Channel-axis, overbank and drift sediment waves in the southern Hikurangi Trough, New Zealand

Marine Geology 192 (2002) 123^151 www.elsevier.com/locate/margeo

Channel-axis, overbank and drift sediment waves in the southern Hikurangi Trough, New Zealand Keith B. Lewis a; , Henry M. Pantin b a

National Institute of Water and Atmospheric Research Ltd (NIWA), Box 14901, Wellington, New Zealand b Earth Sciences Department, University of Leeds, Leeds, UK Received 12 December 2000; received in revised form 1 October 2001; accepted 2 August 2002

Abstract In the southern apex of the Hikurangi Trough east of New Zealand, swath bathymetry and backscatter, seismic profiles and cores from the proximal part of the 2000 km long Hikurangi Channel show different types of sediment waves in the channel axis and on the levee backslopes. These represent different responses to large turbidity currents. Locally, backslope sediment waves merge into sediment waves formed by a deep geostrophic flow. In the channel axis and on channel walls, sediment waves that have wavelengths of 1^4 km and heights of less than 6 m occur where cores show gravelly and coarse sandy turbidites. They are not apparent in layering below the channel axis. They are inferred to be thin, transitory, surface sand bodies migrating across coarser substrates and may have formed in a recent, perhaps the last, turbidity current. Their geometry suggests that they probably formed as antidunes in high-velocity flows that were much thicker than the channel depth. Overbank sediment waves also have similar wavelengths of 1^4 km but heights of up to 45 m. They have conspicuous sub-seabed layering to a depth of at least 400 m beneath the seabed. Seismic stratigraphy indicates that they have developed for most of the last 2 Ma, during near-vertical aggradation of the channel, its levees and overbank slopes, at average rates of 0.4^0.7 m/ka. The sediment waves migrate up levee backslopes towards the levee crest at rates of 3^20 m/ka, and are inferred to be formed beneath lee waves initiated in overbank flows by the linear perturbation of the levee. They are best developed on the outsides of bends, where centrifugal outflow of upper, ‘overspill’ parts of large turbidity currents might be concentrated. They are particularly well developed at right-hand bends on the left bank, where centrifugal and southern hemisphere Coriolis effects combine. Poorly developed sediment waves on the inside of some left-hand bends may be the result of Coriolis effects only. Overbank sediment waves decrease in height and wavelength away from the channel and continue into a depression along the slope-toe deformation front at the landward edge of the turbidite plain. Where the channel intersects the slope-toe depression, the absence of a left-bank levee suggests that overbank flows re-enter the channel. Along the toe of the Chatham Rise, sediment waves up to 5 m high and 2^5 km in wavelength migrate away from the channel and towards a slope-toe ‘moat’. They are inferred to be part of a drift deposit, perhaps associated with a branch of the Pacific Deep Western Boundary Current. 9 2002 Elsevier Science B.V. All rights reserved. Keywords: sediment waves; deep-sea channel; turbidity current

* Corresponding author. Fax: +64-4-386-2153. E-mail addresses: [email protected] (K.B. Lewis), [email protected] (H.M. Pantin).

0025-3227 / 02 / $ ^ see front matter 9 2002 Elsevier Science B.V. All rights reserved. PII: S 0 0 2 5 - 3 2 2 7 ( 0 2 ) 0 0 5 5 2 - 2

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1. Introduction Sediment waves are a feature of turbidite channel systems, with long wavelength bedforms reported from channel axes and climbing up the backslopes of channel levees (Wynn and Stow, 2002). Turbidity currents are rare on a human timescale and are extremely di⁄cult to measure directly. However, the form and orientation of sediment waves provide important clues to the nature of the sediment-laden £ows that form channels, levees and turbidite plains. Deep-ocean channels, which range from many hundreds to several thousand kilometres long, are much less frequently studied than the much shorter and more common fan-channel systems (Carter, 1988; Klaucke et al., 1998). The full length of the 2000 km long Hikurangi Channel, east of New Zealand (Fig. 1), has only recently been discovered (Lewis, 1994, 1999), and it is one of only a small number of deep-ocean channels that is known to be presently active (Lewis and Barnes, 1999). In this paper, we examine channel and

overbank sediment waves at the proximal end of the Hikurangi Channel, in the southern Hikurangi Trough, east of central New Zealand (Fig. 1), as an aid to understanding the nature of turbidity £ows that travel enormous distances. The southern Hikurangi Trough is at the southern extremity of the Hikurangi^Kermadec^Tonga subduction trench system, in an area where highly oblique subduction of the Paci¢c Plate merges into the intra-continental oblique collision and transform, associated with the Alpine Fault of South Island (Fig. 1) (Lewis et al., 1998; Barnes et al., 1998). Onshore compression and uplift results in an abundant supply of sediment from rapidly rising mountains. This is supplied, via northward nearshore drift and the Kaikoura Canyon, to the collision foredeep trough and sediment-£ooded structural trench that together constitute the southern Hikurangi Trough. There, in an area devoid of typical fan structures, sediment waves record the nature of £ows near the start of a s 2000 km long conduit for sediment from rising mountains to deep ocean basin.

Fig. 1. Location of the southern Hikurangi Trough (rectangular box, lower left, shows location of Figs. 2 and 3) o¡ the southern tip of North Island and northeastern South Island, New Zealand. Bathymetric contours at 500 m intervals. Route of the 2000 km long Hikurangi Channel shown by thick broken line. Convergent plate boundary shown by broken £agged line. Large open arrows mark the route of the Paci¢c Deep Western Boundary Current (DWBC). Smaller open arrows indicate an inferred shallow branch of the DWBC along the northern margin of the Chatham Rise.

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2. Methods The data set used in this study includes swath and narrow-beam sounding bathymetry, swath backscatter, high and low frequency seismic pro¢les (Fig. 2), and cores (Fig. 3) collected over several decades. One of the primary data sets is swath backscatter and bathymetry obtained in August 1994 using an MR1 system chartered from the University of Hawaii and towed from the New Zealand vessel Giljanes. Along the continental margin, MR1 data are supplemented by poorer resolution backscatter but higher resolution bathymetry collected in December 1993 by the multibeam Simrad EM12 system hull-mounted on the French research vessel Atalante. Both data sets were collected for the joint France^New Zealand GeodyNZ study of the o¡shore plate boundary zone through New Zealand (Collot et al., 1996). Extensive sets of 3.5 kHz and single-channel seismic data, using 40 and 120 in3 airguns, were collected from 1988 to 1994 for studies of tectonic^sedimentary interactions at the southern end of the Hikurangi subduction system and adjacent Chatham Rise (Barnes, 1992; Barnes, 1994a; Barnes and Mer-

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cier de Lepinay, 1997). These data are supplemented by archived seismic pro¢les collected for earlier scienti¢c studies and commercial reconnaissance (Lewis and Bennett, 1985; Lewis and Pettinga, 1993). The 3D view of the seabed obtained from backscatter, bathymetry and seismic pro¢les is ‘ground-truthed’ using piston and gravity cores.

3. Source canyon sediments and bedforms Of the several large canyons that feed into the apex of the Hikurangi Trough, only the Kaikoura Canyon presently approaches close enough to shore to intercept a vigorous, near-shore sediment transport system (Fig. 3) (Lewis, 1994). An estimated 1 500 000 m3 of mud, sand and gravel, moving northwards along the shelf of northeastern South Island, are trapped in the canyon head each year (Lewis and Barnes, 1999). The accumulated sediment is stored in the canyon head until it collapses under severe earthquakes stress associated with dislocation of nearby plate-boundary faults. The steepness (25^30‡) of the canyon headwall (Fig. 4A,B) facilitates transformation of

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Fig. 2. Seismic tracks from the southern Hikurangi Trough that were used in this study. Hikurangi Channel marked by hatched walls. Thickened lines indicate swath bathymetry and backscatter images also available on these tracks.

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MARGO 3224 18-11-02 Fig. 3. Bathymetric terrain model of the southern Hikurangi Trough showing distances along Hikurangi Channel and overbank trough at 50 km intervals from source at head of Kaikoura Canyon. Location of cores (Fig. 5) and detailed maps (Figs. 6, 9 and 12) are indicated.

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Fig. 4. Plots against distance from Kaikoura in kilometres of (a) depth, (b) slope, (c) sediment wave wavelength, for both channel-axis (solid lines) and overbank areas (broken lines). Note that 400 km on the overbank route is the same place as 560 km along the channel (see Fig. 3).

failed sediments into major turbidity currents. Radiocarbon dates of wood fragments in a core (Q312) (Fig. 5) recovered from strongly backscattering gravel turbidites in the lower Kaikoura Canyon (Figs. 3 and 6) suggest that the last two major turbidity currents may correlate with rupture of the nearby Hope Fault about 170 years ago and about 300 years ago (Carter et al., 1982; Lewis and Barnes, 1999). The rest of the dozen or so canyons around the southern end of the Hikurangi Trough generally incise only the outer continental shelf and it is inferred that they intercept signi¢cant shelf sediment transport systems only during glacial ages (Herzer, 1979; Carter, 1992). Small canyons im-

mediately north of Kaikoura Canyon and on either side of Cook Strait Canyon (Fig. 6) have moderate to strong backscatter but they originate, not at a single ‘point’ source like Kaikoura Canyon, but at rows of small, strongly backscattering gullies, which are inferred to be a product of erosion and headward sapping at £uid seeps (Lewis and Marshall, 1996). However, even with slumping of canyon walls and mass £uid expulsion during severe earthquake stress (Normark and Piper, 1991), it is unlikely that modern input from these is signi¢cant compared with input from Kaikoura Canyon. The Cook Strait Canyon system, which has a con£uence with the Hikurangi Channel about 180

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km from its canyon-head source (Fig. 6), is presently blocked by slumps in two of its three branches (Carter, 1992), and the unblocked branch does not have access to an abundant supply of sediment. However, a core (U632) from the lower Cook Strait Canyon (Figs. 3 and 5) has graded granule gravel and coarse sand to mud turbidites. Also, swath images of the lower Cook Strait Canyon show moderate to strong backscatter, with crescentic or chevron-shaped bedforms that are concave downslope and have a wavelength of about 2 km (Figs. 6 and 7A). Inferred input from Cook Strait has been reported from turbidites near the distal end of the Hikurangi Channel (McCave and Carter, 1997), but it is probable that this contribution occurs mainly during periods of glacially lowered sea-level. Much of the seabed in the area is draped by inferred ‘hemipelagic’ mud (Barnes, 1994b). The inference is that this is derived both from slow settling through the water column, and from dilute, slow-moving, bottom, nepheloid-type layers that deposit preferentially in topographic lows (Lewis, 1973). In the Hikurangi Trough, hemipelagic mud forms bands that are lighter in colour and less silty than the muddy tops of turbidites, and may form up to about one third of the sediment in cores from overbank areas (Lewis and Kohn, 1973; Lewis, 1985).

4. Channel-axis sediment waves

Kaikoura Canyon and proximal foredeep Hikurangi Trough (Figs. 6 and 7B) are in an area where cores (e.g. core Q312; Figs. 3 and 5) and strong backscatter indicate a gravelly substrate. The bands, which have not been sampled, could possibly indicate gravel waves or, more probably if the sonar imaging is mainly of textural di¡erences, they may represent ¢ner-grained, probably sandy, sediment waves moving over gravel. At several places in the foredeep trough, particularly near the edges of areas with strong backscatter and at the lower end of tributary canyons that begin at rows of gullies, faint transverse bands of weaker backscatter suggest bedforms with a wavelength of about 0.5^1.0 km (Fig. 7B). These may also indicate degraded gravel bedforms, or more probably thin waves of ¢ner sediment moving over the gravel. In the distal part of the foredeep trough, which extends to about 130 km from the canyon-head source, the trough/channel is con¢ned between a high left bank, with inferred seep-generated gullies and amphitheatre-like avalanche scars (Fig. 7C), and a right bank formed by the current-scoured £ank of the Chatham Rise (Barnes, 1992) (Fig. 6). The channel is generally £at-£oored with an axial slope that decreases gradually from 0.6‡ to 0.3‡ (Fig. 4B). Backscatter is mottled but generally stronger than on the adjacent slopes, with strong, speckled backscatter from blocky debris-£ow deposits derived from both banks (Figs. 6 and 7C) (Barnes, 1992). Downstream from the debris-£ow deposits, faint transverse backscatter bands might represent degraded sediment waves.

4.1. Bedforms in the con¢ning foredeep trough At 50 km from its headwall source, the erosional Kaikoura Canyon turns into the structural foredeep between obliquely overthrusting continental slope on the western or left-bank (looking downstream) and Chatham Rise to the right (Figs. 1, 3 and 6) (Barnes, 1994a). This foredeep, which forms the apex of the Hikurangi Trough, has an almost £at £oor with an axial slope of 0.6‡ (Fig. 4B) that is only 5^10 km wide and has no separate, incised channel. Trough and channel are the same. Faint, transverse backscatter bands with a wavelength of 0.5^1.0 km in the lower

4.2. Channel-axis bedforms in the foredeep to trench transition At about 130 km from the head of the Kaikoura Canyon, and about 80 km from its toe, the intra-continental foredeep trough merges into a turbidite-¢lled structural trench (Figs. 3 and 6). Beyond this point, Hikurangi Channel and Hikurangi Trough are distinct morphological features. For a further 50 km, to the con£uence with the Cook Strait Canyon, the incised channel continues with its former N60‡E trend, but the axial slope is reduced to 0.15‡ (Fig. 4B). Bathy-

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Fig. 5. Diagrammatic representation of cores from channel-axis and overbank areas: left bank above the channel-axis line and right bank from below. Horizontal scale is estimated percentage of silt in the mud fraction. Positions shown in Fig. 3.

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Fig. 6. Bathymetric, structural and backscatter features, including sediment waves, between Kaikoura Canyon and Cook Strait Canyon. Contours at 50 m intervals. Areas of strong backscatter in the canyon and foredeep trough indicate areas of gravel and debris avalanche deposits. Transverse bands of medium backscatter indicate sediment waves. Rectangles and lines show location of images and pro¢les in Figs. 7 and 8.

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metric, seismic and backscatter data show that the 7 km wide channel £oor is characterised by transverse sediment waves. Those near the foredeep blocky debris-£ow deposits have wavelengths of about 1 km, but those near the junction with Cook Strait Canyon have wavelengths of 3^4 km (Figs. 4C, 6 and 7D). Core S880, from the channel axis in the proximal part of the foredeep to trench transition zone (Figs. 3 and 5), and core E7, from just beyond it (Figs. 3 and 5), contain medium to coarse sand turbidites. Swath bathymetry indicates that the long-wavelength sediment waves are only 4^6 m high, with a steeper upstream face and stronger backscatter coming from the troughs and lower parts of the downstream faces (Fig. 7D). The channel axis sediment waves are low, surface features, with no subseabed layering indicative of growth over an extended period of time (Fig. 8A), unlike overbank sediment waves, which are discussed later. 4.3. Channel-axis bedforms along the foot of the Chatham Rise At 180 km from its canyon-head source, the Hikurangi Channel reaches its con£uence with the Cook Strait Canyon system (Figs. 3 and 9). There, the channel turns 55‡ to the right, almost bisecting the angle between the foredeep trough and Cook Strait Canyon, and the channel £oor is 12 km wide, almost the same as the combined width of the Cook Strait Canyon and the precon£uence Hikurangi Channel. This geometry suggests a long-term equilibrium between channel and input from both sources. From 180 to 310 km from source, the channel is almost straight and heads eastwards towards the outer edge of the 70 km wide turbidite plain with an axial slope of about 0.1‡ (Figs. 3, 4B and 9). The £at channel £oor narrows from about 12 km to about 4 km, whereas the depth of the channel below the higher left-bank levee increases from 230 to 290 m. Channel-axis backscatter is generally weaker than surrounding areas, but transverse banding indicates channel-£oor sediment waves that have a wavelength of 2^3 km near to Cook Strait, reducing to 0.6 km further away (Figs. 4C, 9 and 10A). Bathymetric data

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suggest that these sediment waves are less than 5 m high, and seismic pro¢les show no evidence beneath the channel £oor of long-term development of sediment waves. Channel-axis core E7 (Figs. 3 and 5) contains graded coarse sand to silty ¢ne sand turbidites up to 0.8 m thick, overlain by a 0.15 m thick mud layer (Pantin, 1972). Core U660 contains medium to ¢ne sand turbidites up 0.4 m thick (Figs. 3 and 5). Despite low backscatter at the seabed, seismic pro¢les show high-amplitude re£ectivity to a depth of at least 400 m (450 ms TWT (two-way travel time)) beneath the channel, with no indication of growing bedforms (Figs. 8B, 11). Dated horizons, traced throughout the southern trough (Barnes and Mercier de Lepinay, 1997), indicate that these highamplitude, channel-£oor re£ectors extend back in time more than 800 ka and almost to the 2 Ma re£ector. Thus, the channel axis has remained in almost the same geographical position throughout all or most of the Quaternary. During that time, it has aggraded and remained close to the foot of the Chatham Rise, although in places there has been southward migration of up to 5 km as turbidites onlapped the Chatham Rise. Backscatter banding, similar to that in the channel axis, extends obliquely up the walls of the channel to a height of more than 200 m above the channel £oor (Figs. 9 and 10A). The banding, which has wavelengths of 0.8^2 km, is generally continuous with that on the channel £oor. It is inferred to represent the same type of low sediment waves that occur in the channel axis. Sediment waves are best de¢ned on the left (northern) wall after the sharp right-hand bend at the con£uence with Cook Strait Canyon between 190 and 230 km from the canyon-head source. Here the channel wall slopes at 2^4‡. Transverse banding also occurs on the right bank just beyond this, on the inside of the bend at 225^240 km from source. Between 260 and 310 km from source, the Hikurangi Channel narrows, starts to become sinuous and undercuts the Chatham Rise (Fig. 9). Faint backscatter banding in the channel axis is mainly sub-parallel with the channel walls (and inferred £ow direction) (Fig. 9), and generally downstream from £ow-modi¢ed debris deposits. The largest debris deposit or slide occurs where

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a bend cuts into the toe of the Chatham Rise at about 290 km from source (Fig. 9) (Barnes, 1992). It is about 16 km wide and projects about three quarters of the way across the channel, leaving a gap less than 1 km wide between its toe and the northern wall, which presumably constricts turbidity currents. 4.4. Channel meandering back to the deformation front From 310 to 470 km from source, the channel forms a continuous series of meander bends, each with a radius of 6^10 km (Figs. 3, 10B and 12), and swings from the outer edge of the turbidite plain back towards the deformation front. Its £oor, which narrows to 1.5^2.0 km wide, has weak backscatter (Fig. 10B) but strong re£ectivity (Fig. 8C and 11C), with no indication of channelaxis bedforms either at the seabed or beneath it. Seismic pro¢les show the channel to a depth of at least 400 m (450 ms TWT) beneath the present seabed, in more or less its present position, with no evidence of channel migration or avulsion (Fig. 11C). Trough stratigraphy (Fig. 8C) (Barnes and Mercier de Lepinay, 1997) suggests that the aggrading meanders have hardly moved geographically since latest Pliocene or early Pleistocene times despite 400 m of aggradation.

5. Overbank sediment waves: morphology and £ow directions 5.1. Sediment waves beyond the right bank Just before the con£uence with Cook Strait Canyon, seismic pro¢les show sediment waves

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up to 45 m high, which have migrated up the backslope from a narrow overbank plain towards the right-bank (looking downsteam) levee crest (Figs. 6 and 8A). Bands of stronger backscatter correlate with the wave crests and upslope faces and show that the sediment waves have a wavelength of 3^4 km and are aligned slightly oblique to the channel (Fig. 6). The levee crest is about 150 m above the channel £oor (Fig. 8A) and parallel-layered turbidites onlap the Chatham Rise to about 200 m above the channel £oor. Seismic pro¢les show that these sediment waves have migrated by more than 7 km towards the levee crest while aggrading by more than 400 m (450 ms TWT). This represents about 17.5 m lateral advance for each 1 m of aggradation. Tentatively extrapolating seismic stratigraphy in the southern Hikurangi Trough (Barnes and Mercier de Lepinay, 1997), the 400 m of aggradation represents at least 1 Ma of deposition, so that the largest sediment wave has migrated towards the levee crest (and channel) at up to 7 m/ka, whilst aggrading at up to 0.4 m/ka. As the channel skirts the foot of the Chatham Rise, the right overbank area is narrow and no overbank sediment waves are developed until after the ¢rst left-hand bend near the outer edge of the turbidite plain at 310^320 km from source. There, sediment waves on the outside of the bend beyond the right-bank levee are 5^20 m high and 1.5^3 km in wavelength, and migrate up the back of the levee (Fig. 9, 11B). They extend to the channel’s right bank on the far side of the meander loop at 330^340 km from source, implying that centrifugal out£ows ‘cut the corner’ across the meander loop and back into the channel. On the lower £ank of the Chatham Rise immediately to the south, this levee-back, sediment-wave ¢eld

Fig. 7. Backscatter images of the proximal Hikurangi Trough with dark zones indicating high backscatter and arrows indicating £ow direction (positions shown in Fig. 6). (A) Lower Cook Strait Canyon showing transverse, crescentic backscatter bands (I) that may represent sediment waves with a wavelength of about 2 km. (B) Small, faint backscatter bands at bottom of Kaikoura Canyon (I), at the bottom of seep-fed canyon (II), and at the edge of gravel zone (III); the bands probably represent sediment waves either in gravel or in sand moving over gravel. (C) Edge of the foredeep trough showing an avalanche scar and seep gullies on northwest wall (I), and strongly backscattering debris avalanche deposits on the channel axis (II). (D) Transverse sediment waves immediately before the con£uence with the Cook Strait Canyon (contours at 50 m intervals); sediment wave crests marked by opens triangles and troughs by solid triangles, with steeper faces generally facing up-channel and stronger backscatter on downslope faces.

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merges into sediment waves (described later) that migrate away from the channel and may be part of a drift deposit. At 440^470 km from source, bands of backscatter on the right bank are aligned normal to a line tangential to the outside of a left-hand bend (Figs. 10B and 12). Their orientation is compatible with centrifugal out£ow from the right bank. 5.2. Sediment waves on the left-bank levee From the point where the channel turns right along the toe of the Chatham Rise, large overbank sediment waves migrate up the backslope of a high left-bank levee that is 210^280 m above the channel £oor and typically 50 m higher than a poorly developed levee on the right bank. These sediment waves range from 10 to 45 m high, and have a wavelength of 2^5 km and a crest length of up to 10 km (Figs. 8^13). Their height typically decreases down the 0.2^0.3‡ gradient (Fig. 4B) away from the levee crest. Such bedforms are a common feature of channel levee systems elsewhere (Normark et al., 1980; Carter et al., 1990; Normark and Piper, 1991; Nakajima et al., 1998). Seismic pro¢les show that the largest sediment waves near the levee crest developed during the aggradation of about 400 m (450 ms TWT) of sediment (Figs. 8B,C), commonly by the amalgamation of several smaller sediment waves at depth (Fig. 8B). They show that the buried sedimentwave crests migrate towards the modern levee crest (and channel) at rates of 4^50 m (typical values of 12^30 m) for each 1 m of aggradation (Figs. 8B,C and 13A,B). Using the seismic stratigraphy of Barnes and Mercier de Lepinay (1997) (Fig. 8B,C), the left-bank levee begins to form

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just above a 2.03 Ma re£ector and, at several places, the ¢rst small sediment waves begin to develop at the back of the levee soon after this, perhaps about 1.6^1.8 Ma (Fig. 8C). At other places, development of levee sediment waves is delayed until about the time of a 800 ka re£ector (Fig. 8B). Since then, sediment waves have migrated towards the levee crest at rates of 3^20 m/ka during vertical aggradation of 0.4^0.7 m/ ka. Such rates are high compared with estimates for sediment waves beneath bottom currents and unchannellised turbidite systems (Wynn et al., 2000a) They are comparable with, but generally higher than, those for overbank sediment waves of the Bounty Channel, which is almost inactive in interglacial periods (Carter et al., 1990). A few backscatter images and correlations between all available sounding tracks suggest that left-bank sediment waves are typically oriented sub-parallel to the channel and to the regional bathymetric contours. They are best developed on the outside of right-hand bends, notably at 180, 260m, 330 and 380 km from source (Figs. 8^13). At these places, well-formed sediment waves are oriented transverse (perpendicular) to a line that is tangential from the outside of the bend, implying centrifugal out£ow at the bend (Fig. 10B). In the sinuous section of the channel, at 280, 360and 410 km from source, small sediment waves and faint backscatter liniations also occur along the levee on the insides of bends (Figs. 8C, 9, 10B and 12). 5.3. The overbank plain beyond the left-bank levee The overbank plain between the channel and the deformation front is up to about 60 km

Fig. 8. Seismic pro¢les showing aggradation of the Hikurangi Channel and growth of overbank sediment waves (for locations see Figs. 6, 9 and 12) with stratigraphy from Barnes and Mercier de Lepinay (1997) where (1) is 190 ka, (2) is 400 ka, (3) is 800 ka and (4) is 2.03 Ma. (A) Before the con£uence with Cook Strait Canyon showing (I) low channel-axis sediment waves and large, overbank sediment waves on the right bank advancing towards the channel. (B) At the toe of Chatham Rise (right) showing vertical aggradation of high-amplitude channel-axis re£ectors, and well developed overbank sediment waves on the left-bank levee and backslope plain. (C) In the meander section showing near vertical channel aggradation and levee growth, with development of sediment waves and channel, soon after 2.03 Ma. Note development of backslope plain sediment waves about 400 ka and again after 190 ka. Vertical scale on left in seconds of two-way travel time and on right in metres (assuming velocity of sound of 1500 m s31 ).

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wide and inclined towards the northeast at slopes ranging from 0.3‡ near Cook Strait to about 0.1‡ near where the channel again reaches the deformation front (Figs. 4B and 12). The overbank plain is slightly deeper along the deformation front, which is at the toe of a well developed accretionary margin (Figs. 3 and 12) (Lewis and Pettinga, 1993; Barnes and Mercier de Lepinay, 1997). On the plain beyond the left-bank levee crest, sediment waves gradually reduce in size from typically 8^12 m high and 2^4 km in wavelength just beyond the left-bank levee (Figs. 8B,C) to only 3^7 m high and 1^2 km in wavelength where they occur along the deformation front (Fig. 11A). Backscatter images show them trending almost parallel with the contours, and perpendicular both to the deformation front (Fig. 9) and, presumably, to any £ow along the deformation-front depression. A core, E12, from the depression close to the deformation front has a sequence of 0.1^0.4 m thick, graded silt turbidites (Figs. 3 and 5) (Pantin, 1972). Seismic pro¢les show that parallel-bedded overbank turbidites are up to nearly 5 km thick ( s 4 s TWT) at the deformation front in the southern Hikurangi Trough and have accumulated since mid to late Miocene times (Barnes and Mercier de Lepinay, 1997; Lewis et al., 1998). They also indicate that sediment waves on the plain beyond the left-bank levee have developed in two main phases (Figs. 8B,C). The earlier phase occurred at about the time of the 400 ka re£ector identi¢ed by Barnes and Mercier de Lepinay (1997). The later one occurred after a 190 ka re£ector but are somewhat degraded at the seabed. Where the Hikurangi Channel is about 460^470 km from source, a meander loop incises the low part of the turbidite plain along the deformation front (Fig. 12). Despite the left-bank levee being well-developed everywhere else beyond the con£uence with the Cook Strait Canyon, the levee is absent from this section (Fig. 12). The pattern of a meander, with no left-bank levee, intersecting

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the slope-toe trough of the turbidite plain is repeated where the channel approaches a volcanic knoll 550^560 km from source. Any overbank £ows that were focussed along the deformationfront depression must re-enter the channel at one of these points.

6. ‘Wrong way’ sediment waves Beyond the large avalanche deposit in the channel 290^300 km form source (Fig. 9), sediment waves up to 5 m high and 2^5 km in wavelength appear to migrate up the £ank of the Chatham Rise to a height of 320 m above the channel axis. (Figs. 11B,C and 13C). Unlike nearby overbank sediment waves, they migrate away from the channel, not towards it. Instead, they migrate towards a 2^5 km wide, 10^20 m deep ‘moat’ along the base of the Chatham Rise. As discussed later, the ‘moat’ is most probably formed by a deepwater geostrophic (bottom) current. We suggest that the sediment waves migrating towards it are not related to the Hikurangi Channel, but form part of a slope-toe drift or contourite deposit associated with the geostrophic £ow. Tentative correlations between sounding lines suggest that the sediment waves trend east^west, sub-parallel with the slope-toe ‘moat’. Down the slope towards the Hikurangi Channel, perhaps at a water depth of about 2750 m, the drift deposit with its associated sediment waves merges into right-bank turbidites with sediment waves (Fig. 11B,C). Airgun pro¢les show that the ‘moat’ and associated sediment waves occur to a depth of about 500 m (600 ms) beneath the seabed (Fig. 11C). The pro¢les also indicate that it was in¢lled several times by debris £ows from the Chatham Rise slope. Age correlations (Barnes and Mercier de Lepinay, 1997) suggest that the ‘moat’ and its associated sediment waves date back to at least the deposition of a regional 2.03 Ma re£ector.

Fig. 9. Distribution of channel-axis, levee and overbank sediment waves, marked by wave crests (solid lines) and strong backscatter in the apex of the Hikurangi Trough subduction trench. Contours at 50 m intervals. Legend as Fig. 6. Rectangles and lines show location of images and pro¢les in Figs. 8, 10, 11 and 13).

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MARGO 3224 18-11-02 Fig. 10. Backscatter images of the Hikurangi Channel (position and interpretations shown in Figs. 9 and 12) showing stronger and weak backscatter banding from (A) transverse channel-axis and oblique wall sediment waves just after con£uence with Cook Strait Canyon, (B) overbank sediment waves on left (upper) and right (lower) banks related to centrifugal and Coriolis over£ow from channellised turbidity currents at meander bends. Arrowed boxes show distance along channel axis from source at head of Kaikoura Canyon.

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Therefore, it is at least as old as the Hikurangi Channel in this area.

7. Flow conditions from sediment waves 7.1. Channel-axis £ow As the conduit for the densest, coarsest and fastest moving part of turbidity currents, the channel axis has sediment waves that are signi¢cantly di¡erent from classic, high, muddy, upslope migrating, levee-overbank sediment waves. Channel-axis sediment waves are low and essentially non-depositional, with no detectable subsurface history. Thus, there is no evidence of persistence in either space or time. The techniques available permit no resolution of their internal structure. Despite their long wavelengths, the sediment waves of the channel axis are only a few metres in height, in contrast to the very much larger overbank sediment waves. Also, they are surface features only, so they may be representative of only very recent events, unlike overbank sediment waves which appear to be in very long-term equilibrium with over£ow from a channel (Hesse, 1995). Below a sur¢cial mud layer, sediment in the channel axis reduces gradually in grain size from gravel turbidites at the foot of the Kaikoura Canyon, to coarse sand turbidites in the distal foredeep trough and proximal subduction trough, to very ¢ne sand turbidites where the channel meanders back to the deformation front. Deposition from turbidity currents occurs where the £ow becomes overloaded, with deposition re£ecting a loss in £ow ‘capacity’, which is sediment volume past a point in a given time, rather than the ‘competence’, which is the £ow speed required to transport the particles being deposited (Hiscott, 1994). Nevertheless, the reduction in particle size does indicate a signi¢cant reduction in £ow velocity along the channel. For instance, gravels require velocities of over 10 m/s to be transported in suspension (Piper et al., 1985), with wide variations depending on the nature of the £ow transporting them (Normark and Piper, 1991), whereas

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the very ¢ne sand can be carried in suspension at speeds as slow as 0.35 m/s (Heathershaw, 1981). These velocity estimates based on particle size are minimum values (Komar, 1985). The transverse bands of high and low backscatter in the foredeep trough (Fig. 7B) have been tentatively inferred to represent ¢ne-grained sediment waves moving over gravel, their 0.5^1 km wavelengths (Fig. 7B) being generally greater than published wavelengths for coarse-grained waves (Morris et al., 1998; Piper et al., 1985; Piper and Kontopoulos, 1994; Wynn et al., 2000a,b). Elsewhere, channel-axis backscatter banding is also inferred to represent textural di¡erences, although stronger backscatter does not necessarily correlate with the troughs between low sediment waves (Fig. 7D). The channel-axis sediment waves are unlikely to be formed by undular bores, which represent a relatively weak hydraulic jump occurring downsteam of a transformation from supercritical to subcritical £ow (Fisher, 1983), because there would need to be a signi¢cant decrease in gradient, which is not observed in the Hikurangi Channel. Nor are they likely to be caused by lee waves, because lee waves (in the lee of an obstruction) are generally associated with £ows considerably slower than those to be expected of a channellised turbidity current (Flood, 1988). It is inferred that the channel-axis sediment waves are likely to form in relatively strong £ows as either dunes or antidunes. Dunes are progressive, transverse, sediment waves formed in subcritical £ows (Leeder, 1982), whereas antidunes are transverse, typically regressive, sediment waves formed in £ows where the overall Froude number Fro s W0.84 (Alexander and Morris, 1994). Antidunes are in phase with the waves at the £ow surface and, although they typically move upstream, they may also remain stationary or even move downstream. Dunes, on the other hand, are out of phase with the water waves and only move downstream. Both types develop where there are low-amplitude, random perturbations in bed height or roughness. The wavelength (L) of sediment waves formed by such £ows is dependent on £ow thickness (H). The relationship varies but is approximated by LW2ZH for both

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dunes (Leeder, 1982) and antidunes (Normark et al., 1980; Zeng et al., 1991). Discussion on the conditions for growth of these bedforms is beyond the scope of this paper. However, the Hikurangi Channel axis sediment waves, particularly those with steeper up-channel faces, are more likely to be antidunes if they are generated by the main (strongest) £ows. If they are formed or modi¢ed by dilute, tail-end £ows, they could possibly be a product of lee waves passing over a low-amplitude, local obstruction that generates a ‘ship-wave’ pattern in the £ow. Flows are likely to be subcritical in such situations and then the shipwave consists of two sets of waves, one oblique and diverging symmetrically downstream of the obstruction, and the other curved transverse and convex upstream (Scorer, 1978). In supercritical £ows only the oblique set of waves will form (Lamb, 1932). In the case of the crescentic and V-shaped bedforms in lower Cook Strait, it may be that the centre of the sediment waves has migrated upslope more rapidly in the centre of the channel where £ow velocities are highest (Wynn et al., 2000b). This possibility is consistent with these bedforms being oblique antidunes, perhaps generated by local obstructions such as pre-existing bedforms, each of which gives rise to an oblique ‘ship-wave’ pattern associated with supercritical £ow (Lamb, 1932). These aligned, V-shaped bedforms are super¢cially similar to the scattered ‘chevrons’ formed by uncon¢ned turbidity currents down-fan from the Valencia Channel mouth (Morris et al., 1998), but we are unable to compare their form and origins. The great length of the Hikurangi Channel strongly suggests that turbidity currents £owing along its proximal part are generally in a state of autosuspension (Pantin, 1979). Calculations

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and theoretical arguments based on the model by Parker et al. (1986) indicate that the autosuspension £ows must also be supercritical. If so, then the sediment waves in the Hikurangi Channel are probably antidunes, or at least originated as such. The steeper up-current faces of the sediment waves immediately before the con£uence with the Cook Strait Canyon support this interpretation. If the relationship LW2ZH applies in the proximal Hikurangi Channel, then the 3^4 km wavelength sediment waves immediately before Cook Strait Canyon formed beneath a £ow that was about 500 m thick. Beyond the sharp bend at the con£uence with the Cook Strait Canyon, a reduction in the wavelength of channel-axis sediment waves to an average of about 2.2 km is consistent with £ow-stripping of the top of the £ows to overbank areas and reduction in £ow thickness to about 350 m. The sediment waves on the left wall of the channel immediately after the con£uence appear to be essentially continuations of the low, channel-axis sediment waves and, despite some obliquity of trend, presumably formed in a similar £ow regime. Further downstream reductions in the wavelength of sediment waves may result from loss of speed, concentration, or strati¢cation within £ows that were still capable of overtopping a left-bank levee that is 250^300 m above the channel axis. Some small sediment waves that are convex downstream may be former straight-crested antidunes that subsequently behaved as dunes when modi¢ed by later, weaker currents. They may be products of the last major turbidity current, with reworking at a late stage of this event, as demonstrated on the Laurentian Fan (Piper et al., 1985). The longitudinal bedforms that occur between 260 and 310 km from source in the Hikurangi

Fig. 11. Pro¢les across the southern Hikurangi subduction trough showing horizontal and vertical extent of overbank sedimentwave ¢elds, relative to the Hikurangi Channel (cross in circle) and to the ‘moat’ at the toe of the Chatham Rise. See Figs. 9 and 12 for locations. (A) 3.5 kHz pro¢le showing overbank plain sediment waves extending to the depression along the toe of the deformation front (left). (B) 3.5 kHz pro¢le showing right and left bank sediment waves and the low sediment waves on drift deposit associated with slope-toe ‘moat’. (C) Single-channel airgun pro¢le showing growth, mainly since 2.03 Ma (re£ector 4), of ‘moat’, drift deposits, channel, levees (see Fig. 8C) and overbank plain. Vertical scale in seconds TWT (left) and metres (right) assuming velocity of sound of 1500 m s31 .

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Channel may indicate seabed scour. However, the Kaikoura Canyon^Hikurangi Channel is devoid of deep, local scour structures that would indicate major, stationary, £ow transformation or hydraulic jump, such as occur at the bottom of some canyons (Kenyon et al., 1995). The reason may be that there is no sharp change in slope that would trigger repeated £ow transformations in the same locality (Fisher, 1983). 7.2. Overbank £ows In contrast to sediment waves in the channel axis, overbank sediment waves are depositional, composed of relatively ¢ne-grained sediment and persistent in both space and time. Their relatively high amplitude permits good resolution of their form and internal structure. Southern Hikurangi Trough overbank deposits consist mainly of ¢ne sand^silt turbidites that suggest £ow-stripping away from the channel of the dilute, supra-channel ‘overspill’, which is carried along with the more concentrated, channellised part of large, strati¢ed turbidity currents (Klaucke et al., 1998; Peakall et al., 2000), and which helps to build up the channel-parallel levees. The well developed left-bank levee and overbank sediment waves show that loss from the supra-channel part of large turbidity currents is predominant to the left (looking downstream), with the overbank £ows continuing down the slope towards the lowest part of the turbidite plain along the deformation front (Figs. 9, 11 and 12). This implies a southern hemisphere Coriolis e¡ect. An alternative mechanism involving a consistently northwestward £owing deep ocean current is unlikely in the southern Hikurangi Trough, where deeply penetrating geostrophic eddies, which would produce a less consistent £ow, are inferred (Roemmich and Sutton, 1998). The increased development of large sediment waves in lines tangentially away from the outsides of

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bends implies that loss over the left-bank levee is concentrated tangentially from the outside of right-hand bends, where Coriolis and centrifugal e¡ects combine. Between right-hand bends, low, channel-parallel sediment waves (Figs. 10 and 12) suggest Coriolis e¡ects only. On the right-bank, concentric, tangentially aligned, sediment waves at left-hand bends show where centrifugal e¡ects exceed Coriolis forces. For deep-ocean channel systems in long-term equilibrium, the loss to overbank plains on both sides can be remarkably constant along their length, despite lateral variations in sinuosity and levee height (Hesse, 1995; Klaucke et al., 1998). The overbank sediment waves adjacent to the Hikurangi Channel (Fig. 13A,B) are inordinately persistent in both time and space. They have developed over a nearly 2 Ma period, so they are a very long-term, equilibrium response to turbidity currents of both glacial and interglacial ages (Carter et al., 1990). They are quite distinct from the channel-axis sediment waves, which have little or no sub-seabed history. They migrate upslope and into the £ow because of preferential deposition of ¢ne-grained sediment on the their up£ow side. They retain their waveform despite a signi¢cant proportion of their thickness being ‘hemipelagite’ or ‘hemiturbidite’, deposited from near-stationary, near-bottom turbid layers (Howe, 1996). The Hikurangi Channel overbank sediment waves are interpreted as the product of lee waves, which can be generated in subcritical overbank £ows, downstream (in the lee) of a linear obstruction (Flood, 1988; Howe, 1996). In this case, the linear obstruction is the channel edge or levee. Lee waves are essentially standing waves, in which the wave velocity relative to the current is directed upstream, and exactly opposed to the velocity of the current itself. Such a regime would tend to impose a periodic pattern of sedimentation and, thus, the development of sediment waves. For lee waves to exist, the £ow must be ‘subcritical’, but

Fig. 12. Distribution of overbank sediment waves, indicated by wave crests (solid lines) and strong backscatter adjacent to the meandering segment of channel. Legend as Fig. 6. Contours at 50 m intervals. Arrowed boxes show distance from Kaikoura Canyon head along channel (solid) and overbank plain (grey). Rectangle shows location of backscatter image in Fig. 10B and lines of pro¢les in Figs. 8 and 11.

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the de¢nition of this depends on the density pro¢le of the £ow. In the case of a ‘top-hat’ pro¢le (with vertically uniform density and £ow velocity), the overall Froude number Fro must be 6 1 (Fro being de¢ned as U/(gPH)0:5 , where gP is reduced gravity, H is £ow thickness, and U is velocity). However, for a strati¢ed £ow, the ‘strati¢cation’ Froude number Frs must be 6 1/Z (Frs being de¢ned as U/NH where N is buoyancy frequency (Turner, 1973)). Although in free £ow, lee waves can form oblique to the current direction (Blumsach, 1993), where there is a linear obstruction such as the channel edge or levee, the wave crests will generally lie parallel to that obstruction, even if the current direction is oblique. The lee-wave interpretation for overbank sediment waves is preferred to an antidune interpretation for two reasons. Firstly, their orientation is related to that of the channel levees, which are likely to produce lee waves in any suitable current £owing over them. Secondly, overbank sediment surfaces, even at depth in seismic pro¢les, are generally smooth and devoid of signi¢cant perturbations in bed height or roughness, which might have initiated the growth of antidunes. Although low Froude numbers (however de¢ned) favour the formation of lee waves, and high Froude numbers favour an antidune interpretation, realistic Froude numbers are di⁄cult to estimate for the Hikurangi overbank slopes on the basis of the presently available data, even for a ‘top-hat’ pro¢le. A frequently used formula gives Fro = S/CD , where S is the slope and CD the combined (top + bottom) drag coe⁄cient (Bowen et al., 1984; Wynn et al., 2000a). The average slope normal to the overbank sediment waves is of the order of 0.004^0.005, within the range of widely favoured values of the drag coe⁄cient, but there are not su⁄cient data to estimate possible values for CD . In the case of a strati¢ed £ow, no value for Frs can be obtained without some estimate of the density gradient.

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Elsewhere, some broad sediment-wave ¢elds on the continental rise have been interpreted as antidunes formed by turbidity currents. Those around the base of the volcanic Selvage Islands in the North Atlantic have been interpreted as antidunes formed by full-thickness, unchannellised (never con¢ned) turbidity currents (Wynn et al., 2000a), rather than dilute overbank £ows from the tops of channellised £ows. The Selvage sediment waves are related to somewhat irregular obstructions formed by rock outcrops and debris-£ow deposits, but signi¢cant irregularities, which might promote the formation of antidunes, are rare on the backslopes of deep-ocean channel levees. In addition, some fan-channel sediment waves, in an environment of channel avulsion, may also be antidunes (Normark et al., 1980; Piper and Savoye, 1993). However, on the gentle, levee backslopes of deepocean channels, where there is no channel avulsion, it is di⁄cult to derive the observed sediment waves from a smooth bed, except in the presence of lee waves in the £ow. Beneath lee waves, regular sediment waveforms are imposed on the seabed by the £ow regime itself. Turbidity currents form sediment waves in very di¡erent environments and £ow regimes. It is not necessary that one explanation ¢ts all situations. The orientation of some overbank sediment waves implies that turbidity currents that over£ow the channel may re-enter it lower down. At many meander loops, the sediment waves indicate that overbank £ows simply ‘cut the corner’ and re-enter the channel at the next bend. On a much larger scale, the sediment waves on the turbidite plain, particularly those in the depression along the deformation front, indicate that over£ow from a considerable length of the channel may re-enter it where the channel approaches the deformation front. The absence of a left-bank levee where the channel intersects the inferred overbank £ow may be support for this theory. However, there are no recognisable overbank tributary

Fig. 13. High-resolution (3.5 kHz) pro¢les of sediment waves in the southern Hikurangi Trough (vertical exaggeration about 38U). Positions shown in Fig. 9. (A) Sediment waves with heights up to 40 m and wavelengths of about 3 km near the top of the overbank slope, size decreasing down overbank slope. (B) Sediment waves with a height of about 20 m and an apparent wavelength of 3^4 km migrating up an overbank slope at the ¢rst tight right-hand bend. (C) Low sediment waves with a height of about 5 m and a wavelength of about 4 km migrating towards a shallow moat at the foot of the Chatham Rise slope.

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channel systems, as recognised for the Monterey Channel (McHugh and Ryan, 2000) 7.3. Long-term equilibrium system The proximal Hikurangi Channel system is in very long-term equilibrium. Although similar in surface morphology to £uviatile systems, the channel, levees and overbank slopes have not developed piecemeal with frequent meander cut-o¡ and switches of position. Instead, the channel has aggraded in more or less the same geographical position for nearly 2 million years. This complete lack of downstream translation of meander bends may be a characteristic feature of deep-ocean channels (Peakall et al., 2000) and of the longterm growth of channel overbank sediment waves. The consistently higher left-bank levee indicates strong tilting of the upper, overspilling part of each large turbidity current by the southern hemisphere Coriolis e¡ects (Carter and Carter, 1988; Klaucke et al., 1998), there being no evidence of consistent northwestward geostrophic £ow in the southern Hikurangi Trough. The overbank sediment waves show that at least some of the overspill has been subject to ‘£ow-stripping’ away from the channel, notably at right-hand bends (Piper and Normark, 1983).

8. Bottom current sediment waves Sediment waves that migrate upslope away from the Hikurangi Channel and towards a scour ‘moat’ along the toe of the Chatham Rise are inferred to be the product of bottom currents, not turbidity currents. The sediment waves are part of a drift deposit, which merges downslope into turbidite-dominated deposits. Although the slope-toe ‘moat’ indicates a signi¢cant geostrophic £ow along the base of the slope, there is no direct evidence of £ow direction. Oblique scour and drift deposition at mid-slope depths at the far western end of the Chatham Rise has been attributed to northeastward £ow, mainly during periods of glacially lowered sea-level, and these features end on the lower slope (Barnes, 1994b) near to the slopetoe ‘moat’. Seismic pro¢les across the northern

Chatham Rise several hundred kilometres east of the Hikurangi Trough also show a slope-toe channel and associated drift deposit (Wood and Davy, 1994). These features have been inferred to result from a westward-£owing branch of the Deep Western Boundary Current, the main £ow of the Deep Western Boundary Current being much deeper and con¢ned by the northeast scarp of the Hikurangi Plateau (Fig. 1) (McCave and Carter, 1997). A branch of the Deep Western Boundary Current is likely to form a more laterally and temporally consistent £ow than glacial age £ow through the gap at the western end of the Chatham Rise. If such a slow-moving boundary current reached the apex of the Hikurangi Trough, it might be diverted by deeply penetrating geostrophic eddies (Roemmich and Sutton, 1998). At this stage, the nature of the £ow along the northern toe of the Chatham Rise is unclear. These drift-deposit sediment waves may be due to lee waves, like those inferred for upslope-migrating sediment waves in other abyssal geostrophic £ows (Flood et al., 1993; Manley and Caress, 1994); lee waves typically form in £ow velocities greater than about 0.1 m s31 (Weatherly, 1993; Howe, 1996). Their orientation, which is based only on tentative correlation between pro¢les, is not clear enough to de¢ne the nature and direction of £ow, sediment waves being typically oriented oblique and anticlockwise to £ow in southern latitudes (Flood and Shor, 1988; Flood et al., 1993).

9. Conclusions Sediment waves along the axis of the proximal Hikurangi Channel are 1^4 km in wavelength but less than about 6 m high. They are surface features with no discernible long-term aggradational history. They are probably formed of sand moving over a coarse sand or gravel substrate, and probably originated as antidunes in turbidity currents that were much thicker than the 230^290 m deep channel, perhaps in the order of 500 m thick. Overbank sediment waves, which occur on the backslopes of channel levees, are 1^4 km in wavelength and up to 45 m high. They are conspicuous

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in seismic pro¢les for up to 400 m below the seabed, suggesting an equilibrium with turbidity currents for almost all of the Quaternary Period. They migrate towards levee crests at rates of 3^20 m/ka horizontally, during aggradation of 0.4^0.7 m/ka. They are inferred to form beneath lee waves, generated in the relatively dilute, suprachannel, ‘over-spill’ parts of large, strati¢ed turbidity currents in the lee of the channel edge or levee. They form nearly parallel with the channel edge and levee, and at a high angle to £ow away from the channel, which is commonly tangential to the outsides of bends. They are best developed on the outside of bends on the left bank, where centrifugal and Coriolis e¡ects combine. Sediment waves on the insides of bends on the left bank may be related to Coriolis e¡ects only. Overbank sediment waves are a very long-term equilibrium response to £ow-stripping during both glacial and interglacial ages. Sediment waves that migrate away from the Hikurangi Channel but towards a slope-toe boundary channel along the northern edge of the Chatham Rise are probably formed by lee waves associated with a west-£owing branch of the Paci¢c Deep Western Boundary Current. In the sediment-¢lled subduction trench of the Hikurangi Trough, the Hikurangi Channel, with its meanders and overbank sediment waves, has remained in almost the same geographic location for nearly 2 million years, despite aggradation of nearly half a kilometre of turbidites. The orientation and distribution of overbank sediment waves suggests that the stripped-o¡ tops of turbidity currents £ow towards, into, and then along the deepest part of the turbidite plain at the slope-toe deformation front, before re-entering the Hikurangi Channel where it approaches the deformation front.

Acknowledgements We are indebted to Phil Barnes, NIWA, for generously supplying the data and detailed draft interpretation for his southern Hikurangi margin plate boundary studies, for discussion on long term processes in the southern trough, and for

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helpful suggestion at many stages. We are also very grateful to Lionel Carter of NIWA, David Piper of Bedford Institute of Oceanography, Canada, Russell Wynn of SOC, UK, and an anonymous referee for their constructive, if somewhat drastic, editing of a lengthy manuscript that included the text of this paper. We hope that our theories on the long distance propagation of Hikurangi Channel turbidity currents may be published elsewhere. We also thank Alan Orpin of NIWA for editing a late draft of this paper, Hamish Saunders and Andy Hill for drafting the maps, and Wendy St George for drafting the other ¢gures.

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