ENZOSS REVISITED: THE HISTORY OF PACIFIC DEEP FLOW

Leg 181 drilling confirmed the utility of delineating the Eastern New Zealand Sedimentary System (ENZOSS) as a self-contained, dynamic sedimentary system. Like all regional stratigraphies and as summarized below, that of ENZOSS was at times affected by major events for which the cause lay outside the immediate area. Leg 181 drilling confirmed the timing of several of these influential regional events and additionally provided us with a detailed record of many hitherto undocumented changes in Neogene climate and in oceanographic flows for the ENZOSS region of the Southern Ocean (Figs. F23, F24, F25; also see the "Supplementary Material" contents list).

The New Zealand continental fragment split from the Gondwana supercontinent in the Late Cretaceous (Anomaly 33; ~83 Ma) (Cande and Kent, 1995) by rifting and subsequent seafloor spreading in the South Pacific and Tasman Oceans (Molnar et al., 1975; Cande et al., 2000). The onland geology of New Zealand, especially in South Island, contains an excellent and well-understood record of these events as they affected the western margin of the proto-Pacific (Carter, R., and Norris, 1976). Scattered fault-controlled Late Cretaceous rift basins with a fanglomeratic and immature fluviatile fill (e.g., Bishop and Laird, 1976; Laird, 1993) are succeeded by quartzose coastal plain coal measures and shallow-marine transgressive sediments (e.g., Macpherson, 1949; Wilson, 1956; Gair, 1959; Carter, R., 1988a, 1988b; Warren and Speden, 1978; Bishop and Turnbull, 1996). In eastern South Island, where the Canterbury Basin represents the very inner edge of the rift-stretched Pacific margin, total subsidence was small, resulting in a thin marine transgressive succession, generally <500 m thick, which is entirely of shallow-water origin and characterized by condensed sediment facies and many paraconformities (e.g., McMillan and Wilson, 1997).

Leg 181 drilling was primarily targeted at the paleoceanographic record contained in Neogene current drifts. Nonetheless and because of the rudimentary knowledge of the offshore stratigraphy prior to Leg 181, drilling at Sites 1121 and 1124 unexpectedly penetrated through important Late Cretaceous-late Eocene successions that represent the deep-sea counterpart to the well-understood postrift transgressive sections onland. The offshore sediments are mostly fine-grained claystones rich in biopelagic microfossils and represent part of the deep-marine peri-continental apron that accumulated along the western fringe of the Pacific Ocean after rifting. Drilling at Site 1124 penetrated the Cretaceous/Tertiary boundary, though regrettably, the boundary succession itself, which was imaged on downhole logs, lay between two successive cores and was therefore not retrieved (Carter, R., McCave, Richter, Carter, L., et al., 1999). Late Cretaceous siliceous, nannoplankton-rich claystones below the boundary change above to Paleocene nannofossil chalks with intervals of claystone. Similar nannofossil chalks with diatom-rich interbeds characterize the middle-late Paleocene section at Site 1121, farther south along the margin. The Paleocene chalks are overlain at Site 1124 by a 10-m-thick interval of middle Eocene dark brown mudstone, similar to sediment facies described from Gondwanan immediate postrift margins by Andrews (1977) and similarly inferred to represent a period of low oxygenation and sluggish bottom water movement. At intermediate depths, dysaerobic conditions became established during the late Paleocene, as marked by the deposition of the organic carbon-rich Waipawa ("Black Shale") Formation (Lillie, 1953), which Killops et al. (2000) suggested was formed between ~59 and 55.5 Ma on the outer shelf and upper slope during episodes of upwelling of warm, saline, nutrient-rich deep water (WSDW). A dysaerobic peak also occurred at a major extinction of intermediate-water benthic foraminiferal species that accompanied a 2 decrease in ¹³C at the ~55.5-Ma PETM, as recorded within siltstones of the Tawanui Formation in eastern North Island (Kaiho et al., 1996) and siliceous marl of the Amuri Limestone in Marlborough (Hancock et al., 2003). More vigorous Paleogene bottom flows are, however, suggested by the presence of Paleocene-early Eocene and middle-late Eocene paraconformities at Site 1124. Nonetheless and as may also be true globally (Moore et al., 1978; Wright and Miller, 1993), a period of stagnant circulation and low surface productivity characterized the southwest Pacific Ocean margin in the late middle Eocene (~37-39 Ma), after which the deposition at Site 1124 of a thin (8 m) interval of late Eocene (~34-37 Ma) nannochalk represents the offshore resumption of more normal biopelagic sedimentation prior to the dramatic ocean circulation changes marked by the Marshall Paraconformity at ~33 Ma.

Late Cretaceous and Paleogene fine-grained terrigenous and biopelagic sediments of deep marine origin occur extensively in northern South Island (Herring-Woolshed siltstone and Amuri chalk/chert formations) (Hollis et al., 1995; Field and Browne, 1989; Field et al., 1997) and eastern North Island (Whangai, Waipawa, and Wanstead Formations) (Lillie, 1953; Field et al., 1997). Characteristic Late Cretaceous transgressive shelf facies (Reay and Strong, 1992; Isaac et al., 1991) pass up into this fine-grained, deepening-upward, hemipelagic and biopelagic succession (Moore and Morgans, 1987; Field et al., 1997; Rogers et al., 2001), which accumulated as a regional sediment apron during the postrift foundering of the continental margin to abyssal depths. Today, these sediments occur within the tectonically disturbed East Coast Fold Belt, where they were emplaced by collision or accretion since the establishment of the New Zealand plate boundary in the late Oligocene (~25 Ma).

Leg 181 drilling yielded important in situ samples of the postrift sediment apron and thereby provides insight into the deepwater processes that operated then along the southwest Pacific margin. Earlier studies, including deterministic computer models (Barron and Peterson, 1991; Mikolajewicz et al., 1993), have suggested that the Paleogene Pacific Ocean circulation was dominated by a large anticlockwise gyre that distributed warm water southward from low western latitudes into the New Zealand area (Berggren and Hollister, 1977; Frakes, 1979; Kennett and Stott, 1991). The sediments drilled during Leg 181, and their onland equivalents in New Zealand, are consistent with these earlier hypotheses regarding Pacific circulation, containing warm-water faunas and yielding no particular evidence for the operation of a northward-flowing DWBC prior to the early Oligocene.

The ENZOSS commenced at the Eocene/Oligocene boundary, when eastward-propagating rifting through the Great Australian Bight and Tasmanian Gateway formed the first ocean passage between the Indian and Pacific Oceans and thereby created the proto-Southern Ocean (Molnar et al., 1975; Cande et al., 2000; Exon, Kennett, Malone, et al., 2001; Exon et al., 2002) (Fig. F24; see the "Supplementary Material" contents list). This event had two immediate, climatically profound consequences. First, it resulted in the wind forcing of strong eastward-flowing seabed to sea-surface currents through the gap, creating the precursor circum-Antarctic current. Second, the commencement of this energetic zonal current system began the process of the thermal isolation of Antarctica. The ensuing climatic deterioration may have been accentuated, or even primarily caused, by a marked decline in atmospheric CO₂ content that occurred over the same time period (Pearson and Palmer, 2000; cf. DeConto and Pollard, 2003). Together, these events were a prelude to the Oi-1 oxygen isotope event that marks the Eocene/Oligocene boundary (Devereux, 1967; Shackleton and Kennett, 1975; Miller et al., 1991), which is associated with a spike in ice-rafted debris at Leg 120 Site 748 on the Kerguelen Plateau (Zachos et al., 1992) and also with an associated short-term boost in productivity at Southern Ocean sites (Diester-Haass and Zahn, 1996; Robert et al., 2002). The Oi-1 event has therefore been widely interpreted as marking the formation of the first southern-sourced deep cold-water flows, which were followed by stronger cooling and ice-cap development in the Neogene (e.g., Barrett, 1996). However, recently published benthic foraminiferal Mg/Ca paleotemperatures (Lear et al., 2000) do not show a significant temperature fall at the Eocene/Oligocene boundary, suggesting that the Oi-1 isotope event may correspond entirely to the growth of new ice on Antarctica without concomitant generation of cold deep water. Irrespective of this ocean temperature ambiguity, however, at 33.7 Ma the center of the New Zealand Plateau lay at paleolatitude ~55°S in the southwest Pacific Ocean, immediately downstream from the throat of the developing Tasmania-Antarctica Gateway (Lawver et al., 1992; Exon, Kennett, Malone, et al., 2001) where "circumpolar...bottom waters were restricted to a relatively narrow channel and, hence, must have had very high current velocities," thereby producing "numerous and widespread disconformities" (Watkins and Kennett, 1971, p. 817).

Sites 1123 and 1124 were cored through relatively thick Neogene successions but penetrated beyond to terminate in late Eocene and Late Cretaceous sediment, respectively. Drilling at Site 1121 also encountered a Paleogene succession, which, in contrast, lay below thin, condensed Neogene deposits (Graham et al., in press). At these three sites and also at all other deepwater sites in the southwest Pacific that penetrate to the Eocene (Kennett, Houtz, et al., 1975; Exon, Kennett, Malone, et al., 2001), a paraconformity or complete facies change always separates Oligocene or younger sediments above from Eocene or older sediments below. This, the Marshall Paraconformity, has long been recognized to be present in all onland New Zealand sedimentary sections and to mark the inception of southern-sourced current flows into the Pacific consequent upon the opening of the Tasmanian Gateway (Carter, R., and Landis, 1972). Subsequent DSDP (Kennett, Houtz, et al., 1975; Kennett and von der Borch, 1986) and ODP (Carter, R., McCave, Richter, Carter, L., et al., 1999; Exon, Kennett, Malone, et al., 2001) drilling established the ubiquitous presence of the paraconformity in deep marine sections throughout the southwest Pacific. Deepwater unconformities of similar stratigraphic significance have also been identified in the Indian Ocean (Edwards, 1973), off South Africa (Tucholke and Embley, 1984), and on the southern Kerguelen Plateau (Diester-Haas et al., 1993; Diester-Haas and Zahn, 1996). These Southern Ocean unconformities mark the establishment of strong flows around the Antarctic perimeter (cf. Watkins and Kennett, 1971; Kennett et al., 1974) and probably also the start of the modern thermohaline ocean circulation system.

The deepwater manifestation of the Marshall Paraconformity is illustrated in Figure F25 and by Exon, Kennett, Malone, et al. (2001; frontispiece 2). As is the case for its onland shallow-water manifestation (Carter, R., 1985, fig. 3), the paraconformity at Sites 1123 and 1124 is subtle and at first glance it is hard to believe that it represents a sedimentation hiatus of many million years. The break is ubiquitously marked by Chondrites (or, in shallow water, Thalassinoides) burrows that penetrate into underlying nannofossil chalk and that are filled with the conspicuously younger greenish gray nannofossil chalk that overlies the unconformity. Both below and above the unconformity the sediments are extensively burrow mottled, and their shallow-water counterparts are also richly glauconitic, all pointing to low sedimentation rates. Estimated periods of ~13 and ~6 m.y. are unrepresented by sediment at Sites 1123 and 1124, respectively, the break commencing in the earliest Oligocene (~33.6 Ma) (Carter, R., McCave, Richter, Carter, L., et al., 1999; McGonigal and Di Stefano, this volume). On land, at the type locality of the Marshall Paraconformity at Squires Farm, Sr isotope dating established the duration of the gap as a shorter but still significant ~3.4 m.y. between 32.40 and 29.00 Ma (Fulthorpe et al., 1996). Elsewhere on land, however, the gap across the unconformity may extend for 10 m.y. or longer because of (1) minor marginal tilting and erosion of the sediments beneath the unconformity (Benson, 1969; McLennan and Bradshaw, 1984), which probably relates to the late Eocene-Oligocene propagation of the plate boundary through western South Island (e.g., Carter, R., and Norris, 1976; Turnbull, 1991; Sutherland, 1995); or, as exemplified at Site 1121, (2) extensive synparaconformity erosion under the influence of powerful and probably corrosive seabed currents.

Three agents are capable of causing a profound interruption to the accumulation of biopelagic sediments in both deep and shallow water. These are a complete switch-off of pelagic productivity, the introduction of cold and corrosive water (including possibly a shift in the CCD), or an increase in bottom currents to a level that precludes the deposition of fine-grained sediment. Changes to productivity almost certainly occurred along with the profound ocean reorganization that occurred at the Eocene/Oligocene boundary (e.g., Cifelli, 1969; Benson, 1975; Burns, 1977; Miller, 1992) and at deep Site 1124 corrosion of foraminifers is clearly intensified in sediments from above the Marshall Paraconformity compared with those from below (Carter, R., McCave, Richter, Carter, L., et al., 1999), which is strong evidence for a new cold bottom water source. However, abundant evidence exists that seafloor current activity was the primary cause of the Marshall Paraconformity in shallow water (Carter, R., 1985; Fulthorpe et al., 1996; Carter, R., et al., 1996; see also discussion above) and we infer a similar origin for its deepwater manifestations (cf. Kennett, 1977). Situated as it was, as a shallowly submergent plateau just beyond the Tasmanian Gateway, the ENZOSS region was subjected to the impact of the partial ACC system as soon as it developed. The abyssal seabed to sea-surface character, and the strength, of early Oligocene current activity, which is a key feature of the modern ACC, is demonstrated by the fact that sediment accumulation was inhibited for a minimum of several million years at all depths between a few tens of meters on the interior of the continental platform and several thousand meters on the peri-continental abyssal sediment apron.

Three major features that reflect far-reaching oceanographic change characterize late Eocene to mid-Oligocene sediments of the southwest Pacific. They are

1. The cooling of shelf and oceanic waters, as marked by the sharp Oi-1 oxygen isotope enrichment at the Eocene/Oligocene boundary (Devereux, 1967; Benson, 1975; Shackleton and Kennett, 1975; Wei, 1991);
2. The ubiquitous presence of the regional "mid-Oligocene" Marshall Paraconformity (Carter, R., and Landis, 1972; see also the appendix in Carter, R., et al., 1982) and other probably equivalent "early Oligocene" oceanic and continental margin unconformities (Edwards, 1973; Carter, A., 1978); and
3. A claimed global sea level fall, also of "mid-Oligocene" (~32 Ma) age (Haq et al., 1987; Miller et al., 1985).

Establishing the exact relationships between these events has been bedeviled by problems of accurate dating, with added uncertainty coming from ambiguity in the age of the Eocene/Oligocene boundary itself. For example, Boersma and Shackleton (1977) suggested that the apparently later age of the Oi-1 enrichment in the southwest Pacific (early Oligocene) was caused by diachroneity of zone fossils between tropical and southern locations, a discrepancy that was later partly resolved by adopting the isotope shift as itself marking the Eocene/Oligocene boundary (e.g., Keigwin, 1980; Miller, 1992). Later, Kamp et al. (1990) confirmed the existence of such problems by demonstrating that the traditional late Eocene indices Globigerapsis index and Subbotina linaperta continued well into the early Oligocene, as judged by their occurrence above the Oi-1 event in southern Australian sections. More generally, correlation using New Zealand Oligocene stages (e.g., Morgans et al., 1996) has long been unsatisfactory because of an inadequate type locality for the Duntroonian Stage, which has an unconformable base at the Marshall Paraconformity and comprises a condensed greensand containing reworked microfossils (e.g., Hornibrook, 1966). There is also a lack of any alternative onland sections that are known to be continuous through the Oligocene (Waghorn, 1981).

Resolution of these problems began when the estimate of 32.4-29.0 Ma was established for the age gap across the Marshall Paraconformity at its type locality (Fulthorpe et al., 1996), shortly after the revision of the age of the Eocene/Oligocene boundary to 33.7 Ma (Berggren and Aubry, 1995), but ambiguities nonetheless remain. Leg 181 and 189 drilling demonstrated that the youngest sediments below the paraconformity offshore are 33.5 Ma (Site 1123) and 33.0 Ma (Site 1124) at Leg 181 sites (McGonigal and Di Stefano, this volume), ~33.0 Ma at Leg 189 Sites 1170-1172 (Exon, Kennett, Malone, et al., 2001), and ~33-32 Ma at DSDP Leg 29 Site 277 (Jenkins, 1974; Murphy and Kennett, 1986). Thus the "early Oligocene" isotope event (Shackleton and Kennett, 1975), widespread Southern Ocean "Eocene-Oligocene" deep marine paraconformities (Edwards, 1973; Tucholke and Embley, 1984), and the "mid-Oligocene" Marshall Paraconformity (Carter, R., and Landis, 1972) are actually closely similar in age, with the isotope shift lying at the Eocene/Oligocene boundary at 33.7 Ma and the basal surface of the Marshall Paraconformity following in the early Oligocene no more than 1-2 m.y. later. The indications after Leg 181 drilling are that in the southwest Pacific the paraconformity developed in deep water ~1 m.y. before it spread to shallow water, which might suggest that the proto-DWBC preceded the origin of the proto-ACC. However, the age difference of ~0.6 m.y. for the youngest sediments below the unconformity in deep and shallow water is of the same order as the likely error on the biostratigraphic dating, and alternatively, such a difference could also result from the occurrence of greater seabed erosion at the deepwater sites.

Two new Eocene-Oligocene stratigraphies for the Tasmanian Gateway region have been published since the appearance of the Leg 189 cruise report (Exon, Kennett, Malone, et al., 2001). The first, by Exon et al. (2002) omits the early Oligocene (32-33.5 Ma) hiatus depicted at Sites 1170-1172 by Exon, Kennett, Malone, et al. (2001). Instead, Exon et al. (2002) show continuous deposition across the Eocene/Oligocene boundary at the three Tasman Plateau sites, followed by a middle Oligocene (~27-31 Ma) paraconformity at Sites 1170 and 1171; a similar pattern of Eocene-earliest Oligocene continuity followed by middle Oligocene hiatus occurs at Site 277 on the western edge of the Campbell Plateau (Shackleton and Kennett, 1975) and Site 744 on the Kerguelen Plateau (Barron, Larsen, et al., 1989; Robert et al., 2002). The second reinterpretation, by Pfuhl and McCave (2003), recognizes the Marshall Paraconformity at all Leg 189 sites except Site 1168 (see additional comments in the caption to Fig. F6). Different regional interpretations will apply, depending upon the correctness of these differing stratigraphies on the Tasman Plateau and particularly upon the amount of early Oligocene seafloor erosion that actually occurred at particular sites throughout the Southern Ocean. Overall, however, the Marshall Paraconformity (1) is present in all, or all but one, oceanic drill sites in the region and (2) in its type area is centered in the early to middle Oligocene in shallow water (~29-32.5 Ma) (Fulthorpe et al., 1996) and appears to start slightly earlier offshore (33.0-33.5 Ma) (McGonigal and Di Stefano, this volume). At some locations (perhaps including Sites 1123 and 1124), seafloor erosion prior to resumed late Oligocene sediment deposition may have extended the apparent length of the hiatus down to the Eocene/Oligocene boundary, or beyond (i.e., a nondepositional, as opposed to erosional, hiatus of early Oligocene age may never have existed). Alternatively, two separate unconformities may be involved, the older "Eocene-Oligocene" one centered in the earliest Oligocene (~33 Ma) (Edwards, 1973; Exon, Kennett, Malone, et al., 2001) and the younger Marshall Paraconformity centered in the middle of the Oligocene (~29-32 Ma) (Fulthorpe et al., 1996). The existing data are inadequate to resolve the issue, especially because most of the relevant sections lack good magnetostratigraphic, isotopic, or astrochronologic control.

Two other significant stratigraphic conclusions follow from this discussion. First, the Vail-EXXON "mid-Oligocene" sea level fall, where it is present at all, may, like the base of the Marshall Paraconformity, prove to be of early Oligocene age on the revised timescale (Pekar, in Prothero et al., 2000). Second, within the local New Zealand stage classification, the early Whaingaroan, which was assigned formerly to the early Oligocene and for which Globigerina angiporoides is a key indicator species, is actually partly late Eocene. Furthermore, Globigerapsis index, the extinction of which has traditionally been taken as marking the end of the Eocene in Australasia, is now known to persist well into the Oligocene (Kamp et al., 1990). These facts may help resolve the apparent enigma of alleged "early Oligocene" macrofossils of tropical affinities (including Cocos nuts, Lingula, and cypraeid and coniid gastropods) (cf. Beu and Maxwell, 1990; Hornibrook, 1992; Edwards, 1991) and the conflicts between faunal and oxygen isotope data reported by Burns and Nelson (1981), Adams et al. (1990), and Buening et al. (1998). The isotope data, once displayed against an accurate timescale, probably accurately reflect the global oceanic climatic trend. Many—though not necessarily all (since shallow-water warm refugia could certainly have existed)—of the tropical taxa reported from "early Oligocene" localities in New Zealand may be miscorrelated and actually have lived during the latest Eocene warm interval. Similarly, the widespread occurrence of fossils of modern subantarctic affinities (cetacean whales, penguins, and Notorotalia) in the condensed greensands and limestones that overlie the Marshall Paraconformity (Marples, 1952; Fordyce, 1977, 1981; Carter, R., 1985) is a clear indication of the regional change in oceanography and of post-Eocene cold water influence.

Seismic evidence indicates that the first deepwater ENZOSS drifts began to accumulate along the path of the DWBC north of Chatham Rise (Carter, L., and McCave, 1994). The oldest sediments above the Marshall Paraconformity at nearby Site 1124 comprise cyclic alternations of darker- (terrigenous) and lighter- (carbonate rich) colored, greenish gray nannoplankton-rich chalk of late Oligocene (22.4-27.0 Ma) age (Carter, R., McCave, Richter, Carter, L., et al., 1999). The darker mudstones contain fewer and more corroded foraminifers, an increased siliceous microfauna of cold-water affinity, and reworked Eocene-Oligocene diatoms. These materials are inferred to have been derived partly by erosion from upstream sources such as the Paleogene sediment apron cored at Site 1121, and their concentration in the darker mudstones suggests increased cold water inflow at such times.

Clearly, if the Marshall Paraconformity at this site resulted from current erosion, then that activity must have waned sufficiently by ~27 Ma to allow the accumulation of these cyclic chalks. By comparison of their lithologies with the similar but younger sediments at Site 1123 and given their relatively high accumulation rates of >20 cm/k.y., the Site 1124 late Oligocene chalks represent the earliest DWBC deposits in the Leg 181 record and their cyclicity indicates that, like the younger drifts, they were deposited under the influence of the 41-k.y.-long astronomical cycle. Toward their top, these chalks show decreased rates of accumulation (Handwerger and Jarrard, in press) (Fig. F20), perhaps indicative of an increasing current speed toward the overlying paraconformity. The marked increase in the proportion of terrigenous material after 25 Ma (Joseph et al., in press) (Fig. F22) also indicates the arrival of "new" sediment into the path of the DWBC. This sediment may have been derived from the newly uplifting plate boundary in the west, but no definite pathways are known. Alternatively, therefore, the sediment may be derived from the seabed farther south, as evidenced by the eroded sediment apron and moat at Site 1121 and reworked Eocene-Oligocene microfossils at Site 1124.

Unfortunately, because of the great depth and the inferred influence of corrosive bottom water, post-Marshall Paraconformity microfossils at Site 1124 are poorly preserved. Nonetheless, the late Oligocene section at the site, together with the equivalent parts of Sites 1168 and 1172 (Exon, Kennett, Malone, et al., 2001), will be of particular importance for future research into the earlier history of ACC-DWBC evolution, for which they represent both the earliest and the only records available.

The modern mean flow of water passing through Drake Passage above 3000 m is 97 ± 13 Sv (Orsi et al., 1995). Sensitivity modeling indicates that this vigorous ACC flow depends upon the presence of the passage (Gill and Bryan, 1971), which therefore clearly plays a pivotal role in the maintenance of a fully circum-Antarctic, Southern Ocean circulation. Despite contributions by many authors, it has been surprisingly difficult to achieve a consensus regarding the date at which the passage opened to oceanic circulation.

1. Eocene South American marsupial fossils on the Antarctic Peninsula indicate a connection across Drake Passage up until that time (Woodburne and Zinsmeister, 1982). Relationships between marine benthic organisms on either side of the passage have been used in support of opening ages as widely different as late Eocene (~37-40 Ma) (echinoid distribution, Foster, 1974; increased biopelagic productivity at Maude Rise, Weddell Sea, Diester-Haas and Zahn, 1996) and late Miocene (~6 Ma) (benthic foraminifers; Boltovskoy, 1980).
2. As an increasing number of ocean cores became available from the region, estimates of the age of the passage based on sediment facies distributions or paleoceanographic reasoning mostly came to lie in the range of Oligocene to early Miocene (~30-20 Ma) (e.g., Tucholke et al., 1976; Ciesielski and Wise 1977; Kennett, 1978).
3. About the same time, the first detailed analyses of ocean floor magnetic anomalies near the passage also suggested an opening age for shallow waters in the late Oligocene (~29 Ma) but with continental fragments impeding deepwater flow until the early Miocene (~23.5 Ma) (Barker and Burrell, 1977, 1982). Most recently, Barker (2001) reviewed this evidence and inferred an ACC origin between 17 and 22 Ma.
4. It has been suggested that somewhat earlier passage opening occurred through the Powell Basin, a marginal basin that lies immediately south of the modern Drake Passage, along the Antarctic Peninsula (Lawver et al., 1994; Lawver and Gahagan, 1998). Based on the regional geology, combined with heat flow measurements and age/depth calculations, these authors concluded that a middle- to deepwater passage opened by ridge spreading in the Powell Basin in latest Eocene to Oligocene time (~37-30.5 Ma). This interpretation was refined by Eagles and Livermore (2002), who document continental margin rifting along the Antarctic Peninsula between ~40 and 29.7 Ma and identify seafloor magnetic lineations that indicate the creation of the Powell Basin by spreading between 29.7 and 21.8 Ma.

The best current estimate, therefore, is that a deepwater marine gap opened up between South America and Antarctica, initially through Powell Basin, at ~29.7 Ma (i.e., ~4 m.y. after the opening of the Tasmanian Gateway).

Because of the control it exercises over the relative strengths of meridional (thermohaline) and circum-Antarctic zonal (ACC) flows, the effects of the opening of Drake Passage can be expected to be recognizable in southwestern as well as southeastern Pacific Ocean successions. In that light, one of the most striking aspects of the Marshall Paraconformity is the gap of 3-4 m.y. (32.7-29.0 Ma) or longer that occurs across it, which we take to indicate strong corrosive and erosive seabed current flows. When sedimentation resumed at ~29-28 Ma, sediment drifts were deposited ubiquitously in the New Zealand region in both shallow-water (Ward and Lewis, 1975; Anastas et al., 1997) and deepwater (Carter, R., McCave, Richter, Carter, L., et al., 1999) locations. Noting this, Carter, R. (1985) and Fulthorpe et al. (1996) argued that the early Oligocene proto-ACC-DWBC flowed into the Pacific directly across and along the eastern edge of the New Zealand microcontinent, where it precluded the deposition of biopelagic sediment for a period of several million years, forming the paraconformity in the process. The resumption of sedimentation in the mid-Oligocene (local Duntroonian stage) reflects lessening current flows, possibly caused by a southward migration of the ACC core into its more strictly circum-Antarctic path consequent upon the opening of Drake Passage. As discussed above, the best-estimate mid-Oligocene date of opening of Drake (Powell) Passage is consistent with these previous interpretations, which were based on sediment facies evidence. Indeed, if this line of reasoning is correct, the conspicuous regional sedimentation changes that occur at ~28 Ma (resumption of greensand and calcarenite deposition above the Marshall Paraconformity) and ~23 Ma (change from coarse-grained calcareous drifts to fine-grained terrigenous drifts) in shallow-marine sediments in eastern New Zealand may represent our most accurate indication of the start and the complete establishment of true circum-Antarctic flow. A similar major sub-early Miocene unconformity overlain by terrigenous sediment drifts also occurs on both the upstream and downstream side of Drake Passage in the southeast Pacific and South Atlantic, respectively (Tucholke et al., 1976; Wright and Miller, 1993).

Prior to the opening of Drake Passage, enhanced northward meridional flow is expected to have occurred in the Southern Hemisphere, with a greater outflow of AABW to the world ocean and a weaker ACC, both factors that today act to suppress NADW production (Mikolajewicz et al., 1993). Thus, prior to ~5 Ma, which approximates to the closure of the Isthmus of Panama and the start of NADW production (Warren, 1983; Wright et al., 1991; Haug and Tiedemann, 1998; Collins et al., 1996; Haug et al., 2001; Lear et al., 2003), the primary source of global ocean deepwater production was probably cooling in the high latitudes of the Southern Hemisphere (Katz and Miller, 1991). Accordingly, during the period ~33.7-5 Ma, the Pacific DWBC constituted an even more important limb of the thermohaline circulation than it does today and variations in heat flux, modulated by its changing flow, may have been a primary determinant of global climate change.

Late Oligocene drift deposits at Site 1124 are interrupted by a 5.4-m.y.-long paraconformity that starts at ~22.4 Ma. Above this break, cyclic nannofossil chalks similar to those below accumulated from 17.6 Ma onward, with two further paraconformities at ~16.5-15.0 Ma and ~14-11 Ma (Carter, R., McCave, Richter, Carter, L., et al., 1999). Meanwhile at nearby Site 1123, North Chatham Drift sedimentation commenced at ~21 Ma and continued almost unbroken to today (Carter, R., McCave, Richter, Carter, L., et al., 1999; Wilson et al., 2000b). At both Sites 1123 and 1124, deposition occurred in concert with a prominent 41-k.y. climatic beat, as shown by the cyclic lithologic logs and the orbitally tuned grain size record of Hall et al. (2003). Intervals with reduced biogenic carbonate correspond to enhanced corrosive bottom-water flow from southern sources, as inferred from the increased size of terrigenous sortable silt and the presence of reworked subantarctic diatoms. Like the diatoms, some of the terrigenous component was probably derived from seafloor erosion farther south rather than from the main New Zealand landmass.

The regular Milankovitch obliquity cycles are superimposed upon other episodic shifts that are manifest in the drift record, for instance, the sharp increase in carbonate flux at ~13.5 Ma (Sites 1123 and 1124), which Handwerger and Jarrard (in press) associate with increased productivity and which occurs ~1 m.y. after the marked middle Miocene cooling step recorded in the global oxygen isotope curve (Zachos et al., 2001). Latest Miocene to early Pliocene (~7.0-4.6 Ma) warming and resumed cooling between ~4.0 and ~2.5 Ma are marked at Leg 181 core sites in manifold ways (see the "Supplementary Material" contents list), yet none of these events disrupted the background functioning of the DWBC. From at least 27.0 Ma and probably from 33.7 Ma, the Pacific DWBC current continued its metronomic supply of deep cold water to the Pacific Ocean; waxing—occasionally to the point of seabed erosion—and waning awhile, but always present.

Calc-alkaline volcanic activity commenced in New Zealand in the mid-Cenozoic at the same time as regional tectonism related to the inception of the North Island plate boundary (Stoneley, 1968; Delteil et al., 1996). An active andesitic volcanic chain became established in the Auckland-Northland region at ~25 Ma in the late Oligocene (Ballance et al., 1985), associated with other dramatic geological changes including the emplacement of the Onerahi Chaos Breccia (Kear and Waterhouse, 1967) as part of the Northland Allochthon (Ballance and Sporli, 1979). This, the first of the three active chains of arc volcanoes related to subduction beneath the North Island, now lies far to the west because it was rotated away from the modern trench by intra-arc extension along the younger Coromandel and Taupo Volcanic Zones (cf. Walcott, 1984). Fragmental ejecta from the Northland arc are widely preserved in the nearby Oligocene-Miocene Waitemata Basin (e.g., Hayward, 1993), and altered green clay layers of late Oligocene-late Miocene age from Leg 90 drill sites on the Lord Howe Rise can be interpreted as altered ashes derived from the Northland arc (Gardner et al., 1985). Northland arc ashes may occur in the ENZOSS record as similar altered green clay layers at Site 1124.

Early remnants of ash notwithstanding, the record of unaltered macroscopic tephra derived from the Coromandel and Taupo Volcanic Zones starts at Site 1124 at ~12 Ma, near the middle/late Miocene boundary. Thereafter, the offshore ash record is semicontinuous, with 134 macroscopic tephra layers punctuating the background hemipelagic sediments. As outlined in more detail earlier in this review, these tephra provide a significant source of particulate sediment into the ENZOSS and represent an excellent record of major volcanic eruptions from the North Island arc.

We described earlier the evidence for the early Oligocene (~33.7 Ma) initiation of a strong southern-sourced current system across the southwest Pacific region. Ubiquitous seafloor erosion in both shallow and deep waters suggests that this protothermohaline system resembled the modern ACC, in that it occupied the entire seabed to sea-surface water column. Burns (1977), using microfossil evidence from DSDP Legs 21, 28, and 29, gave the first summary of the development of Southern Ocean water masses and fronts that followed the pivotal Eocene/Oligocene change.

Prior to the Oligocene, an undifferentiated calcareous, open ocean microfossil assemblage was regionally widespread. During the Oligocene, a cool nannoflora (including Isthmolithus recurvus) and siliceous cool-water microfauna developed in coastal locations around Antarctica, followed by marked change in the early Miocene when the first siliceous assemblages characteristic of modern Antarctic waters appeared near Antarctica (Leg 28, Site 267); farther north, away from Antarctica, the undifferentiated oceanic microfauna was replaced by an early Miocene assemblage similar to that of modern cool subtropical water (Leg 21, Site 281). Burns (1977) identified a precursor fauna to that of modern Circumpolar Subantarctic water at Leg 29 Site 278 during the early Miocene, but not until the end of the Miocene did this zonal water mass become regionally extensive. Summarizing, the first proto-Polar Front probably appeared during the Oligocene (cf. Kaneps, 1975) and signaled the start of the regional differentiation of water masses away from a cooling Antarctica. A precursor southwest Pacific Ocean water circulation similar to today's appeared in the early Miocene, and the differentiation of this system took place by northward expansion and strengthening of oceanographic gradients, including an intensification of subtropical gyral circulation (Kennett et al., 1985), during the middle and late Miocene. By the Pliocene, the belt of Subantarctic water intermediate between Antarctic and Subtropical waters was established as a circum-Antarctic feature and the direct predecessor water masses and fronts of the modern circulation were all in place (Kennett, Houtz, et al., 1975; Burns, 1977; Lazarus and Caulet, 1993).

Leg 181 results are consistent with earlier onland studies regarding New Zealand paleoclimate (Hornibrook, 1992) (cf. Fig. F3), which suggests the occurrence of a middle Miocene subtropical climatic optimum at ~16-17 Ma followed by climatic deterioration into the Pliocene-Pleistocene. As discussed earlier, from the Oligocene onward phases of enhanced DWBC flow generally correspond to marked climatic coolings. Increased biopelagic productivity and foraminiferal census counts also demonstrate the existence of early Pliocene climatic warming, albeit followed by further cooling at ~4 Ma. After 4 Ma, all proxy indicators reveal short-period, often Milankovitch-scale, alternations of warmer and colder climate, though only small parts of the record have yet been examined in detail (e.g., Hall et al., 2002, 2003; Sabaa et al., in press; Scott and Hall, in press). Oxygen isotope results suggest that significant surface cooling affected nearshore ENZOSS waters over the last 0.7 m.y. (Site 1119) (Carter, R., unpubl. data), consistent with the intensified winnowing reported for this period for intermediate-depth waters on Lord Howe Rise by Kennett and von der Borch (1986). However, the average peak glacial and interglacial temperatures of oceanic surface waters over the same time interval remained essentially unchanged (Site 1123) (Hall et al., 2001).

Regarding the initiation of oceanic frontal zones in the southwest Pacific, we agree with earlier writers that the number of drill sites available is insufficient to tightly constrain the origins of the STF and SAF. Murphy and Kennett (1986) argued from an analysis of the Eocene-early Oligocene isotope records at Leg 29 Site 277 and Leg 90 Sites 592 and 593 that no frontal structure can have lain between these sites prior to the Oligocene. They inferred from diverging intermediate and surface water isotope records between Site 277 and sites farther north that, after ~33 Ma, intermediate water flows had commenced as part of the enhancement of the latitudinal thermal gradients that eventually led to the formation of the STF and SAF. However, Murphy and Kennett (1986, p. 1359) also noted that " ... biogeographic similarities in calcareous microfossils suggest that the Subtropical Convergence (= STF) had not yet formed by the end of the Paleogene" and that "biostratigraphic evidence suggests formation of this convergence in the middle Miocene." In contrast, Kamp et al. (1990), Hornibrook (1992), and Buening et al. (1998) used shallow-marine faunal distributions and isotope data to infer that a proto-STF had already developed off Southern Australia and eastern South Island by the early Oligocene and Nelson and Cooke (2001) recently argued for the existence of this front since the latest Eocene. Any such early-established front would have initially marked a simple separation between increasingly cold circumpolar water and the major warm tropical-subtropical gyre water, and its modern descendant may therefore be the more northerly Tasman Front that separates warm from cool subtropical water (20°C summer isotherm) rather than the later-developing STF (15°C summer isotherm).

Data from Leg 181 and other available data do not closely constrain the time of formation of either the STF or the SAF in the southwest Pacific but are consistent with the formation of the AAPF in the early Oligocene and the STF and SAF some time between the early Oligocene and the early Miocene. Relevant information includes

1. Oligocene and younger shallow and intermediate-depth sediment drifts occur throughout eastern New Zealand, changing from carbonate facies in the late Oligocene to terrigenous facies in the early Miocene (Carter, R., 1985; Carter, R., et al., 1996). Drilling at Site 1119 confirmed the Pliocene-Pleistocene age of the upper part of the terrigenous drift succession, which shows a consistent (seismic) drift facies back to the early Miocene. We infer that the subduction of AAIW water at the ancestral AAPF was occurring from at least the early Miocene, after which time its northward passage caused the deposition of the Canterbury terrigenous drifts (cf. Fulthorpe and Carter, R., 1991; Lu et al., 2003).
2. Flower and Kennett (1993, 1994) used isotope data to infer that Tethyan-sourced warm, saline deep water was competing with cold southern-sourced deep water between ~16 and 14.9 Ma in the southwest Pacific, just prior to the major increase in oxygen isotope values that represents Antarctic ice sheet growth from 14.5 to 14.1 Ma. Leg 181 data, which come from rather more southerly sites, are consistent with a fluctuating but continuing southern source through the middle and late Miocene (e.g. Fig. F21A). However, bolboforms, which are reliable cold-water indicators and may also specifically mark the presence of AAIW, first appear at Site 1123 at ~13 Ma and have a regular presence after 11.7 Ma (Crundwell, in press) (cf. Fig. F12). Leg 90 Site 593 in the Tasman Sea contains a similar bolboform stratigraphy and also exhibits a slightly older (15-18 Ma) condensed and oxygenated "orange interval," which Kennett and von der Borch (1985) and Nelson (1986) interpreted as marking the start of northward-flowing AAIW in the Tasman. Site 1120 on the Campbell Plateau contains a small hiatus at this time (16.7-15.8 Ma). These various events may mark the site-specific strengthening of a general regional increase in AAIW flow strengths, and AAPF frontal structure, during the middle Miocene.
3. Detailed census studies of foraminifers at Sites 1123 and 1124 (Hayward et al., in press) show that there are significant differences between lower bathyal faunas north and south of Chatham Rise, suggesting the presence of a proto-STF in eastern New Zealand since at least the early Miocene. Major microfaunal changes also occur between ~5.0 and 6.5 Ma (cf. Fig. F14), when the planktonic foraminiferal fragmentation index shows a marked decline at Site 1123 and benthic foraminiferal assemblages at both sites show a marked shift of key species. Hayward et al. interpret these changes as marking a deepening of the CCD, a decrease in dissolution, a decrease in water oxygen content, and an increase in food supply; all these features are consistent with a latest Miocene restructuring of water masses into essentially the modern system (Kennett and von der Borch, 1985; Nelson and Cooke, 2001), after which both the SAF and STF were strong and permanent frontal features.
4. Detailed isotope and MAR studies at Site 1125 (Grant and Dickens, 2003) have demonstrated the presence there of a strong late Miocene-Pliocene "biogenic bloom" starting at ~5.8 Ma, consistent with the initiation of strong frontal (STF) structure at that time; a biogenic bloom of similar age, the "Chron 6 carbon shift," is widespread in the world ocean (Haq et al., 1980; Kennett and von der Borch, 1985; Hodell and Kennett, 1986; Wei, 1998; Dickens and Owen, 1999; Grant and Dickens, 2003), consistent with the ocean restructuring and the global strengthening of frontal zones at that time.
5. After a period of steady middle-late Miocene nannofossil ooze accumulation, the Site 1120 record displays a major paraconformity at 1.9-5.6 Ma, above which occurs only 12 m of Pleistocene sediment, itself punctuated by further paraconformities (Carter, R., McCave, Richter, Carter, L., et al., 1999). These features are consistent with the commencement of steady intermediate-depth cool-water flows across the Campbell Plateau from the latest Miocene onward.

We conclude that southern-sourced intermediate-depth water flows deriving from a proto-AAPF traversed the New Zealand region from the early Oligocene onward, strengthening after the climatic cooling that followed the ~16-Ma climatic optimum (cf. Flower and Kennett, 1993, 1995). Somewhat later, in the latest Miocene (~7-5 Ma), the direct antecedents of today's SAF and STF developed, stimulated by a marked climatic deterioration (e.g., Zachos et al., 2001) and as part of the global ocean restructuring that is inferred to have occurred at that time (e.g., Wright et al., 1991).

The Solander, Bounty, and Hikurangi Channels are the three main conduits through which sediment is today delivered from the New Zealand landmass into the path of the DWBC (Carter, L., et al., 1996).

The modern Bounty Channel is fed by tributary canyons and channels in the head of the Bounty Trough (Carter, R., and Carter, L., 1996), a basin that was created during the Late Cretaceous rifting of the margin (Davey, 1993). The bulk of the trough remains unfilled, despite the advance of the South Island shelf sediment prism into its western head since the early Miocene (Carter, R., and Norris, 1976). The trough center has undoubtedly been present as a marked bathymetric low since the Late Cretaceous, and evidence exists for the presence since the early Eocene of a proto-Bounty Channel in a now-filled subsidiary headward rift (the Great South Basin) (Carter, R., and Carter, L., 1987). Nonetheless, seaward of the Late Cretaceous-Eocene shore-connected, transgressive sediment prism, the ENZOSS deep seafloor continued to receive mainly biopelagic detritus until the late Oligocene-early Miocene. Then, at ~23.4 Ma and almost simultaneously with the beginning of terrigenous regression in the South Island (Finlay, 1953; Lewis et al., 1980; Fulthorpe et al., 1996) and the inception of subduction in the North Island (Ballance, 1976; Field et al., 1997), rhythmic terrigenous-carbonate DWBC sedimentation commenced at Site 1124 (Fig. F22; also see the "Supplementary Material" contents list). The route by which the terrigenous sediment reached the DWBC is unclear, though some was undoubtedly provided from seafloor erosion along the edge of Campbell Plateau. Another possibility is that of transport through a paleo-Bounty Channel, but the presence at Site 1122 of early Pliocene contourites containing reworked diatoms from farther south suggests that a Bounty source only became significant during and after the late Pliocene.

The presence of 41-k.y.-long cycles of clay-rich and clay-poor layers throughout the post-Oligocene successions at Sites 1123 and 1124 indicates that fine-grained terrigenous sediment was continually fed into the DWBC from the late Oligocene onward. In addition, the gradual post-Oligocene enrichment of illite + chlorite at the expense of smectite at Site 1123 (Winkler and Dullo, this volume) is indicative of a "steady-state" system undergoing a progressive change in provenance lithology. Such background flux to the DWBC notwithstanding, the Leg 90 Site 594 record (Dersche and Stein, 1991) shows that during the Miocene-early Pliocene the southern slopes of Chatham Rise continued to accumulate mainly biopelagic chalk, after which the record shows the same enrichment of illite + chlorite that occurs over a longer period at Site 1123. Site 594 clearly lay outside the main zone of terrigenous supply to the offshore until ~4 Ma (age model of Grant and Dickens, submitted [N5]) (Fig. F26), when terrigenous material first started to become a significant sediment component at ~170 mbsf (Kennett, von der Borch, et al., 1986). This date coincides with the start of the strong global Pliocene oxygen isotope enrichment trend (e.g., Zachos et al., 2001) and therefore probably represents lowered sea levels and the direct delivery of terrigenous material to the Bounty Trough. Consequently, overspilling turbidites may have reached Site 594 for the first time then. Sediment movement away from a Bounty Channel source would have been accentuated by the accelerated northward-flowing AAIW contour currents and STF frontal currents that developed within the Bounty gyre as climatic deterioration developed. Terrigenous sediment transported along-slope to Site 594 becomes even more marked at ~2.5 Ma (145 mbsf), above which are strong rhythms of alternating nannochalk/ooze and hemipelagic mudstone that correspond to the established Pliocene-Pleistocene climatic and oxygen isotope cycles (Nelson et al., 1985). McMinn et al. (2001) show that the 145-mbsf facies change at Site 594 corresponds to MIS 100 (~2.53 Ma) and that diatom and dinoflagellate assemblages change at that time in a way that "may reflect the fundamental reorganization of the pelagic ecosystem in response to...major climate change." These interpretations, and the inference of latest Pliocene climatic deterioration, are consistent also with data from Site 1122, on the left-bank levee of the abyssal Bounty Fan. There, a paraconformity straddles the period of ~5.0-2.2 Ma, pointing to enhanced DWBC/ACC flow at this time. The paraconformity separates early Pliocene DWBC contourite sediments below from higher sedimentation-rate turbidites above, within which fan turbidites come to dominate the depositional facies from ~1.7 Ma onward.

Whereas the course of the Bounty Channel through the Neogene was largely determined by the location of the rift axis of the Bounty Trough, the location of the Hikurangi Channel was strongly influenced by the position of the axis of the Hikurangi Trough, which the headward channel course still follows (Lewis, 1994). It is likely, therefore, that a proto-Hikurangi Channel has existed along the eastern North Island margin ever since subduction commenced at ~25 Ma in the late Oligocene. Like other known trench channels, however, the terminus of the paleo-Hikurangi channel was probably "captured" by the trench (e.g., Surveyor Channel) (Ness and Kulm, 1973), to which it delivered an aggradational turbidite fill (Lewis et al., 1998). For the Hikurangi Channel to provide sediments into the DWBC as it does today requires that its path was diverted out of the trench toward the east. This is unlikely to have occurred by simple channel avulsion both because of the height of the oceanward slope up onto the Hikurangi Plateau and because Coriolis-induced turbidity current overspill, and therefore avulsion is preferentially concentrated on left channel banks in the Southern Hemisphere (e.g., Carter, L., and Carter, R., 1988; cf. Menard, 1955). Rather, Lewis (1994) and Lewis et al. (1998) show that diversion of the channel was caused by the catastrophic collapse of a large portion of the inner trench slope, which Hall et al. (2002) estimated to have occurred in the early Pleistocene, at ~1.7 Ma. The resulting Ruatoria seaslide blocked the axis of the Hikurangi Trough and diverted the north-flowing Hikurangi Channel 400 km east across the Hikurangi Plateau to Rekohu Drift, which diverted it northeastward for a further 200 km to Rapuhia Scarp and into the path of the DWBC. Sediment flux records at Site 1124 (Carter, R., McCave, Richter, Carter, L., et al., 1999), on the crest of Rekohu Drift, suggest that overspilling distal turbidites from Hikurangi Channel first reached the drift at ~1.65 Ma and thereby confirm the earlier geological estimates of the timing of channel diversion from the Hikurangi Trough.

Farther south, the Solander Channel/Fan complex prograded toward the DWBC/ACC flow from about the late Miocene onward. However, it was a change in deep circulation that instigated the Solander contribution to the abyssal sediment budget, when the creation of a tectonic seaway through Macquarie Ridge in the Pliocene allowed the ACC to enter the lower Solander Trough, erode the Solander Channel/Fan complex (Schuur et al., 1998), and shift material east and north out of the basin. This eroded sediment contributed to drift deposition in the Emerald Basin (Carter, L., and McCave, 1997).

The establishment of these three major channel sediment injection points, the increased tectonic (e.g., Sutherland, 1996; Mortimer et al., 2001) and volcanic (e.g., Carter, L., et al., 2003, in press a) activity along the New Zealand plate boundary, and powerful paleoclimatic oscillations, all contributed to the major changes that occurred throughout the ENZOSS during the Pliocene-Pleistocene. The combined effect of these changes was to add a prominent terrigenous component to the previously biopelagic-dominant drift deposits of the DWBC (Hall et al., 2001, 2002; Joseph et al., in press; Handwerger and Jarrard, in press). In contrast to the older carbonate-rich Oligocene-Pliocene DWBC drifts at Sites 1123 and 1124, the Hikurangi Fan-Drift and Bounty Fan are dominantly terrigenous (McCave and Carter, L., 1997; Carter, R., McCave, Richter, Carter, L., et al., 1999; Carter, L., and McCave, 2002), though they do contain a small percentage of biopelagic material concentrated in interglacial intervals. In the case of Bounty Fan, at least, the DWBC redistributes some of its sediment north along the DWBC path, as shown by mineralogical tracer studies (Carter, L., and Mitchell, 1987).

Despite the dominance of channel sediment supply, flux studies by Hall et al. (2002), Joseph et al. (in press), and Carter, L., et al. (2000) as well as palynological analysis of Site 1123 (Mildenhall et al., in press) show that the ENZOSS also has a number of nonchannel sediment sources. In particular, Site 1123, which is located well above the discharge depth of Bounty Channel, received pulses of sediment over the past 3 m.y. that Hall et al. (2002) relate to periodic remobilization of sediment upstream by an invigorated ACC. Site 1123 also contains water-borne pollen derived from the North Island, which perhaps indicates transport to the site via the warm, south-flowing East Cape Current (Mildenhall et al., in press), and a minor eolian component was identified also by Thiede (1979) and Stewart and Neall (1984). Finally, the increasing tectonic tempo at the New Zealand plate boundary during the Quaternary was associated with an increase in the frequency of large volcanic eruptions from the Taupo Volcanic Zone (Carter, L., et al., 2003, in press a). Such activity contributed much air fall material to northern ENZOSS, and at Site 1124 macroscopic tephra make up 6% of the late Cenozoic sequence.

In addition to evidence for the longer-term changes summarized above, the Leg 181 record preserves a striking record of cyclic glacial-interglacial change in the ENZOSS area for the last ~2.6 m.y. and of parallel 41-k.y.-modulated cold-warm climate cycles back to the late Oligocene (cf. see the "Supplementary Material" contents list).

During late Cenozoic glacial lowstands, the seaward extension of New Zealand river mouths to feed directly into the heads of canyons and submarine channels introduced a flush of sediments to ENZOSS fans and drifts. Not surprisingly, therefore, glacial sediment fluxes can exceed interglacial counterparts by factors of 2-3 (e.g., Lean and McCave, 1998; Carter, L., et al., 2000; Hall et al., 2001). In addition, the glacial ENZOSS received more eolian detritus under the strong winds of the times (e.g., Stewart and Neall, 1984), more hemipelagic detritus under an enhanced wind-forced surface circulation, and possibly even more volcanic ash. All this sediment interacted with an invigorated glacial DWBC/ACC. Accordingly, sortable silt (Hall et al., 2001) and Antarctic diatoms, including reworked forms (Stickley et al., 2001), increased in abundance, whereas calcium carbonate content decreased because of both reduced productivity and corrosion under the expanding and accelerating Antarctic bottom waters (Hall et al., 2001). In contrast, interglacial sediment intervals record a reduction of terrigenous input and a reversal of the other factors listed above.

Detailed glacial-interglacial studies completed to date have been based mainly on the Site 1119 and 1123 records. The upper part of Site 1119 contains an expanded succession of the last ~250 k.y. (MIS 1-7), which confirms that topographic steering was a controlling influence on the location of the STF in the late Pleistocene and shows that strong along-front current flows were developed during glacial periods (Carter, R., et al., in press). Studies of the orbital sedimentary signal at Site 1123 show that the strength of DWBC flow also varied through time, with stronger flows during glacial periods back to 1.2 Ma (Hall et al., 2001). An alternating flow pattern applies between warm and cold period cycles in the early middle Miocene at Site 1123, which also contains a clear record of longer-term (>41 k.y.) changes in DWBC flow strength (Hall et al., 2003). Undoubtedly, more research will be accomplished in the future on specific aspects of glacial-interglacial cycling in the Leg 181 cores.