The South Westland Basin: seismic stratigraphy, basin geometry and evolution of a foreland basin within the Southern Alps collision zone, New Zealand

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Tectonophysics 300 (1998) 359–387

The South Westland Basin: seismic stratigraphy, basin geometry and evolution of a foreland basin within the Southern Alps collision zone, New Zealand Keith N. Sircombe 1 , Peter J.J. Kamp * Department of Earth Sciences, The University of Waikato, Private Bag 3105, Hamilton 2001, New Zealand Accepted 20 September 1998

Abstract This paper develops further the case for a foreland basin origin of South Westland Basin, located adjacent to the Southern Alps mountain belt. Geohistory analyses show Middle Miocene initiation of subsidence in the basin, with marked increases at 5–6 Ma. Five seismic reflection horizons, including basement, Middle Miocene (top Awarua Limestone), top Miocene, mid-Pliocene (PPB) and mid-Pleistocene (PPA) have been mapped through the grid of seismic data. A series of five back-stripped structure contour maps taken together with five isopach maps show that prior to the Middle Miocene, subsidence and sedimentation occurred mainly along the rifted continental margin of the Challenger Plateau facing the Tasman Sea; subsequently it shifted to a foredeep trending parallel to the Southern Alps and located northwest of them. Through the Late Miocene–Recent this depocentre has progressively widened, and the loci of thickest sediment accumulation have moved northwestward, most prominently during the Late Pliocene and Pleistocene with the progradation of a shelf–slope complex. At the northern end of the basin the shelf–slope break is currently located over the forebulge, which appears not to have migrated significantly, probably because the mountain belt is not advancing significantly northwestwards. Modelling of the lithospheric flexure of the basement surface normal to the trend of the basin establishes values of 3.1 to 9:8 ð 1020 N m for the flexural rigidity of the Australia Plate. This is at the very low end of rigidities for plates, and 1–2 orders of magnitude less than for the Australia Plate beneath the Taranaki Basin. Maps of tectonic subsidence where the influence of sediment loading is removed also clearly identify the source of the loading as lying within or beneath the mountain belt. The basin fill shows a stratigraphic architecture typical of underfilled ancient peripheral foreland basins. This comprises transgressive (basal unconformity, thin limestone, slope-depth mudstone, flysch sequence) and regressive (prograding shelf–slope complex followed by molasse deposits) components. In addition the inner margin of the basin has been inverted as a result of becoming involved in the mountain building, as revealed earlier by fission track thermochronological data. The timing and degree of inversion fits well with the geometrical and stratigraphic development of the basin. That the inversion zone and the coastal plain underlain by molasse deposits are narrow, and most of the basin is beneath the sea, highlights this as an underfilled active foreland basin. The basin is geodynamically part of the Southern Alps collision zone.  1998 Elsevier Science B.V. All rights reserved. Keywords: Southern Alps; New Zealand; Australia–Pacific plate boundary; stratigraphy; tectonics

1

Present address: Geological Survey of Canada, 601 Booth Street, Ottawa, ON K1A OE8, Canada. author. Fax: C64-7-856-0115; E-mail: [email protected]

Ł Corresponding

0040-1951/98/$ – see front matter  1998 Elsevier Science B.V. All rights reserved. PII: S 0 0 4 0 - 1 9 5 1 ( 9 8 ) 0 0 2 5 4 - 6

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1. Introduction The central part of South Island is regarded widely as comprising a zone of continent–continent collision involving the Australia and Pacific plates

(Fig. 1). Certain elements of this collision zone have historically been investigated in some detail, including particularly the Alpine Fault, which is the boundary between the Australia and Pacific plates, and the Southern Alps, which are the topographic ex-

Fig. 1. Map of South Island showing the main tectonic features of the Southern Alps collision zone, including the South Westland and Canterbury foreland basins. Depth to basement shown as generalised structure contours for the South Westland Basin.

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pression of uplift and erosion of the leading edge of the Pacific Plate. Few previous studies have viewed the two sedimentary basins (Canterbury and South Westland) flanking the central zone of uplift as integral parts of the collisional orogen. The results of recent geodynamical modelling, constrained by what is known about the surface and subcrustal structure of the orogen, showed clearly (Beaumont et al., 1996, fig. 15) that foredeeps are predictable elements on both sides of the Southern Alps. The basins fill space created marginal to the orogen by broad downward lithospheric flexure caused by crustal and subcrustal loading within and beneath the mountain belt (see for example fig. 15, Beaumont et al., 1996). A key feature of the South Westland Basin is that its geometry is asymmetrical, deepening or thickening towards the mountain belt and it has a distinctive stratigraphic architecture. This study documents the geometry, seismic stratigraphy and aspects of the basin history of the South Westland Basin in the context of the Southern Alps collision zone. It builds upon earlier definition of the basin from seismic data (Nathan et al., 1986), and its identification as a foreland basin (Kamp et al., 1992). This had become evident from the results of apatite fission track thermochronology on samples from basement exposed onshore in the narrow coastal strip, the interpretation of which required the former occurrence of a greater thickness (ca. 4 km) of Middle to Late Miocene section onshore than is preserved in the offshore realm, thus defining a fundamental basin asymmetry typical of foreland basins (e.g. Beaumont, 1981).

2. South Westland Basin We use the term South Westland Basin to refer to the sedimentary basin lying between Hokitika and Milford Sound, and west of the Alpine Fault (Fig. 1). It is geographically distinct in terms of including the narrow part of the coastal plain and contiguous offshore shelf and slope bounded to the southeast by the Alpine Fault Zone and Southern Alps. This basin passes southwestward into the Puysegur Trench where the Australia Plate subducts obliquely beneath Fiordland (Davey and Smith, 1983). Northwestward of the basin lies the transition from early

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Palaeozoic continental crust underlying the Challenger Plateau to Late Cretaceous oceanic crust of the Tasman Sea (Fig. 1). Existing information about the character and development of the basin has originated chiefly from the stratigraphy and structure exposed onshore (Nathan et al., 1986). The principal structure in southern Westland is a monoclinal fold, known as the Coastal Monocline (Cotton, 1956). The steeply dipping Late Cretaceous and Cenozoic beds exposed along parts of the coastline constitute the faulted vertical limb of the monocline (Nathan, 1977). The gently dipping upper limb has mostly been eroded, whereas the lower limb is preserved entirely offshore, seismic profiles of which show an almost flat-lying sequence several km thick (McNaughton and Gibson, 1970; Nathan et al., 1986). Basement rocks onshore comprise mostly a Lower Ordovician greenschist facies, thickly bedded greywacke succession (Greenland Group) (Nathan, 1977). Several small granite plutons confined to a zone 2–3 km west of the Alpine Fault have intruded the Greenland Group; a Rb–Sr age of 357 Ma has been reported for one of them (Aronson, 1968). A Late Cretaceous–Cenozoic sedimentary succession with interbedded volcanics unconformably overlies basement and has been described in detail by Nathan (1977). The lowermost unit comprises local occurrences of breccias and sandstones (Otumotu Formation) up to 500 m thick of mid-Cretaceous age (Ngaterian Stage). These are overlain by up to 500 m of a transgressive Late Cretaceous–Eocene succession of (Tauperikaka) coal measures, glauconitic (Whakapohai) sandstone, basalt flows and breccias (Arnott Basalt), foraminiferal limestone and mudstone (Abbey Fm), and further basalt (Otitia) and associated volcaniclastic-derived sediments. In a few localities a thin (10 m) foraminiferal limestone (Awarua Limestone) of poorly determined and presumed Oligocene age overlies basement unconformably. This is followed stratigraphically by the Tititira Formation of Middle Miocene (Lillburian–Waiauan) age consisting of a regressive submarine fan sequence, passing upwards from hemipelagic mudstone through distal to proximal turbidites (Nathan, 1978). This grades upwards into thick-bedded mass-emplaced conglomerates and sandstones inferred to be inner-fan and channel de-

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posits (Nathan, 1978). The Tititira Formation and all underlying units onshore are truncated by a regional unconformity that is overlain by glacial moraines and outwash deposits of Pleistocene age. Along strike to the northeast the Coastal Monocline develops into a steeply dipping reverse fault zone known as the Hohonu Fault (Nathan et al., 1986) with about 2.5 km of offset. Two discrete fault systems, the Fraser Fault and a set of younger thrust faults, lie between the Hohonu Fault and the Alpine Fault. The Fraser Fault juxtaposes a complex of amphibolite facies gneiss and granitoids enveloped by anastomosing mylonitic zones (Rattenbury, 1986) against Early Cretaceous granitoids in the north and Greenland Group greywackes with Devonian– Carboniferous granitoids in the south (Morgan, 1908; Rattenbury, 1986). A set of younger low-angle (20º–30º) east and southeast dipping thrusts displace the Fraser Complex over basement to the west of the Fraser Fault (Rattenbury, 1986). Two highly eroded nappes of alpine schist originating east of the current Alpine Fault trace, but now located about 2 km northwest of it, are included in the younger set of thrusts by Rattenbury (1986), and are considered to represent inactive surface traces of the Alpine Fault. The results of fission track thermochronological analyses on basement rocks from South Westland have indicated the amount and timing of late Neogene uplift and denudation within the southeastern part of the basin (Kamp et al., 1992). In southern South Westland the amount of denudation increases from 3.5 to about 5.0 km between the coastline and the Alpine Fault; this has occurred since 5 Ma, and involved chiefly the erosion of a Middle-to-Late Miocene basin fill succession. In the northern part of South Westland two basement blocks have been differentially inverted since about 10 Ma, involving the erosion of up to 5 km of the sedimentary section. 2.1. Seismic reflection and well data Several seismic reflection surveys have been undertaken by the industry for hydrocarbon exploration in the South Westland Basin and three exploration holes have been drilled. Open-file petroleum reports (Table 1) have been the source of the seismic profiles and well stratigraphy. Some 1715 km of seismic reflection profiles including 104 km onshore, as il-

Table 1 Open-file petroleum reports and seismic line series Seismic line series

Petroleum Rep. No.

Year

Operator

EZDL-, OSMWNZPS-

400 523 548 575 614 629

1968 1971 1967 1971 1973 1974

Esso Exploration N.Z. Petroleum Shell B.P. Magellan Petroleum Australian Gulf N.Z. Petroleum

lustrated in Fig. 2, have been used in this study. The quality of the seismic data varies greatly. Earlier data sets, such as the M-series, are often of poor quality. Fortunately, the prominence of certain seismic reflectors, as discussed below, have meant that some useful information could be gained even from the lesser-quality data sets. The data were analysed in several steps. First, the seismic reflection profiles were interpreted using standard seismic stratigraphic techniques (Vail, 1987; Cross and Lessenger, 1988). Prominent horizons tied into the well records were digitised using an engineering relational database management system (TECHBASE) running on an IBM-PC compatible computer (MINEsoft, 1991). The two-way travel time data were converted to depth using velocity information from well-logs. Structure contour, and isopach maps were drawn and contoured for and between each of the horizons mapped within TECHBASE. 2.2. Lithostratigraphy and geohistory of well successions There are three exploration wells in the basin: Harihari-1 (New Zealand Petroleum Co. Ltd., 1971), Waiho-1 (New Zealand Petroleum Co. Ltd., 1972), which are both onshore, and Mikonui-1 (Diamond Shamrock Oil Company, 1981), located about 50 km offshore. All three wells lie on seismic lines enabling good estimates of geophysical values to be made for the mapped units (Fig. 3). The lithostratigraphy of the three wells and the correlation are shown in Fig. 4. All three wells are dominated by mudstone lithologies, with lesser amounts of sandstone and conglomerate. Two of

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Fig. 2. Map showing the location of seismic reflection profiles and exploration holes used in this study, as well as bathymetry and the extent of the zone of inversion (underlain by Palaeozoic basement) and Late Cretaceous and Cenozoic successions exposed onshore. Location of flexure profiles illustrated in Fig. 12 are also shown. Bathymetry from Eade (1972) and Carter (1981).

the wells (Waiho-1 and Harihari-1) contain a thin limestone (Awarua Limestone) at the base of the sedimentary section which in turn overlies basement. In Mikonui-1 a thin coal measure section overlies basement. The upper units of Mikonui-1 are quite rich in bioclastic sediments, whereas the upper parts of the two onshore wells comprise glacial and=or fluvioglacial deposits. Geohistory analyses were undertaken on all three wells based on stratigraphic data in the respective well reports. The porosity and compaction constants of Falvey and Deighton (1982) and Schmoker and

Halley (1982) were used along with a programme from Wood (1989). The geohistory plots for each well are illustrated in Fig. 5. Mikonui-1 shows clearly a two-stage subsidence history (Fig. 5). A 55 m.y. history of slow subsidence is followed at 5–6 Ma by a dramatic increase in the rate of subsidence. The tectonic subsidence corresponding to the first stage has been modelled using the simple McKenzie (1978) uniform stretching model. This model assumes ‘instantaneous’ and uniform stretching of the lithosphere with passive upwelling of hot asthenosphere to maintain isostatic

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Fig. 3. Summary of the age and nomenclature of the seismic reflection horizons and successions mapped throughout South Westland Basin, together with seismic velocity, compaction coefficients .k/ and original porosity ./ calculated from geophysical and lithologic data in Mikonui-1.

equilibrium. A number of more complex modifications to the McKenzie model exist, but the lack of detailed geologic data needed to support them in this region prevents their application. A summary of the formulas and standard parameters used in subsidence modelling is given in Appendix A. Models have been calculated assuming a pre-rifting lithosphere thickness of 125 km and a stretching factor of 1.20, 1.25 and 1.30, and are compared with the calculated tectonic subsidence for Mikonui-1 (Fig. 6). In this well the overall best-fit line appears to be þ D 1:25, indicating that the lithosphere at Mikonui-1, approximately 200 km from the edge of the Challenger Plateau, was thinned from 125 km to 100 km, probably during the Late Cretaceous preceding the start of Tasman Sea spreading at about 80 Ma. Calculated tectonic subsidence in Mikonui-1 continues to fit the model quite well even after the inception of the Australia–Pacific plate boundary (Alpine Fault) at about 23 Ma. We attribute the rapid increase in subsidence at 5–6 Ma to originate from loading within the collision zone onshore in South Island. The sedimentary sections encountered at Waiho-1 and Harihari-1 both have thin Middle Miocene limestone at their bases, followed by thick late Neogene beds. Subsidence and terrigenous sedimentation started during the Middle Miocene, and in both well sections show increases in the rates of subsidence at 5–6 Ma. The shapes of the tectonic subsidence curves in all three wells are very similar to those constructed for classical foreland basins in North

America, e.g. the Alberta (Kominz and Bond, 1986), Hoback and Utah basins (Cross, 1986). After close examination of the seismic reflection profile data set, six horizons, including the modern seafloor, and five sequences (A upwards to E) were interpreted and mapped (Fig. 3). Twin basal reflectors corresponding to basement and the Awarua Limestone are very prominent in most seismic profiles and relatively easy to interpret and map. A top Miocene reflector tied to the succession encountered in Mikonui-1 has been mapped although it is not everywhere as a prominent horizon. Two prominent horizons were mapped in the thick Pliocene–Pleistocene sequence. Both appear to be reflectors overlain by sediments that contain clinoforms. The lower horizon, named Pliocene–Pleistocene B (PPB), is correlated with the Waipipian–Mangapanian boundary (3.1 Ma) in Mikonui-1. This horizon appears to be regionally extensive. The Upper Pliocene– Pleistocene horizon (PPA), is of more limited extent in the basin. This horizon is correlated with the Nukumaruan–Castlecliffian boundary (1.2 Ma). Unfortunately the PPA horizon often coincides with the first seafloor multiple in the reflection profiles. This limits somewhat the interpretation of the horizon and overlying reflector associations and structures. The basement horizon is a relatively smooth virtually unfaulted surface but it does show some irregularities. It can be mapped easily across the basin because it marks a boundary in the profiles between

K.N. Sircombe, P.J.J. Kamp / Tectonophysics 300 (1998) 359–387 Fig. 4. Stratigraphy and correlation of Mikonui-1, Harihari-1 and Waiho-1 exploration holes in South Westland Basin. After Diamond Shamrock Oil Company (1981), Enclosure 6. 365

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Fig. 5. Geohistory plots for each of Mikonui-1, Harihari-1 and Waiho-1 exploration holes. Note the increase in the rate of subsidence in each of the wells at 5–6 Ma.

continuous reflectors of the overlying cover and discontinuous reflectors in the more indurated and deformed basement (Figs. 7–10). The relatively smooth basement surface reflects the weathered character of the upper 5 to 10 m of the basement as shown by the composite logs for the wells. The Middle Miocene horizon corresponding to the Awarua Limestone is also prominent in reflection profiles (Figs. 7–10) and is more or less parallel to the basement reflector. In places the horizon probably includes limestone of Oligocene age. Limestone of Oligocene age, particularly Late Oligocene age, are common in western New Zealand resulting from the Late Cretaceous–Cenozoic marine transgression of the New Zealand platform (Nelson, 1978).

The top Miocene horizon is difficult to trace across the basin as it is often of low amplitude, although it is associated with several prominent reflectors (EZD-3, Fig. 10). It has been correlated with the Miocene–Pliocene boundary in Mikonui-1. The horizon generally dips toward the southeast (PS-21, Fig. 9; EZD-3, Fig. 10), but it rises near the coastline in some sections (NZ-102, Fig. 7; L-206, Fig. 8). The PPB horizon is the regionally more extensive of two breaks in the basin fill succession. It is well displayed in line L-206 (Fig. 8), where it dips in a series of shallow steps to the northwest. This horizon is correlated with the Waipipian–Mangapanian stage boundary in Mikonui-1 at around 3.1 Ma. The undulose nature of the horizon and the fact that it

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Fig. 6. A plot of calculated tectonic subsidence in Mikonui-1 compared with modelled tectonic subsidence for a lithosphere of original thickness 125 km and stretched instantaneously by a factor of 1.20, 1.25 and 1.30 at 60 Ma. The best-fit model is the 1.25 factor. Note the change in tectonic subsidence at 5–6 Ma when the well data diverge from the subsidence modelled for extension.

truncates underlying reflectors implies an erosional or slump origin possibly associated with relative sea level changes. The PPA horizon is a prominent slope reflector that can be traced into topset beds (e.g. Figs. 7 and 9). By correlation with the succession in Mikonui-1, the horizon is just older than the Nukumaruan–Castlecliffian boundary and is assigned an age of 1.2 Ma. The reflector marks the boundary, at least, between two depositional packages, as discussed below, and its origin is probably associated with relative sea level changes. The continental shelf is very narrow offshore from South Westland, and most of the basin investigated underlies the continental slope at water depths of 250–1000 m (Fig. 2). The break in slope between the shelf and slope is very clear on seismic profiles that transect this region. Several major canyon systems cut the shelf and slope, notably the Hokitika Canyon complex in the northeast, and the Cook Canyon to the southwest, which has two tributaries (Moeraki and Haast).

3. Basin geometry A series of five structure contour maps, drawn for each of the seismic horizons mapped, illustrate the basin geometry and its subsidence history (Fig. 11). Fig. 11a is a traditional structure contour map on

basement. Fig. 11b through Fig. 11e are maps that show for particular times corresponding to each of the horizons mapped, the depth to basement having stripped off the overlying sediments and decompacted those remaining. The structure contour on the basement map (Fig. 11a) shows several features. (1) The minimum depth to basement occurs along the central axis of the basin, where the basin floor is between 1500 and 2100 m below sea level. The Mokonui basement high shows about 500 m of relief in an elliptical pattern, some of which will have been inherited from Late Cretaceous–early Cenozoic topography. (2) To the northwest and southeast of the central axial high, the depth to basement increases. On the northwest side the depth to basement increases, but data coverage runs out; it is expected that the depth to basement would gradually increase towards the edge of the Challenger Plateau and continent–ocean crust transition. On the southeast side of the central high, the depth to basement increases regularly towards the coastline and reaches a maximum of 3600 m, although it is more commonly 3000 m (Fig. 11). This geometrical pattern has been attributed to loading within the continental collision zone (Kamp et al., 1992). Fig. 11b through Fig. 11e are structure contour maps on basement showing the theoretical depth to basement below contemporary sea level for times corresponding to the seismic horizons mapped earlier. The maps have been generated in the same man-

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Fig. 7. Reproduction of seismic reflection line NZ-102, and, below, its interpretation.

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Fig. 8. Reproduction of seismic reflection line L-206 and, below, its interpretation.

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Fig. 9. Reproduction of seismic reflection line PS-21 and, below, its interpretation.

ner and by using the same parameters as applied in the geohistory analysis of the well sections (Fig. 5). The chief limitations on the calculation of the theoretical structure contours for each map lie in the palaeowater depths at the times (horizons) for which the maps are drawn, and the Cenozoic eustatic sea level history. For this exercise the Haq et al. (1987)

sea level curve was used and a uniform palaeowater depth was applied across the basin. Although these assumptions are imprecise, the magnitude of the errors introduced by them (500 m for the Oligocene horizon, and 300 m for subsequent horizons) are unlikely to be substantial and to affect the major conclusions drawn from the analysis.

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Fig. 10. Reproduction of seismic reflection line EZD-3 and, below its interpretation.

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372 K.N. Sircombe, P.J.J. Kamp / Tectonophysics 300 (1998) 359–387 Fig. 11. A series of five structure contour maps: (a) is a traditional structure contour map on basement; (b–e) are maps for the Middle Miocene, top Miocene, mid-Pliocene and mid-Pleistocene horizons, respectively. In these maps the overlying units have been stripped off and the remaining sediments decompacted. This shows the location and magnitude of the depocentres and sequentially, the geometrical evolution of the basin. Contour intervals are 250 m. Note that the southwestern end of the South Westland Fault Zone is poorly constrained.

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Fig. 11 (continued).

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Fig. 11 (continued).

The usefulness of a series of back-stripped structure contour maps is that they illustrate more effectively than traditional structure contour maps, the changes in location and magnitude of the basin depocentre(s), and hence the geometrical evolution of the basin. During the Middle Miocene, after the deposition of the Awarua Limestone and related beds, the depocentre lay distant from the modern shoreline and along the northwest margin of the area mapped. Sediment accumulation was focused along the continental margin, although there appear to have been some small areas near the modern coastline where a Late Cretaceous–Oligocene section up to 1000 m thick had accumulated, which is consistent with the onshore stratigraphy. By the end of the Late Miocene (Fig. 11d) a new axial depocentre located parallel to the Alpine Fault and in the vicinity of the modern coastline had developed. The depth to basement there increased from about 600 m to at least 1500 m and possibly 2000 m in places. By comparison, over the central axial high, a distance of only 45 to 50 km to the northwest, the depth to basement increased

by only 100–200 m. Over the same interval, the depth to basement along the northwest continental margin of the basin increased by about 200 m. This indicates clearly that the subsidence generating the new depocentre during the Middle-to-Late Miocene originated along the southeast margin of the basin. The pattern in the basin geometry established by the Late Miocene was accentuated during the Pliocene and Pleistocene. During the Early and midPliocene, that is, for the interval between Fig. 11c and Fig. 11d, the total amount of subsidence along the southeast margin of the basin doubled to about 3000 m and it was still focused near the modern coastline. During the Late Pliocene and Pleistocene (Fig. 11d,e), however, the depocentre widened and deepened in an offshore direction, which resulted in about 1000 m of subsidence of the axial central high. This was probably driven by contemporaneous uplift and erosion of the inner margin of the basin between the modern coastline and the Alpine Fault (Kamp et al., 1992).

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3.1. Modelling of the lithospheric flexure The geometrical development of the South Westland Basin in the context of other geological features (see below) and the proximity of an active collisional mountain belt, suggests that it may have formed by loading and flexure of the leading edge of the Australia Plate. The flexure revealed by the present structure on the basement horizon can be modelled and information about lithospheric behaviour assessed. In Fig. 12 the profile of the depth to basement in seismic line L-206 and two cross-sections drawn across the total depth to basement map (Fig. 11a) are compared with flexural models (Sircombe, 1993). The South Westland Fault Zone is inferred as the edge of the Australia Plate for flexural modelling purposes. The formulas and standard parameters used in the flexural models are listed in Appendix A, and the observed and derived parameters are listed in Table 2. The results for each of the three cross-sections indicate immediately that if the basement structure is completely the result of flexure, the lithosphere of the Australia Plate under the South Westland Basin is extremely weak with a flexural rigidity 3.1 to Table 2 Lithospheric parameters observed and calculated in the flexural models illustrated in Fig. 12 Observed parameters: Distance to forebulge xb 60 km Maximum deflection !0 1.40 km 1.10 km 0.80 km Calculated parameters: Unbroken lithosphere Flexural parameter Þ Flexural rigidity D Elastic thickness Te Line load V0

Broken lithosphere Flexural parameter Flexural rigidity Elastic thickness Line load

Þ D Te V0

(96K section) (150K section) (L-206 section)

375

9:8 ð 1020 N m depending on the model used. These values occur at the low end of lithospheric rigidities established so far (Karner et al., 1983). Holt and Stern (1991) calculated the rigidity of the lithosphere under the Western Platform of Taranaki Basin to be in the region of 5:6 ð 1022 –1:5 ð 1023 N m, 1–2 orders of magnitude greater than the values for South Westland calculated here. There are several possible explanations for the apparent weakness of the lithosphere beneath the South Westland Basin. The most obvious is that the model adopted is too simple or there is something special about the region. For example, the basin lies at the southern end of the Challenger Plateau and close to the transition with the oceanic crust of the Australia Plate generated during the Late Cretaceous–mid-Cenozoic by spreading in the Tasman Sea and Emerald Basin. Heating associated with rifting and early spreading was therefore more marked and more recent than in Taranaki Basin, which has much higher lithospheric rigidities (Holt and Stern, 1991). Another possible explanation of the apparent weakness of the lithosphere is that our seismic mapping has underestimated the original topography associated with the central high running along the axis of the basin. Part of this ridge may be original topography, and therefore may not originate completely as a forebulge. For the moment we note the apparently weak lithospheric rigidity for the leading edge of the Australia Plate, which is reflected in the narrow depocentre parallel to the mountain belt, and leave any further explanation to a more rigorous analysis.

4. Basin fill 19.10 km 3:10 ð 1020 3.65 km 4:98 ð 1011 3:92 ð 1011 2:85 ð 1011 25.46 km 9:79 ð 1020 5.35 km 3:32 ð 1011 2:61 ð 1011 1:90 ð 1011

Nm Nm Nm Nm

1 1 1

(96K section) (150K section) (L-206 section)

Nm Nm Nm Nm

1 1 1

(96K section) (150K section) (L-206 section)

The character of the basin fill succession as described from outcrop (Nathan, 1977, 1978), recorded in well logs and observed in seismic reflection profiles, complements the basin evolution identified from the structure contour maps. Here we address the stratal patterns evident in the reflection profiles (e.g. Figs. 7–10) and in isopach maps (Fig. 13), and from them describe further aspects of the basin evolution. Fig. 13a illustrates the thickness of the Late Cretaceous–Oligocene succession (E, Fig. 3). Sedi-

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Fig. 12. Flexural models of the 96K, 150K and L-206 sections on basement. The model curves were derived using parameters given in Appendix A. Locations of sections are shown in Fig. 2.

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mentation was focused along the continental margin as evident in the distal sections of NZ-102, L-206 and PS-21 (Figs. 7–9), but there were also depocentres in the vicinity of the modern coastline, which are also known from the outcrop section. These may reflect contemporary crustal mobility along an eastern margin of South Westland now overthrust by the Pacific Plate. The thinness of succession E precludes in many cases interpretation of internal seismic structures, but where this is possible the reflectors are of moderate to high amplitude and continuity. The reflectors terminate by onlap or are concordant; some show toplap with respect to the Oligocene horizon. The thickness distribution of succession E and the reflector terminations are consistent with the stratigraphic pattern in western New Zealand for a change from locally influenced sedimentation to widespread submergence and marine transgression (Nathan et al., 1986). Succession D (Figs. 3 and 13b) shows the opposite direction of thickening (towards the southeast) compared with that of succession E (Fig. 13a). The succession comprises a broad layer across the basin, with a wedge-like thickening along the southeastern margin, organised into several lobes that appear to form the toe of a continental slope. The reflectors terminate mostly by downlap onto the Middle Miocene reflector within the southeastern depocentre and by onlap onto the axial high. Reflectors are generally concordant with the upper boundary. Internally the succession comprises low-amplitude continuous reflectors. Mudstone is the dominant lithology, with a few sandstone and conglomeratic horizons recorded in Waiho-1. This succession correlates with the Tititira Formation onshore. It grades upwards from hemipelagic mudstone through distal and proximal turbidites into a mass-flow conglomerate–sandstone complex (Nathan, 1978). The stratigraphic transition from hemipelagic mudstone to mass-emplaced deposits coincides with the development of slope fans, a more proximal source area supplying the sediments, and increasing rates of subsidence and sedimentation along the southeast margin of the basin in the vicinity of the modern shoreline (Fig. 11c and Fig. 13b). The isopach map for succession C (Fig. 13c) shows marked Early to mid-Pliocene deposition within the basin along the southeast margin, in-

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volving a marked increase in the width of the depositional axis. This represents progradation of a lower slope environment northwestwards across the depocentre and towards the forebulge=axial high. In profile EZD-3 (Fig. 10) a prominent wedge structure shows strong downlap onto the top Miocene reflector. This could represent progradation of a welldeveloped lower slope fan complex. In line L-206 several strong reflectors are interpreted as another fan complex, possibly of low-stand origin. In more distal parts of the basin there is gentle onlap onto the contemporary forebulge, and the reflectors are continuous and weakly hummocky. From the section encountered in Waiho-1 and Harihari-1, the main sedimentary lithology is expected to be mudstone with sandy turbidites. Successions A and B show further evolution of the basin fill in that the locus of deposition shifts progressively further offshore and into the basin (Fig. 13e and Fig. 13d, respectively). This reflects active progradation of the upper slope and shelf, resulting from increased sediment flux. Coinciding with this phase is subsidence of the initial forebulge=axial high, such that its northern part in the vicinity of Mikonui-1 lies beneath the outer shelf–upper slope. Seismic reflection profiles (Figs. 7, 9 and 10) show very clearly a clinoform structure within these (A and B) depositional successions. This structure is very similar to that displayed by deposits of similar age in the Taranaki Basin where they are known as the Giant Foresets Formation (Beggs, 1990). Succession B in general comprises the middle to lower parts of a prograding continental slope where the clinoforms are concave upwards. The contemporary upper slope and shelf would have been very narrow and located in the vicinity of the modern coastline or further inland. By comparison, succession A comprises the whole prograding slope and shelf (compare Figs. 7, 9 and 10). This illustrates the amount of progradation of the shelf and slope since about 1.2 Ma. Another notable feature of both successions is that parts of the reflectors are disrupted in middle to lower slope positions. These are considered to result from collapse of parts of the slope giving rise to large slump structures. The main lithology of both sequences is expected to be variably sandy mudstones, based on well records (Fig. 4). The bioclastic content of the upper part

378 K.N. Sircombe, P.J.J. Kamp / Tectonophysics 300 (1998) 359–387 Fig. 13. A series of five isopach maps, one for each of the stratigraphic units lying between the reflection horizons mapped. (a) Illustrates sequence=succession E (Fig. 3), through to (e), which illustrates sequence=succession A. Note the marked change in the loci of sedimentation between maps (a) and (b), which reflects the shift from deposition on the rifted continental margin of the Challenger Plateau to the new foredeep adjacent to the developing mountain belt. Note also in maps (c), (d) and (e) how the loci of sedimentation shifted progressively into the basin and away from the mountain belt, reflecting progradation of a shelf–slope complex.

379

Fig. 13 (continued).

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380

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Fig. 13 (continued).

of succession A in Mikonui-1 reflects the development of shelf depths following the seaward passage past this location of the shelf–slope break. The accumulation of bioclasts is possible during marked sea level rises in shelf settings with high rates of terrigenous deposition because the advancing shoreconnected wedge temporarily ties up much of the sediment being delivered to the basin (Abbott and Carter, 1994; Kamp and Naish, 1998). Most of the deposition on the outer shelf and slope occurs during times of stable or falling sea level. During the Pleistocene a narrow coastal plain underlain by glacial moraine, outwash and valley fill deposits formed between the modern coastline and the Alpine Fault. These are the molasse deposits of typical foreland basins.

of the topographic load of the Southern Alps (Kamp et al., 1992). The basin is out of sight beneath the Tasman Sea in a comparatively poorly seismically explored part of New Zealand, and, historically has been overlooked as an element of the collisional orogen; most research attention has been focused on the adjacent Alpine Fault and mountain belt. Indeed, that the modern coastline is so close to the Main Divide, a distance of only about 30 km, underscores a crucial feature about the basin: it is an active, underfilled peripheral foreland basin. In this section we attempt to draw together the essential elements of the tectonostratigraphic development of the basin to show that it has many of the characteristics of foreland basins, the best documented examples (e.g. Alpine) being those that are fossilised and have progressed through filled to overfilled depositional states.

5. Tectonostratigraphic evolution of the basin

5.1. Tectonic subsidence and geometrical development of the basin

The concept advanced here is that the South Westland Basin originated as a foreland basin due to loading of the leading edge of the Australia Plate by part

As noted above, for each of the three exploration wells the rates of total subsidence and tectonic

K.N. Sircombe, P.J.J. Kamp / Tectonophysics 300 (1998) 359–387 Fig. 14. Maps showing tectonic subsidence across the basin at the end of the Miocene (a) and at present (b). As the contribution of sediment loading to total subsidence has been removed, the maps reveal the subsidence component due to the tectonic driving force. Note how the amount of subsidence in both maps increases towards the mountain belt, which is the vicinity of the load causing the lithospheric flexure.

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subsidence increase dramatically between 5 and 13 Ma, depending upon proximity of the wells to the plate boundary (Fig. 5), and show similar subsidence patterns to those of wells in other foreland basins (Kominz and Bond, 1986; Cross, 1986). Fig. 14 shows two maps, one drawn for the top Miocene horizon and another for the present seafloor, in which the amounts of tectonic subsidence have been estimated. These maps are derivatives of Fig. 11a,c. They have been constructed by back-stripping the subsidence component caused by sediment loading up to the Late Miocene and the present. The amount of subsidence shown in Fig. 14 is the component due to the tectonic driving force. A limitation of these maps is that they have been calculated assuming Airy isostatic compensation, when it is known that the Australia Plate has some flexural rigidity, albeit weak. Nevertheless, Fig. 14 shows that at the end of the Miocene there was clearly tectonically driven subsidence along the southeast margin of the basin. It amounted to about 800 m, with about 1000 m at the northeastern end of the basin. Subsidence did not extend far into the basin as there was a prominent axial ridge at this time, and continental margin subsidence is evident further to the northwest. The proximity of the zone of obvious tectonic subsidence to the southeast margin of the basin suggests that up until 5 Ma there was limited loading of the Australia Plate by the Pacific Plate, and therefore limited topography and=or subcrustal load within the Southern Alps; a significant length of the plate boundary had probably only just become obliquely compressive. Comparison of Fig. 14a and Fig. 14b shows that most of the tectonic subsidence of the basin, and consequently the occurrence of loading within the collision zone, occurred during the Pliocene and Pleistocene. The contours of tectonic subsidence in Fig. 14b are colinear with the Alpine Fault and mountain belt. This spatial association together with the Plio–Pleistocene timing of most of the tectonic subsidence are key indicators that the Neogene origin and development of the basin were driven by the consequences of convergence within the wider Australia–Pacific plate boundary zone, as opposed to the strike-slip component of the total relative plate motion. Hence we conclude from Figs. 11 and 14 that the basin geometry and the pattern of tectonic subsidence that gave rise to it originated from load-

ing within the orogen, and therefore the basin can be classed as a foreland basin. 5.2. Stratigraphic architecture Previous studies have described how peripheral foreland basins evolve from an underfilled to a filled or overfilled depositional state (Graham et al., 1975; Homewood et al., 1986; Coakley and Watts, 1991). In ancient settings the degree of filling of a foreland basin is assessed from the long-term trends in the sedimentary facies of the basin fill (Sinclair, 1997). Deep-marine facies equate with underfilled, shallow-marine distal-continental facies equate with filled, and fully continental facies equate with overfilled basins. In South Westland the proximity of the shoreline to the mountain belt, the narrow coastal plain and the narrow continental shelf, together indicate that the basin is in an underfilled depositional state. This is perhaps anticipated by the fact that it lies within an active orogen. In a review of the stratigraphic and sedimentation patterns of ancient foreland basins, Sinclair (1997) highlighted a tripartite stratigraphic signature typical of foreland basins. It comprises three lithostratigraphic units, which are commonly diachronous and superimposed on top of one another during cratonward migration of the facies belts. (1) A lower unit that may or may not be underlain by a basal unconformity and comprises a variable thickness of shallow-marine limestone and sandstone. (2) A middle unit composed of mudstone commonly rich in pelagic microfauna. (3) An upper unit comprising sandstone dominated by turbidites and classically termed ‘flysch’. The origin of the character and migration of these facies belts cratonward across the basin can be understood in terms of the evolution of a collisional orogen. Initiation of the foreland basin occurs when a thrust wedge starts to advance towards the craton, and the previous tectonic setting starts to be overprinted. The first expression on the craton margin may be slow rates of uplift and erosion due to migration of the forebulge, forming a basal unconformity. This is followed by flexure-induced subsidence, and carbonate sedimentation on the margin of the forebulge in the absence of siliciclastic sediments because the eroding thrust wedge is distant. The subsequent accumulation of mudstones in

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outer shelf–slope depths reflects the generation of increased accommodation by flexure-induced subsidence and sufficient siliciclastic sedimentation to end carbonate accumulation. The flysch deposits accumulate in the axis of the basin upon and adjacent to the thrust wedge. The migration of the facies belts leading to their superposition is driven by cratonward advance of the thrust wedge as the collision zone widens, and accompanying retreat of the forebulge and basin axis. Ultimately the basin is infilled from the mountain belt side by molasse deposits and the inner margin of the basin becomes involved in the deformation and is inverted and eroded. The South Westland Basin shows this stratigraphic architecture and pattern of development. The Awarua Limestone is a thin carbonate unit corresponding to the lower carbonate unit in Sinclair’s (1997) tripartite subdivision. It is underlain by a regional unconformity. The middle unit corresponds to the mudstone of the lower part of the Tititira Formation as observed in outcrop (Nathan, 1978) and in the lower part of the three well successions (Fig. 4). The upper unit corresponds to the mass-emplaced flysch deposits of the upper part of the Tititira Formation (Nathan, 1978). These pass unconformably upwards into glacial moraines, glacial outwash and coastal plain deposits that represent the infilling of the basin, but this is very limited as shown by the narrow width of these deposits. The Late Pliocene and Pleistocene progradation of the shelf and slope documented above represents a transitional phase between the flysch deposition, which occurs on the slope, and as we have noted includes large-scale slumping, and the coastal plain–fluvial infilling of the basin. Although the stratigraphic architecture of the basin and its history of sedimentation are similar to the patterns in classical foreland basins, aspects of it may be co-incidental. For example, the unconformity beneath the Awarua Limestone may not have originated through uplift and erosion associated with migration of a forebulge. The uplift and erosion may have resulted from earlier (Early Miocene) formation of the Alpine Fault. Also, the extent of the Awarua Limestone suggests that its origin is not necessarily tied to carbonate deposition marginal to a forebulge; this may simply reflect the absence of a supply of siliciclastics (Kamp and Nelson, 1988).

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5.3. Basin inversion As noted earlier, the inner margin of the South Westland Basin has been inverted. Moreover, there is a pattern in the timing and degree of inversion: it started earlier in the north (10 Ma) compared with the south (5 Ma), and the amount of uplift and erosion has been greater in the north compared with the south (Kamp et al., 1992). This ties in with the pattern of basin infilling. The timing of the initiation of uplift in the north follows by about 3 million years the initiation of basin subsidence. In the south it follows by about 8 million years. In Sections 3 and 5.1 it was outlined how for a given horizon the foredeep was wider in the northern compared with the southern parts of the basin, which reflects the earlier and greater degrees of inversion and hence the location of the topographic load. The pattern of basin inversion also complements the pattern of infilling, with displacement of the locus of sedimentation further out into the basin through time, particularly during the Late Pliocene and Pleistocene. This is also expressed in the position and extent of the continental shelf at present (Fig. 2). We conclude that there is a profound unity in the timing and pattern of subsidence, infilling and deformation of the South Westland Basin, which carries the foreland basin signal. 5.4. Influence of oblique convergence and strike-slip displacement The Australia and Pacific plates are converging obliquely (DeMets et al., 1990), having evolved out of a continental transform with purely strike-slip displacement (Norris et al., 1978; Stock and Molnar, 1982; Kamp, 1986). However, in the classical ancient foreland basins the direction of convergence is usually understood to have been more orthogonal (e.g. Sinclair, 1997). The more oblique convergence in the Southern Alps Orogen has affected the evolution of the South Westland Basin in the following way. Because the Alpine Fault has accommodated most of the Neogene relative plate motion, it is a crustal structure that has persistently defined the northwestern flank of the high alps. Consequently, most of the surface and rock uplift in the mountain belt has been confined by this structure (Wellman, 1979; Kamp et al., 1989; Tippett and Kamp, 1993;

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Kamp and Tippett, 1993), which has also fundamentally determined the location of lower crustal thickening and subcrustal loading beneath and to the southeast of the high alps (Stern, 1995; Beaumont et al., 1996). This means that the topographic and subcrustal loads located within and beneath the mountain belt and inferred to have driven the flexural subsidence of the Australia Plate, have effectively been pinned to a particular surface trace (the Alpine Fault) throughout most of the basin formation. The degree of erosion of the rock section uplifted in the mountains is also implicated in limiting the amount of overthrusting of Pacific crust over the Australia Plate across the trace of the Alpine Fault (Norris et al., 1990). The net result is that there has been minimal migration of the thrust wedge (leading edge of the Pacific crust) into the basin. By comparison, in more orthogonal continental collision zones the thrust wedge can advance significantly into the basin, causing cratonward retreat of the forebulge. A feature that compounds the limited migration of loads in the Southern Alps is the occurrence of a thin cover sequence, which has therefore not been able to detach and form a fold thrust belt, which typically advances into the foreland basin and contributes to its subsidence. Nevertheless inversion of the leading edge of the Australia crust has occurred, probably by reverse faulting of the footwall of the Alpine Fault (Kamp et al., 1992). This deformation has been suggested by Beaumont et al. (1996) to be a consequence of the development of topography within the main part of the mountain belt and attempts by the orogen to adjust (flatten) the dip of the retro-shear zone (Alpine Fault) to maintain a balance between the gravitational and buoyancy forces within the mountain belt. The evidence for inversion of the inner margin of the South Westland Basin also means that the Alpine Fault Zone is strictly not the northwest margin of the mountain belt — an observation that needs to be addressed in any model, numerical and otherwise, of the collision zone.

6. Summary and conclusions This study has been concerned with documenting the geometry, seismic stratigraphy and aspects

of the development and history of the South Westland Basin within the context of the Southern Alps collision zone. It has been based chiefly on interpretation of the results of the mapping of five seismic reflection horizons across the basin. The main data sources were open-file industry petroleum reports and especially reflection profiles. The basin contains a Late Cretaceous–Cenozoic sedimentary and volcanic succession (Nathan et al., 1986), and has a polygenetic history. The study has concentrated on identifying the features and patterns that relate to the late Cenozoic development of a peripheral foreland basin within the Southern Alps collision zone. Geohistory analysis of the exploration hole successions show subsidence patterns typical of foreland basins. In Waiho-1 and Harihari-1, located along the southeast mountain front margin of the basin, subsidence started in the Middle Miocene and shows increases in rates at 5–6 Ma; in Mikonui-1, located over the axial high or forebulge, foreland basin subsidence overprinted extension-induced subsidence with þ D 1:25 at 5–6 Ma. Five seismic reflection horizons, including basement, Middle Miocene (top Awarua Limestone), top Miocene, mid-Pliocene (PPB) and mid-Pleistocene (PPA) have been mapped through the grid of seismic data. A structure contour map on basement (Fig. 11a) identifies a central axial high trending parallel to the mountain belt with a depocentre to the northwest that is part of the continental margin of the Challenger Plateau, and one to the southeast being the foreland basin. Back-stripped and decompacted structure contour maps drawn for each of the other horizons (Fig. 11) illustrate sequentially the location and magnitude of the basin depocentres and hence the geometrical evolution of the basin. By the Late Miocene a new axial depocentre located parallel to the Alpine Fault and in the vicinity of the modern coastline had developed. Subsidence of this depocentre was accentuated during the Pliocene and Pleistocene, when the extent of subsidence also widened. Modelling of the lithospheric flexure of the basement surface normal to the trend of the basin establishes values of 3.1 to 9:8 ð 1020 N m for the flexural rigidity of the Australia Plate. This is 1–2 orders of magnitude less than for the same plate beneath the Taranaki Basin (Holt and Stern, 1991). This could be due to some of the structure on the

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basement being inherited topography, or it could reflect the proximity of the foreland basin to a rifted continental margin. Isopach maps constructed for the stratigraphic section between mapped horizons show clearly the switch from pre-Middle Miocene sedimentation mainly along the continental margin to subsequent sedimentation mainly within a foreland depocentre parallel to the mountain belt. Through the Late Miocene to Present the loci of thickest sediment accumulation have moved northwestward within this depocentre, which reflects the progradation of a shelf–slope system into the basin. The South Westland Basin shows a stratigraphic architecture distinctive of ancient peripheral underfilled foreland basins. This comprises in order of a basal unconformity, a thin limestone (Awarua Limestone), a slope-depth mudstone and a submarine fan flysch sequence. The regressive part of the succession involves progradation of a shelf–slope complex followed by the accumulation of molasse deposits, which are confined mainly to the narrow coastal plain. This architecture results from a cycle of initial slow subsidence, which accelerates to create slope depths, which are then infilled by the increasing proximity of a denuding mountain belt, the topographic expression of the developing orogen. Sedimentation is terminated by uplift and erosion of the inner margin of the basin, which becomes involved in deformation at the margin of the mountain belt. This aspect of the basin evolution has been established from fission track thermochronological investigations (Kamp et al., 1992). The timing and degree of inversion fits well with the geometrical and stratigraphic development of the basin as documented here from seismic reflection and exploration hole data. The case for an origin of the South Westland Basin as a foreland basin marginal to the Southern Alps mountain belt, and resulting from lithospheric flexure as a consequence of loading within the orogen, is based upon the sum of many structural and stratigraphic features and their evolution as documented here. In particular, maps of tectonic subsidence (Fig. 14), where the influence of sediment loading is removed, identify the source of the loading as lying within or beneath the mountain belt.

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Acknowledgements We thank Simon Nathan for assistance with the acquisition of reflection profiles during the early stages of this project. P.J.J. Kamp acknowledges gratefully S. Cloetingh and P. Andriessen for the provision of facilities in the Institute of Earth Sciences, Free University of Amsterdam, during a visit in 1997 when this paper was written. We thank R. Stephenson, P. Andriessen and P. Roure for constructive reviews of the manuscript.

Appendix A A.1. Thermal subsidence Parameters used in the McKenzie (1978) type subsidence modelling of Mikonui-1, after values in Parsons and Sclater (1977) and Allen and Allen (1990): – Initial thickness of the lithosphere, 125 km – Density of the mantle, 3330 kg m 3 – Density of the crust, 2800 kg m 3 – Average bulk density of the sediment=water infilling rift, 1030 kg m 3 – Thermal expansion coefficient of crust and mantle, 3:28 ð 10 5 ºC 1 – Temperature of the asthenosphere, 1333ºC A.2. Flexural modelling The general flexural equation for a loaded plate and flexed plate, assuming boundary conditions and that the applied load is at the end of the plate and no horizontal forces are applied, is: d4 ! C .²m ²s /g! D 0 (A1) dx 4 where: D is the flexural rigidity of the plate; x is the horizontal scale; ! is the deflection of the plate; ²m is the density of the mantle; ²s is the bulk density of the basin sediment fill; g is the gravitational acceleration (9.8 m s 2 ). Ignoring the fundamental complexities of continental lithosphere and assuming elastic behaviour, the equation can be applied in continental situations (Allen and Allen, 1990). The flexural parameter, Þ (Walcott, 1970), is given by: ² ¦1=4 4D ÞD (A2) g.²m ²s / Assuming an unbroken lithosphere, and a known maximum deflection the profile of the deflection obeys: D

! D !0 e

x=Þ

.cos x=Þ C sin x=Þ/

(A3)

Alternatively the maximum deflection may also be calculated as: !0 D

V0 Þ 3 8D

(A4)

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where V0 is the line emplaced end load at x D 0. The distance from the line load to the highest part of the forebulge can be found as: xb D ³ Þ

(A5)

Thus if the distance of the forebulge is already known, the flexural parameter .Þ/ can be calculated, which in turn can be used to calculate the flexural rigidity .D/ and the line load .V /. If the lithosphere under the line load is broken, such as the case in rift zones and subduction, then the boundary conditions need to be modified (Walcott, 1970). If the maximum deflection is known, the deflection profile for a broken plate is given by: ! D 0 e

x=Þ

cos x=Þ

Maximum deflection can be found as: V0 Þ 3 !0 D 4D The distance to the crest of the forebulge is given by: 3³ Þ xb D 4

(A6)

(A7)

(A8)

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