Introduction

The fundamental building blocks of all ancient and modern continental margins are three stratal associations that result from the interplay of accommodation and sedimentation, which are downbuilding progradation, upbuilding progradation, and retrogradation (Fig. 1; e.g., Hunt and Tucker 1992; Posamentier and Morris 2000; Catuneanu 2006; Catuneanu et al. 2009; Helland-Hansen and Hampson 2009). Downbuilding progradation results from a base-level fall and the seaward migration of the shoreline independent of sediment supply; upbuilding progradation is driven by sediment supply, in which the sedimentation rate outpaces rates of base-level rise at the shoreline; the retrogradation of depositional systems results from a base-level rise, in which the rates of base-level rise outpace sedimentation rates at the shoreline (Catuneanu 2006). Today, state-of-the-art 3D seismic-reflection data allow these stratal associations to be mapped and analysed over hundreds of square kilometres in 3D detail down to a resolution of a few metres, potentially revealing spectacular views of past landscapes and depositional environments (e.g., Cartwright and Huuse 2005). 3D seismic-reflection information, ideally supported by borehole data, can thus provide key evidence for the interpretation of palaeo-depositional environments (Catuneanu 2006).

Fig. 1
figure 1

Key stratal associations that result from the interplay of accommodation and sedimentation: downbuilding progradation, upbuilding progradation, and retrogradation (see Catuneanu 2006; Catuneanu et al. 2009). a Downbuilding progradation results from a base-level fall and the seaward migration of a shoreline independent of sediment supply. b Upbuilding progradation is driven by sediment supply, in which the sedimentation rate outpaces the rates of base-level rise at the shoreline. c Retrogradation of depositional systems results from a base-level rise, in which the rates of base-level rise outpace the sedimentation rates at the shoreline. d Erosion of clinoform topsets hampering the differentiation between downbuilding and upbuilding progradation

Full size image

The previous 2D and 3D seismic-based research on the prograding Neogene sequences of the Taranaki Basin offshore New Zealand (Fig. 2) includes the studies of Bussell (1994) and Baur et al. (2011), who investigated Miocene deposits in the 3D seismic Maui survey and described internal sedimentary structures and depositional elements of submarine channels. Shumaker et al. (2017) provided a detailed analysis of submarine gullies of latest Miocene-to-Pliocene age in the 3D seismic Tui survey, linking these to large, dilute, sheet-like turbidity currents. The pioneering work of Beggs (1990) and Hansen and Kamp (2002, 2006) described the development and architecture of the Pliocene–Pleistocene Giant Foresets Formation in the northern Taranaki Basin focusing on 2D seismic interpretation and sequence stratigraphy. Salazar et al. (2016) combined 2D and 3D seismic-reflection data to study 3D clinoform morphologies in the northern Taranaki Basin, and Anell and Midtkanal (2015) quantified different types, shapes, and geometries of clinoforms in 2D in the northern and central Taranaki Basin.

Fig. 2
figure 2

Geological setting of New Zealand, stratigraphy, and location of the studied data. a Crustal thickness of present-day emergent New Zealand and submerged Zealandia. Blue: crustal thickness around 5 km; green: crustal thickness between 10 and 25 km; red: crustal thickness up to 50 km. Crustal thickness data from NZEP (2015, 2017). b Miocene to recent chronostratigraphy, lithostratigraphy, and tectonic regime of the southern and central Taranaki Basin (after Roncaglia et al. 2010; Strogen 2011; Reilly et al. 2015)

Full size image

This study uses three overlapping 3D seismic-reflection surveys (Maui, Tui, Kokako; Fig. 3) of the central Taranaki Basin, supported by well data, to exemplarily document the key architectural elements and sedimentary features of the rapidly prograding continental margin of western New Zealand. Palaeo-sedimentary systems are analysed in terms of depositional environment and, through time, with respect to the migration of facies belts. The use of specific sequence-stratigraphic terminology (for review, see, e.g., Catuneanu et al. 2009; Burgess 2016) was deliberately avoided in this study to emphasise seismic observations and palaeo-sedimentary interpretations. The sequence-stratigraphic terms regressive and transgressive are used in the following solely for eustatic variations; the sedimentary system analysis presented exclusively uses the descriptive stratigraphic terms downbuilding progradation, upbuilding progradation, and retrogradation (Fig. 1).

Fig. 3
figure 3

Location map of the Maui, Tui, and Kokako 3D seismic-reflection surveys and ten boreholes in the central Taranaki Basin. Faults and magmatic features simplified after Stagpoole and Nicol (2008), Giba et al. (2010, 2012) and Mouslopoulou et al. (2012). 2D seismic line SUNZ-91-119 was used to tie borehole information from well Witiora-1 to the 3D Kokako seismic volume. Wells Maui-1, -3, -5, -6, and -7 are located in the footwall of Cape Egmont Fault. Black line X–X’ indicates the location of 2D seismic overview section shown in Fig. 4

Full size image

Geological framework

The present-day emergent New Zealand depicts around 5% of the largely submerged palaeo-continent of Zealandia (Fig. 2a). Submergence was a result of rifting and the associated thinning of most of the continental crust to an average of 25 km (Campbell et al. 2012). With the onset of extensional tectonics in the Cretaceous, Zealandia separated from Gondwana and the Tasman Sea rift formed (Holt and Stern 1994; King and Thrasher 1996; Strogen et al. 2017); dated by the oldest basalts found in the Tasman Sea, spreading started at c. 83 Ma and lasted for a period of c. 30 Ma (Campbell et al. 2012). Zealandia rifted in an NE direction and subsided continuously due to stretching and subsequent lithospheric cooling. The Miocene saw a fundamental change in Zealandia, with the onset of intense active margin tectonics caused by the collision of the Australian and Pacific plate splitting the continent into northern and southern parts (e.g. King 2000; Cande and Stock 2004). As a result, today’s emergent New Zealand got uplifted and increased erosion rates led to vast amounts of sediment entering the surrounding basins (Sutherland 1999; Wood and Stagpoole 2007).

The Taranaki Basin (Figs. 2a, 3) is located mostly offshore of the west coast of North Island, New Zealand, extending over an approximate area of 100,000 km2 (King and Thrasher 1996). It initiated as rift basin (Middle–Late Cretaceous to Paleocene), developed into a passive continental margin (Eocene to Early Oligocene), then into a foreland basin (Miocene) and finally into a back-arc basin (Middle/Late Miocene to recent). Collision-related uplift and erosion led to the establishment of a major progradational sedimentary sequence in the Early Miocene. Due to major sediment influx from rising land areas, the continental margin rapidly prograded northwards across the Taranaki Basin (summarized in Strogen 2011). In the southern parts of the basin, deep-marine mudstones with minor amounts of siltstone (Manganui Formation) and sandy–muddy submarine fans (Moki and Mount Messenger formations and informal ‘Sw-Sands’) were initially deposited (Fig. 2b), successively overridden by a shelf sedimentary wedge (Hayward 1990; Bussel 1994; King and Thrasher 1996; King et al. 1999).

The southern part of the Taranaki Basin started to invert during the Late Miocene, ultimately involving emergence of much of the previous southern basin areas. The study area (Figs. 2a, 3) is on the margin of this emergent region and responded by the accommodation of progradational clinoforms (Fig. 4). The inversion started with a change in the sense of displacement on Late Cretaceous faults, resulting in the growth of antiforms with synforms subsequently being uplifted, as well (King and Thrasher 1996). The progradation of the Late Miocene section was driven by this inversion, with the erosion of inverted parts of the basin succession sourcing the prograding wedge material. A further significant source for the prograding wedge was erosional material of the Southern Alps, which started to be uplifted during the Late Miocene as part of the wider development of the plate boundary across New Zealand (Kamp et al. 1989; Kamp and Green 1990; Tippett and Kamp 1993; Crowhurst et al. 2002). Following the Late Miocene inversion phase, the southern and central Taranaki Basin subsided (Bull et al. 2016), and sediments onlapped the prior Late Miocene erosion surface during the Pliocene, followed by significant Plio-Pleistocene progradation. The Plio-Pleistocene Giant Foresets Formation (Fig. 2b; Pilaar and Wakefield 1978; Beggs 1990; Scott et al. 2004; Hansen and Kamp 2002) has been studied in detail by Beggs (1990), Scott et al. (2004), Hansen and Kamp (2002, 2006), Anell and Midtkanal (2015), Salazar et al. (2016), and Shumaker et al. (2017).

Fig. 4
figure 4

Seismic overview section across the 3D Maui, 3D Tui, and 3D Kokako seismic-reflection volumes. For line location, see Fig. 2c and inset box. a Un-interpreted seismic section and location of borehole Maui-6. b Interpreted section with key marker horizons 1–8 (old to young) subdividing the studied clinoform succession into 9 seismic units (+ acoustic basement). Dashed lines indicate areas in which downbuilding progradational, upbuilding progradational, and retrogradational units are separated by zones of reflection termination instead of a single sharp, precisely defined bounding reflection. c Interpreted section with seismic units 1–9 bound by key marker horizons (colour coding corresponds to Fig. 4). Also shown are intra-unit horizons (dashed black lines) that correspond to spectral-decomposition horizon slices and well data on Figs. 6, 7, 8, 9, 10, and 11

Full size image

Given the situation that the southeastern edge of the study area is the Cape Egmont Fault (Fig. 3), this fault needs to be considered in detail having important implications for accommodation development and sedimentation immediately to the north. The Cape Egmont Fault has ~ 2.5 km displacement, separating in the Maui field reservoir and potential source rocks (Reilly et al. 2016). The fault has been reactivated multiple times with initial normal faulting in the Late Cretaceous–Paleocene, Eocene-to-Miocene (c. 11 Ma) tectonic quiescence, reverse displacement between 11 and 5.5 Ma, an inactive interval between 5.5 and 3.7 Ma, and a final reactivation as a normal fault from 3.7 Ma until recent (King and Thrasher 1996; Nicol et al. 2005; Reilly et al. 2015, 2016). The series of volcanoes around the northern termination of the Cape Egmont Fault Zone (Fig. 3) is dated younger than 2 Ma (Giba et al. 2013).

Data and methods

The overlapping 3D seismic-reflection surveys Maui-3D, Tui-3D, and Kokako-3D are located about 40 km west of the Taranaki Peninsula (Fig. 3). The Maui-3D survey covers an approximate area of 1000 km2. It was acquired in 1991 using a dual-vessel system, four sources, and four streamers (Lutz et al. 1993). The final 3D seismic-reflection volume has a 25-by-25 m bin size. The data set is in zero-phase European polarity, in which a downward acoustic impedance increase is negative, displayed on all figures of this study as red reflection. The Tui-3D survey (extent c. 350 km2) was acquired in 2003 using a dual-source/two-streamer configuration (Veritas 2003). The final 3D migrated data set has a bin size of 12.5 by 12.5 m and is phase-rotated by 120°. For comparability with the neighbouring Maui-3D and Kokako-3D surveys, a phase rotation to zero-phase European polarity was applied to the Tui-3D volume. The Kokako-3D survey (extent: 593 km2) was shot in 2013 using a system of two sources and ten streamers (DownUnder 2013). The final 3D migration bin size is 25 by 12.5 m, and the survey is in zero-phase European polarity. For ground truthing of the seismic interpretation results, 11 wells were integrated in this study (Fig. 3) enabling seismic-well ties based on checkshots (NZPE 2015, 2017), log-based lithology interpretations, and borehole-lithology information from cuttings (e.g. STOS 1970; Stroud et al. 2003, 2004; NZOP 2004; AWE 2006).

The 3D seismic interpretation methodology of this study differs from a classical interpretation approach: the first interpretation step comprised a fully automated 3D interpretation of all positive (peaks) and negative reflections (troughs) in all the seismic volumes in PaleoScan. In a subsequent workstep, a 3D geomodel was calculated by horizon interpolation, assigning a relative geological position to each horizon pixel within the model volume. The 3D geomodel was then interpolated assigning a relative geological position to each datapoint within the model volume. Series of horizons were then extracted from the 3D geomodel, matched against the original seismic data, and—for marker horizons defined by reflection terminations—snapped to a precise position constrained by distinct amplitude signature and reflection configuration. Seismic-well ties, biostratigraphic analyses of wells Maui-1, Maui-2 and Witiora-1 (Hornibrook and Hoskins 1969; Schröder 1970; Morgans 2006), and chronostratigraphic interpretation provided rounded ages that reflect the inherent biostratigraphic age uncertainties (Table 1). All ages presented are internally consistent, although associated with some error. A final comparison of the seismic horizons and ages of this study with regional seismic interpretation in the Taranaki Basin (summarized in Bull et al. 2016; Table 1) shows a reasonable correlation between the different interpretation approaches.

Table 1 Eight seismic marker horizons interpreted in this study in the central Taranaki Basin
Full size table

3D seismic-sedimentological interpretations were carried out by frequency-decomposition colour blending along horizon slices (Partyka et al. 1999; Henderson et al. 2007, 2008; McArdle and Ackers 2012) using a combination of Ricker-wavelet central frequencies of 10 Hz (red), 20 Hz (green), and 30 Hz (blue), which yielded optimum results in terms of contrast and sharpness of seismic-sedimentological features. A strength of frequency decomposition is the identification of anomalous seismic features that are normally covered by the broadband seismic signal (Li et al. 2015), which helps to produce a clearer image of the feature of interest (e.g., Fahmy et al. 2005). Particularly, in combination with borehole-based subsurface lithological interpretations, 3D seismic spectral-decomposition horizon-slice analysis, in places complemented by reflectivity, RMS amplitude, and variance attribute analysis, provided robust predictions of subsurface depositional environments.

Seismic overview

The Miocene-to-recent depositional system of the central Taranaki Basin is dominated by kilometre-scale progradational clinoforms (Fig. 4). Eight seismic marker horizons (1–8 from old to young; Table 1) were defined based on reflection terminations, subdividing the succession into nine reflection units that formed between > 19 Ma and today (Fig. 2b).

Unit 1 (> c. 19 Ma) is a reflection package above the acoustic basement with a thickness of around 300 ms two-way time (TWT) in the southeast and 900 ms (TWT) in the northwest. Reflections within this unit are generally of low-to-medium amplitudes and high continuity (Fig. 4). Horizon 2 at the top of unit 1 is close to the regional marker N05 of Bull et al. (2016). Unit 2 (c. 19–11 Ma) forms a basinward thinning wedge with a maximum thickness of c. 800 ms (TWT) in the southeast. Internal reflections are of medium-to-high continuity (with local truncations) and low-to-medium amplitude. Horizon 3 locally truncates the top of unit 2 in the southeast, and forms a downlap surface for succeeding unit 3. Horizon 3 is just above the regional marker N40 (intra-late Waiauan; c. 11 Ma) of Bull et al. (2016). Unit 3 (c. 11–7.5 Ma) is characterized by pronounced clinoform foresets that are truncated at the top by marker horizon 4 (Fig. 4). The foresets of the unit 3 clinoforms have a concave, oblique shape with a highly variable reflection configuration from partly chaotic in the southeast to almost parallel in the northwest. The reflections within this unit are mostly of low, locally of medium amplitude. In regional comparison, horizon 4 is between markers N40 and N50 (Tongaporutuan) of Bull et al. (2016).

Unit 4 (c. 7.5-6 Ma) is characterized by concave-to-convex, sigmoidal, downbuilding clinoform reflections. In the southeast, reflections show repeated onlaps and local truncations (Fig. 4). Horizon 5 at the top of unit 1 is close to the regional marker N50 (c. 7.2 Ma) of Bull et al. (2016). Unit 5 (c. 6–5.5 Ma) has a lens-shaped form and onlaps onto unit 4. The internal reflections are of a slightly aggradational, convex clinoform shape, and of low-to-medium strength, with a few high amplitudes around local incisions (Fig. 4). Units 4 and 5 are not recorded in the southeast of the study area (Fig. 4). Unit 6 (c. 5.5–4.5 Ma) has an overall sheet form with a rather constant thickness of ca. 200 ms (TWT) in the southeast, pinching out in northwestern direction on the topsets of unit 5. The internal reflection configuration shows multiple stacked packages of shingled reflections in the southeast, and more subparallel reflections locally cut by incisions in the northwest. Due to its basinward termination, unit 6 lacks in the northwest of the study area (Fig. 4). In regional comparison, horizons 6 and 7 are between markers N50 and N68 of Bull et al. (2016).

Unit 7 (c. 4.5–4 Ma) extends across the entire study area with a relatively constant thickness of 100–150 ms (TWT). It is characterized by slightly aggradational, convex clinoforms of low-to-medium amplitude with low (southeast)-to-medium (northwest) continuity (Fig. 4). In places, unit 7 shows internal incisions and distorted reflections. Horizon 8 at the top of unit 7 is close to the regional marker N68 (intra Opotian; c. 4 Ma) of Bull et al. (2016). Unit 8 (c. 4–2 Ma) consists of progradational, downbuilding, complex sigmoid-oblique clinoforms that onlap onto unit 7. The base of unit 8 shows deformed reflections of low-to-medium amplitude and high continuity; the upper part is characterized by steeply inclined reflections that downbuild and repeatedly onlap onto their respective predecessor. Unit 8 is not recorded in the central and southeastern parts of the study area (Fig. 4). The top of unit 8 is close to the regional marker N82 (Mangapanian-Nukumaruan bounday; c. 2 Ma) of Bull et al. (2016). Unit 9 (< 2 Ma) finally onlaps onto units 7 and 8 and shows an overall thickening towards the northwest. The reflection configuration is parallel to subparallel with medium amplitudes and medium-to-high reflection continuity.

Stratigraphic architecture and internal 3D seismic unit characteristics

The pronounced downbuilding of the complex sigmoid-oblique clinoforms of units 4 and 8 is associated with numerous repeated onlaps below the breakpoint of preceding clinoform reflections, which indicates that progradation in these intervals was driven by base-level fall (Fig. 1a). In contrast, the more convex, sigmoidal clinoforms of units 5 and 7 show topset aggradation (upbuilding), indicating that sedimentation rates outpaced the rates of base-level rise (Fig. 1b). The basinward pinch out and landward thickening, and backstepping geometry of unit 6 document retrogradation of the sedimentary system, in which the rates of base-level rise must have outpaced sedimentation rates (Fig. 1c). The topsets of the prograding clinoforms of unit 3 are lacking either due to erosion or non-deposition; it, therefore, cannot be unambiguously established whether this prograding unit was originally downbuilding or upbuilding (Fig. 1d). Due to the data set limits in the southeast and northwest, the genetic stratigraphic architecture of units 1, 2, and 9 cannot be established. In the following, the description and interpretation of 3D seismic-sedimentological and geomorphological features within all seismic units (Fig. 4) is presented on spectral-decomposition horizon slices at selected horizon levels (Fig. 5) using a red–green–blue (RGB) blend of frequencies, in places ground-truthed by well-log data. As Unit 1 contains all the pre-Early Miocene strata and Unit 9 is located distal of the shelf break without sedimentary features of interest, they are not considered for this study.

Fig. 5
figure 5

Depth maps (in ms TWT) of seismic horizons between seismic units 2 and 7 selected for detailed frequency-decomposition display (Figs. 6, 7, 8, 10). Note major northward progradation of sedimentary system and general upward-shallowing trend

Full size image

Unit 2 (19–11 Ma)

Horizon 2.1 (Figs. 5a, 6a) of unit 2 (for location, see Fig. 4) shows a large channel system trending in northwestern direction across the Maui and Tui areas, changing its direction to the north in the Kokako area. The horizon is located between 1.6 s (TWT) in the southeast and 2.25 s (TWT) in the northwest. The approximately 70 km long-channel system starts to be tracable in the southeast at the Cape Egmont Fault Zone. In the proximal Maui area, four different feeder channels with variable sinuosity, width, and incision depth can be recognized each trending to the northwest (Fig. 6a). A highly sinuous channel system (width c. 400 m; incision depth c. 100 ms (TWT); sinuosity c. 2.3) with extensive meander loops and lateral migration patterns characterizes the east of the study area. Less sinuous, in their upper reaches partly straight channels (width c. 400–800 m; incision depth c. 70–100 ms (TWT); sinuosity c. 1.1–1.5) which characterize the central and western Maui area. Wells Maui-3 and Maui-5 (Fig. 6b) document a generally shaly background lithology around horizon 2.1, encompassing up to 20 m thick, sandstone-rich intervals at channel intersections.

Fig. 6
figure 6

RGB blend of spectral frequencies on horizons 2.1 and 3 (for horizon position, see Fig. 4c), well data (for well location, see Fig. 2c) and vertical seismic data. a Horizon-slice 2.1 of central unit 2 showing large, sinuous, NW-oriented channels that merge in the 3D Tui and Kokako areas into a big, north-oriented master channel flanked by extensive overbank deposits. b GR measurements at wells Maui-3 and Maui-5 document a generally shaly background lithology around horizon 2.1, encompassing up to 20 m thick, sandstone-rich intervals at channel intersections. c Vertical seismic section x–x’ documenting complex cut-and-fill geometry of proximal channels. d GR measurements at wells Tieke-1 (T1), Tui-1, Amokura-1 (A-1) and Pateke-2 (P-2) document several sandstone-rich intervals at channel intersections in the interval below horizon 2.1, and shalier lithologies towards horizon 3 and further up. e Horizon-slice 3 of the base of unit 3 showing NNW-oriented channel systems flanked by extensive overbank deposits and sediment-wave fields. e Vertical seismic section y–y’ across westernmost channel and downslope overbank with sediment waves. CH channel, SWF sediment-wave field

Full size image

Downslope in the northern Tui area, the channels merge into a large, north-oriented master channel flanked by extensive overbank deposits. Numerous channel bends that are frequently cut by adjacent and succeeding channels form a complex cut-and-fill geometry with variable seismic-reflection strength and orientation (Fig. 6c). The subsurface lithology encountered by the distal wells Tieke-1, Tui-1, Amokura-1, and Pateke-2 is generally shaly, containing between horizons 2 and 2.1 at channel intersections several 10–20 m thick sandstone units (Fig. 6d).

Unit 3 (c. 11–7.5 Ma)

Horizon 3 (Figs. 5b, 6e) defines the base of unit 3 (Fig. 3) in depths around 1.2 s (TWT) in the southeast and 2.1 s (TWT) in the northwest. In the Maui area, three separate channel systems occur, all trending in northern direction (Fig. 6e). The easternmost channel has a high sinuosity of 1.90 and probably continues into the Kokako area with a slight directional change to the northwest. The central and western channels merge in the Tui area, continuing into the Kokako area in a northwestern direction. Sinuosity of these systems is c. 1.2 (central) and 1.6 (western) for the proximal channels, and c. 1.1 for the joined system. Extensive sediment-wave fields occur at meander loops and channel bends. These sediment-wave fields (Fig. 6e, f) show a distinct pattern of interchanging high and low frequencies, extending in places over a length of 12 km and a width of 6 km. The geomorphological wavelength of the sediment waves ranges between 800 and 1100 m, with an average wave height of 40 ms (TWT). Wells in the Maui area (Maui-3, Maui-5 and Maui-6) show relatively uniform GR values between 80 and 94 API around horizon 3, increasing downslope to values of 100–120 API in wells in the Tui area (Fig. 6d) with well cuttings documenting claystone–dominated lithologies (Stroud et al. 2003, 2004; NZOP 2004; AWE 2006).

Horizon 3.1 (Figs. 5c, 7a) is located in the top part of unit 3 (Fig. 4) at depths between 1.0 s (TWT) (south) and at 1.9 s (TWT) (north). In its most proximal part, it corresponds to an erosional surface (Figs. 4, 7a). Topsets in the uppermost part of unit 3 close to horizon 3.1 are absent. Its steep, oblique foresets are characterized by a series of slope gullies (straight, regularly spaced channels) that incise between 130 and 170 ms (TWT) in northwestern direction into the reflections below (Fig. 7b). The gullies can be traced over approximately 23 km into the Tui area and are accompanied on either side by sediment-wave fields that extend into the distal northwestern Kokako area. The geomorphological wavelength of the sediment waves is between 800 and 1100 m in the Tui area, and between 1200 and 1800 m in the Kokako area. The sediment-wave fields in the Kokako area are cut-off in the east by a mass-transport complex (MTC).

Fig. 7
figure 7

RGB blend of spectral frequencies on horizons 3.1 and 4.1 (for horizon position, see Fig. 4c) and vertical seismic data. a Horizon-slice 3.1 in the top part of progradational unit 3 showing series of slope gullies (straight, regularly spaced channels) accompanied on their sides by sediment-wave fields. In the Kokako area, the sediment-wave fields are cut-off by a MTC. b Vertical seismic section x–x’ perpendicular to gullies. c Horizon-slice 4.1 of the lower part of downbuilding progradational unit 4 showing several small, downdip oriented, branching feeder channels and incisions merging into gullies further downslope. Inhomogeneous, laterally varying frequency spectrum in the Tui and Kokako areas interpreted as a mass-transport complex. d Zoom-in on topset channels. e Vertical seismic section y–y’ immediately beyond the clinoform breakpoint showing proximal feeder channels. Note difference in size between feeder channels on horizon 4.1 and gullies of horizon 3.1. G gully, SWF sediment-wave field, MTC mass-transport complex

Full size image

Unit 4 (c. 7.5–6 Ma)

Horizon 4.1 (Fig. 5d, 7c–e) is located in the lower part of downbuilding unit 4 (Fig. 4) in depths between 1.0 s (TWT) in the southeast and 2.1 s (TWT) in the northwest. Most of the Maui area corresponds at this horizon level to an erosional surface. Yet, Horizon 4.1 shows some preservation of topset beds landward of the clinoform breakpoint, revealing several small, downdip oriented, branching feeder channels, and incisions merging into gullies further downslope (Fig. 7d). Individual branches of the feeder channels are 40–80 m wide, up to 40 ms (TWT) deep (Fig. 7e), and can be traced for a length of 400–2500 m. On the sigmoidal foresets, linear-to-slightly sinuous gullies (Fig. 7c) occur with a spacing of 400 to 800 m, are up to 240 m wide, and incise about 50 ms (TWT) into the substratum. Several of these gullies merge in the eastern Tui area and can be traced for up to 20 km downslope. In the western Tui and Kokako areas, an inhomogeneous, laterally strongly varying frequency spectrum with a complex rugose geomorphology is interpreted as a mass-transport complex (Fig. 7c). Gamma-ray logs of wells Tui-1, Amokura-1, and Pateke-2 around horizon 4.1 show generally high values (100–130 API) suggesting a shale-dominated lithology; well cuttings are claystone dominated with minor siltstone and sandstone contents (Stroud et al. 2003, 2004; NZOP 2004).

Unit 5 (c. 6–5.5 Ma)

Horizon 5.1 (Figs. 5e, 8a–c) is located in the upper part of upbuilding unit 5 (Fig. 4), onlapping in the proximal Tui area onto unit 4. In the Maui area, unit 5 is lacking. Contemporaneous unconformity development due to uplift and erosion can be, e.g., interpreted in well Maui-3 by a downward decrease of average sonic values (DTC; compressional slowness) of around 20 µs/ft, and in wells Maui-5 and Maui-7 by a decrease of 10–15 µs/ft (Fig. 9).

Fig. 8
figure 8

RGB blend of spectral frequencies on horizons 5.1 and 6.1 (for horizon position, see Fig. 4c) and vertical seismic data. a Horizon-slice 5.1 of the upper part of upbuilding progradational unit 5 showing medium-to-high frequencies with smooth transitions between different slope-parallel frequency bands in the Tui area, and gullies flanked by sediment-wave fields in the Kokako area. b Branching networks of small incisions characterize the distal topsets, buried predecessors of these systems further landward indicated by chains of small pockmarks. c Vertical seismic section x–x’ across gullies. d Horizon-slice 6.1 of the lower part of retrogradational unit 6 showing WSW-ENE-oriented, elongate zones of variable frequency. e Incised valleys in the Maui area containing numerous upward and sideward backstepping, anastomosing channels (fluvial?). f Vertical seismic section y–y’ across distal gullies of horizon 6.1. CH channel, G gully, SWF sediment-wave field

Full size image
Fig. 9
figure 9

Gamma-ray (GR) and sonic (DTC; compressional slowness) logs of wells Maui-3, Maui-5, and Maui-7 around the unconformable contact between progradational seismic unit 3 (c. 11–7.5 Ma; Late Waiauan to Late Tongaporutuan) and retrogradational seismic unit 6 (c. 5.5–4.5 Ma; Late Kapitean to Middle Opoitian). Note change in sonic trend at unconformity. King and Thrasher (1996) calculated up to 300 m of missing sediment around well Maui-3. Horizon 6 is located in zone characterized by change from retrogradation (unit 6) to upbuilding progradation (unit 7; see Fig. 4)

Full size image

In the Tui and Kokako areas, horizon 5.1 occurs in depths between of 1.1 s (TWT) and 1.8 s (TWT). Seismic frequencies in the Tui area are medium to high with smooth transitions between different slope-parallel frequency bands (Fig. 8a) interpreted as shelf- to lower shoreface succession. Cuttings of wells Tieke-1 and Tui-1 document a siltstone-claystone succession (Stroud et al. 2003; NZOP 2004), the more distal well Amokura-1 encountered dominantly claystone. In the proximal Kokako area, branching networks of small incisions characterize the distal topsets (Fig. 8b), and buried predecessors of these systems further landward are indicated by chains of small pockmarks. At the clinoform breakpoint, the incisions merge into gullies up to 700 m wide, 100 ms (TWT) deep (Fig. 8c), and > 20 km long. The distal Kokako area is characterized by merging of the upper foreset gullies into two large gully systems that are flanked by sediment-wave fields.

Unit 6 (c. 5.5–4.5 Ma)

Horizon 6.1 (Figs. 5f, 8d) is located a few ms (TWT) above the base of the retrograding unit 6 (Fig. 4) at a depth of 0.6 s (TWT) near the Cape Egmont Fault Zone, and at 1.2 s (TWT) in the eastern Kokako area. Seismic frequencies in the Maui area show several 2–3 km wide, WSW-ENE-oriented elongate zones of variable frequency that are > 30 km in lateral extent (Fig. 8d). In section view, several stacked packages of small-scale, shingled clinoforms form these zones (Fig. 4). Spectral-decomposition imaging additionally shows incised valleys in the Maui area (Fig. 8e), and channels and gullies in the Tui area (Fig. 8f). The two northward trending valleys in the western Maui area (Fig. 8e) are each up to 3 km wide, > 10 km long, and less than 40 ms (TWT) incised, and contain numerous upward and laterally backstepping, anastomosing channels. Wells Maui-5 and -7 are located very close to the upper reaches of these valleys and record an upward-fining gamma-ray pattern around horizon 6.1 (Fig. 9). The gullies of the Tui area (up to 135 m wide, c. 30 ms (TWT) deep; Fig. 8f) are in their lower part straight, and in their upper part fed by networks of small, slope-parallel, and slope-perpendicular, moderately sinuous tributaries originating in the WSW-ENE-oriented patches of variable frequency in the overlapping area of the Maui, Tui, and Kokako surveys.

Unit 7 (c. 4.5–4 Ma)

Horizon 7.1 (Figs. 5g, 9, 10) is located in the lower part of upbuilding progradational unit 7 (Fig. 4) at a depth of 0.5 s (TWT) near the Cape Egmont Fault, and at 1.6 s (TWT) in the distal Kokako area. The spectral-composition horizon-slice (Fig. 10a) shows several arcuate, SW–NE to E–W trending, subparallel zones of variable frequency in the Maui and Tui areas, corresponding in vertical display to shingled clinoform units (Fig. 10b). In the northern Maui area, small channels (< 1 km in length) incise into these zones. Basinward in the Tui area, several northwest trending, straight to slightly sinuous gullies (c. 140 m wide, up to 50 ms (TWT) deep) occur. These gullies branch in their upper reaches into the networks of small tributaries (Fig. 10c) that originate at the basinward edge of the arcuate variable-frequency zones. The northern part of the study area contains major NW-trending gullies (around 180 ms (TWT) deep) of > 30 km length, a northward trending sinuous channel system (> 10 km length), and localized frequency distortions interpreted as mass-transport deposits on the middle slope.

Fig. 10
figure 10

RGB blend of spectral frequencies on horizons 7.1 and 7.2 (for horizon position, see Fig. 4c) and vertical seismic data. a Horizon-slice 7.1 of the lower part of upbuilding progradational unit 7 showing arcuate, subparallel zones of variable frequency landward of the offlap break interpreted as shelf to shoreface sediment (coarsening upward GR pattern in wells Maui-5 and Maui-7; fining upward in Maui-3). b Variable-frequency pattern corresponding in vertical display (section x–x’) to shingled clinoforms. c Gullies of the Tui area branching in their upper reaches into networks of small tributaries originating at the basinward margin of the arcuate, variable-frequency zones. Note the occurrence of mass-transport deposits in the Kokako area. d Horizon-slice 7.2 near the top of unit 7 showing inhomogeneous, partly arcuate patches of variable frequency interpreted as nearshore to coastal plain (fining upward GR pattern in well Maui-7). e Vertical seismic section y–y’ showing the location of horizon 7.2 at the top of shingled clinoform succession. Note difference between horizons 7.1 and 7.2 on both, vertical seismic data and horizon slices. CH channel, G = gully, SWF sediment-wave field, MTC mass-transport complex

Full size image

Horizon 7.2 (Figs. 5h, 10d) is located near the top of unit 7 (Fig. 4) at a depth of 0.4 s (TWT) along the Cape Egmont Fault, and at 1.5 s (TWT) in the northwestern Kokako area. Seismic frequencies in the proximal Maui area are inhomogeneous, organized in partly arcuate patches containing the complete frequency spectrum. In the distal Maui area, elongate SW–NE-trending zones of variable frequency correspond to shingled clinoforms (Fig. 10e). The Tui area, in turn, shows a distinct change of frequencies in NNW direction, with a medium-frequency spectrum in the south and very low frequencies in the north (Fig. 10d) corresponding to a clinoform topset–foreset transition. The southwest exhibits a channel system 500–800 m wide, up to 200 ms (TWT) deep and c. 15 km long. The Kokako area in the northwest shows several slope gullies, localized mass-transport complexes, and small sediment-wave fields.

Unit 8 (c. 4–2 Ma)

Horizon 8 (Fig. 11) at the base of unit 8 can be only observed in the Kokako area (0.7–1.5 s TWT). It marks the onset of a second set of downbuilding, with complex sigmoid-oblique clinoforms prograding to the northwest (Figs. 4, 11a). The topset areas in the southeast are characterized by homogeneous frequencies in the lower spectrum (Fig. 11b). The upper foresets of horizon 8 show a broad, chaotic frequency pattern, 5–12 km wide in the SE–NW direction. Reflections in this zone have a poor continuity, which show large amplitude differences and major frequency and RMS amplitude contrasts (Fig. 11b). Small, sinuous channels branching near the clinoform breakpoint enter the discontinuity zone and merge downstream into larger gully systems.

Fig. 11
figure 11

Vertical reflectivity section, RGB blend of spectral frequencies on horizons 8–8.5, and root-mean-square (RMS) amplitude extractions (close ups) along horizons. a Vertical reflectivity section x–x’ across Pliocene–Pleistocene downbuilding clinoforms. Black arrows mark onlap reflection terminations. b Horizon-slice 8 showing topset areas with homogeneous frequency distribution, and foresets with a chaotic frequency signature interpreted as MTC. c Horizon-slice 8.1 showing lobate feature on topsets, grading distally into elongate patches of variable frequency. d Horizon-slice 8.2 showing several gully systems fed in the topset areas by a network of partly sinuous, branching, shoreline-parallel channels. e Horizon-slice 8.3 with shoreline-parallel and oblique channels landward of the clinoform breakpoint, merging into a large, straight foreset gully offshore. f Horizon-slice 8.4 documenting abandonment of horizon 8.3 gully (Fig. 11e) and development of new, large foreset gully in the further north. Branching topset channels extend laterally for up to 11 km, and are joined from on their sides by numerous smaller tributaries. g Horizon-slice 8.5 showing partial coverage of the topsets of the preceding horizons 8.1–8.4, coinciding with the development of a partly slope-parallel, partly slope-perpendicular, dendritic channel system on the clinoform foresets. CH channel, G gully, MTC mass-transport complex

Full size image

Horizon 8.1 is up to 230 ms (TWT) above horizon 8, with its clinoform breakpoint shifted c. 10 km basinward (Fig. 11a, c). The slope angles of horizon 8.1 are steeper than those of horizon 8, and only minor frequency variations occur on the upper foresets. Linear gullies [200 m wide, 40 ms (TWT) deep] incise the foresets and continue downslope over c. 12 km. In the southeast of horizon 8.1, a distinct lobate feature of medium seismic frequency occurs on the topsets. This feature is incised at its front by numerous small channels (< 1.8 km long), and grades distally into elongate patches of variable frequency (Fig. 11c). Horizon 8.2 (Fig. 11a, d) onlaps onto horizon 8.1. The foresets of horizon 8.2 exhibit several gully systems [main system in the north: 600 m wide, 100 ms (TWT) deep], which are fed in the topset areas by pairs of shoreline-parallel, partly sinuous, branching channels that extend laterally over up to 10 km (Fig. 11d).

Horizon 8.3 (Fig. 11e) is up to 120 ms (TWT) above horizon 8.2, with its clinoform breakpoint another 2.5 km to the northwest. Dendritic networks of shoreline-parallel and a few oblique channels occur landward of the clinoform breakpoint, merging into a large, straight foreset gully in the centre of the Kokako-3D survey. The clinoform breakpoint of succeeding horizon 8.4 (Fig. 11f) is another 4 km to the northwest. This horizon shows one large foreset gully in the north fed by two shoreline-parallel topset channels joining from either side the gully at the clinoform breakpoint. The topset channels extend laterally for > 10 km, and are joined on their sides by numerous smaller tributaries (Fig. 11f, RMS amplitude slice). The branched network is in total up to 3 km wide and located approximately 500 m landward of the clinoform breakpoint. Finally, horizon 8.5 (Fig. 11g) shows sedimentary upbuilding and partial coverage of the topsets of the preceding horizons 8.1–8.4 (Fig. 11a), coinciding with another basinward shift of the clinoform breakpoint by 6 km. The proximal topset beds of horizon 8.5 show elongate patches of variable frequencies (Fig. 11g). In the northeast, a dendritic topset channel trends in northern direction, in the northwest, a partly shoreline-parallel, partly shoreline-perpendicular, dendritic channel system commences landward of the clinoform breakpoint.

Geological interpretation

The clinoform system under study prograded more than 60 km across the central Taranaki Basin between c. 19 and 2.4 Ma in northwesterly direction (Fig. 4). The clinoform breakpoint trajectories (Fig. 12a), often corresponding to sites of channel or gully initiation, record a major progradation of units 3–5, the retrogradation of unit 6, and a renewed progradation of units 7 and 8. The upper clinoform foresets, topsets, and inferred palaeo-shorelines of units 1 and 2 are several tens of kilometres south of the study area (Strogen 2011); the lower foresets of these units are characterized by large submarine channels (e.g., Fig. 6a).

Fig. 12
figure 12

Clinoform breakpoint migration and depositional elements of units 3 to 8. a Clinoform breakpoint trajectories and approximate stratigraphic ages. Coloured lines indicate maximum progradation or retrogradation in each unit. Purple dashed lines in the Maui area show truncated clinoform foresets and brown dashed lines in the Kokako area indicate offlap breaks of individual clinoform packages. b–g Key depositional elements of each seismic unit. Note differences of depositional systems in units characterized by the same genetic architecture (downbuilding progradation, upbuilding progradation, and retrogradation)

Full size image

Around c. 11 Ma, the clinoforms of unit 3 approached the study area, with lower foresets marked by large sinuous deepwater channels flanked by extensive sediment-wave fields (Fig. 6e). This pattern changed with the slope proper reaching the study area to increasing foreset angles to a dominance of gullies and mass-transport complexes (Figs. 7a, 12b). It can be speculated that the transition from sinuous channels to gullies and mass-wasting corresponded to a change from upbuilding to downbuilding progradation. This assumption, however, cannot be proven due to erosion of the complete topset record of unit 3 (Fig. 4), with the amount of missing strata previously estimated as c. 300 m (near well Maui-3) based on compaction-trend analysis from sonic logs (King and Thrasher 1996; their Fig. 3.6; also see Fig. 9).

In contrast, the late Miocene (< c. 9 Ma) downbuilding clinoforms of unit 4 show some limited topset preservation (Figs. 4, 7d). Downslope-oriented feeder channels established landward of the clinoform breakpoint, supplying relatively small upper foreset gullies with topset sediment (Fig. 12c). Shumaker et al. (2017) interpreted slope gullies in this unit as primarily resulting from large, dilute, sheet-like turbidity currents stemming from autogenic slope instabilities. The observation of topset feeder channels of these gullies likely relates to backstepping erosion of the initial slope channels through time. At the base of unit 4, the lower clinoform foresets contain extensive mass-transport deposits.

The topsets of upbuilding progradational unit 5 are preserved over relatively large areas (> 10 km between the offlap break and landward horizon termination) that are interpreted to comprise shelf to shoreface deposits (Figs. 8a, 12d). Landward of the clinoform breakpoint, the topsets are incised by a multitude of small, downslope-oriented channels connected to large gullies beyond the offlap break (Fig. 8b). In comparison to unit 4, upstream incision of the feeder channels reaches further inland, implying that larger sediment quantities (and maybe more constant supply) led to the development of larger slope systems.

The c. 5 Ma retrogradation of unit 6 into the proximal Maui area covered the major erosional peneplain at the top of units 3 and 4, enlarging the shelf area to > 30 km in width. Relatively wide and shallow valleys of possible fluvial origin formed in the southwest (Fig. 12e), and few, small submarine channels and gullies formed on the slope (Figs. 8d, 12e). The transition between retrograding unit 6 and upbuilding prograding unit 7 corresponds on vertical seismic data to a rather diffuse zone of multiple reflection terminations (Fig. 4), interpreted to record the development of a large shelf (Figs. 8d, 10a) changing within unit 7 to an extensive coastal plain (Figs. 10d, 12f). At the clinoform breakpoint, small gullies are superseded through time by larger channels, and the lower clinoform foresets contain mass-transport deposits (Fig. 12f).

Around 4 Ma, downbuilding progradational unit 8 shows foresets that record with increasing slope angles a change from mass-wasting to gully development (Fig. 11). Series of extensive, shoreline-parallel valleys on narrow topsets are interpreted to have developed in backshore environments sheltered by beach ridges (Figs. 11c–f, 12f). Such seismic-sedimentological, shoreline-parallel topset features have—to our knowledge—not been previously documented on downbuilding clinoforms.

Discussion

A key observation of this study is that the Miocene-to-Pleistocene clinoforms of the central Taranaki Basin show a highly variable topset preservation: the topsets of Miocene units 3 and 4 either lack (unit 3) or are only partially preserved (unit 4), which contrasts the extensive topset records of Pliocene-to-Pleistocene units 5, 7 and 8. Upbuilding progradational clinoforms (e.g., units 5 and 7) can be assumed to show a relatively high potential topset preservation, since forming under base-level rise conditions providing space for topset accumulation. In contrast, downbuilding clinoforms (e.g., units 4 and 8) should be rather prone for the cannibalization of former topsets (Fig. 1). Yet, there are marked differences in the appearance of the two downbuilding clinoform units 4 and 8, with a significant downcutting and little topset preservation in unit 4 (Figs. 4, 7c), and less downcutting associated with high topset preservation in unit 8 (Figs. 4, 11a–g). Since downbuilding progradation exclusively results from a base-level fall independent of sediment supply (Fig. 1), the contrasting development of clinoform units 4 and 8 must have been controlled either by different tectonic (subsidence/uplift) or different eustatic conditions. The location of the southern study area on the footwall of the Cape Egmont Fault (Fig. 3), which was active in a reverse sense during unit 4 deposition but during unit 8 in a normal sense, and the location of the northern study area (Kokako survey) on the hanging wall of a young (< 4 Ma), minor (< 200 m offset), normal fault (Fig. 3), suggest an important tectonic contribution to accommodation and clinoform development.

Figure 4 shows that the amount of topset downbuilding for the Miocene units 3 and 4 is in total > 300 ms (TWT; >300 m, assuming an average velocity of 2000 m/s for the overburden), measured between the erosive top of unit 3 and the most basinward clinoform breakpoint in unit 4 (Fig. 4). In contrast, downbuilding within the Pliocene–Pleistocene unit 8 is in total around 120 ms (TWT; c. 120 m), measured between horizons 8 and 8.4 (Fig. 11a). For both units, the base-level fall shifting shorelines and clinoforms downward and basinward could be induced by either (1) a tectonic uplift in the proximal parts of the study area, (2) a fall in global sea level (eustatic fall), or (3) a combination of tectonics and eustasy. Figure 13 shows that global sea level was rather stable in the Late Miocene during the deposition of units 3 and 4, even if taking rather large age uncertainties for the studied units into account. The base-level fall of over 300 m recorded by the basinward shifted onlaps between units 3 and 4, therefore, cannot be explained by a eustatic change. Yet, Reilly et al. (2015, 2016) have documented contractional tectonics at the Cape Egmont Fault (Figs. 3, 6, 8, 10) and neighbouring structures with significant reverse faulting and folding in the Late Miocene (c. 11–5.5 Ma). Inversion-driven uplift related to this contractional phase in the south of the study area seems, therefore, a very likely trigger for the > 300 m Late Miocene base-level fall forcing the basinward migration and stratal downbuilding of the unit 3 and 4 clinoforms.

In contrast, downbuilding unit 8 developed between c. 4 and 2 Ma in times when the Cape Egmont Fault (~ 30 km south of the unit) was acting as a normal fault (Reilly et al. 2016), as was the smaller normal fault at the eastern edge of the Kokako Survey (Fig. 3). This tectonic configuration supported the subsidence of the study area with tectonics contributing to accommodation development. However, the late Pliocene was characterized by major high-frequency eustatic variations cumulatively lowering sea level (Fig. 13). The pronounced topset downbuilding of the Plio-Pleistocene unit 8 can be, therefore, best explained by a cumulative eustatic fall of around 100 m between c. 4 and 2 Ma. The exceptional topset preservation in this clinoform succession likely results from the recurrent, high-frequency, short-term ups and downs in Plio-Pleistocene sea level, with individual intervals of accommodation loss and gain often only lasting between few 10’s (Naish and Kamp 1997) to c. 100 kyears (Miller et al. 2005). These exposure times were likely too short for the establishment of more widespread erosive systems capable of removing a significant amount of rock from onshore coastal areas and further inland. Recurrent flooding of backshore areas might have contributed to the spectacular preservation of (tidally influenced?) palaeo-shorelines and coastal-plain environments (Fig. 11) by sealing. In turn, the low preservation of the topsets of the Miocene units 3 and 4 probably resulted from longer and constant, tectonically triggered exposure times lasting at least several 100,000 up to a few millions of years (Fig. 13), producing a major erosional peneplain and depositional hiatus at the top of units 3 and 4 (Fig. 4).

Fig. 13
figure 13

Summary chart of seismic units, approximate chronology, short-term eustasy (Hardenbol et al. 1998), base-level change, unit architecture, and tectonic activity along the Cape Egmont Fault (King and Thrasher 1996; Nicol et al. 2005, Reilly et al. 2015, 2016). Red arrows correspond to mean eustatic change and black arrows to total base-level change measured on clinoform breakpoints. Note pronounced differences in base-level change and eustatic variations between downbuilding progradational units 4 and 8. Progradational units 3 and 4 are interpreted as forced by tectonic uplift. Downbuilding progradational unit 8 is interpreted as predominantly forced by eustasy

Full size image

Topset erosion, topset deposition, and shoreline migration likely exert an important control on depositional processes beyond the clinoform breakpoint, and their understanding can be potentially used to predict slope sedimentary processes. Figures 10, 11 indicate that the transition between upbuilding and downbuilding progradation (unit 7 to unit 8; APD pattern sensu Neal and Abreu 2009; Neal et al. 2016) is marked by the widespread occurrence of mass-transport deposits on the clinoform foresets, which seems not necessarily tied to foreset steepness. This observation can be possibly set in relation with a decreasing water depth in distal shelf environments unloading the sedimentary column and lowering the storm-wave base, leading to increased sediment reworking and instability at the clinoform breakpoint triggering mass failure. The resulting mass flows can involve fine-grained outer shelf deposits that “freeze” on decelaration (Stow et al. 1996; Catuneanu 2006), with a little suspended sediment exiting the mass-transport deposits downstream in gullies (Fig. 11b). Figure 7 images the clinoform foresets of units 3 and 4 that also show a significant amount of mass-transport deposits. Following the above line of arguments, this occurrence could also relate to deposition in a transitional phase between upbuilding and downbuilding progradation; however, the lack of complete foreset–topset profiles in these units makes this interpretation more speculative.

Above the mass-transport deposits, the downbuilding clinoform units 4 and 8 are marked by series of gullies on the upper foresets, which seem to grow landward through time onto the clinoform foresets (e.g., Fig. 11c–e). Shumaker et al. (2017) have interpreted slope gullies near the transition between downbuilding unit 4 and upbuilding unit 5 as resulting from sheet-like turbidity currents stemming from autogenic slope instabilities that promoted regularly spaced variations in erosion and deposition. Many of the gullies documented in this study also seem to initiate on the slope (see Fig. 11c), but often extend upward with time firstly into and then across the clinoform breakpoint by headward erosion (Figs. 7d, 11d to g). At and landward of the clinoform breakpoint, which was in the downbuilding clinoform units 4 and 8 likely in the shoreface depositional environment, foreshore processes (tidal currents and waves) as well as backshore and coastal-plain erosional systems seem to have controlled the respective gully position, and, therefore, sediment supply to the clinoform foresets. Yet, although similar in size, the gullies of the eustatically forced downbuilding unit 8 show mainly shoreline-parallel feeder systems restricted to the coastal zone (Fig. 11d–f), while the feeder systems of the tectonically forced downbuilding unit 4 are generally downslope oriented (Fig. 7d). This observation indicates that genetically similar stratigraphic units can originate from different sediment supply systems, with slope depositional systems potentially carrying a footprint of tectonics or eustasy as superordinate controls.

A final discussion point concerns the stratigraphic approach and terminology used in this study. The omission of “classic” sequence-stratigraphic terms and nomenclature (i.e., relative sea level, forced regression, sequences and sequence boundaries, systems tracts, key stratigraphic surfaces) was deliberately chosen (1) to emphasise the importance of model-independent 3D seismic–stratigraphic analysis for depositional systems interpretation; (2) to avoid complex and potentially misleading terminology; one example of many would be naming tectonically forced sedimentary downbuilding a “forced regression”, with nothing in the studied system “regressing” unless defined as such; (3) to avoid the choice between a multitude of sequence-stratigraphic models (see Catuneanu et al. 2009 for review), that—if used in a strict sense—might force unwanted aspects into the interpretation. An example for this is the separation of upbuilding progradational unit 5 and retrogradational unit 6 by a rather diffuse zone of reflection terminations up to a few tens of ms (TWT) thick. All sequence stratigraphic models require a bounding surface instead of a bounding zone for systems tract separation; however, such a distinct surface might not necessarily exist. The development of a bounding zone instead of a surface could, e.g., result from to higher order variations in either sediment input or accommodation rates (e.g., higher order sequences of Plint and Nummedal 2000), or—under marine conditions—current-driven sedimentation or slope processes independent of base-level change. This indicates, in either case, that a sharp separation of genetic stratigraphic units by single bounding surfaces, as inherent to most sequence-stratigraphic models, might not be in all the cases applicable.

Conclusions

  1. 1.

    The central Taranaki Basin offshore New Zealand is characterized by a spectacular Miocene-to-Pleistocene clinoform succession that developed immediately north of a major fault zone under highly variable tectonic conditions: the Miocene part of the dominantly progradational succession was strongly influenced by contractional tectonics and uplift at and south of the Cape Egmont Fault. In contrast, the Pliocene part of the succession developed in a tectonic environment dominated by extension and basin subsidence. The study area is extensively covered by industrial 3D seismic data and boreholes. The favourable data situation supported a largely automated 3D seismic-reflection interpretation of the palaeo-sedimentary record.

  2. 2.

    The studied clinoform succession reveals two distinct intervals of downbuilding progradation between c. 7.5 and 6 Ma and 4 and 2 Ma, respectively. The Late Miocene interval is interpreted as solely controlled by tectonics, with sedimentary progradation and downbuilding of the shoreline reflecting inversion-related uplift at the Cape Egmont Fault and along neighbouring structures. Pliocene–Pleistocene downbuilding is interpreted as dominantly controlled by a cumulative eustatic fall (overlaid by short-term, high-amplitude sea-level fluctuations) against a local tectonic background characterized by subsidence.

  3. 3.

    The preservation of clinoform topsets in large parts of the studied succession provides unique insights into depositional systems at and above the shelf break, including exceptional views into shoreline systems formed during sedimentary downbuilding. The topsets of Pliocene–Pleistocene downbuilding clinoforms are particularly well preserved, revealing shoreline-parallel sediment supply systems controlling the location of slope gullies at and beyond the clinoform breakpoint. The preservation of these features likely results from short Plio–Pleistocene exposure times (few 10’s to 100 kyears?), which were likely too short for erasing the topset sedimentary record.

  4. 4.

    Linkage of depositional systems between topsets and foresets allows analysing deepwater sedimentary patterns dependent on the respective clinoform architecture and trajectory trend. Periods of significant mass-wasting seem to fall preferentially into the transition between upbuilding and downbuilding progradation. Deepwater gullies and channels can be observed during both progradation and retrogradation. Upbuilding progradational units seem to develop larger deepwater gullies and channels than downbuilding progradational and retrogradational units.