Marine Geology 229 (2006) 159 – 178
www.elsevier.com/locate/margeo
Late Quaternary (marine isotope stages 6-1) seismic
sequence 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, includes
fluvially incised sequence boundaries eroded during marine isotope stages 2, 4 and 6, bounding three sequences subdivided into
lowstand, 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 of
independent global sea-level curves. A second chronostratigraphic model that lacks a stage 4 sequence boundary and distinct stage
3 highstand/forced regressive deposit is discounted because its subsidence-corrected paleoshoreline depths do not match lowstands
in 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, but
may be erroneously deep owing to erosion of paleoshoreline indicators. Clastic, shallow marine wedges and back-barrier valley-fill
deposits accumulated during the stage 6–5, 4–3 and 2–1 marine transgressions. Contrary to the traditional sequence stratigraphic
model, deltaic and strandline units deposited during highstand through falling sea level (stages 5–4, 3–2, 1) volumetrically
dominate Otago Late Quaternary sequences. Discrete units and minor marine erosion surfaces within the transgressive and
regressive systems tract deposits are interpreted as strong evidence for seventh- (∼ 20 ka) or higher-order sea-level fluctuations
influencing 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, New
Zealand, 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 active
sediment source regions to abyssal fans in the Pacific Ocean (Fig. 1a). Terrigenous sediment is shed from
the rapidly uplifting Southern Alps, transported to the
coast via large rivers, and deposited either on the continental shelf or in the submarine canyons feeding the
Bounty Channel and Fan Complex (Carter et al., 1985).
The sedimentological evolution of the margin since its
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E.C. Osterberg / Marine Geology 229 (2006) 159–178
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 of
ODP Site 1119 and the Galleon-1 and Endeavor-1 petroleum exploration boreholes is also shown. b) Study area bathymetric map showing the grid of
high-resolution seismic reflection profiles and the location of figures in this paper (bold lines).
inception ∼ 80 Ma (Molnar et al., 1975) has consequently depended on the transient interplay of the uplift
and erosion rates of the Southern Alps, the competence
and location of rivers draining the Alps, marine currents
associated with the Southland Front, and relative sea
level, all of which depend on global and local climate
conditions. Like most passive margins, Milankovitchorder (20–120 ka) sea-level fluctuations induced by the
waxing and waning of continental ice sheets have been
the primary control on the Late Quaternary (last
∼ 500 ka) evolution of the Otago margin. In this paper,
the sedimentary response to these eustatic changes is
investigated from a sequence stratigraphic perspective
based on newly acquired high-resolution seismic
reflection data from the Otago shelf.
Late Quaternary shelf sequences have increased our
understanding of the relationship between sedimentation, erosion and sea level change because they were
deposited during a period for which sea level has been
well constrained by independent proxy records (e.g.
Chappell and Shackleton, 1986; Shackleton et al., 1990).
High-resolution seismic reflection studies have found
that contrary to the traditional sequence stratigraphic
model, units deposited during falling relative sea level
often volumetrically dominate Late Quaternary shelf
sequences because of the longer duration of fifth(∼ 120 ka) and sixth-order (∼ 40 ka) glacio-eustatic
regressions compared to rapid transgressions (e.g.
Posamentier et al., 1992; Naish and Kamp, 1997;
Haywick, 2000; Browne and Naish, 2003). Such studies
E.C. Osterberg / Marine Geology 229 (2006) 159–178
161
have also described discrete sediment units and minor
erosion surfaces attributed to rapid (1–10 ka), climatedriven sea-level cycles superimposed on the dominant
Milankovitch-order glacio-eustatic signal (HernandezMolina et al., 2000; Hamberg and Nielsen, 2000; Plint
and Nummedal, 2000). The shelf sequences described
here provide examples of both patterns in a region that
has not been studied extensively.
While sea-level proxies can be valuable for understanding sedimentary responses to base level changes,
they can also provide chronological constraints on Late
Quaternary sequences in lieu of datable material from
cores or boreholes. Using sequence stratigraphic principles, highstand and lowstand shorelines are identifiable
on seismic profiles and can be compared to eustatic
curves after taking account of subsidence. This method
has been used to chronologically constrain shelf
sequences in New Zealand (Browne and Naish, 2003)
and Spain (Hernandez-Molina et al., 2000), and is used
here to define two chronostratigraphic models for Otago
shelf evolution over the past 250 ka, incorporating a
detailed analysis of subsidence in the study area. The two
models differ primarily in their interpretation of sea level
during oxygen isotope stages 3 and 4, which is not as
tightly constrained as during other periods of the Late
Quaternary, 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 model
requires that the stage 4 lowstand and stage 3 highstand
were significant sea-level fluctuations capable of
producing a complete sequence on the Otago shelf.
Waihemo fault system (Allan, 1990), and has been
relatively passive during the Late Quaternary. The outer
Otago shelf is incised by seven major submarine
canyons, including Papanui Canyon within the study
area (Fig. 1b), which have periodically channeled terrigenous sediment to the abyssal Bounty Channel and Fan
complex since the Early Miocene (Fig. 1a; Carter and
Carter, 1987).
Geophysical surveys and petroleum exploration boreholes on the Otago and Canterbury (located north of the
study area) shelves have revealed a 2–5 km-thick transgressive–regressive megasequence spanning ∼ 80 Ma,
the deposition of which was controlled primarily by the
evolution of the Pacific–Australian plate boundary
(Molnar et al., 1975; Carter, 1988; Fulthorpe and Carter,
1991; Fulthorpe et al., 1996). Jurassic schist basement
underlies a Late Cretaceous to Oligocene marine
transgressive sequence consisting of basal fluvial deposits
and coal measures, overlain by a deepening succession of
marine sandstone, limestone and mudstone (Wilson,
1985; Carter, 1988). Maximum transgression in the
Oligocene (∼ 30 Ma) is represented by a limestone unit,
followed by an eastward thickening, Miocene-to-present
regressive marine sediment wedge consisting predominantly 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 focus
of this paper.
2. Regional setting
The dominant hydrologic feature on the Otago continental margin is the Southland Front, which is an
extension of the global Subtropical Front separating
relatively warm and salty Subtropical Water on the
continental shelf from relatively cold and fresh Subantarctic Water offshore (Heath, 1972). Associated with the
Southland Front is the Southland Current, a northward
flow (∼ 8.3 Sv; ∼ 0.15–0.30 m/s) of dominantly
Subantarctic water (Chiswell, 1996; Sutton, 2003) that
has been responsible for the formation of the Canterbury
Drifts north of the study area since at least the Pliocene
(Fulthorpe and Carter, 1991; Carter et al., 2004b; Lu and
Fulthorpe, 2004). The local semidiurnal tide averages
1.5 m in height with an average velocity over the Otago
shelf of 0.1 to 0.2 m/s (Andrews, 1973; Carter et al.,
1985). Southerly gales mobilize sand on the middle and
outer shelf up to several times a year by accentuating
the Southland Current and producing storm swell capable of stirring sand at N 100 m depth (Andrews, 1973;
Carter and Herzer, 1979; Carter et al., 1985). Bed-load
2.1. Otago shelf description and Tertiary evolution
The modern Otago continental shelf stretches from
Shag 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 east
of the Otago Peninsula (a remnant Miocene volcanic
complex; Fig. 1b). South of the peninsula, most of the
Otago coastline is cliffed as a result of uplift along
northeast-trending (roughly shore-parallel), eastwarddipping reverse faults within the Otago reverse fault
province (Litchfield and Norris, 2000). North of the
peninsula, the northwest-trending, reverse Waihemo
fault system has uplifted Shag Point and sediments
further eastward on the continental shelf (Allan, 1990).
The study area discussed here lies northeast of the Otago
reverse fault province (Johnstone, 1990) and south of the
2.2. Hydraulic regime
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E.C. Osterberg / Marine Geology 229 (2006) 159–178
mobilization during calm weather is believed to be restricted to the surf zone where active longshore processes 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 sourced
primarily from the Clutha (pre-dam sediment flux of
∼ 3.2 million tons per annum (Mt a− 1)), Taieri (∼ 0.6 Mt
a− 1) and Tokamairiro (∼ 0.1 Mt a− 1) Rivers (Carter,
1986). The Clutha River drains Lakes Wanaka, Hawea
and Wakatipu nested within the high-relief outboard
(eastern) portion of the Southern Alps (Fig. 1a). All three
rivers drain catchments enriched in Mesozoic, low-grade
prehnite–pumpellyite and greenschist facies, and consequently have a predominantly quartzofeldspathic bedload with a high mica content (Andrews, 1973; Carter et
al., 1985).
Four, roughly shore-parallel surface sediment belts
have been identified in previous Otago shelf studies: a
fine, well-sorted, modern quartz sand facies from 0 to
∼ 55 m depth; a rounded, iron-stained, relict quartz gravel
facies from ∼ 55 to ∼ 70 m 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 m
depth (Andrews, 1973; Carter et al., 1985; Orpin et al.,
1998). The modern sand and biogenic gravel facies have
been interpreted as Holocene highstand deposits, while
the relict terrigenous sand and gravel facies have been
interpreted as post-glacial transgressive units deposited in
response to episodic changes in the rate of post-glacial
sea-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, highresolution seismic reflection data were collected in a grid
pattern over a 220 km2 area of the Otago continental
shelf and Papanui Canyon during 14 individual day-long
cruises on the University of Otago research vessel
Munida between March 9, 2000 and February 1, 2001
(Fig. 1b). Position fixes were acquired at a 5-s interval
using a Fugro OmniSTAR™ differential global positioning system. The seismic data were collected using a
Ferranti-ORE Geopulse Sub-bottom Profiling System™, and recorded digitally in raw form for reprocessing. In general, power levels of 175 and 350 J were used
over the continental shelf and Papanui Canyon, respectively, with a 0.5–2 pulse/s firing rate. Vessel survey
speed ranged from approximately 3–5 knots, resulting in
a ∼ 1.5–3 m seismic shot spacing interval. A FerrantiORE 5210A Receiver™ was used to process the raw
seismic data with flat gain, time-varied gain, bandpass
filter (500–2000 Hz) and swell filter. Seismic profiles
were all printed on an EPC 4800™ graphic recorder, for
which a constant seismic velocity of 1470 m/s was
utilized for all time–depth conversions.
3.2. Sequence stratigraphic interpretation
Seismic reflection data are interpreted in a sequence
stratigraphic framework using established procedures
(e.g. Mitchum and Vail, 1977; Vail, 1987). Regionally
extensive seismic horizons are interpreted as sequence
boundaries (SB), ravinement surfaces (RS), downlap
surfaces (DLS) and internal truncation surfaces (ITS)
based on their orientation, stratal relationship (lapout,
truncation), amplitude and continuity. Four systems
tracts are identified in each complete sequence: the
lowstand (LST), transgressive (TST), highstand (HST)
and regressive (RST) system tracts in ascending
stratigraphic order. The regressive systems tract (“falling
stage systems tract” of Hart and Long (1996) and Plint
and Nummedal (2000); “forced regressive systems tract”
of Hunt and Tucker (1992)), comprises sedimentary
units deposited during falling relative sea level (Naish
and Kamp, 1997; Browne and Naish, 2003). Where the
boundary between HST and RST deposits within a single
sequence is indiscernible, the HST and RST have been
combined into the highstand/regressive systems tract
(HRST). Sea-level lowstand paleoshorelines (PS) are
identified at the landward pinch-out depths of LST
deltaic wedges, at the crests of LST barrier islands, and at
the inflection of LST paleobeach faces (Section 4.1.1).
The sequence stratigraphic classification of forced
regressive deposits and their position relative to the
sequence boundary have been somewhat controversial
(Posamentier et al., 1992; Hunt and Tucker, 1992; Naish
and Kamp, 1997; Plint and Nummedal, 2000; Tesson et
al., 2000). Otago forced regressive deposits most closely
resemble the strongly progradational Wanganui basin
and Canterbury Shelf deposits assigned to the regressive
systems tract by Naish and Kamp (1997) and Browne
and Naish (2003), and the same terminology is adopted
here for simplicity. Under this labeling scheme, sequence
boundaries immediately overlie and may represent
significant erosion of regressive systems tract deposits,
which are therefore the youngest deposits in each
sequence (Naish and Kamp, 1997; Haywick, 2000;
E.C. Osterberg / Marine Geology 229 (2006) 159–178
Browne and Naish, 2003). Sequences are labeled with
Arabic numerals by increasing age (1 = youngest). Each
systems tract, unconformity and paraconformity is
denoted by its acronym with its sequence number in
subscript (e.g. SB1, TST2).
3.3. Chronostratigraphy
Developing a chronostratigraphic model for the
Otago continental shelf is difficult owing to the paucity
of Late Quaternary borehole and sediment core data
from the region. The Plio-Pleistocene portion of the
sediment column was discarded by petroleum borehole
contractors who drilled on the shelf (e.g. Galleon-1;
Wilson, 1985), and the few piston cores that were
collected by previous researchers only penetrated 0.5–
1 m into Holocene shelf sediments (Carter et al., 1985).
163
ODP Site 1119, located ∼ 175 km north of the study
area on the South Canterbury upper continental slope
(395 m below mean sea level (bmsl); Fig. 1a), is the
most accurately dated Quaternary sediment record on
the Otago or Canterbury margins (Carter and Gammon,
2004; Carter et al., 2004a,b; Lu and Fulthorpe, 2004; Lu
et al., 2005). Although the Site 1119 record provides a
valuable analog to Otago shelf sequences, its location
∼ 15 km beyond the shelf break and proximity to
sediment-laden rivers (e.g. Waitaki River) draining the
eastern flanks of the Southern Alps limit its usefulness
for developing a high-resolution chronostratigraphic
model for the Otago continental shelf.
Given these limitations, two chronostratigraphic
models of Otago shelf evolution are proposed by
correlating Otago shelf sequences to a compilation of
sea-level data from raised coral terraces at Huon
Table 1
Otago shelf stratigraphic column and chronostratigraphy under Models 1 and 2
Underlying Unit b
horizon a
DLS1
HST1
RS1
TST1a
TST1b
SB1
LST1a,b
DLS2
HRST2
RS2
TST2a
TST2b
SB2
DLS3
TST3
SB3
LST3a,b
HRST4
b
Thickness
Stratal
relationship
Depositional
environment
Seq. Model 1
age
Model 2
age
Shore-connected,
prograding wedge
Three mid-shelf
back-stepping
sediment wedges
Inner to mid-shelf
incised
channel fill
Outer shelf
prograding wedge
(a) and fill (b)
Inner to outer shelf
prograding wedge
Mid-shelf stranded
sediment lens
Inner to mid-shelf
incised channel fill
15–30 m
Low amp. clinoforms
downlapping DLS1
Transparent wedges
onlapping RS1
Holocene inner shelf
1
Stage 1
LST2a,b Outer shelf
prograding wedge
(a) and mound (b)
HRST3 Inner to outer shelf
prograding wedge
RS3
a
Modern
description
Three mid-shelf
back-stepping
sediment wedges
Outer shelf
prograding wedge
(a) and fill (b)
Inner to outer
shelf prograding wedge
5–8 m
5–15 m
Chaotic fill bounded
by SB1 below and RS1
above
a) 60+ m Onlapping SB1
b) 20 m
landward and offlapping
it seaward
10–20 m Low amp. clinoforms
downlapping DLS2
1–5 m
Transparent lens
onlapping RS2
5–15 m
Chaotic fill bounded
by SB2 below
and RS2 above
a) 5–30 m a) On/offlapping SB2
b) 60+ m b) fill capped by RS2
20–30 m
5–15 m
Moderate amp.
clinoforms
downlapping
DLS3; ITSsc
Transparent wedges
onlapping RS3
Inner shelf during
higher order sea-level
fluctuations
Estuary
Stage 2-1 Stage 2-1
Stage 2-1 Stage 2-1
Marine delta (a)
and barrier (b)
Strandline or inner shelf
Stage 2
2
Inner shelf during higher
order sea-level fluctuations
Estuary
Stage 2
Stage 3-2 Stage 5-2
Stage 4-3 Stage 6-5
Stage 4-3 Stage 6-5
Marine delta (a)
and canyon fill (b)
Marine delta
Stage 1
Stage 4
3
Inner shelf during
higher order sea-level
fluctuations
Marine delta (a)
and canyon fill (b)
a) 40+ m
b) 60 m
a) On/offlapping SB3
b) fill capped by RS3
20+ m
Low amp. clinoforms;
Marine delta or strandline? 4
downlap surface not seen
Stage 6
Stage 5-4 Stage 7-6
Stage 6-5 Stage 8-7
Stage 6
Stage 8
Stage 7-6 Stage 9-8
SB: sequence boundary, DLS: downlap surface, RS: ravinement surface. c ITSs: internal truncation surfaces.
HST: highstand systems tract, TST: transgressive systems tract, LST: lowstand systems tract, HRST: highstand/regressive systems tract.
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E.C. Osterberg / Marine Geology 229 (2006) 159–178
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.
E.C. Osterberg / Marine Geology 229 (2006) 159–178
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 tract deposits. Locations of lowstand paleoshorelines (PS)
are also noted. Note the massive canyon fill deposit LST2b.
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E.C. Osterberg / Marine Geology 229 (2006) 159–178
Peninsula, Papua New Guinea (PNG; Chappell and
Shackleton, 1986; Chappell et al., 1996) and Malakula,
Vanuatu (Cabioch and Ayliffe, 2001), sedimentary
deposits from the Bonaparte Gulf, Australia (Yokoyama
et al., 2001a), planktonic foraminiferal δ18O ratios from
ODP sites 768 and 769 in the Sulu Sea (Linsley, 1996),
benthic foraminiferal δ18O ratios from ODP site 677 in
the equatorial Pacific (Shackleton et al., 1990), and the
sea-level component of the V19–30 benthic foraminiferal δ18O signal (determined by extracting the temperature component through comparison with the Vostok
air δ18O record; Shackleton, 2000). The total sea-level
range indicated by the proxies through time defines a
“sea-level envelope” (SLE), and it is assumed that sea
level on the Otago shelf was within the SLE at any time
during the Late Quaternary.
Using the SLE as an independent sea-level proxy,
Otago shelf sedimentary sequences are counted back
from a temporally-constrained starting point by assuming that 1) sea level was the predominant control on
sedimentation and erosion, and 2) every significant sealevel cycle created a corresponding sequence that can be
identified on high-resolution seismic profiles. These
assumptions are justified for shelf sequences deposited
during the Late Quaternary, when the amplitudes of
dominant (∼ 120 ka period) sea-level cycles were greater
than 80 m (e.g. Imbrie et al., 1984) and instantaneous
rates of sea-level change often exceeded 10 mm/yr
(Lambeck et al., 2002). Though not a substitute for
absolute dating methods, this dating technique is useful
for developing a robust chronostratigraphic model when
absolute dates are sparse, and for providing a working
model to aid in the collection of future data.
4. Results
4.1. Sequence architecture
Sequence stratigraphic interpretation of the seismic
reflection data reveals portions of four sequences
separated by three regionally extensive sequence boundaries on the Otago shelf (Table 1, Figs. 2 and 3). Sequence
1 is presently being deposited, and will not be complete
until its upper sequence boundary is created during the
next major glacio-eustatic regression and lowstand.
Sequences 2 and 3 are both visible in full, but only the
uppermost portion of Sequence 4 is visible and will not be
discussed further. Sequence boundaries on the Otago shelf
are interpreted as diachronous unconformities that
progressively formed through erosive fluvial and subaerial processes during sea-level falls and lowstands.
Ravinement surfaces, created by shoreface wave erosion
during marine transgressions, are generally amalgamated
with subjacent sequence boundaries on seismic profiles,
creating a readily identifiable high-amplitude seismic
reflector. Sequence boundaries and ravinement surfaces
diverge, however, at fluvially incised valleys and channels
on the inner and middle shelf. There is no evidence on the
seismic profiles of active tectonism that would affect
sedimentary deposition.
Within the study area, Sequences 1, 2 and 3 display
uniform stratal architecture and thickness along strike,
with three exceptions. First, an infilled fossil Taiaroa
Canyon head visible on line 22 (Fig. 3) has eroded a
significant portion of Sequence 3 at mid-shelf depths in
the northern part of the study area. Multichannel (120channel) seismic reflection profiles collected by BP
Fig. 4. Uninterpreted (a) and interpreted (b) portion of seismic line 22 on the outer shelf showing architecture of lowstand deposits (medium gray),
including deltaic wedges (LST3a, LST2a), a paleobeach (LST1a) and paleobarrier deposit (LST1a). Note the erosional truncation of the LST2a wedge
by SB1, and the similarity in depth of the paleobeach inflection and barrier crest (PS1).
E.C. Osterberg / Marine Geology 229 (2006) 159–178
Shell Todd Services, Ltd. in 1982 and 1984 (not shown
here) reveal this feature to be the edge of a large (N4 km
wide, ∼ 400 m deep) region that has systematically been
eroded and infilled at the head of Taiaroa Canyon, with
the most recent cut and fill cycle occurring during
deposition of Sequence 2 (Fig. 3). Second, lowstand
deposits in the study area have been modified by the
presence of Papanui Canyon, with different LST
geometries north and south of the canyon (see 4.1.1
below), and the absence of LST1a at the head of the
canyon on lines 9 and 10 (Fig. 6). Third, the fluvially
incised channels show considerable variation in thickness and width, and an increased abundance towards the
northern part of the study area (e.g. line 22).
4.1.1. Lowstand systems tract deposits
Lowstand systems tract features on the outer Otago
shelf include three landward-pinching wedges, a
stranded barrier-like structure and infilled fossil canyon
heads. The three wedges (LST1a,2a,3a) display highangle (up to 5°), low-amplitude, oblique parallel clinoforms that offlap the subjacent sequence boundary
seaward and onlap it landward (Table 1; Figs. 2–4). This
aggredational/progradational architecture is indicative
of marine–deltaic deposition during peak lowstand and
earliest transgression and is commonly observed on
outer 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
167
located ∼ 4 km basinward from an interpreted
paleobeach deposit overlying SB1, which truncates
LST2a (Figs. 3 and 4). The stranded barrier and
paleobeach are assigned to LST1b. Line 22 also includes
a massive (60+ m thick), infilled, fossil Taiaroa Canyon
head within Sequence 2 (LST2b; Fig. 3), while an
infilled, fossil Papanui Canyon head is visible on several
profiles within Sequence 3 (LST3b; not shown). Sealevel regressions and lowstands promoted the incision of
Papanui and Taiaroa Canyon heads several kilometers
into 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 have
accumulated during late lowstand through early transgressive periods (Chiocci et al., 1997; Tesson et al.,
2000), and are assigned to the LST here for simplicity.
The identification of lowstand paleoshoreline positions on the Otago shelf is an essential part of the
correlation between shelf sequences and the sea-level
envelope. The three LST wedges belong to different
sequences (i.e. they were deposited during different
lowstands) and pinch out at different depths, but their
depths do not simply increase with age as might be
expected. LST2a, the intermediately-aged wedge (as
determined by stratigraphic principles), pinches out at
121–122 m depth (PS2), while the younger and older
wedges (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
Fig. 5. Uninterpreted (a) and interpreted (b) portion of seismic line 22 on the inner shelf showing channels incised during lowstand sequence boundary
(SB) formation, infilled with interpreted transgressive estuarine sediment (TST, light gray), capped by ravinement surfaces (RS), and overlain by
transgressive shallow marine wedges (dark gray) and/or subsequent highstand and forced regressive deposits (HRST, no fill).
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E.C. Osterberg / Marine Geology 229 (2006) 159–178
within 1–2 m on multiple seismic profiles. The crest of
the interpreted stranded barrier and the top of the
paleobeach face (LST1b) north of Papanui Canyon are
both found at 133–134 m depth (Fig. 4), equivalent to
the pinch-out depth of wedge LST1a south of the canyon
(Fig. 2). Thus, it appears that the most recent lowstand
reached an uncorrected (for subsidence) depth of 133–
134 m (PS1). During this lowstand a deltaic system
south of Papanui Canyon produced wedge LST1a, and a
barrier island system north of the canyon eroded
portions of LST2a and left remnants of the barrier and
back-barrier beach (LST1b).
4.1.2. Transgressive systems tract deposits
Transgressive systems tract deposits are preserved as
incised valley fill and back-stepping, middle-shelf
wedges (Table 1; (Figs. 2, 3 and 5)). Fluvially incised
valleys on the inner shelf generally range from 150 to
3000 m in width and 5–15 m in thickness, although the
width values are maximums because seismic lines are
most likely oblique to individual valleys. TST1b and
TST2b are interpreted as back-barrier, incised valley fill
deposits capped by transgressive ravinement surfaces
(RS1,2; Table 1, Fig. 5). The transgressive, back-barrier
origin of at least the uppermost portion of TST1b was
confirmed by piston core P152 collected by Carter et al.
(1985), which penetrated estuarine mud complete with
an in situ Austrovenus stutchburyi.
Thin (5–15 m thick), discrete, backstepping units
within TST1a, TST2a and TST3 are interpreted as
shallow marine deposits that accumulated as shoreconnected wedges during brief periods of increased
sediment flux and/or periods of slower transgression,
stillstand or temporary regression during lower-order
transgressions (Table 1, (Figs. 2, 3, 5 and 6); Carter et
al., 1985, 1986; Hernandez-Molina et al., 2000). The
units were probably reworked significantly by shelf
currents and storm waves during their submergence
(Carter et al., 1985, 1986; Goff et al., 2005). The
youngest backstepping units (TST1a) thin considerably
to the south and most likely represent the source
deposits for the relict gravel and sand facies (Carter et
al., 1985) on the shelf surface today.
4.1.3. Highstand/regressive systems tract
Highstand and regressive systems tract deposits
volumetrically dominate Sequences 1, 2 and 3 with
thicknesses of 10–30 m (Table 1; (Figs. 2, 3 and 6)).
Forced regressive deposits are characterized by the
progressively lower elevation of depositional clinoforms
Fig. 6. (a) Seismic Line 9 with sequence boundaries (SB), paleo-shorelines (PS), 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 identified. (b) Expanded section
of Line 9 showing prograding clinoforms and internal truncation surfaces (ITS) within HRST3, interpreted as a forced-regressive marine deltaic unit.
Note the lack of high-amplitude clinoforms within HRST2, interpreted as a strandline or inner shelf forced-regressive deposit.
E.C. Osterberg / Marine Geology 229 (2006) 159–178
towards the shelf break, the seaward-dipping orientation
of the overlying sequence boundary, and their stratigraphic position as strongly progradational deposits on
the middle to outer shelf (Posamentier and Morris, 2000;
Roberts et al., 2004). Otago shelf highstand deposits
gradually transition to forced regressive deposits with
no discernable boundary, and are consequently grouped
into the combined highstand/regressive systems tract
(HRST). The only HST unit in the study area without a
corresponding RST component is HST1, the shoreconnected Holocene sand wedge interpreted to have
prograded to a depth of 55–60 m during the present
highstand (∼ 6.5 ka to present; Carter et al., 1985; Carter
and Carter, 1986; (Figs. 2, 3 and 5)). Seaward of HST1
(60 m to at least 150 m depth), terrigenous sediment
starvation of the modern shelf is currently producing a
condensed section with a mid-cycle shell bed consisting
of modern biogenic growth amidst reworked relict shell
hash (Carter et al., 1985; Orpin et al., 1998).
HRST3 is characterized by a succession of oblique
tangential clinoforms and internal truncation surfaces
(Fig. 6) that exhibit a complex mounded structure on
strike profiles, implying a deltaic origin (Hart and Long,
1996; Tesson et al., 2000; Roberts et al., 2004). A deltaic
interpretation is supported by the presence of incised
fluvial valleys shoreward. The high-angle (1–2°)
portions of clinoforms are interpreted as sand- and
gravel-rich delta-front foresets, grading downdip to
mud-rich prodelta bottomset facies that downlap at very
low 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 erosion
surfaces within the deltaic unit created by either lobe
switching (Hart and Long, 1996; Kolla et al., 2000) or
an abrupt lowering of wave base (Plint and Nummedal,
2000).
In contrast to the high-amplitude ITSs, mounded
clinoforms and mud-rich pro-delta facies of deltaic
HRST3, HRST2 displays only sporadic, low-amplitude
oblique clinoforms (Figs. 2 and 3). This stratal architecture is characteristic of forced regressive marine deposition 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 the
former seaward edge of the shore-connected wedge or
lower shoreface as it prograded across the shelf over
time (Dominguez and Wanless, 1991). Thus, HRST2
probably formed in a depositional regime resembling
that of the study area today, with longshore currents
transporting sediment northwards from a proximal delta
located to the south.
169
4.2. Chronostratigraphic models
Two chronostratigraphic models displayed schematically in Table 1 comply with the interpretation of
seismic data from the Otago shelf. Sequence 1 is well
constrained to oxygen isotope stages 2 and 1 (oxygen
isotope stages (hereafter “stages”) referred to here are
those of Imbrie et al., 1984) by radiocarbon-dated faunal
samples (Carter et al., 1985), sedimentological evidence
from dredges and core samples (Andrews, 1973; Carter
et al., 1985), and stratigraphic principles, specifically:
1. A fossil A. stutchburyi recovered from estuarine mud
(TST1b) in core P152 has a radiocarbon age of
12,150 ± 300 yr BP (Carter et al., 1985), constraining
deposition to the stage 2/1 transgression;
2. shallow-water faunal specimens associated with the
surficial relict/palimpsest sand facies on the middleouter shelf similarly yield a post-glacial transgressive
radiocarbon age of 10,050 ± 250 yr BP (Carter et al.,
1985);
3. SB1 is the youngest sequence boundary within the
study area, and most likely represents the stage 2
lowstand erosion surface;
4. The 133–134 m pinch-out depth of the LST1a wedge
correlates well with stage 2 sea-level data from the
Bonaparte Shelf, PNG, foraminiferal δ18O ratios, and
with the estimated stage 2 lowstand depth of 135 m
along the Canterbury shelf break (Chappell and
Shackleton, 1986; Linsley, 1996; Shackleton et al.,
1990; Shackleton, 2000; Yokoyama et al., 2001a;
Browne and Naish, 2003).
Sequence 1 is the temporally-constrained starting
point for the countback method in both models. Carter et
al. (1985) further refined the stage 2–1 transgressive
Otago sequence by assigning radiocarbon ages to
specific surficial shelf deposits, and Gibb (1986)
published a New Zealand Holocene sea-level curve
based on radiocarbon dated coastal deposits. The seismic
data collected in this study support the interpretations of
Carter et al. (1985) and Gibb (1986), though it was not
possible to collect additional radiocarbon samples to test
their transgressive models more rigorously.
The fundamental difference between chronostratigraphic Model 1 and Model 2 is the interpretation of sea
level during stages 4 and 3 from approximately 70 to
40 ka. Model 1 is based on the assumptions that the
stage 4 lowstand was sufficient in duration and
amplitude to produce an identifiable sequence boundary
and lowstand wedge within the study area, and that the
stage 3 highstand produced a distinct HRST deposit.
170
E.C. Osterberg / Marine Geology 229 (2006) 159–178
Model 2, however, is based on the assumption that the
stage 4 lowstand and stage 3 highstand were only minor
sea-level excursions within the overall stage 5–2 marine
regression. Thus, SB2 was created during either the stage
4 (Model 1) or stage 6 (Model 2) lowstand (Table 1).
The two models are scrutinized and evaluated below
through stratigraphic analyses and a comparison of
subsidence-corrected paleoshorelines to the sea-level
envelope.
4.3. Quantifying subsidence to correct lowstand
paleoshorelines
Lowstand paleoshorelines are assigned at the crests
of interpreted barrier bars, at the inflection of paleobeach faces, and at the landward pinch-out depths of
lowstand wedges (Figs. 2–4). However, in order to
compare lowstand paleoshorelines to sea-level curves,
present-day paleoshoreline depths must be corrected for
subsidence. No comprehensive accounting of subsidence 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 tectonically
stable. Subsidence on the Otago margin is the result of
lithospheric cooling since rifting, sediment compaction,
and loading by sediment and water. Each component of
subsidence is examined separately below to quantify
subsidence corrections under both models.
4.3.1. Thermal cooling
The amount of subsidence attributable solely to
lithospheric cooling since initial rifting ∼ 90 Ma can
be estimated with the 1-D thermal model of McKenzie
(1978), which depends largely on the amount of syn-rift
lithospheric stretching. Wilson (1985) determined a
Canterbury/Otago margin extension factor (β) of 1.8
from the subsidence history of the 3100 m-deep Galleon1 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 the
thermal expansion coefficient of the mantle and crust,
and 1333 °C for asthenospheric temperature in the
sediment-filled McKenzie model yields a prediction of
only ∼ 165 m of thermal subsidence over the past 10 Ma,
or a rate of 0.017 m/ka (Table 2). This subsidence rate is
considered a maximum since elasticity and 2-D thermal
effects significantly reduce the 1-D approximation.
Thus, lithospheric cooling contributed b2 m of accommodation per Late Quaternary 4th-order (∼ 120 ka)
sequence. Such low thermal subsidence values are
expected (McKenzie, 1978) since the age of the Otago
margin exceeds 60 Ma (e.g. Molnar et al., 1975).
Table 2
Otago shelf subsidence rate estimates separated by component for each
chronostratigraphic model
Subsidence component
Subsidence rate
(m/ka)
Model 1
Model 2
Thermal cooling
Compaction
Sediment loading
Airy isostasy
Periodic load flexure
Line load flexure
Total subsidence rate
Total w/ periodic flexure
Total w/ line flexure
Total w/ airy Isostasy
0.017
0.08–0.13
0–0.20
0.17–0.20
∼ 0.001
0.08–0.09
0.10–0.34
0.10–0.14
0.18–0.23
0.27–0.34
0.017
0.04–0.07
0–0.10
0.09–0.10
∼ 0.0004
0.04–0.05
0.06–0.19
0.06–0.08
0.10–0.13
0.15–0.19
4.3.2. Sediment compaction
Compaction of the Cretaceous–Pleistocene sediment
column caused by the deposition of an overlying 40 mthick Late Pleistocene sequence (equivalent to Sequence
2 + 3) can be estimated with the general decompaction
equation (Eq. (17) in Sclater and Christie, 1980):
z2V−z1V¼ z2 −z1 −
f0 −cz1 −cz2
f0
½e −e þ ½e−cz1V−e−cz2V
c
c
where z1′ and z2′ are the compaction-corrected upper and
lower depths of a sediment block, z1 and z2 are the
uncorrected depths, f0 is the lithology-dependent surface
porosity (0.5–0.6), and c is a constant representing the
slope of the depth–porosity curve (0.3–0.5). For these
calculations, the Cretaceous–Holocene shelf sequence
was divided into four temporal–lithologic units based on
data from the Galleon-1 borehole (Wilson, 1985):
Cretaceous breccia, Paleocene–Oligocene sandstone
and mudstone, Miocene sandstone and limestone, and
Plio-Pleistocene sand and mud. By using porosity data
from the Galleon-1 borehole (Wilson, 1985) and the
ranges for c and f0 listed above, ∼ 10–15 m of
compaction-induced subsidence is modeled from the
combined deposition of Sequences 2 and 3. Not
surprisingly, more than 75% of this compaction occurs
in the Miocene–Pleistocene portion of the sediment
column in this model. The Galleon-1 well log shows that
the Miocene–Recent sediment is not overpressured
(Wilson, 1985), suggesting that compaction does not
significantly lag burial. Thus, compaction-induced
subsidence rates of ∼ 0.1 and ∼ 0.05 m/ka are estimated
for Models 1 and 2, respectively (Table 2). These rates
are considered minimums because underlying sediment
may have been irreversible compacted before Sequences
2 and 3 were subaerially eroded (during SB formation) to
their present-day reduced thicknesses.
E.C. Osterberg / Marine Geology 229 (2006) 159–178
171
Table 3
Otago shelf paleoshoreline subsidence corrections under chronostratigraphic Model 1
PaleoModern depth Model 1 age
shoreline (m bmsl)
(ka)
PS1
PS2
PS3
133–134
121–122
157–159
Total subsidence rate range Subsidence correction range Corrected paleoshoreline depth range
(m/ka)
(m)
(m bmsl)
15–25 (stage 2)
0.1–0.34
60–70 (stage 4)
0.1–0.34
135–145 (stage 6) 0.1–0.34
1.5–8.5
6.0–23.8
13.5–49.3
124.5–132.5
97.2–116.0
107.7–145.5
bmsl = below mean sea-level.
4.3.3. Sediment, ice and water loading
Subsidence induced by sediment loading is more
difficult to quantify because of uncertainty about the
specific mechanisms, spatial distribution and timing of
lithospheric compensation to an applied load. If Airy
isostasy is assumed then ∼ 20–24 m of load-induced
subsidence is modeled under a 40 m-thick sediment load
(equivalent to Sequences 2 + 3) when the mantle and
sediment have a density contrast of 900–1100 kg/m3
(Steckler and Watts, 1978). Subsidence of 20–24 m
during deposition of Sequences 2 and 3 is equivalent to a
subsidence rate of 0.17–0.2 m/ka under Model 1 and
0.09–0.1 m/ka under Model 2 (Table 2). If an infinite
elastic lithosphere is assumed (Te = 20–30 km, Young's
modulus = 7 × 1010 Pa, Poisson's ratio = 0.25; Holt and
Stern, 1991) then lithospheric flexure beneath a 40 mthick, 25 km-wide sediment load ranges from b 0.2 m in a
periodic load flexure model (∼ 0 m/ka) to ∼ 9–11 m in a
line load flexure model (∼ 0.8 and ∼ 0.4 m/ka under
Models 1 and 2; Table 2). Although these 1-D loading
models do not account for the flexure induced by longwavelength onshore topography (e.g. Southern Alps),
they are useful for defining a potential range of
subsidence rates caused by sediment loading. A more
comprehensive 2-D finite-difference flexural model (e.g.
Holt and Stern, 1991) is beyond the scope of this paper.
New Zealand alpine glaciers expanded considerably
during glacial periods (Suggate, 1990; Carter and
Gammon, 2004), but they were too small to produce a
significant isostatic response on the Otago margin
(Porter, 1975). Coupled ice–ocean load models suggest
that glacial-age continental ice sheets created a minimal
isostatic signal at far-field sites (Lambeck et al., 2000,
2002). However, contrasting water loads over the
world's ocean basins and continental shelves may
cause far-field Late Quaternary lowstand paleoshorelines to appear shallower than predicted from global sealevel curves (Lambeck et al., 2002). This apparent
shallowing discrepancy at far-field sites decreases
offshore from a maximum of ∼ 20 m along modern
coastlines to ∼ 5 m along continental shelf breaks
(Yokoyama et al., 2001a). Because Otago paleoshoreline
depths are measured on the outer shelf where the water
load correction is minimal, they are not corrected for
glacio-hydro-isostatic effects. The sea-level envelope
constituents derived from Bonaparte Gulf sediments,
Huon Peninsula terraces and Vanuatu terraces (all farfield sites) are likewise uncorrected, and those derived
from foraminiferal oxygen isotope ratios require no such
correction (see Yokoyama et al., 2001a,b for glaciohydro-isostatic corrections for the Huon Peninsula and
Bonaparte Gulf sea-level curves).
4.3.4. Subsidence summation
The total estimated subsidence rate is heavily
dependent on which isostatic model is favored in the
sediment loading calculations (Table 2), with Airy
isostasy providing the upper limit and periodic load
flexure providing the lower limit of the total range. The
subsidence rate estimates include both conservative and
extreme values, and therefore the true subsidence rate is
unlikely to fall outside of these ranges. These results,
combined with Gibb's (1986) assessment of tectonic
stability along the shoreline west of the study area,
suggest that the Otago continental shelf is tiling beneath
the weight of the marine sediment column with a pivot
point near the present-day shoreline. Consequently, the
subsidence rates estimated here are probably only valid
Table 4
Otago shelf paleoshoreline subsidence corrections under chronostratigraphic Model 2
PaleoModern depth Model 2 age
shoreline (m bmsl)
(ka)
PS1
PS2
PS3
133–134
121–122
157–159
Total subsidence rate range Subsidence correction range Corrected paleoshoreline depth range
(m/ka)
(m)
(m bmsl)
15–25 (stage 2)
0.06–0.19
135–145 (stage 6) 0.06–0.19
250–275 (stage 8) 0.06–0.19
bmsl = below mean sea-level.
0.9–4.7
8.1–27.5
15.0–52.2
128.3–133.1
93.5–113.9
104.8–144.0
172
E.C. Osterberg / Marine Geology 229 (2006) 159–178
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 paleoshorelines are located. Each lowstand paleoshoreline is
corrected for subsidence by multiplying its age under
each chronostratigraphic model by the range of potential
subsidence rates, and then subtracting those values from
its present-day depth (Tables 3 and 4). The resulting
paleoshoreline depth ranges (under both models) are
compared graphically to the composite sea-level envelope in Fig. 7.
data substantially favor Model 1 over Model 2. Even if
absolute paleoshoreline depths are ignored, Model 2 requires that the stage 6 lowstand on the Otago shelf was an
average of 25 m shallower than those of stages 2 and
8 (Table 4; Fig. 7). The SLE, however, reveals that the stage
8 lowstand was the least extreme of the three (Fig. 7).
5. Discussion
5.1. Comparison of lowstand paleoshorelines to sealevel curves
The global sea-level proxies comprising the sea-level
envelope reveal lowstands during stages 2 and 6 that each
reached ∼ 115–140 m depth, with more moderate sea level
falls to ∼ 95–110 m depth during stage 4 and 85–120 m
during stage 8 (Fig. 7). Stage 4, 6 and 8 lowstand depths in
the SLE are exclusively defined by foraminiferal δ18O data
from deep-sea cores, while that of stage 2 additionally
includes data from Bonaparte Gulf sediments and a single
Papua New Guinea coral terrace. Subsidence-corrected
Otago paleoshorelines consistently fall within the SLE
under Model 1, with lowstand depth ranges of 125–133 m
(PS1), 97–116 m (PS2) and 108–146 m (PS3) for stages 2, 4
and 6, respectively (Table 3; Fig. 7). Under Model 2,
however, the lowstand paleoshoreline range assigned to
stage 6 (PS2, 94–114 m) is considerably shallower than
stage 6 lowstands in the SLE (120–140 m), and that assigned to stage 8 (PS3, 105–144 m) only partially overlaps
stage 8 SLE lowstands (90–120 m) (Table 4; Fig. 7). These
Fig 8. Total range of sea-level envelope lowstands (shaded) compared
to Otago paleoshoreline depths calculated under Model 1 (a) and
Model 2 (b) using the entire range of possible subsidence rates (lines).
Note scale changes in both axes.
E.C. Osterberg / Marine Geology 229 (2006) 159–178
Fig. 8 displays the entire range of subsidencecorrected Otago paleoshoreline depths under Models 1
and 2 compared to SLE lowstand depths for each stage.
This figure emphasizes how much closer Otago
paleoshorelines follow global sea-level proxies under
Model 1 than Model 2. None of the Model 2 subsidence
rate estimates provide stage 6 paleoshoreline depths that
fall within the SLE (Fig. 8b). The best fit between
Model 1 and the SLE (all three paleoshorelines within
the SLE) is achieved when subsidence rates of 0.18–
0.28 m/ka are used in depth corrections (Fig. 8a). This
range includes the upper end of subsidence rates
estimated by assuming line load flexure and the lower
end of those calculated by assuming Airy isostasy
(Table 2, Fig. 8a). Using this reduced, best-fit subsidence
rate range in chronostratigraphic Model 1, subsidencecorrected stage 2, 4 and 6 lowstand paleohorelines on the
Otago 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 the
0.2 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) for
the Canterbury shelf based on the Resolution-1 petroleum exploration well and ODP Site 1119 (Fig. 1a).
However, the average sedimentation rate over the past
250 ka at Site 1119 is 0.34 m/ka, closely matching the
0.33 m/ka sedimentation rate for Otago Sequences 2 + 3
under Model 1 (using an average thickness of 40 m and a
depositional period of 120 ka), while the sedimentation
rate over the past 4.2 Ma at the Galleon-1 well is 0.17 m/
ka (Wilson, 1985). Although the sedimentation and
tectonic characteristics of the Canterbury shelf differ
from those in the study area, these data nevertheless
suggest that the Model 1 sedimentation and subsidence
rates are reasonable for this region.
Otago lowstand paleoshorelines under Model 1
consistently fall at the deep limit of the SLE (Figs. 7
and 8a). This is most likely attributable to errors in
determining the present-day depth of the paleoshorelines, although a systematic error in isolating the sealevel component of the foraminiferal δ18O signal from
the temperature component is also possible. An effort
was made to minimize the latter contingency by
including data in the SLE from both benthic and
planktonic forams, and by including the Shackleton
(2000) sea-level curve for which the temperature
component was removed through a different analysis
than the other δ18O curves. Submerged barrier islands,
stranded paleobeach faces and deltaic wedges, however,
are all susceptible to erosion by submarine and subaerial
processes, potentially providing paleoshoreline depth
173
assignments that are too deep. Thus, Otago paleoshorelines represent maximum depths for each sea-level
lowstand. Water loading effects could not be responsible
for the offset because water loading at far-field sites
causes paleoshorelines to appear shallower than would
be expected from eustatic curves (Yokoyama et al.,
2001a; Lambeck et al., 2002). Subsidence rate errors are
also unlikely to be the cause because the subsidence rate
estimate would have to be increased to cause a
shallowing of paleoshoreline depths, and it is doubtful
that subsidence rates would exceed estimates based on
an Airy isostatic model.
5.2. Stage 4 sea level
The major difference between the two models is that
Model 1 includes a stage 4 sequence boundary while
Model 2 does not. Although the Sulu Sea eustatic curve
is the only component of the SLE that includes a stage 4
lowstand deeper than –100 m, the V19-30 and ODP-677
curves both indicate a stage 4 lowstand at 95 m depth,
which is only 2 m shallower than the top of the Otago
stage 4 paleoshoreline range in Model 1 and could be
accounted for by erosion of the stage 4 paleoshoreline as
discussed above. Thus, the interpretation of a stage 4
lowstand delta and sequence boundary in Model 1 is
reasonable based on several components of the SLE.
A further way to evaluate the two models is to investigate whether a stage 4 subaerial unconformity exists on
continental shelves in New Zealand and beyond. Evidence
supporting both chronostratigraphic models can be found
within other New Zealand sequences. Nodder (1995)
inferred the existence of a stage 4 unconformity on the
Taranaki shelf (North Island) based on seismic data, dated
samples and predicted accumulation rates. However,
chronological interpretations of the north (Barnes, 1995)
and south (Browne and Naish, 2003) Canterbury shelf
Quaternary sequences do not include a stage 4 sequence
boundary. Browne and Naish (2003) utilize estimated subsidence rates and lowstand paleoshoreline positions to
match the Canterbury shelf sequences to a sea-level curve
based 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 minor
internal truncation surface within the stage 5-2 RST
deposit, and they attribute an overlying slight landward
shift in onlap to the stage 3 highstand (Browne and Naish,
2003). The difference in preservation of stage 4 LST and
stage 3 HRST deposits on the Canterbury and Otago shelves may be due to the much steeper (3–4 times) gradient of
the Otago shelf (Browne and Naish, 2003), or its lower
sediment flux (Carter, 1986; Carter et al., 2004a).
174
E.C. Osterberg / Marine Geology 229 (2006) 159–178
A stage 4 sequence boundary has been frequently
identified in shelf sequences abroad, supporting chronostratigraphic Model 1. Seismic data and microfossil
samples from boreholes in the Lagniappe Delta complex
on the Mississippi–Alabama shelf indicate that the two
youngest sequence boundaries and deltaic lobes formed
during stages 2 and 4 (Kolla et al., 2000; Roberts et al.,
2004). Extensive stage 4 incision associated with the
deposition of a shelf-edge delta has also been documented on the Louisiana shelf (Suter et al., 1987). Seismic
profiles from the Spanish Gulf of Cadiz shelf sequence
(Hernandez-Molina et al., 2000) display remarkably
similar stratal architecture to the Otago shelf sequence.
Although the Cadiz sequence has not been temporally
constrained, correlation with dated Mediterranean
sequences and sea-level curves provide evidence for
deposition of forced regressive wedges during stages 5
and 3, separated by a fluvially incised stage 4 sequence
boundary (Hernandez-Molina et al., 2000).
The New Jersey continental shelf has been the
subject of extensive study over the past several decades,
and interpretations of the chronostratigraphy have
evolved over time. The present consensus is that
lowstands during stages 2 and 6 created erosive shelfwide sequence boundaries and thick shelf-edge deposits,
while a thinner stage 3 progradational deposit is
confined to the middle shelf and underlain by an erosive
stage 4 sequence boundary (Duncan et al., 2000; Carey
et al., 2005). Taken together, the examples discussed
above suggest that the stage 4 lowstand was associated
with extensive shelf exposure and the deposition of midouter shelf deltas globally. This clearly supports Otago
chronostratigraphic Model 1 over Model 2.
Shackleton, 2000; Cabioch and Ayliffe, 2001). This total
range of stage 3 highstand depths (22–78 m) is the largest
uncertainty associated with any highstand or lowstand in
the SLE over the last 120 ka (Fig. 7).
Additional evidence for a stage 3 highstand above
40 m depth is found in North America, Europe and
Australia. Temporally-constrained (δ18O, biostratigraphy and radiocarbon dates) Quaternary sequences on the
Texas continental shelf indicate a ∼ 15 m bmsl stage 3
maximum highstand (Rodriguez et al., 2000). The same
− 15 m stage 3 sea-level height was determined through
luminescence and biostratigraphic dating of raised
terraces in southern Italy (Mauz and Hassler, 2000).
On the south coast of Australia, a maximum stage 3 sealevel elevation of − 22 m is inferred from shallow marine
sequences chronologically constrained by radiocarbon
and amino acid racemization ages from fossil mollusks
(Murray-Wallace et al., 1993; Cann et al., 2000). Wellner
et al. (1993) identified a stage 3 barrier complex on the
inner New Jersey shelf, providing a stage 3 maximum
highstand elevation of ∼ 20 m bmsl. This estimation may
be significantly influenced by the New Jersey Shelf's
proximity to the Laurentide Ice Sheet during stage 3
(Potter and Lambeck, 2004), but it nonetheless matches
estimations from the far-field sites described above. This
evidence provides a precedent for stage 3 highstands 15–
20 m above the level required under Model 1. It is
possible that an extensive seismic survey north and south
of the study area would reveal a better location for
tracing SB1 and SB2 landward to where they become
amalgamated, providing a maximum stage 3 highstand
shoreline on the Otago shelf to complement the global
stage 3 paleoshoreline dataset.
5.3. Stage 3 sea level
5.4. Forced regressive and lowstand deltaic deposits
HRST2 is interpreted as a shallow marine or strandline
unit, and is constrained to stage 3 in Model 1. If these
interpretations are correct, the stage 3 highstand must
have been sufficiently high to allow marine deposition of
HRST3. HRST3 is discernible within the study area at
depths as shallow as ∼ 40 m bmsl, but cannot be traced
further landward due to the thickness of the overlying
Holocene sand wedge and acoustic interference from
multiples. Model 1 therefore requires a stage 3 highstand
shoreline at least above the present-day −40 m isobath. A
stage 3 highstand above 40 m depth is indicated by Sulu
Sea δ18O data (∼27 m bmsl) and Vanuatu coral terrace
data (22–40 m bmsl), but PNG terraces and foraminiferal
δ18O ratios (V19–30 and site 677) suggest a lower stage 3
highstand of 38–78 m bmsl (Fig. 6; Chappell and
Shackleton, 1986; Chappell et al., 1996; Linsley, 1996;
Highstand/regressive systems tract unit HRST3 and
lowstand outer shelf wedges LST1a, 2a and 3a are interpreted as deltaic deposits based on their stratal architecture. The presence of fluvially incised valleys and
channels at each sequence boundary confirms that rivers
meandered across the exposed continental shelf during
periods of lowered sea level, supporting the interpretation of the deposits as deltaic. Unfortunately, the grid of
seismic reflection data is not fine enough to definitively
link incised valleys from one profile to the next and trace
them back to their origin. Waitati River and Careys Creek
currently flow into Blueskin Bay estuary located north of
the Otago Peninsula directly up-dip from the study area
(Fig. 1b). It is hypothesized that the Late Quaternary
equivalents of these rivers, perhaps combined with a
river flowing out of present-day Dunedin Harbor, were
E.C. Osterberg / Marine Geology 229 (2006) 159–178
the source of the deltaic deposits seen in the study area.
Seismic Line 16, the closest profile to the Otago Peninsula, shows a higher concentration of fluvially incised
channels and valleys at its northern end, supporting this
hypothesis.
Browne and Naish (2003) interpret widespread RSTand
LST deltaic deposition on the Canterbury shelf, but do not
observe LST fluvial incision due to the lower gradient of
the shelf relative to the adjacent Canterbury Plains. The
offlapping, strongly progradational delta deposits on the
middle to outer Canterbury shelf (Browne and Naish, 2003)
closely resemble those on the Otago shelf (e.g. HRST3).
The Otago and Canterbury shelf deltaic deposits are fundamentally different from the massive (up to 20 km wide and
1 km thick), long-lived (3–7 Ma) Canterbury Drifts, which
were deposited at mid-slope (300–750 m) depths during
the Miocene through Pliocene by the Southland Current (or
its 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 interpreted as marine erosion surfaces created by either lobe
switching (Hart and Long, 1996; Kolla et al., 2000; Plint
and Nummedal, 2000) or an abrupt lowering of wave
base (Plint and Nummedal, 2000). Lobe switching only
occurs during relative sea-level rise caused by either
eustatic transgression or sediment dewatering and
compaction during eustatic stillstand (Hart and Long,
1996; Kolla et al., 2000). Fluvial systems incise and
become entrenched in their channels during relative sealevel lowering (if shelf has a steeper gradient than the
coastal plain; Browne and Naish, 2003), preventing lobe
switching (Hart and Long, 1996; Posamentier and
Morris, 2000). Thus, regardless of whether ITSs
represent autocyclic lobe switching surfaces or allocyclic
sea-level lowering surfaces, their presence suggests that
HRST3 was deposited during a series of higher-order
eustatic oscillations during a lower-order sea-level fall
(Plint and Nummedal, 2000). In the preferred chronostratigraphic model, HRST3 was deposited during stage 5,
when seventh-order (20 ka, stages 5d,c,b,a) oscillations
punctuated the sixth-order (40 ka) sea-level fall from
stage 5e to stage 4 (Fig. 7).
Clastic, backstepping wedges assigned to TST1a and
TST3 are interpreted as shallow marine units deposited
during brief periods of increased sediment flux and/or
periods of slower transgression, stillstand or temporary
regression during lower-order transgressions (Carter et
al., 1985, 1986; Hernandez-Molina et al., 2000). Similar
deposits at corresponding isostatically-corrected depths
175
on Australian, North American, European and Asian
continental shelves suggests that global eustatic fluctuations, rather than local variations in sediment supply,
may have been responsible for their formation (Carter et
al., 1986; Hernandez-Molina et al., 2000). Brief (1–
5 ka) but dramatic climate change events during Late
Quaternary glacial and deglaciation periods are well
known from paleoclimate proxy records in ice and
sediment cores (e.g. Bond et al., 1993; Dansgaard et al.,
1993). Although such events are potentially linked to
the deposition of discrete sedimentary deposits, more
rigorous age control of the Otago sequences is required
before a correlation can be investigated.
6. Conclusions
Two chronostratigraphic models of Late Quaternary
Otago shelf evolution are proposed through correlation
with sea-level curves from far-field sites around the globe.
Model 1 includes a stage 4 sequence boundary (erosive
unconformity) and a stage 3 HST/RST deposit, while
Model 2 is based on the assumption that the stage 3 highstand and stage 4 lowstand were only minor sea-level
fluctuations during the overall stage 5-2 regression. In both
models, Sequence 1 is constrained to stages 2 and 1 by
radiocarbon dated shells from piston cores previously
collected within the study area (Carter et al., 1985). Model 1
is considered more favorable for the following reasons:
1. Subsidence-corrected lowstand paleoshorelines on
the Otago shelf more closely match global sea-level
curves under Model 1 than under Model 2.
2. A stage 4 sequence boundary (subaerial unconformity), as required under Model 1, has been identified
in chronologically constrained shelf sequences worldwide, implying that stage 4 included a significant sealevel lowstand.
3. Stage 3 highstand shoreline elevations of 15–22 m
bmsl reported from North America, Europe and
Australia provide precedents for stage 3 sea-level
above 40 m bmsl as required by Model 1.
Under this preferred model, lowstands during stages 2, 4
and 6 reached 126–131, 101–111 and 116–135 m bmsl,
respectively on the Otago shelf when corrected for subsidence with a best-fit range of 0.18–0.28 m/ka. During
each lowstand, landward-pinching deltaic wedges were
deposited on the outer Otago shelf, fluvially incised sequence boundaries were produced on the subaerially exposed shelf, and submarine canyon heads incised further
landward through mass wasting. Incised channels and
valleys were infilled with interpreted back-barrier deposits
176
E.C. Osterberg / Marine Geology 229 (2006) 159–178
during stage 6-5, 4-3 and 2-1 transgressions, while backstepping, shallow marine wedges were deposited during
high-order stillstands, regressions or periods of slower
transgression within these transgressive intervals. Voluminous, progradational units are interpreted as delta and
strandline deposits that accumulated on the shelf during
highstand through falling sea level associated with stages 54 and 3-2, respectively. The older deltaic deposit includes
internal truncation surfaces created during seventh-order
sea-level fluctuations within stage 5 by either lobe switching during brief stillstands or transgressions, or wave
erosion during abrupt forced regressions.
Acknowledgements
This research was funded by a graduate scholarship
from the J. William Fulbright Scholarship Board and the
University 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 first
mate Keith Murphy of the R.V. Munida were invaluable
during the seismic data collection. Many thanks also to
Mike Trinder and Damian Walls for equipment assistance.
Discussions with Dr. Tim Naish, Dr. Lionel Carter and Dr.
Robert Carter, and suggestions from an anonymous
reviewer, helped to develop concepts in this paper.
References
Allan, T.M.H., 1990. A high resolution seismic study of the offshore
extension of the Waihemo fault system. MSc. Thesis, Department of
Geology, University of Otago.
Andrews, P.B., 1973. Late Quaternary continental shelf sediments off
Otago Peninsula, New Zealand. N.Z. J. Geol. Geophys. 16, 793–830.
Barnes, P.M., 1995. High-frequency sequences deposited during Quaternary sea-level cycles on a deforming continental shelf, north Canterbury, New Zealand. Sediment. Geol. 97, 131–156.
Berryhill Jr., H.L., Suter, J.R., Hardin, N.S., 1986. Late Quaternary facies
and structure, northern Gulf of Mexico; interpretations from seismic
data. AAPG Stud. Geol. 23.
Bond, G., Broecker, W.S., Johnsen, S.J., McManus, J., Labeyrie, L.D.,
Jouzel, J., Bonani, G., 1993. Correlations between climate records from
North Atlantic sediments and Greenland ice. Nature 365, 143–147.
Browne, G.H., Naish, T.R., 2003. Facies development and sequence
architecture of a Late Quaternary fluvial–marine transition, Canterbury
plains and shelf, New Zealand: implications for forced regressive
deposits. Sediment. Geol. 158, 57–86.
Cabioch, G., Ayliffe, L.K., 2001. Raised coral terraces at Malakula, Vanuatu, southwest Pacific, indicate high sea level during marine isotope
stage 3. Quat. Res. 56, 357–365.
Cann, J.H., Belperio, A.P., Murray-Wallace, C.V., 2000. Late Quaternary
paleosealevels and paleoenvironments inferred from foraminifera, northern Spencer Gulf, South Australia. J. Foraminiferal Res. 30, 29–53.
Carey, J.S., Sheridan, R.E., Ashley, G.M., Uptegrove, J., 2005. Glaciallyinfluenced Late Pleistocene stratigraphy of a passive margin; New
Jersey's record of the North American ice sheet. Mar. Geol. 218,
155–173.
Carter, L., 1986. A budget for modern-Holocene sediment on the south
Otago continental shelf. N.Z. J. Mar. Freshw. Res. 20, 665–676.
Carter, R.M., 1988. Plate boundary tectonics, global sea-level changes and
the development of the eastern South Island continental margin, New
Zealand, southwest Pacific. Mar. Petrol. Geol. 5, 90–103.
Carter, L., Carter, R.M., 1986. Holocene evolution of the nearshore sand
wedge, south Otago continental shelf, New Zealand. N.Z. J. Geol.
Geophys. 29, 413–424.
Carter, R.M., Carter, L., 1987. The Bounty Channel System: a 55-million
year old sediment conduit to the deep sea, southwest Pacific Ocean.
Geo. Mar. Lett. 7, 183–190.
Carter, R.M., Gammon, P., 2004. New Zealand maritime glaciation: millennial-scale southern climate change since 3.9 Ma. Science 304,
1659–1662.
Carter, L., Herzer, R.H., 1979. The hydraulic regime and its potential to
transport sediment on the Canterbury continental shelf. N.Z. Oceanog.
Inst. Mem. 83.
Carter, R.M., Carter, L., Williams, J.J., Landis, C.A., 1985. Modern and
relict sedimentation on the south Otago continental shelf, New Zealand.
N.Z. Oceanog. Inst. Mem. 93.
Carter, R.M., Carter, L., Johnson, D.P., 1986. Submergent shorelines in the
SW Pacific: evidence for an episodic post-glacial transgression. Sedimentology 33, 629–649.
Carter, R.M., Gammon, P., Millwood, L., 2004a. Glacial–interglacial (MIS
1–10) migrations of the Subtropical Front across ODP Site 1119,
Canterbury Bight, southwest Pacific Ocean. Mar. Geol. 205, 29–58.
Carter, R.M., Fulthorpe, C.S., Lu, H., 2004b. Canterbury Drifts at Ocean
Drilling Program Site 1119, New Zealand: climatic modulation of
southwest Pacific intermediate water flows since 3.9 Ma. Geology 32,
1005–1008.
Chappell, J., Shackleton, N.J., 1986. Oxygen isotopes and sea level. Nature
324, 137–140.
Chappell, J., Omua, A., Esat, T., McCulloch, M., Pandolfi, J., Ota, Y.,
Pillans, B., 1996. Reconciliation of Late Quaternary sea levels derived
from coral terraces at Huon Peninsula with deep sea oxygen isotope
records. Earth Planet. Sci. Lett. 141, 227–236.
Chiocci, F.L., Ercilla, G., Torres, J., 1997. Middle–Late Pleistocene stratal
architecture of western Mediterranean margins as the result of the
stacking of lowstand deposits below ‘glacio-eustatic fluctuation baselevel’. Sediment. Geol. 112, 195–217.
Chiswell, S.M., 1996. Variability in the Southland Current, New Zealand.
N.Z. J. Mar. Freshw. Res. 30, 1–17.
Dansgaard, W., Johnsen, S.J., Clausen, H.B., Dahl-Jensen, D.,
Gundestrup, N.S., Hammer, C.U., Hvidberg, C.S., Steffensen, J.P.,
Sveinbjornsdottir, A.E., Jouzel, J., Bond, G., 1993. Evidence for
general instability of past climate from a 250-kyr ice-core record.
Nature 364, 218–220.
Dominguez, J.M.L., Wanless, H.R., 1991. Facies architecture of a falling
sea-level strandplain, Doce River coast, Brazil. In: Swift, D.J.P.,
Oertel, G.F., Tillman, R.W., Thorne, J.A. (Eds.), Shelf Sand and
Sandstone Bodies: Geometry, Facies and Sequence Stratigraphy.
Int. Assoc. Sediment. Spec. Pub., vol. 14, pp. 259–281.
Duncan, C.S., Goff, J.A., Austin, J.A., Fulthorpe, C.S., 2000. Tracking the
last sea-level cycle: seafloor morphology and shallow stratigraphy of
the latest Quaternary New Jersey middle continental shelf. Mar. Geol.
170, 395–421.
Fulthorpe, C.S., Carter, R.M., 1991. Continental shelf progradation by
sediment drift accretion. Bull. Geol. Soc. Am. 103, 300–309.
Fulthorpe, C.S., Carter, R.M., Miller, K.G., Wilson, J., 1996. Marshall
paraconformity; a Mid-Oligocene record of inception of the
E.C. Osterberg / Marine Geology 229 (2006) 159–178
Antarctic circumpolar current and coeval glacio-eustatic lowstand?
Mar. Petrol. Geol. 13, 61–77.
Gibb, J.G., 1986. A New Zealand regional Holocene eustatic sea-level
curve and its application to determination of vertical tectonic movements: a contribution to IGCP-Project 200. R. Soc. N.Z. Bull. 24,
377–395.
Goff, J.A., Austin Jr., J.A., Gulick, S., Nordfjord, S., Christensen, B.,
Sommerfield, C., Olson, H., Alexander, C., 2005. Recent and modern
marine erosion on the New Jersey shelf. Mar. Geol. 216, 275–296.
Hamberg, L., Nielsen, H.L., 2000. Shingled, sharp-based shoreface
sandstones: depositional response to stepwise forced regression in a
shallow basin, Upper Triassic Gassum Formation, Denmark. In:
Hunt, D., Gawthorpe, R.L. (Eds.), Sedimentary Responses to Forced
Regression. Geol. Soc., London, Spec. Pubs., vol. 172, pp. 69–89.
Hart, B.S., Long, B.F., 1996. Forced regressions and lowstand deltas:
Holocene Canadian examples. J. Sediment. Res. 66, 820–829.
Haywick, D.W., 2000. Recognition and distinction of normal and forced
regression in cyclothemic strata: a Plio-Pleistocene case study from
eastern North Island, New Zealand. In: Hunt, D., Gawthorpe, R.L.
(Eds.), Sedimentary Responses to Forced Regression. Geol. Soc.,
London, Spec. Pubs., vol. 172, pp. 193–215.
Heath, R.A., 1972. The Southland current. N.Z. J. Mar. Freshw. Res. 6,
497–533.
Hernandez-Molina, F.J., Somoza, L., Lobo, F., 2000. Seismic stratigraphy of the Gulf of Cadiz continental shelf: a model for Late
Quaternary very high-resolution sequence stratigraphy and response
to sea-level fall. In: Hunt, D., Gawthorpe, R.L. (Eds.), Sedimentary
Responses to Forced Regression. Geol. Soc., London, Spec. Pubs.,
vol. 172, pp. 329–362.
Holt, W.E., Stern, T.A., 1991. Sediment loading on the Western Platform
of the New Zealand continent: implications for the strength of the
continental margin. Earth Planet. Sci. Lett. 107, 523–538.
Hunt, D., Tucker, M.E., 1992. Stranded parasequences and the forced
regressive systems tract: deposition during base-level fall.
Sediment. Geol. 81, 1–9.
Imbrie, J., Hayes, J.D., Martinson, D.G., McIntyre, A., Mix, A.C.,
Morley, J.J., Pisias, N.G., Prell, W.L., Shackleton, N.J., 1984. The
orbital theory of Pleistocene climate: support from a revised
chronology of the marine d18O record. In: Berger, A., Imbrie, J.,
Hayes, J.D., Kukla, G., Satzmann, B.D. (Eds.), Milankovitch and
Climate — Part 1. Reidel Publishing Company, Dordrecht,
Holland, pp. 269–305.
Johnstone, T., 1990. A high resolution seismic study of offshore
southeast Otago. MSc Thesis, Department of Geology, University
Of Otago.
Kolla, V., Biondi, P., Long, B., Fillon, R., 2000. Sequence stratigraphy
and architecture of the Lake Pleistocene Lagniappe delta complex,
northeast Gulf of Mexico. In: Hunt, D., Gawthorpe, R.L. (Eds.),
Sedimentary Responses to Forced Regression. Geol. Soc., London,
Spec. Pubs., vol. 172, pp. 291–327.
Lambeck, K., Yokoyama, Y., Johnston, P., Purcell, A., 2000. Global
ice volumes at the Last Glacial Maximum and Early Lateglacial.
Earth Planet. Sci. Lett. 181, 513–527.
Lambeck, K., Yokoyama, Y., Purcell, A., 2002. Into and out of the Last
Glacial Maximum: sea-level change during oxygen isotope stages
3 and 2. Quat. Sci. Rev. 21, 343–360.
Linsley, B.K., 1996. Oxygen-isotope record of sea level and climate
variations in the Sulu Sea over the past 150,000 years. Nature 380,
234–237.
Litchfield, N., Norris, R.J., 2000. Holocene motion on the Akatore
Fault, south Otago coast, New Zealand. N.Z. J. Geol. Geophys. 43,
405–418.
177
Lu, H., Fulthorpe, C.S., 2004. Controls on sequence stratigraphy of a
Middle Miocene–Holocene, current-swept, passive margin: offshore Canterbury Basin. N.Z. Geol. Soc. Am. Bull. 116,
1345–1366.
Lu, H., Fulthorpe, C.S., Mann, P., Kominz, M.A., 2005. Miocene–
Recent tectonic and climate controls on sediment supply and
sequence stratigraphy: Canterbury basin, New Zealand. Basin Res.
17, 311–328.
Mauz, B., Hassler, U., 2000. Luminescence chronology of Late
Pleistocene raised beaches in southern Italy: new data of relative
sea-level changes. Mar. Geol. 170, 187–203.
McKenzie, D., 1978. Some remarks on the development of
sedimentary basins. Earth Planet. Sci. Lett. 40, 25–32.
Mitchum Jr., R.M., Vail, P.R., 1977. Seismic stratigraphy and global
changes of sea level. Part 7. Seismic stratigraphic interpretation
procedure. In: Payton, C.E. (Ed.), Seismic Stratigraphy; Applications to Hydrocarbon Exploration. Mem. Am. Assoc. Petrol. Geol.,
vol. 96, pp. 135–143.
Molnar, P., Atwater, T., Mammerycx, J., Smith, S.M., 1975. Magnetic
anomalies, bathymetry and the tectonic evolution of the South Pacific
since later Cretaceous. R. Astron. Soc. Geophys. J. 40, 383–420.
Murray-Wallace, C.V., Belperio, A.P., Gostin, V.A., Cann, J.H., 1993.
Amino acid racemization and radiocarbon dating of interstadial
marine strata (oxygen isotope stage 3), Gulf St. Vincent, South
Australia. Mar. Geol. 110, 83–92.
Naish, T., Kamp, P., 1997. Sequence stratigraphy of sixth-order (41 k.y.)
Pliocene–Pleistocene cyclothems, Wanganui Basin, New Zealand: a
case for the regressive systems tract. Geol. Soc. Am. Bull. 109,
978–999.
Nodder, S.D., 1995. Late Quaternary transgressive/regressive
sequences from Taranaki continental shelf, western New Zealand.
Mar. Geol. 123, 187–214.
Orpin, A.R., Gammon, P.R., Naish, T.R., Carter, R.M., 1998. Modern
and ancient Zygochlamys delicatula shellbeds in New Zealand, and
their sequence stratigraphic implications. Sediment. Geol. 122,
267–284.
Plint, A.G., Nummedal, D., 2000. The falling stage systems tract:
recognition and importance in sequence stratigraphic analysis. In:
Hunt, D., Gawthorpe, R.L. (Eds.), Sedimentary Responses to Forced
Regression. Geol. Soc., London, Spec. Pubs., vol. 172, pp. 1–17.
Porter, S.C., 1975. Glaciation limit in New Zealand's Southern Alps.
Arct. Alp. Res. 7, 33–37.
Posamentier, H.W., Morris, W.R., 2000. Aspects of the stratal
architecture of forced regressive deposits. In: Hunt, D., Gawthorpe,
R.L. (Eds.), Sedimentary Responses to Forced Regression. Geol.
Soc., London, Spec. Pubs., vol. 172, pp. 19–46.
Posamentier, H.W., Allen, G.P., James, D.P., Tesson, M., 1992. Forced
regressions in a sequence stratigraphic framework: concepts,
examples and exploration significance. Am. Assoc. Petrol. Geol.
Bull. 76, 1687–1709.
Potter, E.K., Lambeck, K., 2004. Reconciliation of sea-level observations in the western Atlantic during the last glacial cycle. Earth
Planet. Sci. Lett. 217, 171–181.
Roberts, H.H., Fillon, R.H., Kohl, B., Robalin, J.M., Sydow, J.C.,
2004. Depositional architecture of the Lagniappe Delta; sediment
characteristics, timing of depositional events, and temporal
relationship with adjacent shelf-edge deltas. Soc. Sediment.
Geol. 79, 143–188.
Rodriguez, A.B., Anderson, J.B., Banfield, L.A., Taviana, M.,
Abdulah, K., Snow, J.N., 2000. Identification of a −15 m middle
Wisconsin shoreline on the Texas inner continental shelf.
Palaeogeogr. Palaeoclimatol. Palaeoecol. 158, 25–43.
178
E.C. Osterberg / Marine Geology 229 (2006) 159–178
Sclater, J.G., Christie, P.A.F., 1980. Continental stretching: an
explanation of the post-Mid-Cretaceous subsidence of the central
North Sea basin. J. Geophys. Res. 85, 3711–3739.
Shackleton, N.J., 2000. The 100,000-year ice-age cycle identified and
found to lag temperature, carbon dioxide, and orbital eccentricity.
Science 289, 1897–1902.
Shackleton, N.J., Berger, A., Peltier, W.A., 1990. An alternative
astronomical calibration of the Lower Pleistocene timescale based
on ODP Site 677. Trans. R. Soc. (Edinburgh) Earth Sci. 81, 251–261.
Steckler, M.S., Watts, A.B., 1978. Subsidence of the Atlantic-type
continental margin off New York. Earth Planet. Sci. Lett. 41, 1–13.
Suggate, R.P., 1990. Late Pliocene and Quaternary glaciations of New
Zealand. Quat. Sci. Rev. 9, 175–197.
Suter, J.R., Berryhill, H.L., Penland, S., 1987. Late Quaternary sea-level
fluctuations and depositional sequences, southwest Louisiana
continental shelf. In: Nummedal, D., Pilkey, O.H., Howard, J.D.
(Eds.), Sea-Level Fluctuation and Coastal Evolution. Spec. Pub. —
Soc. Econom. Paleontol. Mineral., vol. 41, pp. 199–219.
Sutton, P.J.H., 2003. The Southland Current: a subantarctic current. N.Z.
J. Mar. Freshw. Res. 37, 645–652.
Tesson, M., Posamentier, H.W., Gensous, B., 2000. Stratigraphic
organization of Late Pleistocene deposits of the western part of the
Golfe du Lion shelf (Languedoc shelf), western Mediterranean
Sea, using high-resolution seismic and core data. Am. Assoc.
Petrol. Geol. Bull. 84, 119–150.
Vail, P.R., 1987. Seismic stratigraphy interpretation procedure. In:
Bally, A.W. (Ed.), Atlas of Seismic StratigraphyAAPG Studies in
Geology, vol. 27, pp. 1–11.
Wellman, H.W., 1979. An uplift map for the South Island of New
Zealand, and a model for uplift of the Southern Alps. In: Walcott,
R.I., Cresswell, M.M. (Eds.), The Origin of the Southern Alps.
Roy. Soc. New Zeal. Bull., vol. 18, pp. 13–20.
Wellner, R.M., Ashley, G.M., Sheridan, R.E., 1993. Seismic
stratigraphic evidence for a submerged middle Wisconsin barrier:
implications for sea-level history. Geology 21, 109–112.
Wilson, I.R., 1985. Well completion report Galleon-1. New Zealand
Geological Survey Unpublished Open-File Petroleum Report,
vol. 1146.
Yokoyama, Y., De Deckker, P., Lambeck, K., Johnston, P., Fifield, L.K.,
2001a. Sea-level at the last glacial maximum: evidence from
northwestern Australia to constrain ice volumes for oxygen isotope
stage 2. Palaeogeogr. Palaeoclimatol. Palaeoecol. 165, 281–297.
Yokoyama, Y., Esat, T.M., Lambeck, K., 2001b. Coupled climate and
sea-level changes deduced from Huon Peninsula coral terraces of
the last ice age. Earth Planet. Sci. Lett. 193, 579–587.