hz) and swell filter. Seismic profileswere all printed on an EPC 4800™ graphic recorder, forwhich a constant seismic velocity of 1470 m/s wasutilized for all time–depth conversions.

3.2. Sequence stratigraphic interpretation

Seismic reflection data are interpreted in a sequencestratigraphic framework using established procedures(e.g. Mitchum and Vail, 1977; Vail, 1987). Regionallyextensive seismic horizons are interpreted as sequenceboundaries (SB), ravinement surfaces (RS), downlapsurfaces (DLS) and internal truncation surfaces (ITS)based on their orientation, stratal relationship (lapout,truncation), amplitude and continuity. Four systemstracts are identified in each complete sequence: thelowstand (LST), transgressive (TST), highstand (HST)and regressive (RST) system tracts in ascendingstratigraphic order. The regressive systems tract (“fallingstage systems tract” of Hart and Long (1996) and Plintand Nummedal (2000); “forced regressive systems tract”of Hunt and Tucker (1992)), comprises sedimentaryunits deposited during falling relative sea level (Naishand Kamp, 1997; Browne and Naish, 2003). Where theboundary between HSTand RST deposits within a singlesequence is indiscernible, the HST and RST have beencombined into the highstand/regressive systems tract(HRST). Sea-level lowstand paleoshorelines (PS) areidentified at the landward pinch-out depths of LSTdeltaic wedges, at the crests of LST barrier islands, and atthe inflection of LST paleobeach faces (Section 4.1.1).

The sequence stratigraphic classification of forcedregressive deposits and their position relative to thesequence boundary have been somewhat controversial(Posamentier et al., 1992; Hunt and Tucker, 1992; Naishand Kamp, 1997; Plint and Nummedal, 2000; Tesson etal., 2000). Otago forced regressive deposits most closelyresemble the strongly progradational Wanganui basinand Canterbury Shelf deposits assigned to the regressivesystems tract by Naish and Kamp (1997) and Browneand Naish (2003), and the same terminology is adoptedhere for simplicity. Under this labeling scheme, sequenceboundaries immediately overlie and may representsignificant erosion of regressive systems tract deposits,which are therefore the youngest deposits in eachsequence (Naish and Kamp, 1997; Haywick, 2000;


Browne and Naish, 2003). Sequences are labeled withArabic numerals by increasing age (1=youngest). Eachsystems tract, unconformity and paraconformity isdenoted by its acronym with its sequence number insubscript (e.g. SB1, TST2).

3.3. Chronostratigraphy

Developing a chronostratigraphic model for theOtago continental shelf is difficult owing to the paucityof Late Quaternary borehole and sediment core datafrom the region. The Plio-Pleistocene portion of thesediment column was discarded by petroleum boreholecontractors who drilled on the shelf (e.g. Galleon-1;Wilson, 1985), and the few piston cores that werecollected by previous researchers only penetrated 0.5–1 m into Holocene shelf sediments (Carter et al., 1985).

Table 1Otago shelf stratigraphic column and chronostratigraphy under Models 1 an

Underlyinghorizon a

Unit b Moderndescription

Thickness Stratalrelationship

DLS1 HST1 Shore-connected,prograding wedge

15–30 m Low amp. clindownlapping D

RS1 TST1a Three mid-shelfback-steppingsediment wedges

5–8 m Transparent weonlapping RS1

TST1b Inner to mid-shelfincisedchannel fill

5–15 m Chaotic fill boby SB1 belowabove

SB1 LST1a,b Outer shelfprograding wedge(a) and fill (b)

a) 60+ mb) 20 m

Onlapping SB1

landward and oit seaward

DLS2 HRST2 Inner to outer shelfprograding wedge

10–20 m Low amp. clindownlapping D

RS2 TST2a Mid-shelf strandedsediment lens

1–5 m Transparent lenonlapping RS2

TST2b Inner to mid-shelfincised channel fill

5–15 m Chaotic fill boby SB2 belowand RS2 above

SB2 LST2a,b Outer shelfprograding wedge(a) and mound (b)

a) 5–30 mb) 60+ m

a) On/offlappinb) fill capped b

DLS3 HRST3 Inner to outer shelfprograding wedge

20–30 m Moderate ampclinoformsdownlappingDLS3; ITSs

c

RS3 TST3 Three mid-shelfback-steppingsediment wedges

5–15 m Transparent weonlapping RS3

SB3 LST3a,b Outer shelfprograding wedge(a) and fill (b)

a) 40+ mb) 60 m

a) On/offlappinb) fill capped b

HRST4 Inner to outershelf prograding wedge

20+ m Low amp. clindownlap surfac

a SB: sequence boundary, DLS: downlap surface, RS: ravinement surface.b HST: highstand systems tract, TST: transgressive systems tract, LST: low

ODP Site 1119, located ∼ 175 km north of the studyarea on the South Canterbury upper continental slope(395 m below mean sea level (bmsl); Fig. 1a), is themost accurately dated Quaternary sediment record onthe Otago or Canterbury margins (Carter and Gammon,2004; Carter et al., 2004a,b; Lu and Fulthorpe, 2004; Luet al., 2005). Although the Site 1119 record provides avaluable analog to Otago shelf sequences, its location∼ 15 km beyond the shelf break and proximity tosediment-laden rivers (e.g. Waitaki River) draining theeastern flanks of the Southern Alps limit its usefulnessfor developing a high-resolution chronostratigraphicmodel for the Otago continental shelf.

Given these limitations, two chronostratigraphicmodels of Otago shelf evolution are proposed bycorrelating Otago shelf sequences to a compilation ofsea-level data from raised coral terraces at Huon

d 2

Depositionalenvironment

Seq. Model 1age

Model 2age

oformsLS1

Holocene inner shelf 1 Stage 1 Stage 1

dges Inner shelf duringhigher order sea-levelfluctuations

Stage 2-1 Stage 2-1

undedand RS1

Estuary Stage 2-1 Stage 2-1

fflappingMarine delta (a)and barrier (b)

Stage 2 Stage 2

oformsLS2

Strandline or inner shelf 2 Stage 3-2 Stage 5-2

s Inner shelf during higherorder sea-level fluctuations

Stage 4-3 Stage 6-5

unded Estuary Stage 4-3 Stage 6-5

g SB2

y RS2

Marine delta (a)and canyon fill (b)

Stage 4 Stage 6

. Marine delta 3 Stage 5-4 Stage 7-6

dges Inner shelf duringhigher order sea-levelfluctuations

Stage 6-5 Stage 8-7

g SB3

y RS3

Marine delta (a)and canyon fill (b)

Stage 6 Stage 8

oforms;e not seen

Marine delta or strandline? 4 Stage 7-6 Stage 9-8

c ITSs: internal truncation surfaces.stand systems tract, HRST: highstand/regressive systems tract.

Fig. 2. Uninterpreted (a) and interpreted (b) seismic line 12 showing three Late Quaternary sequences delineated by sequence boundaries (SB) and sub-divided into lowstand (LST; medium gray),transgressive (TST) estuarine (light gray) and marine (dark gray), highstand (HST; no fill) and highstand/regressive (HRST; no fill) systems tract deposits. Locations of lowstand paleoshorelines (PS)are also noted.

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Fig. 3. Uninterpreted (a) and interpreted (b) seismic line 22 showing three Late Quaternary sequences delineated by sequence boundaries (SB and sub-divided into lowstand (LST; medium gray),transgressive (TST) estuarine (light gray) and marine (dark gray), highstand (HST; no fill) and highstand/regressive (HRST; no fill) systems tra deposits. Locations of lowstand paleoshorelines (PS)are also noted. Note the massive canyon fill deposit LST2b.

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)ct


Peninsula, Papua New Guinea (PNG; Chappell andShackleton, 1986; Chappell et al., 1996) and Malakula,Vanuatu (Cabioch and Ayliffe, 2001), sedimentarydeposits from the Bonaparte Gulf, Australia (Yokoyamaet al., 2001a), planktonic foraminiferal δ18O ratios fromODP sites 768 and 769 in the Sulu Sea (Linsley, 1996),benthic foraminiferal δ18O ratios from ODP site 677 inthe equatorial Pacific (Shackleton et al., 1990), and thesea-level component of the V19–30 benthic foraminif-eral δ18O signal (determined by extracting the temper-ature component through comparison with the Vostokair δ18O record; Shackleton, 2000). The total sea-levelrange indicated by the proxies through time defines a“sea-level envelope” (SLE), and it is assumed that sealevel on the Otago shelf was within the SLE at any timeduring the Late Quaternary.

Using the SLE as an independent sea-level proxy,Otago shelf sedimentary sequences are counted backfrom a temporally-constrained starting point by assum-ing that 1) sea level was the predominant control onsedimentation and erosion, and 2) every significant sea-level cycle created a corresponding sequence that can beidentified on high-resolution seismic profiles. Theseassumptions are justified for shelf sequences depositedduring the Late Quaternary, when the amplitudes ofdominant (∼ 120 ka period) sea-level cycles were greaterthan 80 m (e.g. Imbrie et al., 1984) and instantaneousrates of sea-level change often exceeded 10 mm/yr(Lambeck et al., 2002). Though not a substitute forabsolute dating methods, this dating technique is usefulfor developing a robust chronostratigraphic model whenabsolute dates are sparse, and for providing a workingmodel to aid in the collection of future data.

Fig. 4. Uninterpreted (a) and interpreted (b) portion of seismic line 22 on theincluding deltaic wedges (LST3a, LST2a), a paleobeach (LST1a) and paleobarby SB1, and the similarity in depth of the paleobeach inflection and barrier

4. Results

4.1. Sequence architecture

Sequence stratigraphic interpretation of the seismicreflection data reveals portions of four sequencesseparated by three regionally extensive sequence bound-aries on the Otago shelf (Table 1, Figs. 2 and 3). Sequence1 is presently being deposited, and will not be completeuntil its upper sequence boundary is created during thenext major glacio-eustatic regression and lowstand.Sequences 2 and 3 are both visible in full, but only theuppermost portion of Sequence 4 is visible and will not bediscussed further. Sequence boundaries on theOtago shelfare interpreted as diachronous unconformities thatprogressively formed through erosive fluvial and subaer-ial processes during sea-level falls and lowstands.Ravinement surfaces, created by shoreface wave erosionduring marine transgressions, are generally amalgamatedwith subjacent sequence boundaries on seismic profiles,creating a readily identifiable high-amplitude seismicreflector. Sequence boundaries and ravinement surfacesdiverge, however, at fluvially incised valleys and channelson the inner and middle shelf. There is no evidence on theseismic profiles of active tectonism that would affectsedimentary deposition.

Within the study area, Sequences 1, 2 and 3 displayuniform stratal architecture and thickness along strike,with three exceptions. First, an infilled fossil TaiaroaCanyon head visible on line 22 (Fig. 3) has eroded asignificant portion of Sequence 3 at mid-shelf depths inthe northern part of the study area. Multichannel (120-channel) seismic reflection profiles collected by BP

outer shelf showing architecture of lowstand deposits (medium gray),rier deposit (LST1a). Note the erosional truncation of the LST2a wedgecrest (PS1).


Shell Todd Services, Ltd. in 1982 and 1984 (not shownhere) reveal this feature to be the edge of a large (N4 kmwide,∼ 400 m deep) region that has systematically beeneroded and infilled at the head of Taiaroa Canyon, withthe most recent cut and fill cycle occurring duringdeposition of Sequence 2 (Fig. 3). Second, lowstanddeposits in the study area have been modified by thepresence of Papanui Canyon, with different LSTgeometries north and south of the canyon (see 4.1.1below), and the absence of LST1a at the head of thecanyon on lines 9 and 10 (Fig. 6). Third, the fluviallyincised channels show considerable variation in thick-ness and width, and an increased abundance towards thenorthern part of the study area (e.g. line 22).

4.1.1. Lowstand systems tract depositsLowstand systems tract features on the outer Otago

shelf include three landward-pinching wedges, astranded barrier-like structure and infilled fossil canyonheads. The three wedges (LST1a,2a,3a) display high-angle (up to 5°), low-amplitude, oblique parallel clino-forms that offlap the subjacent sequence boundaryseaward and onlap it landward (Table 1; Figs. 2–4). Thisaggredational/progradational architecture is indicativeof marine–deltaic deposition during peak lowstand andearliest transgression and is commonly observed onouter continental shelves worldwide (Berryhill et al.,1986; Tesson et al., 2000). North of Papanui Canyon,Line 22 reveals a 20 m thick barrier-like structure

Fig. 5. Uninterpreted (a) and interpreted (b) portion of seismic line 22 on the in(SB) formation, infilled with interpreted transgressive estuarine sediment (Ttransgressive shallow marine wedges (dark gray) and/or subsequent highstan

located ∼ 4 km basinward from an interpretedpaleobeach deposit overlying SB1, which truncatesLST2a (Figs. 3 and 4). The stranded barrier andpaleobeach are assigned to LST1b. Line 22 also includesa massive (60+ m thick), infilled, fossil Taiaroa Canyonhead within Sequence 2 (LST2b; Fig. 3), while aninfilled, fossil Papanui Canyon head is visible on severalprofiles within Sequence 3 (LST3b; not shown). Sea-level regressions and lowstands promoted the incision ofPapanui and Taiaroa Canyon heads several kilometersinto the shelf through retrogressive mass wasting (e.g.Berryhill et al., 1986; Chiocci et al., 1997; Tesson et al.,2000). Canyon fill deposits are interpreted to haveaccumulated during late lowstand through early trans-gressive periods (Chiocci et al., 1997; Tesson et al.,2000), and are assigned to the LST here for simplicity.

The identification of lowstand paleoshoreline posi-tions on the Otago shelf is an essential part of thecorrelation between shelf sequences and the sea-levelenvelope. The three LST wedges belong to differentsequences (i.e. they were deposited during differentlowstands) and pinch out at different depths, but theirdepths do not simply increase with age as might beexpected. LST2a, the intermediately-aged wedge (asdetermined by stratigraphic principles), pinches out at121–122 m depth (PS2), while the younger and olderwedges (LST1a and LST3a) pinch out at 133–134 m(PS1) and 157–159 m (PS3) depth, respectively (Figs.2–4). The pinch-out depth of each wedge is consistent

ner shelf showing channels incised during lowstand sequence boundaryST, light gray), capped by ravinement surfaces (RS), and overlain byd and forced regressive deposits (HRST, no fill).


within 1–2 m on multiple seismic profiles. The crest ofthe interpreted stranded barrier and the top of thepaleobeach face (LST1b) north of Papanui Canyon areboth found at 133–134 m depth (Fig. 4), equivalent tothe pinch-out depth of wedge LST1a south of the canyon(Fig. 2). Thus, it appears that the most recent lowstandreached an uncorrected (for subsidence) depth of 133–134 m (PS1). During this lowstand a deltaic systemsouth of Papanui Canyon produced wedge LST1a, and abarrier island system north of the canyon erodedportions of LST2a and left remnants of the barrier andback-barrier beach (LST1b).

4.1.2. Transgressive systems tract depositsTransgressive systems tract deposits are preserved as

incised valley fill and back-stepping, middle-shelfwedges (Table 1; (Figs. 2, 3 and 5)). Fluvially incisedvalleys on the inner shelf generally range from 150 to3000 m in width and 5–15 m in thickness, although thewidth values are maximums because seismic lines aremost likely oblique to individual valleys. TST1b andTST2b are interpreted as back-barrier, incised valley filldeposits capped by transgressive ravinement surfaces(RS1,2; Table 1, Fig. 5). The transgressive, back-barrierorigin of at least the uppermost portion of TST1b was

Fig. 6. (a) Seismic Line 9 with sequence boundaries (SB), paleo-shorelines (Pgray) and marine (dark gray), highstand (HST; no fill) and highstand/regressivof Line 9 showing prograding clinoforms and internal truncation surfaces (ITNote the lack of high-amplitude clinoforms within HRST2, interpreted as a

confirmed by piston core P152 collected by Carter et al.(1985), which penetrated estuarine mud complete withan in situ Austrovenus stutchburyi.

Thin (5–15 m thick), discrete, backstepping unitswithin TST1a, TST2a and TST3 are interpreted asshallow marine deposits that accumulated as shore-connected wedges during brief periods of increasedsediment flux and/or periods of slower transgression,stillstand or temporary regression during lower-ordertransgressions (Table 1, (Figs. 2, 3, 5 and 6); Carter etal., 1985, 1986; Hernandez-Molina et al., 2000). Theunits were probably reworked significantly by shelfcurrents and storm waves during their submergence(Carter et al., 1985, 1986; Goff et al., 2005). Theyoungest backstepping units (TST1a) thin considerablyto the south and most likely represent the sourcedeposits for the relict gravel and sand facies (Carter etal., 1985) on the shelf surface today.

4.1.3. Highstand/regressive systems tractHighstand and regressive systems tract deposits

volumetrically dominate Sequences 1, 2 and 3 withthicknesses of 10–30 m (Table 1; (Figs. 2, 3 and 6)).Forced regressive deposits are characterized by theprogressively lower elevation of depositional clinoforms

S), lowstand (LST; medium gray), transgressive (TST) estuarine (lighte (HRST; no fill) systems tract deposits identified. (b) Expanded sectionS) within HRST3, interpreted as a forced-regressive marine deltaic unit.strandline or inner shelf forced-regressive deposit.


towards the shelf break, the seaward-dipping orientationof the overlying sequence boundary, and their strati-graphic position as strongly progradational deposits onthe middle to outer shelf (Posamentier andMorris, 2000;Roberts et al., 2004). Otago shelf highstand depositsgradually transition to forced regressive deposits withno discernable boundary, and are consequently groupedinto the combined highstand/regressive systems tract(HRST). The only HST unit in the study area without acorresponding RST component is HST1, the shore-connected Holocene sand wedge interpreted to haveprograded to a depth of 55–60 m during the presenthighstand (∼ 6.5 ka to present; Carter et al., 1985; Carterand Carter, 1986; (Figs. 2, 3 and 5)). Seaward of HST1

(60 m to at least 150 m depth), terrigenous sedimentstarvation of the modern shelf is currently producing acondensed section with a mid-cycle shell bed consistingof modern biogenic growth amidst reworked relict shellhash (Carter et al., 1985; Orpin et al., 1998).

HRST3 is characterized by a succession of obliquetangential clinoforms and internal truncation surfaces(Fig. 6) that exhibit a complex mounded structure onstrike profiles, implying a deltaic origin (Hart and Long,1996; Tesson et al., 2000; Roberts et al., 2004). A deltaicinterpretation is supported by the presence of incisedfluvial valleys shoreward. The high-angle (1–2°)portions of clinoforms are interpreted as sand- andgravel-rich delta-front foresets, grading downdip tomud-rich prodelta bottomset facies that downlap at verylow angles (∼ 0.1°) or fade out without downlapping(Kolla et al., 2000; Tesson et al., 2000; Roberts et al.,2004). Internal truncation surfaces are marine erosionsurfaces within the deltaic unit created by either lobeswitching (Hart and Long, 1996; Kolla et al., 2000) oran abrupt lowering of wave base (Plint and Nummedal,2000).

In contrast to the high-amplitude ITSs, moundedclinoforms and mud-rich pro-delta facies of deltaicHRST3, HRST2 displays only sporadic, low-amplitudeoblique clinoforms (Figs. 2 and 3). This stratal archi-tecture is characteristic of forced regressive marine de-position in a strandline or shallow marine environment(Dominguez and Wanless, 1991; Posamentier et al.,1992; Naish and Kamp, 1997; Hamberg and Nielsen,2000; Haywick, 2000). Clinoforms may define theformer seaward edge of the shore-connected wedge orlower shoreface as it prograded across the shelf overtime (Dominguez and Wanless, 1991). Thus, HRST2

probably formed in a depositional regime resemblingthat of the study area today, with longshore currentstransporting sediment northwards from a proximal deltalocated to the south.

4.2. Chronostratigraphic models

Two chronostratigraphic models displayed schemat-ically in Table 1 comply with the interpretation ofseismic data from the Otago shelf. Sequence 1 is wellconstrained to oxygen isotope stages 2 and 1 (oxygenisotope stages (hereafter “stages”) referred to here arethose of Imbrie et al., 1984) by radiocarbon-dated faunalsamples (Carter et al., 1985), sedimentological evidencefrom dredges and core samples (Andrews, 1973; Carteret al., 1985), and stratigraphic principles, specifically:

1. A fossil A. stutchburyi recovered from estuarine mud(TST1b) in core P152 has a radiocarbon age of12,150±300 yr BP (Carter et al., 1985), constrainingdeposition to the stage 2/1 transgression;

2. shallow-water faunal specimens associated with thesurficial relict/palimpsest sand facies on the middle-outer shelf similarly yield a post-glacial transgressiveradiocarbon age of 10,050±250 yr BP (Carter et al.,1985);

3. SB1 is the youngest sequence boundary within thestudy area, and most likely represents the stage 2lowstand erosion surface;

4. The 133–134 m pinch-out depth of the LST1a wedgecorrelates well with stage 2 sea-level data from theBonaparte Shelf, PNG, foraminiferal δ18O ratios, andwith the estimated stage 2 lowstand depth of 135 malong the Canterbury shelf break (Chappell andShackleton, 1986; Linsley, 1996; Shackleton et al.,1990; Shackleton, 2000; Yokoyama et al., 2001a;Browne and Naish, 2003).

Sequence 1 is the temporally-constrained startingpoint for the countback method in both models. Carter etal. (1985) further refined the stage 2–1 transgressiveOtago sequence by assigning radiocarbon ages tospecific surficial shelf deposits, and Gibb (1986)published a New Zealand Holocene sea-level curvebased on radiocarbon dated coastal deposits. The seismicdata collected in this study support the interpretations ofCarter et al. (1985) and Gibb (1986), though it was notpossible to collect additional radiocarbon samples to testtheir transgressive models more rigorously.

The fundamental difference between chronostrati-graphic Model 1 and Model 2 is the interpretation of sealevel during stages 4 and 3 from approximately 70 to40 ka. Model 1 is based on the assumptions that thestage 4 lowstand was sufficient in duration andamplitude to produce an identifiable sequence boundaryand lowstand wedge within the study area, and that thestage 3 highstand produced a distinct HRST deposit.

Table 2Otago shelf subsidence rate estimates separated by component for eachchronostratigraphic model

Subsidence component Subsidence rate(m/ka)

Model 1 Model 2

Thermal cooling 0.017 0.017Compaction 0.08–0.13 0.04–0.07Sediment loading 0–0.20 0–0.10Airy isostasy 0.17–0.20 0.09–0.10Periodic load flexure ∼ 0.001 ∼ 0.0004Line load flexure 0.08–0.09 0.04–0.05

Total subsidence rate 0.10–0.34 0.06–0.19Total w/ periodic flexure 0.10–0.14 0.06–0.08Total w/ line flexure 0.18–0.23 0.10–0.13Total w/ airy Isostasy 0.27–0.34 0.15–0.19


Model 2, however, is based on the assumption that thestage 4 lowstand and stage 3 highstand were only minorsea-level excursions within the overall stage 5–2 marineregression. Thus, SB2 was created during either the stage4 (Model 1) or stage 6 (Model 2) lowstand (Table 1).The two models are scrutinized and evaluated belowthrough stratigraphic analyses and a comparison ofsubsidence-corrected paleoshorelines to the sea-levelenvelope.

4.3. Quantifying subsidence to correct lowstandpaleoshorelines

Lowstand paleoshorelines are assigned at the crestsof interpreted barrier bars, at the inflection of paleo-beach faces, and at the landward pinch-out depths oflowstand wedges (Figs. 2–4). However, in order tocompare lowstand paleoshorelines to sea-level curves,present-day paleoshoreline depths must be corrected forsubsidence. No comprehensive accounting of subsi-dence has previously been published for the study area,although Gibb (1986) determined that Blueskin Bay(20 km west of the study area; Fig. 1b) is tectonicallystable. Subsidence on the Otago margin is the result oflithospheric cooling since rifting, sediment compaction,and loading by sediment and water. Each component ofsubsidence is examined separately below to quantifysubsidence corrections under both models.

4.3.1. Thermal coolingThe amount of subsidence attributable solely to

lithospheric cooling since initial rifting ∼ 90 Ma canbe estimated with the 1-D thermal model of McKenzie(1978), which depends largely on the amount of syn-riftlithospheric stretching. Wilson (1985) determined aCanterbury/Otago margin extension factor (β) of 1.8from the subsidence history of the 3100m-deep Galleon-1 borehole located ∼ 100 km north of the study area.Using 1.8 for β, 125 km for lithospheric thickness,3300 kg/m3 for mantle density, 3.28×10−5 °C for thethermal expansion coefficient of the mantle and crust,and 1333 °C for asthenospheric temperature in thesediment-filled McKenzie model yields a prediction ofonly∼ 165m of thermal subsidence over the past 10Ma,or a rate of 0.017 m/ka (Table 2). This subsidence rate isconsidered a maximum since elasticity and 2-D thermaleffects significantly reduce the 1-D approximation.Thus, lithospheric cooling contributed b2 m of accom-modation per Late Quaternary 4th-order (∼ 120 ka)sequence. Such low thermal subsidence values areexpected (McKenzie, 1978) since the age of the Otagomargin exceeds 60 Ma (e.g. Molnar et al., 1975).

4.3.2. Sediment compactionCompaction of the Cretaceous–Pleistocene sediment

column caused by the deposition of an overlying 40 m-thick Late Pleistocene sequence (equivalent to Sequence2+3) can be estimated with the general decompactionequation (Eq. (17) in Sclater and Christie, 1980):

z2V−z1V¼ z2−z1−f0c½e−cz1−e−cz2 � þ f0

c½e−cz1V−e−cz2V�

where z1′ and z2′ are the compaction-corrected upper andlower depths of a sediment block, z1 and z2 are theuncorrected depths, f0 is the lithology-dependent surfaceporosity (0.5–0.6), and c is a constant representing theslope of the depth–porosity curve (0.3–0.5). For thesecalculations, the Cretaceous–Holocene shelf sequencewas divided into four temporal–lithologic units based ondata from the Galleon-1 borehole (Wilson, 1985):Cretaceous breccia, Paleocene–Oligocene sandstoneand mudstone, Miocene sandstone and limestone, andPlio-Pleistocene sand and mud. By using porosity datafrom the Galleon-1 borehole (Wilson, 1985) and theranges for c and f0 listed above, ∼ 10–15 m ofcompaction-induced subsidence is modeled from thecombined deposition of Sequences 2 and 3. Notsurprisingly, more than 75% of this compaction occursin the Miocene–Pleistocene portion of the sedimentcolumn in this model. The Galleon-1 well log shows thatthe Miocene–Recent sediment is not overpressured(Wilson, 1985), suggesting that compaction does notsignificantly lag burial. Thus, compaction-inducedsubsidence rates of ∼ 0.1 and ∼ 0.05 m/ka are estimatedfor Models 1 and 2, respectively (Table 2). These ratesare considered minimums because underlying sedimentmay have been irreversible compacted before Sequences2 and 3 were subaerially eroded (during SB formation) totheir present-day reduced thicknesses.

Table 3Otago shelf paleoshoreline subsidence corrections under chronostratigraphic Model 1

Paleo-shoreline

Modern depth Model 1 age Total subsidence rate range Subsidence correction range Corrected paleoshoreline depth range(m bmsl) (ka) (m/ka) (m) (m bmsl)

PS1 133–134 15–25 (stage 2) 0.1–0.34 1.5–8.5 124.5–132.5PS2 121–122 60–70 (stage 4) 0.1–0.34 6.0–23.8 97.2–116.0PS3 157–159 135–145 (stage 6) 0.1–0.34 13.5–49.3 107.7–145.5

bmsl=below mean sea-level.


4.3.3. Sediment, ice and water loadingSubsidence induced by sediment loading is more

difficult to quantify because of uncertainty about thespecific mechanisms, spatial distribution and timing oflithospheric compensation to an applied load. If Airyisostasy is assumed then ∼ 20–24 m of load-inducedsubsidence is modeled under a 40 m-thick sediment load(equivalent to Sequences 2+3) when the mantle andsediment have a density contrast of 900–1100 kg/m3

(Steckler and Watts, 1978). Subsidence of 20–24 mduring deposition of Sequences 2 and 3 is equivalent to asubsidence rate of 0.17–0.2 m/ka under Model 1 and0.09–0.1 m/ka under Model 2 (Table 2). If an infiniteelastic lithosphere is assumed (Te=20–30 km, Young'smodulus=7×1010 Pa, Poisson's ratio=0.25; Holt andStern, 1991) then lithospheric flexure beneath a 40 m-thick, 25 km-wide sediment load ranges from b0.2m in aperiodic load flexure model (∼ 0 m/ka) to∼ 9–11 m in aline load flexure model (∼ 0.8 and ∼ 0.4 m/ka underModels 1 and 2; Table 2). Although these 1-D loadingmodels do not account for the flexure induced by long-wavelength onshore topography (e.g. Southern Alps),they are useful for defining a potential range ofsubsidence rates caused by sediment loading. A morecomprehensive 2-D finite-difference flexural model (e.g.Holt and Stern, 1991) is beyond the scope of this paper.

New Zealand alpine glaciers expanded considerablyduring glacial periods (Suggate, 1990; Carter andGammon, 2004), but they were too small to produce asignificant isostatic response on the Otago margin(Porter, 1975). Coupled ice–ocean load models suggestthat glacial-age continental ice sheets created a minimalisostatic signal at far-field sites (Lambeck et al., 2000,2002). However, contrasting water loads over the

Table 4Otago shelf paleoshoreline subsidence corrections under chronostratigraphic

Paleo-shoreline

Modern depth Model 2 age Total subsidence rate range(m bmsl) (ka) (m/ka)

PS1 133–134 15–25 (stage 2) 0.06–0.19PS2 121–122 135–145 (stage 6) 0.06–0.19PS3 157–159 250–275 (stage 8) 0.06–0.19

bmsl=below mean sea-level.

world's ocean basins and continental shelves maycause far-field Late Quaternary lowstand paleoshore-lines to appear shallower than predicted from global sea-level curves (Lambeck et al., 2002). This apparentshallowing discrepancy at far-field sites decreasesoffshore from a maximum of ∼ 20 m along moderncoastlines to ∼ 5 m along continental shelf breaks(Yokoyama et al., 2001a). Because Otago paleoshorelinedepths are measured on the outer shelf where the waterload correction is minimal, they are not corrected forglacio-hydro-isostatic effects. The sea-level envelopeconstituents derived from Bonaparte Gulf sediments,Huon Peninsula terraces and Vanuatu terraces (all far-field sites) are likewise uncorrected, and those derivedfrom foraminiferal oxygen isotope ratios require no suchcorrection (see Yokoyama et al., 2001a,b for glacio-hydro-isostatic corrections for the Huon Peninsula andBonaparte Gulf sea-level curves).

4.3.4. Subsidence summationThe total estimated subsidence rate is heavily

dependent on which isostatic model is favored in thesediment loading calculations (Table 2), with Airyisostasy providing the upper limit and periodic loadflexure providing the lower limit of the total range. Thesubsidence rate estimates include both conservative andextreme values, and therefore the true subsidence rate isunlikely to fall outside of these ranges. These results,combined with Gibb's (1986) assessment of tectonicstability along the shoreline west of the study area,suggest that the Otago continental shelf is tiling beneaththe weight of the marine sediment column with a pivotpoint near the present-day shoreline. Consequently, thesubsidence rates estimated here are probably only valid

Model 2

Subsidence correction range Corrected paleoshoreline depth range(m) (m bmsl)

0.9–4.7 128.3–133.18.1–27.5 93.5–113.915.0–52.2 104.8–144.0

Fig 8. Total range of sea-level envelope lowstands (shaded) comparedto Otago paleoshoreline depths calculated under Model 1 (a) andModel 2 (b) using the entire range of possible subsidence rates (lines).Note scale changes in both axes.

Fig. 7. Independent sea-level proxies from raised coral terraces, foraminiferal δ18O values and marine sediments comprising the sea-level envelope,compared to subsidence-corrected Model 1 and Model 2 lowstand paleoshoreline depths on the outer Otago shelf.


for the outer continental shelf where lowstand paleoshor-elines are located. Each lowstand paleoshoreline iscorrected for subsidence by multiplying its age undereach chronostratigraphic model by the range of potentialsubsidence rates, and then subtracting those values fromits present-day depth (Tables 3 and 4). The resultingpaleoshoreline depth ranges (under both models) arecompared graphically to the composite sea-level enve-lope in Fig. 7.

5. Discussion

5.1. Comparison of lowstand paleoshorelines to sea-level curves

The global sea-level proxies comprising the sea-levelenvelope reveal lowstands during stages 2 and 6 that eachreached∼ 115–140m depth, withmoremoderate sea levelfalls to ∼ 95–110 m depth during stage 4 and 85–120 mduring stage 8 (Fig. 7). Stage 4, 6 and 8 lowstand depths inthe SLE are exclusively defined by foraminiferal δ18O datafrom deep-sea cores, while that of stage 2 additionallyincludes data from Bonaparte Gulf sediments and a singlePapua New Guinea coral terrace. Subsidence-correctedOtago paleoshorelines consistently fall within the SLEunder Model 1, with lowstand depth ranges of 125–133 m(PS1), 97–116m (PS2) and 108–146m (PS3) for stages 2, 4and 6, respectively (Table 3; Fig. 7). Under Model 2,however, the lowstand paleoshoreline range assigned tostage 6 (PS2, 94–114 m) is considerably shallower thanstage 6 lowstands in the SLE (120–140 m), and that as-signed to stage 8 (PS3, 105–144 m) only partially overlapsstage 8 SLE lowstands (90–120m) (Table 4; Fig. 7). These

data substantially favor Model 1 over Model 2. Even ifabsolute paleoshoreline depths are ignored, Model 2 re-quires that the stage 6 lowstand on the Otago shelf was anaverage of 25 m shallower than those of stages 2 and8 (Table 4; Fig. 7). The SLE, however, reveals that the stage8 lowstand was the least extreme of the three (Fig. 7).


Fig. 8 displays the entire range of subsidence-corrected Otago paleoshoreline depths under Models 1and 2 compared to SLE lowstand depths for each stage.This figure emphasizes how much closer Otagopaleoshorelines follow global sea-level proxies underModel 1 than Model 2. None of the Model 2 subsidencerate estimates provide stage 6 paleoshoreline depths thatfall within the SLE (Fig. 8b). The best fit betweenModel 1 and the SLE (all three paleoshorelines withinthe SLE) is achieved when subsidence rates of 0.18–0.28 m/ka are used in depth corrections (Fig. 8a). Thisrange includes the upper end of subsidence ratesestimated by assuming line load flexure and the lowerend of those calculated by assuming Airy isostasy(Table 2, Fig. 8a). Using this reduced, best-fit subsidencerate range in chronostratigraphic Model 1, subsidence-corrected stage 2, 4 and 6 lowstand paleohorelines on theOtago shelf are located at 126–131, 101–111 and 116–135 m bmsl, respectively.

A subsidence rate of 0.18–0.28 m/ka brackets the0.2 m/ka subsidence rate proposed by Wellman (1979)for the Canterbury Plains, but is approximately half that(0.55 m/ka) proposed by Browne and Naish (2003) forthe Canterbury shelf based on the Resolution-1 petro-leum exploration well and ODP Site 1119 (Fig. 1a).However, the average sedimentation rate over the past250 ka at Site 1119 is 0.34 m/ka, closely matching the0.33 m/ka sedimentation rate for Otago Sequences 2+3under Model 1 (using an average thickness of 40 m and adepositional period of 120 ka), while the sedimentationrate over the past 4.2 Ma at the Galleon-1 well is 0.17 m/ka (Wilson, 1985). Although the sedimentation andtectonic characteristics of the Canterbury shelf differfrom those in the study area, these data neverthelesssuggest that the Model 1 sedimentation and subsidencerates are reasonable for this region.

Otago lowstand paleoshorelines under Model 1consistently fall at the deep limit of the SLE (Figs. 7and 8a). This is most likely attributable to errors indetermining the present-day depth of the paleoshor-elines, although a systematic error in isolating the sea-level component of the foraminiferal δ18O signal fromthe temperature component is also possible. An effortwas made to minimize the latter contingency byincluding data in the SLE from both benthic andplanktonic forams, and by including the Shackleton(2000) sea-level curve for which the temperaturecomponent was removed through a different analysisthan the other δ18O curves. Submerged barrier islands,stranded paleobeach faces and deltaic wedges, however,are all susceptible to erosion by submarine and subaerialprocesses, potentially providing paleoshoreline depth

assignments that are too deep. Thus, Otago paleoshor-elines represent maximum depths for each sea-levellowstand. Water loading effects could not be responsiblefor the offset because water loading at far-field sitescauses paleoshorelines to appear shallower than wouldbe expected from eustatic curves (Yokoyama et al.,2001a; Lambeck et al., 2002). Subsidence rate errors arealso unlikely to be the cause because the subsidence rateestimate would have to be increased to cause ashallowing of paleoshoreline depths, and it is doubtfulthat subsidence rates would exceed estimates based onan Airy isostatic model.

5.2. Stage 4 sea level

The major difference between the two models is thatModel 1 includes a stage 4 sequence boundary whileModel 2 does not. Although the Sulu Sea eustatic curveis the only component of the SLE that includes a stage 4lowstand deeper than –100 m, the V19-30 and ODP-677curves both indicate a stage 4 lowstand at 95 m depth,which is only 2 m shallower than the top of the Otagostage 4 paleoshoreline range in Model 1 and could beaccounted for by erosion of the stage 4 paleoshoreline asdiscussed above. Thus, the interpretation of a stage 4lowstand delta and sequence boundary in Model 1 isreasonable based on several components of the SLE.

A further way to evaluate the two models is to inves-tigate whether a stage 4 subaerial unconformity exists oncontinental shelves in New Zealand and beyond. Evidencesupporting both chronostratigraphic models can be foundwithin other New Zealand sequences. Nodder (1995)inferred the existence of a stage 4 unconformity on theTaranaki shelf (North Island) based on seismic data, datedsamples and predicted accumulation rates. However,chronological interpretations of the north (Barnes, 1995)and south (Browne and Naish, 2003) Canterbury shelfQuaternary sequences do not include a stage 4 sequenceboundary. Browne and Naish (2003) utilize estimated sub-sidence rates and lowstand paleoshoreline positions tomatch the Canterbury shelf sequences to a sea-level curvebased on foraminiferal δ18O values (similar to this study),but they see no evidence of a stage 4 SB or lowstand delta.Rather, they interpret the stage 4 lowstand as a minorinternal truncation surface within the stage 5-2 RSTdeposit, and they attribute an overlying slight landwardshift in onlap to the stage 3 highstand (Browne and Naish,2003). The difference in preservation of stage 4 LST andstage 3 HRST deposits on the Canterbury and Otago shel-ves may be due to the much steeper (3–4 times) gradient ofthe Otago shelf (Browne and Naish, 2003), or its lowersediment flux (Carter, 1986; Carter et al., 2004a).


A stage 4 sequence boundary has been frequentlyidentified in shelf sequences abroad, supporting chron-ostratigraphic Model 1. Seismic data and microfossilsamples from boreholes in the Lagniappe Delta complexon the Mississippi–Alabama shelf indicate that the twoyoungest sequence boundaries and deltaic lobes formedduring stages 2 and 4 (Kolla et al., 2000; Roberts et al.,2004). Extensive stage 4 incision associated with thedeposition of a shelf-edge delta has also been documen-ted on the Louisiana shelf (Suter et al., 1987). Seismicprofiles from the Spanish Gulf of Cadiz shelf sequence(Hernandez-Molina et al., 2000) display remarkablysimilar stratal architecture to the Otago shelf sequence.Although the Cadiz sequence has not been temporallyconstrained, correlation with dated Mediterraneansequences and sea-level curves provide evidence fordeposition of forced regressive wedges during stages 5and 3, separated by a fluvially incised stage 4 sequenceboundary (Hernandez-Molina et al., 2000).

The New Jersey continental shelf has been thesubject of extensive study over the past several decades,and interpretations of the chronostratigraphy haveevolved over time. The present consensus is thatlowstands during stages 2 and 6 created erosive shelf-wide sequence boundaries and thick shelf-edge deposits,while a thinner stage 3 progradational deposit isconfined to the middle shelf and underlain by an erosivestage 4 sequence boundary (Duncan et al., 2000; Careyet al., 2005). Taken together, the examples discussedabove suggest that the stage 4 lowstand was associatedwith extensive shelf exposure and the deposition of mid-outer shelf deltas globally. This clearly supports Otagochronostratigraphic Model 1 over Model 2.

5.3. Stage 3 sea level

HRST2 is interpreted as a shallow marine or strandlineunit, and is constrained to stage 3 in Model 1. If theseinterpretations are correct, the stage 3 highstand musthave been sufficiently high to allow marine deposition ofHRST3. HRST3 is discernible within the study area atdepths as shallow as ∼ 40 m bmsl, but cannot be tracedfurther landward due to the thickness of the overlyingHolocene sand wedge and acoustic interference frommultiples. Model 1 therefore requires a stage 3 highstandshoreline at least above the present-day −40 m isobath. Astage 3 highstand above 40 m depth is indicated by SuluSea δ18O data (∼27 m bmsl) and Vanuatu coral terracedata (22–40 m bmsl), but PNG terraces and foraminiferalδ18O ratios (V19–30 and site 677) suggest a lower stage 3highstand of 38–78 m bmsl (Fig. 6; Chappell andShackleton, 1986; Chappell et al., 1996; Linsley, 1996;

Shackleton, 2000; Cabioch and Ayliffe, 2001). This totalrange of stage 3 highstand depths (22–78 m) is the largestuncertainty associated with any highstand or lowstand inthe SLE over the last 120 ka (Fig. 7).

Additional evidence for a stage 3 highstand above40 m depth is found in North America, Europe andAustralia. Temporally-constrained (δ18O, biostratigra-phy and radiocarbon dates) Quaternary sequences on theTexas continental shelf indicate a ∼ 15 m bmsl stage 3maximum highstand (Rodriguez et al., 2000). The same−15 m stage 3 sea-level height was determined throughluminescence and biostratigraphic dating of raisedterraces in southern Italy (Mauz and Hassler, 2000).On the south coast of Australia, a maximum stage 3 sea-level elevation of −22 m is inferred from shallow marinesequences chronologically constrained by radiocarbonand amino acid racemization ages from fossil mollusks(Murray-Wallace et al., 1993; Cann et al., 2000).Wellneret al. (1993) identified a stage 3 barrier complex on theinner New Jersey shelf, providing a stage 3 maximumhighstand elevation of∼ 20m bmsl. This estimationmaybe significantly influenced by the New Jersey Shelf'sproximity to the Laurentide Ice Sheet during stage 3(Potter and Lambeck, 2004), but it nonetheless matchesestimations from the far-field sites described above. Thisevidence provides a precedent for stage 3 highstands 15–20 m above the level required under Model 1. It ispossible that an extensive seismic survey north and southof the study area would reveal a better location fortracing SB1 and SB2 landward to where they becomeamalgamated, providing a maximum stage 3 highstandshoreline on the Otago shelf to complement the globalstage 3 paleoshoreline dataset.

5.4. Forced regressive and lowstand deltaic deposits

Highstand/regressive systems tract unit HRST3 andlowstand outer shelf wedges LST1a, 2a and 3a are inter-preted as deltaic deposits based on their stratal archi-tecture. The presence of fluvially incised valleys andchannels at each sequence boundary confirms that riversmeandered across the exposed continental shelf duringperiods of lowered sea level, supporting the interpreta-tion of the deposits as deltaic. Unfortunately, the grid ofseismic reflection data is not fine enough to definitivelylink incised valleys from one profile to the next and tracethem back to their origin.Waitati River and Careys Creekcurrently flow into Blueskin Bay estuary located north ofthe Otago Peninsula directly up-dip from the study area(Fig. 1b). It is hypothesized that the Late Quaternaryequivalents of these rivers, perhaps combined with ariver flowing out of present-day Dunedin Harbor, were


the source of the deltaic deposits seen in the study area.Seismic Line 16, the closest profile to the Otago Penin-sula, shows a higher concentration of fluvially incisedchannels and valleys at its northern end, supporting thishypothesis.

Browne andNaish (2003) interpret widespreadRSTandLST deltaic deposition on the Canterbury shelf, but do notobserve LST fluvial incision due to the lower gradient ofthe shelf relative to the adjacent Canterbury Plains. Theofflapping, strongly progradational delta deposits on themiddle to outer Canterbury shelf (Browne andNaish, 2003)closely resemble those on the Otago shelf (e.g. HRST3).The Otago and Canterbury shelf deltaic deposits are funda-mentally different from the massive (up to 20 km wide and1 km thick), long-lived (3–7Ma) Canterbury Drifts, whichwere deposited at mid-slope (300–750 m) depths duringtheMiocene through Pliocene by the Southland Current (orits Miocene–Pliocene equivalent; Fulthorpe and Carter,1991; Carter et al., 2004b; Lu and Fulthorpe, 2004).

5.5. High-order sea-level and climate signals

Internal truncation surfaces within HRST3 are inter-preted as marine erosion surfaces created by either lobeswitching (Hart and Long, 1996; Kolla et al., 2000; Plintand Nummedal, 2000) or an abrupt lowering of wavebase (Plint and Nummedal, 2000). Lobe switching onlyoccurs during relative sea-level rise caused by eithereustatic transgression or sediment dewatering andcompaction during eustatic stillstand (Hart and Long,1996; Kolla et al., 2000). Fluvial systems incise andbecome entrenched in their channels during relative sea-level lowering (if shelf has a steeper gradient than thecoastal plain; Browne and Naish, 2003), preventing lobeswitching (Hart and Long, 1996; Posamentier andMorris, 2000). Thus, regardless of whether ITSsrepresent autocyclic lobe switching surfaces or allocyclicsea-level lowering surfaces, their presence suggests thatHRST3 was deposited during a series of higher-ordereustatic oscillations during a lower-order sea-level fall(Plint and Nummedal, 2000). In the preferred chronos-tratigraphic model, HRST3was deposited during stage 5,when seventh-order (20 ka, stages 5d,c,b,a) oscillationspunctuated the sixth-order (40 ka) sea-level fall fromstage 5e to stage 4 (Fig. 7).

Clastic, backstepping wedges assigned to TST1a andTST3 are interpreted as shallow marine units depositedduring brief periods of increased sediment flux and/orperiods of slower transgression, stillstand or temporaryregression during lower-order transgressions (Carter etal., 1985, 1986; Hernandez-Molina et al., 2000). Similardeposits at corresponding isostatically-corrected depths

on Australian, North American, European and Asiancontinental shelves suggests that global eustatic fluctua-tions, rather than local variations in sediment supply,may have been responsible for their formation (Carter etal., 1986; Hernandez-Molina et al., 2000). Brief (1–5 ka) but dramatic climate change events during LateQuaternary glacial and deglaciation periods are wellknown from paleoclimate proxy records in ice andsediment cores (e.g. Bond et al., 1993; Dansgaard et al.,1993). Although such events are potentially linked tothe deposition of discrete sedimentary deposits, morerigorous age control of the Otago sequences is requiredbefore a correlation can be investigated.

6. Conclusions

Two chronostratigraphic models of Late QuaternaryOtago shelf evolution are proposed through correlationwith sea-level curves from far-field sites around the globe.Model 1 includes a stage 4 sequence boundary (erosiveunconformity) and a stage 3 HST/RST deposit, whileModel 2 is based on the assumption that the stage 3 high-stand and stage 4 lowstand were only minor sea-levelfluctuations during the overall stage 5-2 regression. In bothmodels, Sequence 1 is constrained to stages 2 and 1 byradiocarbon dated shells from piston cores previouslycollectedwithin the study area (Carter et al., 1985).Model 1is considered more favorable for the following reasons:

1. Subsidence-corrected lowstand paleoshorelines onthe Otago shelf more closely match global sea-levelcurves under Model 1 than under Model 2.

2. A stage 4 sequence boundary (subaerial unconfor-mity), as required under Model 1, has been identifiedin chronologically constrained shelf sequences world-wide, implying that stage 4 included a significant sea-level lowstand.

3. Stage 3 highstand shoreline elevations of 15–22 mbmsl reported from North America, Europe andAustralia provide precedents for stage 3 sea-levelabove 40 m bmsl as required by Model 1.

Under this preferredmodel, lowstands during stages 2, 4and 6 reached 126–131, 101–111 and 116–135 m bmsl,respectively on the Otago shelf when corrected for sub-sidence with a best-fit range of 0.18–0.28 m/ka. Duringeach lowstand, landward-pinching deltaic wedges weredeposited on the outer Otago shelf, fluvially incised se-quence boundaries were produced on the subaerially ex-posed shelf, and submarine canyon heads incised furtherlandward through mass wasting. Incised channels andvalleys were infilled with interpreted back-barrier deposits


during stage 6-5, 4-3 and 2-1 transgressions, while back-stepping, shallow marine wedges were deposited duringhigh-order stillstands, regressions or periods of slowertransgression within these transgressive intervals. Volumi-nous, progradational units are interpreted as delta andstrandline deposits that accumulated on the shelf duringhighstand through falling sea level associatedwith stages 5-4 and 3-2, respectively. The older deltaic deposit includesinternal truncation surfaces created during seventh-ordersea-level fluctuations within stage 5 by either lobe swit-ching during brief stillstands or transgressions, or waveerosion during abrupt forced regressions.

Acknowledgements

This research was funded by a graduate scholarshipfrom the J. William Fulbright Scholarship Board and theUniversity of Otago. I am grateful to my advisors Dr.Charles Landis and Dr. Peter Koons for their insight,support and assistance. Captain Chris Spears and firstmate Keith Murphy of the R.V. Munida were invaluableduring the seismic data collection. Many thanks also toMike Trinder andDamianWalls for equipment assistance.Discussions with Dr. TimNaish, Dr. Lionel Carter and Dr.Robert Carter, and suggestions from an anonymousreviewer, helped to develop concepts in this paper.

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