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The geologic record provides an opportunity to quantify the timing, rate, amplitude, mechanisms/controls, and effects (stratigraphic response) of eustatic change which, in turn, provide a baseline for predicting future relative sea level changes and assessing anthropogenic influences. However, eustatic effects are complexly intertwined with processes of basin subsidence and sediment supply (e.g., Cloetingh et al., 1985; Karner, 1986; Posamentier et al., 1988; Christie-Blick et al., 1990; Reynolds et al., 1991; Christie-Blick and Driscoll, 1995; Kominz et al., 1998; Kominz and Pekar, 2001). Controversy arises from the application of the sequence stratigraphic model (SSM) (Mitchum et al., 1977; Van Wagoner et al., 1988; Posamentier et al., 1988; Posamentier and Vail, 1988; Vail et al., 1991) to sea level studies. Sequence stratigraphy highlighted the cyclic nature of the continental margin stratigraphic record and led to the theory of eustatic control of sequences and the resultant eustatic cycle chart (Haq et al., 1987). This global sea level model (GSM) (Carter et al., 1991) remains contentious (e.g., Cloetingh et al., 1985; Carter, 1985; Karner, 1986; Christie-Blick et al., 1990; Christie-Blick, 1991; Carter et al., 1991; Karner et al., 1993; Christie-Blick and Driscoll, 1995; Dewey and Pitman, 1988; Miall and Miall, 2001).
To understand the history of eustasy versus subsidence/sediment supply changes, borehole transects across passive continental margins are required (COSOD II, 1987). The long-term strategy developed by ODP-related planning groups (Watkins and Mountain, 1990; Loutit, 1992) involves drilling margins worldwide to evaluate global synchroneity, by correlation among multiple basins and with the oxygen isotopic record, and to document stratigraphic responses in diverse tectonic and depositional settings, including carbonate, siliciclastic, and mixed siliciclastic-carbonate sedimentary systems on both continental and oceanic crust. Initial investigation was to focus on the Neogene "Icehouse" period (Miller et al., 1991) when high-resolution chronological control is available and glacial cycles provide a well-understood mechanism for eustatic change, calibrated by the deep-ocean oxygen isotope record. This approach has guided ODP efforts off New Jersey (Leg 150: Mountain, Miller, Blum, et al., 1994; Miller and Mountain, 1994; Leg 174A: Austin, Christie-Blick, Malone, et al., 1998; Legs 150X and 174AX: Miller et al., 1994; Miller, Sugarman, Browning, et al., 1998), and the Bahamas (Leg 166: Eberli, Swart, and Malone, 1996) and continues to influence IODP planning (e.g., Canterbury Basin IODP Expedition 317).
The passive margin approach integrates seismic profiles and a drilling transect to calibrate the SSM and test the GSM, including investigation of local controls on sequence formation. Seismic profiles provide sequence architecture, seismic facies, and morphologic constraints on depositional processes and tectonism. A drilling transect is required to document (1) ages of sequence stratigraphic surfaces, including sequence-bounding unconformities, or their correlative conformities, and maximum flooding surfaces; (2) facies and lithologies comprising each sequence (stratigraphic response to sea level oscillations); (3) porosity, cementation, and diagenesis; and (4) paleowater depths from benthic biofacies. Two-dimensional (2-D) modeling of these data within the sequence stratigraphic framework allows estimation of eustatic amplitudes because the form of the tectonic component of subsidence is known for passive margins (Kominz and Pekar, 2001).
The ideal approach involves drilling target sequences in at least two locations:
The ideal location for dating is near the clinoform toe to minimize the hiatus at the sequence boundary (Christie-Blick et al., 1998), but locations higher on the slope are necessary to reduce drilling depths (e.g., location of Leg 174A New Jersey margin Site 1072 relative to sequence boundary m1[s], Austin, Christie-Blick, Malone, et al., 1998). In addition, locations on the slope will provide better constrained paleowater depths, which are likely to be poorly constrained at clinoform toes (as is the case at the Clipper well). A further reason for drilling higher on the slope is that seismic correlation from clinoform toes landward to the clinoform front and shelf is difficult on all margins because the section basinward of clinoform toes is condensed and landward divergence of reflections contributes to mis-ties.
The eastern margin of the South Island of New Zealand is part of a continental fragment, the New Zealand Plateau, that rifted from Antarctica beginning at ~80 Ma (Anomaly 33). Rifting between the New Zealand Plateau and Antarctica-Australia was active along a mid-ocean-ridge system passing through the southern Tasman Sea and Pacific basins until ~55 Ma (Anomaly 24). Linking of the Indian Ocean and Pacific spreading centers truncated the spreading ridge in the southern Tasman Sea in the late Eocene. Spreading on the Indian and Pacific segments of this now-continuous Southern Ocean ridge system resulted in the formation of the modern boundary between the Australian and Pacific plates, comprising the Macquarie Ridge, Alpine Fault, and Tonga-Kermadec subduction zone (Molnar et al., 1975).
The Canterbury Basin lies at the landward edge of the rifted continental fragment and underlies the present-day onshore Canterbury Plains and offshore continental shelf. Banks and Otago peninsulas (Fig. F2) are middle–late Miocene volcanic centers. Basin sediments thin toward these features and also westward, where they onlap basement rocks onshore that are involved in uplift and faulting linked to the latest Miocene (8–5 Ma) initiation of the current period of mountain building along the Southern Alps (Adams, 1979; Tippett and Kamp, 1993a; Batt et al., 2000).
The plate tectonic history of the New Zealand Plateau is recorded in the stratigraphy of South Island. The postrift Cretaceous–Holocene sedimentary history of the Canterbury Basin comprises a first-order (80 m.y.) tectonically controlled transgressive–regressive cycle. The basin formed part of a simple passive margin from the Late Cretaceous to some time in the late Eocene when convergence between the Australasian and Pacific plates began to influence the region, eventually leading to formation of the Alpine Fault at ~23 Ma (King, 2000). The marine sedimentary section can be divided into three principal intervals, the Onekakara, Kekenodon, and Otakou groups (Carter, 1988), during which contrasting large-scale sedimentary processes operated (Fig. F4).
The postrift transgressive phase (Onekakara Group) produced ramplike seismic geometries and terminated during the late Eocene when flooding of the land mass was at a maximum (Fleming, 1962). Reduced terrigenous influx during the postrift phase of subsidence and transgression resulted in deposition of regionally widespread siliceous or calcareous biopelagites (Amuri Formation), which range in age up to early Oligocene (~33 Ma). The sequence is then interrupted by a current-induced unconformity, the Marshall Paraconformity (Carter and Landis, 1972), which occurs at the base of mid–late Oligocene cross-bedded glauconitic sand (Concord Formation) and calcarenite limestone (Weka Pass Formation), together comprising the Kekenodon Group (Fig. F4) (Carter, 1985, 1988). Exploration wells reveal that the paraconformity and probable equivalents of the Amuri and Weka Pass formations exist offshore (Wilding and Sweetman, 1971; Milne et al., 1975; Hawkes and Mound, 1984; Wilson, 1985). The paraconformity, which is the deepest target of the proposed drilling, is also recognized at drill sites throughout the region east of the Tasmanian Gateway and is hypothesized to represent the initiation of thermohaline circulation (Deep Western Boundary Current) and the proto-Antarctic Circumpolar Current upon opening of the seaway between Antarctica and Australia (~33.7 Ma), prior to opening of the Drake Passage (Carter et al., 2004c). New Zealand lay directly in the path of the developing current system. The paraconformity has been dated onshore using strontium isotopes as representing a hiatus lasting from 32.4 to 29 Ma (Fulthorpe et al., 1996). Its deepwater representation may have formed 1–2 m.y. earlier (Carter et al., 2004c).
Regression commenced in the late Oligocene or early Miocene in response to an increase in sediment supply provided by the initiation of Alpine Fault movement (Carter and Norris, 1976; Kamp, 1987). The Alpine Fault formed as a dextral strike-slip zone with 500 km displacement since the earliest Miocene (23 Ma) (Kamp, 1987). On eastern South Island, this resulted in the deposition of a widespread shelf siltstone (Bluecliffs Formation), starting in the latest Oligocene. At abyssal depths, in the path of the Deep Western Boundary Current, fine-grained terrigenous–carbonate rhythms at 41 k.y. Milankovitch frequency commenced at almost the same time (~24–23 Ma) (Carter et al., 2004c). This early uplift is distinct from the 5–10 Ma pulse of uplift that has culminated in the present-day Southern Alps because the early uplift is not recognized by fission track dating (e.g., Batt et al., 2000). Uplift of the Southern Alps accelerated at ~8–5 Ma (Tippett and Kamp, 1993a; Batt et al., 2000) or ~10–8 Ma (Carter and Norris, 1976; Norris et al., 1978; Adams, 1979; Tippett and Kamp, 1993b), indicating an increased component of convergence along the fault. Transpression led to an increase in the rate of sediment supply to the offshore Canterbury Basin (Lu et al., 2005). As in New Jersey, this sediment influx was deposited as prograding clinoforms (Otakou Group) (Fig. F4). Currents continued to influence deposition. At present, the core of the northward-flowing Southland Current, inboard of the Southland Front (part of the Subtropical Front [STF]) is over the ~300 m isobath (Chiswell, 1996). In deeper water, to at least 900 m, a local gyre of the Antarctic Circumpolar Current circulates clockwise within the head of Bounty Trough parallel to the Southland Current (Fig. F2) (Morris et al., 2001). Large sediment drifts within the prograding section (Fig. F4) show that similar currents, probably strengthened during glacial periods, existed throughout much of the Neogene (Fulthorpe and Carter, 1991; Lu et al., 2003; Carter et al., 2004c).
Backstripping suggests little tectonic subsidence in the central part of the offshore basin between ~30 and ~8 Ma (Figs. F5, F6) (Browne and Field, 1988), though an increase in subsidence rate beginning at ~8 Ma may be a response to increasing convergence at the Alpine Fault. This could be an in-plane force effect since deformation is locally absent. Figures F5 and F6 show evidence for uplift at ~50–35 Ma, which could reflect the late Eocene reorganization of plate boundaries.
A large amount of multichannel seismic (MCS) data collected in 1982 for a petroleum exploration consortium (B.P. Shell Todd) is available from the offshore Canterbury Basin, most significantly the CB-82 profiles (e.g., Fulthorpe and Carter, 1989). Although covering an extensive area, the data exhibit relatively low vertical resolution (20 m). Therefore, this survey was augmented in 2000 by 2-D high-resolution MCS profiles collected during a 21-day cruise (R/V Maurice Ewing cruise EW00-01) (Figs. F3, F7, F8, F9). The survey grid lies approximately midway between the Banks and Otago peninsulas on the present-day middle to outer shelf and slope, above the late Miocene–Holocene depocenter, and over the area where the largest sediment drifts developed (Fig. F3). The seismic source consisted of two generator-injector (GI) air guns (both 45/45 in3), and the streamer was deployed with 12.5 m groups in 96- and 120-channel configurations. A total of 57 profiles (~3750 km) were collected, covering ~4840 km2. Line spacings are 0.7–3.0 km in the dip direction and 2.0–5.5 km in the strike direction. Penetration, 1.7–2.0 s below seafloor, is sufficient to image the entire Oligocene–Holocene section. Data were processed using Focus software and then loaded into the GeoQuest interpretation system. High resolution was achieved by using reliable high-frequency sources (maximum frequency = 500 Hz), a small sample interval (1 ms), and high fold (48–60). Vertical resolution (~5m for two-way traveltimes [TWT] <1 s) is up to 4–5 times better than that of existing commercial MCS data (Fig. F10). However, seafloor and peg leg multiples are pronounced beneath the shelf. In order to deal with this problem, prestack deconvolution and FK-filtering were applied to some critical sections, yielding some improvements in quality.
The commercial low-resolution MCS grid data, however, were used to extend interpretations beyond the EW00-01 grid. The CB-82 data were particularly useful for ties to exploration wells, determining sediment drift distribution, and locating clinoform breakpoints, onlap, and canyons associated with the oldest sequences. The grid consists of 81 profiles, representing ~6000 km (Fig. F3). Record length is 5 s; sample rate is 4 ms. Digital copies of all stacked profiles were loaded into the GeoQuest system. Paper copies of migrated profiles were also available and 20 of the digital profiles were migrated as part of this project.
The high-resolution EW00-01 MCS data have been interpreted to provide a high-frequency sequence stratigraphic framework for the offshore Canterbury Basin (e.g., Fig. F7). Nineteen regional sequence-bounding unconformities (Unconformities U1–U19) are identified in the middle Miocene–Holocene shelf-slope sediment prism of the offshore Canterbury Basin (Lu and Fulthorpe, 2004). Three larger seismic units are defined based on seismic architecture and facies (Fig. F7):
Ages of Unconformities U3–U10 are based on ties to the Clipper exploration well (Hawkes and Mound, 1984), and the ages of Unconformities U11–U19 are based on ties to ODP Site 1119 (Fig. F11) (Shipboard Scientific Party, 1999b; R.M. Carter, pers. comm., 2002). Endeavour (Wilding and Sweetman, 1971) and Resolution (Milne et al., 1975) exploration wells were tied to the EW00-01 survey using the commercial CB-82 MCS profiles to calibrate the lowermost Unconformities U1, U2, and U3 (Fig. F7). Ages of upper Pliocene–Holocene Unconformities U11–U19 are most reliable since they derive from continuously cored Site 1119. The Site 1119 section is virtually complete; only one downlap unconformity (Unconformity U18 at ~87 meters below seafloor [mbsf], representing the Stage 7/8 boundary, ~252–277 k.y.) has a significant hiatus (~25 k.y.). Unconformity U19 (~48 mbsf) corresponds to the Stage 5/6 boundary (~113 ka), and the age of the deepest sediment recovered is ~3.9 Ma (Carter et al., 2004a). Ages of lower Pliocene and Miocene unconformities are less well constrained.
Correlation with oxygen isotopic records suggests a eustatic origin for the sequence boundaries. The number of seismic sequences is similar to that of coeval cycles on the Miocene–Holocene δ18Osw record of Billups and Schrag (2002) when cycles of comparable frequency are compared (Fig. F11). However, local processes have exerted fundamental control on sequence architecture. Along-strike currents strongly influence sequence development and large sediment drifts dominate parts of the Neogene section.
The presence of long-lived drifts beneath the modern shelf confirms that currents swept the New Zealand Plateau as early as 15 Ma (Fig. F12) (Lu and Fulthorpe, 2004). The STF, Subantarctic Front (SAF), and associated currents may have existed close to their present positions relative to New Zealand by the middle or latest Miocene (Carter et al., 2004c). The STF is represented by the Southland Front along the eastern South Island (Fig. F2). The δ13C record at Site 1119 is interpreted as reflecting glacial–interglacial alternations of subtropical and subantarctic water caused by movement of the STF across the site (Carter et al., 2004a, 2004b). Falling sea level deflects the front basinward during glacial periods, with the counterintuitive result that the site experiences a warmer water mass during glacials; intervals of interbedded sand and mud mark the passage of the front basinward and landward across the site (Carter et al., 2004b). Sand beds occurring near the peak of glaciation are interpreted to represent proximity of STF and SAF, which may coalesce near the site, intensifying current strength (Carter et al., 2004b).
Currents have formed at least 11 large elongate drifts within the lower Miocene–Holocene section (e.g., Fig. F12) (Lu et al., 2003). The drifts were initiated near the slope toe and aggraded to (or nearly to) shelf water depths. Drift deposits can be 1000 m thick and have mounded morphologies with channel-like moats along their landward flanks. Internal geometries define two end-members of elongate drift: simple and complex. Early (middle Miocene) simple drifts are small and concentrated in the southern part of the survey area (Figs. F3, F7). Drift thickness and longevity both increased as the shelf aggraded, increasing accommodation space, while the locus of drift development migrated northeastward through time. Late (late Miocene–Holocene) simple drifts are therefore larger and occur in the northeastern part of the survey area (Figs. F3, F8). Late simple drifts are divided into three parts (base, core, and crest) based on seismic facies. These facies form in response to progressive confinement of current flow within the moat. Complex drifts (multicrested and multistage; Fig. F3) may form as current pathways migrate in response to sea level change, modulated by paleoslope inclination, and as a result of fluctuations in the rate of sediment supply.
Current erosion in drift moats forms diachrononous unconformities, which cut across sequence boundaries. Several sequence boundaries pass through some of the larger drifts, indicating that they existed throughout several cycles of relative sea level change (Figs. F7, F12). Drift deposition controls sequence thickness distributions: sequences are thickest at drift mounds and thin within moats. In addition, currents focus deposition on the slope, reducing the rate of basinward movement of the shelf edge but increasing that of the slope toe. As a result, slope inclination is minimized (cf. Figs. F8 and F9). Cessation of drift development and replacement of along-strike by downslope processes result in increased rates of shelf-edge progradation and slope steepening as the accommodation space over the expanded slope is filled. Termination of large elongate drift development (~3.25 Ma at Site 1119) (Carter et al., 2004a) may have been caused by the initiation of late Pliocene–Pleistocene high-amplitude sea level change, which enhanced downslope processes by exposing the shelf edge. Slope platforms can form above extinct drifts, reducing accommodation and locally accelerating shelf-edge progradation. Along strike from some large elongate drifts, seismic evidence for current activity is lacking and coeval strata are clinoformal (Fig. F9) in spite of the demonstrable presence of a current. Therefore, elongate drift formation is the product of multiple controls, including current intensity, seafloor morphology, and sediment input.