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Geol. Soc. Australia Spec. Publ. 22, and Geol. Soc. America Spec. Pap. 372 (2003), 265–289 Mesozoic–Cenozoic evolution of Australia’s New Guinea margin in a west Pacific context K. C. HILL1* AND R. HALL2 1 2 3D-GEO, School of Earth Sciences, University of Melbourne, Vic. 3010, Australia. SE Asia Research Group, Department of Geology, Royal Holloway University of London, Egham, Surrey TW20 0EX, UK. The northern Australian margin includes the island of New Guinea, which records a complex structural and tectonic evolution, largely masked by Mio-Pliocene orogenesis and the Pleistocene onset of tectonic collapse. In the Palaeozoic, New Guinea contained the boundary between a Late Palaeozoic active margin in the east and a region of extension associated with Gondwana breakup along the western margin of Australia. In the Permian and Early Triassic, New Guinea was an active margin resulting in widespread Middle Triassic granite intrusions. The Mesozoic saw Triassic and Jurassic rifting followed by Cretaceous passive margin subsidence and renewed rifting in the Late Cretaceous and Paleocene. Since the Eocene, New Guinea tectonics have been driven by rapid northward movement of the Australian Plate and later sinistral oblique convergence with the Pacific Plate, resulting in Mio-Pliocene arc–continent collision. Neogene deformation along the margin, however, has been the result of direct interaction with the Philippine and Caroline Plates. Collision with the Philippine–Caroline Arc commenced in the Late Oligocene and orogenesis continues today. We suggest that the New Guinea Mobile Belt comprises a collision zone between a north-facing Cretaceous indented margin and a south-facing Palaeogene accretionary prism, subsequently cut by a Neogene strike-slip fault system with well over 1000 km sinistral displacement that has alternated between extension and compression. The change in character of the lithosphere in New Guinea, from thick and strong in the west to thin and weak north and east of the Tasman Line, was also an important influence on the style and location of Mesozoic and Cenozoic deformation. KEY WORDS Australia, collision zone, New Guinea, plate tectonics, wrench faults INTRODUCTION Australia’s northern margin, in New Guinea, is part of a complex convergent plate boundary between the IndoAustralian Plate, the Pacific Plate, and several other smaller plates including the Philippine Sea and Caroline Plates (Figure 1). Plate-tectonic reconstructions suggest that the relative motions between all of these plates must be considered when attempting to reconstruct the history of Australia’s New Guinea margin during the Cenozoic. These small plates did not exist prior to the Cenozoic and north of Australia there is a region with indications of a long subduction history so that much of the oceanic record of previous events has been lost. Reconstruction of the geological history relies on the geological record from outcrops in New Guinea, radiometric and palaeontological dating, regional palaeomagnetic data and regional plate-tectonic reconstructions. Rare basement outcrops suggest a fundamental difference between east and west New Guinea in the Palaeozoic, consistent with the differences between eastern and western Australia further to the south. During the Palaeozoic there was an important change from the eastern Tasman Orogenic belt to the region of extension associated with Gondwana breakup of an older stable craton along the western margin of Australia. Following Triassic–Jurassic rifting and breakup a passive margin developed in New Guinea which is preserved in the New Guinea Fold Belt (Figure 1). The Cenozoic was characterised by orogenic events, including ophiolite emplacement, arc–continent collision and the development of the major mountains known today. However, despite the importance and youth of these events, their timing and significance remain the subject of much debate. In this paper, we aim to synthesise knowledge of New Guinea geology with a plate-tectonic reconstruction of Australia’s northern margin for the Cenozoic. Furthermore, we consider the increased understanding of Palaeozoic and Mesozoic tectonics along Australia’s northern margin and its implications for the interpretation of Cenozoic plate motions and orogenesis. We suggest that the geology of Australia’s northern margin results from the interaction between the fabric of the Australian continent and the adjacent plate motions. Finally, we review some of the main tectonic issues and uncertainties in New Guinea still to be resolved. Thus, our paper is divided into four parts: (i) a review and discussion of constraints provided by the Palaeozoic, Mesozoic and Cenozoic geology and architecture of the northern Australia continental margin; (ii) a review and discussion of constraints provided by the Cenozoic tectonic history of New Guinea in the context of the plate-tectonic history of the west Pacific and southeast * Corresponding author: [email protected] 266 K. C. Hill and R. Hall Figure 1 (a) Tectonic map of New Guinea and the southwest Pacific region with principal geographical locations referred to in the text. Barbed lines are active subduction zones and thick lines are active spreading centres. The light blue shaded areas, drawn at the 200 m isobath, are the continental shelves of Eurasia and Australia and areas of thickened oceanic crust/arcs. (b) New Guinea showing simplified tectonic belts and the principal tectonic features. AB, Amanab Block; AR, Adelbert Ranges; BB, Bintuni Basin; BG, Bena Bena–Goroka Terrane; B–T, Bewani–Torricelli Mountains; CM, Cyclops Mountains; COB, Central Ophiolite Belt; DF, Derewo Fault; FR, Finisterre Ranges; G, Gauttier Terrane; GM, Grasberg mine; HG, Huon Gulf; HP, Huon Peninsula; In, Indenburg Inlier; K, Kubor Range; La, Landslip Ranges; LF, Lagaip Fault; MB, Meervlakte Basin; Po, Porgera Intrusive Complex and mine; RB, Ramu Basin; SB, Sepik Basin; SG, Strickland Gorge; ST, Sepik Terrane; Wa, Wandaman Peninsula; WT, Weyland Terrane. Tectonic evolution of New Guinea Asia; (iii) a tectonic synthesis of Australia’s northern margin integrating the plate-margin architecture and the plate-tectonic motions; and (iv) tectonic issues still to be resolved. The Cenozoic plate reconstructions are described in detail in Hall (2002), accompanied by computer animations, and are based on a synthesis of a range of data, from spreading histories obtained from small ocean basins to more qualitative information such as that obtained from terrestrial geology. The reconstructions were made with the ATLAS computer program (Cambridge Paleomap Services 1993) using the Indian–Atlantic hotspot frame of Müller et al. (1993). Full details of the approach and the model are provided in Hall (2002). The pre-Cenozoic reconstructions are based mainly on the geological data from New Guinea and the architecture of the margin, as discussed below. NEW GUINEA GEOLOGICAL ARCHITECTURE The island of New Guinea has been divided into four tectonic regions based on the Miocene to Holocene orogenesis affecting the northern part of the Australian Plate 267 (Figure 1) (Dow 1977; Hill et al. 1996). In the south is the Stable Platform, which is the northern continuation of Australian continental crust that preserves Mesozoic and possibly Palaeozoic extensional structures (Cole et al. 2000; Kendrick 2000). The adjacent Fold Belt, to the north, comprises the same crust deformed into fold and thrust structures following Late Oligocene–Miocene arc–continent collision. The fold and thrust deformation occurred in the Late Miocene to Holocene (Hill 1991; Hill & Raza 1999). Most of the northern half of New Guinea is made up of the Mobile Belt, including ophiolites of Mesozoic to Paleocene age (Davies & Jaques 1984; Rogerson et al. 1987). The Mobile Belt also includes distal Mesozoic–Tertiary sediments, abundant Miocene and some Cretaceous volcanic and intrusive igneous rocks, and medium- to high-grade metamorphic rocks with common Early Miocene and some Mesozoic cooling ages. In places, the Mobile Belt has been highly deformed, with a record of local medium- to highgrade metamorphic rocks that record ductile deformation, particularly along strike-parallel shear zones. These have been overprinted by low-angle thrusting and sinistral strikeslip structures (Crowhurst et al. 1996, 1997). The northern Figure 2 Simple chronostratigraphic column for the Fold Belt in New Guinea (after Kendrick 2000; Pigram & Panggabean 1984; Home et al. 1990; Parris 1994), showing the preserved Palaeozoic section west of the Tasman Line and the low-grade Permian to Early Triassic metasediments to the east, intruded by Middle Triassic and Early Jurassic granites. Following Jurassic breakup, the clastic Mesozoic section is relatively uniform. Early Tertiary uplift and denudation in Papua New Guinea led to a significant unconformity in the east prior to deposition of New Guinea Limestone, whilst there was continuous deposition in the west. The Late Miocene marks the onset of deposition of a molasse sequence. 268 K. C. Hill and R. Hall rim of New Guinea is interpreted to consist of accreted Palaeogene arcs and oceanic terranes including the Gauttier and Cyclops mountains, the Bewani–Torricelli mountains, the Finisterre Ranges and the Huon Peninsula (Pigram & Davies 1987). The suture between these terranes and the Mobile Belt is partly covered by the Miocene to Pliocene Meervlakte, Sepik and Ramu successor basins (Francis 1990). The successor basins are reported to be underlain by Mesozoic and Early Tertiary igneous and metamorphic rocks of basic to intermediate composition, inferred to be of island arc or oceanic origin (Doust 1990). Findlay (2003) and Findlay et al. (1997) review recent mapping of the Ramu–Markham part of the suture zone and suggest other interpretations. PROTEROZOIC HISTORY The Arafura Platform to the south of the Irian Jaya Fold Belt, comprises Mesoproterozoic cratonic basement (Pieters et al. 1983) overlain by Silurian–Devonian to Jurassic sedimentary sequences that record rifting of the Gondwana margin (Pigram & Panggabean 1981, 1984). To the north, the 5 km-high frontal ranges of the Fold Belt expose a thick, relatively undeformed uppermost Proterozoic and Palaeozoic sedimentary section that overlies Proterozoic marine rift metabasalts and metavolcanics of the Nerewip Formation exposed near the Grasberg Mine (Figures 1, 2) (Parris 1994). This Neoproterozoic sedimentary section may record erosion of the proposed Rodinian supercontinent from ca 725 Ma, when Laurentia separated from Australia (Powell 1996) giving rise to the Tasman Line, defined by the eastern limit of Proterozoic crust in Australia (Scheibner 1974). If so, then the Tasman Line trends roughly north–south near the Papua New Guinea – Irian Jaya border and then trends approximately west-northwest along the front of the ranges (Figure 1). The change in basement across the north–south part of the Tasman Line in Australia is consistent with tomographic images of the Australian region that indicate a marked change in the character of the lithosphere (Simons et al. 1999; Debayle & Kennett 2000; Hall & Spakman 2003). To the west of Queensland is lithosphere with high velocities beneath much of northern Australia which corresponds to old Precambrian crust, and a cold, very thick lithosphere with great strength. To the east of the western edge of Queensland are much lower lithosphere velocities consistent with younger, hotter and much weaker continental lithosphere east of the Tasman Line. PALAEOZOIC HISTORY The Palaeozoic margin in New Guinea is here interpreted to have been divided into two parts: an eastern active margin and a western rifted margin, with the boundary roughly along the Tasman Line. East of the line in Australia, rocks of the Tasman Orogen are thought to largely overlie Cambrian or Ordovician oceanic crust (Aitchison et al. 1992) and to have been accreted in multiple orogenies throughout the Palaeozoic and earliest Triassic (Veevers 2000). This east-facing edge of Gondwana is believed to Figure 3 Devonian–Carboniferous setting of the New Guinea margin (after Metcalfe 1996 figure 12). NG, New Guinea. have been a long-lived active margin, similar to the westfacing Rocky Mountains of North America in the Mesozoic and Palaeogene (Coney et al. 1990). The youngest Palaeozoic event along the eastern margin was the Permian New England Orogeny recorded in New South Wales and Queensland that continued into the Early Triassic (Korsch in press). Upper Permian to Lower Triassic low-grade metasediments intruded by Middle Triassic granites, recorded in Papua New Guinea suggest a similar history (Van Wyck & Williams 2002; Crowhurst 1999). In contrast the margin west of the Tasman Line underwent substantial extension and rifting, with major microcontinents breaking away in the Early Palaeozoic and the Permo-Carboniferous (Metcalfe 1996, 1998) (Figure 3). The thick, relatively undeformed Palaeozoic section in mainland Irian Jaya is consistent with this rift history (Figure 2). In contrast to mainland Irian Jaya, the widely exposed basement of the Kemum Terrane in the Bird’s Head comprises Silurian–Devonian low-grade metamorphic rocks intruded by Triassic granites. This has led to models suggesting that the Bird’s Head was derived from Papua New Guinea or eastern Australia (Pigram et al. 1985; Struckmeyer et al. 1993). It is suggested here that the Tasman Line continued in a more east–west direction through Irian Jaya and that to the north of it were rocks similar to those of the Tasman Orogen, which rifted off Australia during the Mesozoic. Therefore an alternative to deriving the Bird’s Head from eastern Australia is that during the Palaeozoic the Bird’s Head was north of the Tasman Line in a position relative to the Australian continent similar to the present. In the Papua New Guinea Fold Belt, basement crops out as Triassic granites in the Kubor Ranges and in the Tectonic evolution of New Guinea Strickland Gorge (Page 1976; Van Wyck & Williams 1998, 2002). The Kubor Granites intrude the low-grade Omung metasediments, which were deposited in the Late Permian to Early Triassic and contain Permian to Proterozoic detrital zircons showing affinity with Australia (Van Wyck & Williams 1998, 2002). Lying well to the northeast of the inferred Tasman Line, it is likely that east Papua is underlain by crust no older than late Precambrian. Crowhurst’s (1999) U–Pb age dating of zircons from the Amanab diorite yielded an Ordovician inherited zircon population, consistent with this hypothesis. However, recent U–Pb age dating of zircons in two young intrusive rocks in Papua New Guinea indicates the presence of Archaean grains. The Porgera intrusive complex in the western Papua New Guinea Fold Belt reveals a mixture of Late Miocene and Archaean zircons (Munroe & Williams 1996). Similarly, Pliocene and Archaean U–Pb ages were obtained from zircons from gneiss and granodiorite in the D’Entrecasteaux Islands of eastern Papua (Baldwin & Ireland 1995). These two datasets could indicate underlying Archaean terranes, but alternatively may demonstrate that the underlying Palaeozoic terrane remained close to Australia and received sediments derived from the Archaean of Gondwana. Northern limit of continental crust The present northern limit of intact Australian continental crust is important as it helps to define the terranes accreted during the Cenozoic. In addition, it is important to identify the terranes removed from the margin in the Cenozoic in order to determine the previous margin geometries. Davies et al. (1997) and Pigram and Davies (1987) suggested that in Papua New Guinea continental crust extends to the northern limit of the Fold Belt, defined by the Lagaip Fault and the Kubor Range (Figure 1), and that all terranes to the north were accreted. In contrast, Rogerson et al. (1987) and Simandjuntuk and Barber (1996) proposed that continental crust continues beneath the Mobile Belt as far north as the suture with the accreted arcs. Rogerson et al. (1987) suggested that the ophiolites exposed in the western Papua New Guinea Mobile Belt were transported there as thin nappes. In Irian Jaya continental crust is thought to underlie the Fold Belt as far north as the Derewo Fault, and is bounded to the north by the Central Ophiolite belt and the Mobile Belt (Figure 1). Two terranes constrain the interpretations regarding the limit of Mesozoic continental crust in Papua New Guinea. The first is the pre-Jurassic metasediments, minor metabasic rocks and synkinematic granites of the Bena Bena–Goroka terrane (Rogerson & Hilyard 1990) in the Mobile Belt to the northeast and east of the Kubor Range. Following Crowhurst (1999), the second, composite, terrane in the Mobile Belt is here defined to include the metabasics intruded by Permian–Triassic diorite in the Amanab block of northwest Papua New Guinea, the contiguous Landslip block to the south and adjacent Idenburg inlier across the border in Irian Jaya (Figure 1). Van Wyck and Williams (1998, 2002) used U–Pb dating of detrital zircons from the Goroka terrane to correlate it with the nearby Omung metasediments, i.e. deposited in the Permian to Early Triassic and derived from the 269 Australian continent. Further, they dated the Bena Bena terrane as Permian to Late Triassic and concluded that Australia’s cratonic margin extends at least as far north as the northern edge of the Bena Bena terrane, adjacent to the Finisterre Ranges (Figure 1). Crowhurst (1999) confirmed a Middle Triassic U–Pb age on zircons for the Amanab metadiorite, indicating that it intruded a Palaeozoic or older terrane, consistent with the Ordovician ages from inherited zircons in the metadiorite. They also reported 16 Nd–Sr analyses, which indicate the type of underlying crust and source of the intrusions, and concluded that the Amanab–Idenburg–Landslip terrane was probably a promontory of extended continental crust, possibly rafted in, adjacent to an embayment of oceanic crust to the east. Many tectonic models suggest that some continental crust has been moved along the New Guinea margin since the Early Miocene by strike-slip faulting (Hall 2002). In the island of Bacan, just southwest of Halmahera, there are high-grade continental metamorphic rocks including Barrovian kyanite–staurolite–garnet gneisses (Brouwer 1923; Hall et al. 1988) which yield very young (<<1 Ma) K–Ar and Ar–Ar ages (Malaihollo 1993; Malaihollo & Hall 1996). The chemistry of most of the present-day volcanic rocks of the Halmahera islands, including Bacan, indicates a typical intra-oceanic arc character but at the southern end of the arc there is evidence of contamination by underlying continental basement (Morris et al. 1983). Geochemical work on older volcanic rocks (Forde 1997) similarly shows an intra-oceanic arc character for all preNeogene volcanic rocks, but continental crustal contamination of ?Middle to Late Miocene igneous rocks at the south end of the Neogene arc, suggesting that Australian and Philippine Sea arc crust were brought into contact by Early to Middle Miocene collision and subsequent strikeslip movements. Pb, Nd, Sr and O isotope geochemical studies (Forde 1997) of Neogene volcanic rocks on Bacan show that some of the contamination could have been produced by Permian granites similar to those known from Queensland and Banggai–Sula. However, extreme Pb isotope ratios in some volcanic rocks suggest that an old Precambrian component is also present deeper beneath Bacan (Vroon et al. 1996). The Bacan rocks are within strands of the Sorong Fault zone, bound to the north and south by Mesozoic ophiolites with a Philippine Sea plate origin. They are clearly fault-bounded and must have moved from the east. Similar small fragments of continental crust are known in Obi (Figure 1). In Obi and in the Banggai–Sula islands there are also Mesozoic sediments of Australian margin character. These islands demonstrate the dispersal of continental crust fragments, undoubtedly derived from northern New Guinea. Basement fabric Davies (1991), Corbett (1994) and Hill et al. (1996), amongst others, recognised a northeast structural grain in New Guinea, here interpreted to be inherited from the fabric in the underlying crust. Davies (1991) recognised major north–south or northeast-trending alignments of volcanoes, stocks and normal faults cutting across the Fold Belt. Hill (1991) interpreted similar lineaments as lateral ramps compartmentalising the structures of the Papuan Fold Belt. 270 K. C. Hill and R. Hall Mesozoic margin, the microcontinents derived from northern Australia that are now in southeast Asia must also be considered. Some of these were separated during the Mesozoic rifting but others are thought to have been sliced off the margin during westward shear in the Cenozoic. Triassic orogenesis, igneous activity and rifting EARLY TRIASSIC Figure 4 Schematic Early to Middle Triassic palaeogeography (after Metcalfe 1996 and Charlton 2001). The dashed line through northern New Guinea is the inferred site of Late Triassic to Jurassic rifting. The part of New Guinea shaded red is the Stable Platform (see Figure 1). Both authors inferred the trends to result from underlying crustal or lithospheric features. Corbett (1994) defined northeast-trending transfer structures throughout the Papuan Fold Belt and related them to the copper–gold deposits. Hill et al. (1996) and Kendrick (2000) linked the north–south and northeast structural grain with a similar rectilinear pattern of shear zones in northern Australia recognisable from structural, regional gravity and magnetic mapping (Elliot 1994). Hill et al. (2002a) inferred that the northeast-trending lineaments continued across the Mobile Belt, limiting the amount of strike-slip motion that can be inferred between the Mobile Belt and Fold Belt. MESOZOIC AND PALEOCENE HISTORY The Mesozoic history of the New Guinea margin is well recorded in the sedimentary record. The relatively uniform stratigraphy in the New Guinea Fold Belt (Figure 2) records roughly synchronous rifting and subsidence of the margin. Widespread Middle Triassic igneous activity preceded Late Triassic and Early Jurassic rifting, and Mid-Jurassic breakup, although it is unclear which terrane, if any, separated from New Guinea. The Valanginian–Aptian sequence in New Guinea is mainly shale recording a flooding event as the margin subsided into deep water (Pigram & Panggabean 1984; Home et al. 1990). Similar mudstone deposition associated with subsidence and flooding is recorded along Australia’s North West Shelf to the Carnarvon Basin (Bradshaw et al. 1998), but not down Australia’s east coast (Norvick et al. 2001). Widespread volcanism occurred in the Aptian–Albian and Cenomanian, but there is little evidence of later Cretaceous igneous activity. Late Cretaceous rifting (Home et al. 1990) is inferred in the Gulf of Papua associated with uplift and erosion of southern Papua New Guinea. Pigram and Symonds (1991) inferred that rifting also occurred along the northern New Guinea margin with the possible formation of Late Cretaceous marginal basins. In order to reconstruct the Van Wyck and Williams (2002) demonstrated that the metasediments intruded by Middle Triassic granites were deposited and deformed in the Late Permian to Early Triassic, perhaps during an event correlating with the New England Orogeny along Australia’s eastern margin (Korsch in press). Metcalfe (1996) inferred subduction beneath the eastern Australia and Papua New Guinea margin to account for the orogenesis and volcanism. Early Triassic orogenesis is consistent with the stratigraphic section recorded in the fold belt of Irian Jaya (Kendrick 2000) (Figure 2). There the Palaeozoic section does not show evidence of Triassic deformation, but the Triassic section appears to be absent indicating non-deposition or regional uplift and erosion, recorded by an unconformity within the Tipuma Formation (Figure 2). MIDDLE TRIASSIC Middle Triassic igneous activity has been documented in Papua New Guinea by Page (1976), Crowhurst (1999) and Van Wyck and Williams (1998, 2002). As summarised by Charlton (2001), Late Permian–Triassic igneous activity is also recorded in the Bird’s Head of Irian Jaya and the nearby continental terranes of the greater Sula Spur, such as east Sulawesi and Banggai–Sula as well as Sibumasu (peninsula Myanmar, Thailand, western Malaysia and northwest Sumatra). The precise age of these intrusives is less well constrained than the Papua New Guinea examples. In New Guinea most of the Middle Triassic intrusive rocks are granitic to dioritic and are interpreted to result from the northwestern extension of a volcanic arc along Australia’s east coast related to subduction to the southwest beneath the continent (Figure 4) (Metcalfe 1996). We interpret the Permo-Triassic orogenic belt to have extended from eastern Australia to Papua New Guinea and along the north coast of New Guinea after Metcalfe (1996). LATE TRIASSIC Pigram and Panggabean (1984), Home et al. (1990) and Struckmeyer et al. (1993) inferred rifting in New Guinea in the Middle to Late Triassic, requiring a rapid change from compression to extension in the Middle Triassic. Crowhurst (1999) and Van Wyck and Williams (2002) have dated magmatic events in Papua New Guinea at 240 Ma and ca 220 Ma, consistent with previous dating of Page (1976). We infer that these correspond to a widespread Middle Triassic arc at ca 240 Ma, and Late Triassic extension-related volcanism at ca 220 Ma. Correlating these events with those in the New England Fold Belt, they suggest that Late Permian to Early Triassic orogenesis followed by Middle Triassic arc volcanism and Late Triassic extension and rifting may have Tectonic evolution of New Guinea 271 been common to Australia’s eastern margin and New Guinea. This is consistent with the interpretation of Stampfli and Borel (in press) that renewed rifting along the north Australian margin of Tethys occurred in the Late Triassic, due to subduction of the Palaeo-Tethys ridge beneath Asia. Monnier et al. (1999, 2000) suggested that the Central Ophiolite Belt in Irian Jaya was formed in a Jurassic backarc basin, although no new radiometric dating was presented, whereas ophiolites of the Cyclops Mountains were suggested to have formed in an Oligocene backarc basin, based on a small number of K–Ar ages. SUMMARY Mesozoic rifting An important issue for Australia’s northern margin is the location of the Sula Spur terranes, including the Bird’s Head. Struckmeyer et al. (1993), following Pigram and Panggabean (1984), inferred that, in the Triassic, these terranes lay to the northeast of Queensland in the position of the present Solomon Sea. Recent reflection seismic and drilling data in the Bintuni Basin, immediately south of the Bird’s Head, show a rifted Permian and Upper Palaeozoic section akin to that along Australia’s North West Shelf and in Irian Jaya (Sutriyono 1999). Thus it seems likely that the Sula terranes were adjacent to Irian Jaya as indicated by Metcalfe (1996) and Charlton (2001). Here we follow Metcalfe in placing the Sula terranes northwest of Irian Jaya, accounting for the lack of Triassic igneous rocks recorded in mainland Irian Jaya (Figure 4). The rift history recorded in the sediments of the Papuan Basin was well documented by Pigram and Panggabean (1984) and Home et al. (1990). Both recorded Mid–Late Triassic and Early–Mid Jurassic rifting events followed by breakup and formation of a passive margin. The effect of rifting and breakup in what is now the Mobile Belt has received little attention, mainly due to a paucity of data and uncertain terrane provenance. Hill et al. (in press) compared Australia’s northern margin in the Mesozoic to the extensional margin preserved along Australia’s North West Shelf (Chen et al. 2002). Hill et al. (in press) noted abrupt ~50 km offsets of the ophiolite belts along strike in New Guinea in compartments 125 or 250 km wide and inferred a rectilinear post-rift margin in the Mesozoic, with promontories of extended continental crust separated by embayments of newly formed oceanic crust. This interpretation is consistent with that of Crowhurst (1999) based on isotopic analyses of crust in the Amanab area, discussed above. The nature of the terranes rifted away from New Guinea in the Jurassic is unclear. Many authors show various southeast Asian terranes, such as those of parts of Sumatra, Sulawesi and the Banda arc, situated to the north of New Guinea in the Palaeozoic (Audley-Charles et al. 1988; Audley-Charles 1991; Metcalfe 1996). Struckmeyer et al. (1993) inferred separation of the Sula Spur terranes from northern Papua New Guinea and separation of the Bird’s Head from north Queensland in the Cretaceous. In contrast, Charlton (2001) shows no terranes north of New Guinea in the Triassic and the Sula Spur area lying to the west, in its present position, throughout the Palaeozoic. If the Middle Triassic igneous activity in Papua New Guinea was subduction related, the fact that the ca 240 Ma plutons are preserved at Amanab and in the Kubor Range indicates that only the relatively narrow forearc region can have separated. Allowing a wide arc–trench distance of ~300 km, the departing terranes are likely to have been a narrow strip of up to that width. If, as suggested, the Sula Spur terranes and others like West Sulawesi lay to the north of Irian Jaya in the Triassic (Metcalfe 1996) then some of those terranes would have rifted away in the Mid to Late Jurassic (Figure 4). The Bird’s Head is an exception. Hill et al. (2002) used the shallow water clastic Cretaceous sediments in that area to show that it was attached to northern Australia until the Paleocene. This is further supported by the recognition of Early Palaeozoic detrital zircons in the Paleocene Waripi section of Bintuni Bay (Sutriyono 1999) suggesting that the Bird’s Head was attached to the Australian continent. Ophiolites in New Guinea The ages and time(s) of emplacement of the ultramafic complexes interpreted as ophiolites in northern New Guinea are important in any reconstruction, but are not well known and are based largely upon the work completed by Bureau of Mineral Resource’s (now Geoscience Australia) geologists in the late 1970s and early 1980s. Davies and Jaques (1984) summarised the known geology of the three major ophiolite complexes in Papua New Guinea. Based on K–Ar dating of three gabbro and basalt samples, they inferred that the rocks of the Papuan Ultramafic Belt crystallised in the Jurassic and/or Cretaceous, but that two K–Ar dates on the underlying metamorphic rocks indicated emplacement in the Eocene. The Marum ophiolite yielded similar Early Jurassic and Paleocene ages (Page 1976) and is overlain by a sequence of Eocene argillites (Jaques 1981). Davies and Jaques (1984) interpreted it to be of Late Mesozoic or possibly Early Tertiary age and emplaced after the Middle Eocene but before the Late Oligocene–Early Miocene. Applying apatite fission track analysis to two samples, Hill and Raza (1999) confirmed a Palaeogene or older age for the Marum ultramafic suite and their analysis indicated that the rocks have not been heated to temperatures greater than 100°C since then. They also confirmed thrusting of the ophiolite at about 5 Ma. The April Ultramafics (Figure 1) were thought to be Mesozoic or older (Davies & Hutchison 1982) and to have been emplaced by tectonic stacking in an Eocene subduction system followed by further tectonism during Oligocene arc–continent collision (Davies & Jaques 1984). In contrast, Rogerson et al. (1987) considered the April Ultramafics to be of Early Palaeogene age, by analogy with the Papuan and Marum ultramafic suite. The ages of the ophiolites in Irian Jaya are even less well known. Pigram and Davies (1987) suggested that the ophiolites are Jurassic and Cretaceous and possibly lower Paleocene. Widespread Aptian–Albian volcanism Northern New Guinea was a passive margin through the Late Jurassic and Neocomian (Home et al. 1990; 272 K. C. Hill and R. Hall Struckmeyer et al. 1993), but in the mid-Cretaceous there was volcanic activity, as recorded in the Kondaku Tuff and Kumbruf Volcanics exposed around the Kubor Range. Hill and Gleadow’s (1990) apatite fission track analyses of midCretaceous clastic samples from the Papuan Basin showed common contemporaneous volcanic grains. Since then, zircon fission track analyses of mid-Cretaceous and younger formations across New Guinea have revealed common mid-Cretaceous grains, indicating a widespread and significant volcanic event (Sutriyono 1999; Hill & Raza 1999; Kendrick 2000). This event may have been related to the resumption of volcanism along Australia’s eastern margin (Norvick et al. 2001), which deposited vast amounts of Aptian–Albian volcanogenic sediments in the Surat Basin in Queensland (Exon & Senior 1976; Fielding et al. 1990). The volcanism has been interpreted to be due to subduction beneath Australia’s eastern margin (Jones & Veevers 1983; Elliot 1993; Norvick et al. 2001; Sutherland et al. 2001), but Francis (1990) reported alkaline magmatism associated with rifting in northern New Guinea, and Ewart et al. (1992) and Bryan et al. (1997) argued that the volcanic rocks of east Australia have a rift signature. In an extensive discussion of the origin of the mid-Cretaceous volcanics in New Guinea, Struckmeyer et al. (1993) also argued for a rift-related affinity. If the volcanic activity was associated with renewed extension, the cause may have been events that occurred at the northern edge of the Australian Plate between southeast Asia and the Pacific. For example, Müller et al. (2000) drew attention to a regional tectonic event between 100 and 90 Ma related to plate reorganisation between India and Australia at about 99 Ma which they postulated was due to the subduction of a Neo-Tethyan ridge beneath southeast Asia changing the stresses along Australian margins. Late Cretaceous – Paleocene rifting Home et al. (1990) inferred that rifting of the Coral Sea began in the Cenomanian and continued through the Late Cretaceous. An important coeval event is the widespread uplift and erosion of the southern Papuan Basin, inferred by Home et al. (1990) to be due to uplift of the rift-shoulder. However, similar uplift of the whole eastern margin of Australia occurred in the mid to Late Cretaceous, interpreted by Gurnis et al. (2000) to be due to rebound associated with the breaking off of a long-lived subducted slab. Francis (1990) argued that rifting occurred along the northeastern Australian margin in the Campanian–Maastrichtian, citing as evidence the alkaline basaltic volcanism along the outer shelf and upper slope and common tuffs in the Upper Cretaceous mudstone section. He inferred the formation of marginal basins, but of unknown distribution. This is similar to the interpretation of Pigram and Symonds (1991, 1993) and Davies et al. (1997) that Late Cretaceous rifting in the Coral Sea and along the northern margin led to the formation of marginal basins in the latest Cretaceous and Paleocene with the two systems linked by a transform fault across the western end of the Papuan Peninsula (Figure 1). Oceanic spreading of the Coral Sea occurred in the Paleocene to Eocene (Weissel & Watts 1979; Gaina et al. 1999) but the timing of formation of marginal basins along the north coast is less well known due to poor dating of the ophiolites that record the event, and a Neogene orogenic overprint. On the basis of the evidence for their ages discussed above it is possible some of the ophiolites represent marginal basins formed in the New Guinea Margin in the Late Mesozoic or Early Tertiary. In the Bird’s Head region of Irian Jaya, Brash et al. (1991) recorded an abrupt change in the stratigraphic section in the Paleocene, from a section which deepens to the west to one that deepens to the east towards Cenderawasih Bay. Hill et al. (2002) inferred that this was due to the formation of a Paleocene marginal basin floored by oceanic crust in what is now Cenderawasih Bay, at the same time as extension in other parts of northern New Guinea. The Late Cretaceous and Palaeogene geological history is particularly poorly known in New Guinea and several models are possible. The association of mid to Late Cretaceous volcanism, rifting, and subsequent oceanic spreading is a pattern that is recognisable from eastern Australia to New Guinea. This suggests some regional mechanism but it is not clear what this is. The north and east Australian margins have suffered several episodes of rifting of thin elongate slivers of crust over hundreds of millions of years. This is a feature of all the reconstructions made for the pre-Cenozoic and the best example is the Lord Howe Rise. Gaina et al. (2000) and Müller et al. (2001) suggested that rifting of the Tasman Sea was initiated by a ridge–hotspot interaction, but it is difficult to see how this mechanism could be applied to the north Australian margin, unless there were numerous small plumes. New models for the causes of the rifting are required. CENOZOIC HISTORY Sea-floor spreading in the Tasman and Coral Seas ceased in the Eocene (Gaina et al. 1998, 1999) and Australia began to move rapidly northwards. Ophiolites and arc terranes now found in northern New Guinea are here interpreted to have formed mainly in intra-Pacific oceanic arcs north of New Guinea and to have been accreted to the margin. Davies et al. (1997) have summarised these events, including an Early Eocene collision of the Bena Bena terrane, a Late Eocene to Early Oligocene collision of the Sepik terrane, Late Oligocene collision of a Papuan Peninsula terrane and Early Miocene collision of the Finisterre terrane, although there is disagreement about the nature and timing of all these events (cf. Abbott 1995; Cullen 1996). Eocene and Oligocene tectonic events in New Guinea are poorly constrained. Eocene sediments are largely lacking in the Papua New Guinea portion of the Fold Belt (Figure 2) but sandstone and limestone were deposited in the Irian Jaya Fold Belt. In the Papua New Guinea portion of the Mobile Belt (ST on Figure 1), Davies and Hutchison (1982) interpreted the variously metamorphosed Late Cretaceous to Eocene clastic and limestone strata to have been deposited in a volcanic island arc and trench environment. Around the Sepik Basin, Wilson et al. (1993) reported strongly recrystallised Middle to Upper Eocene limestones, consistent with those reported through much of the Papua New Guinea portion of the Mobile Belt (Davies 1983; Bain & McKenzie 1975). There have been few reports of Lower–Middle Oligocene strata, but Findlay (2003) Tectonic evolution of New Guinea infers that the continentally derived clastics adjacent to the Finisterre Ranges (Figure 1) may range from Eocene to Upper Miocene. In the Late Oligocene to Early Miocene there was widespread shallow-water carbonate deposition across the Platform and Fold Belt of New Guinea (Figure 2), reflecting regional subsidence, and also carbonate deposition in a low-energy environment around the Sepik Basin (Wilson et al. 1993). However, volcanism is indicated in parts of the Mobile Belt. Page (1976) and Rogerson et al (1987) have identified Late Oligocene to earliest Miocene gabbroic to dioritic intrusions in the Sepik Terrane (ST on Figure 1). Page (1976) and Findlay (2003) reported Late Oligocene to earliest Miocene K–Ar ages on a microsyenite and lava, respectively, from the Ramu River immediately south of the Finisterre Ranges (Figure 1). A further 40 km southwest, a quartz diorite yielded a Late Oligocene zircon fission track age (Hill & Raza 1999). The carbonates to the south of the Sepik Terrane and in and around the Sepik Basin to the north of the terrane indicate relatively little volcanogenic detritus. On the basis of mapping of limestone outcrops and interpretation of seismic data, the Sepik Basin (Figure 1) is reported to have been starved of clastic sediment during the Late Oligocene to Early Miocene (Wilson et al. 1993), suggesting limited igneous activity, at least in that area. Metamorphic rocks in the Mobile Belt of northern Papua New Guinea, have previously been dated by K–Ar as Late Oligocene to Early Miocene (Page 1976). Using 40Ar/39Ar analysis, Crowhurst (1999) redated many of the metamorphic rocks and determined that the samples cooled rapidly from temperatures above 300–500°C mainly between 20 and 18 Ma. Based partly on the presence of adjacent and coeval starved graben and limestone deposits in the Sepik Basin (Wilson et al. 1993; Hill et al. 1993), Crowhurst (1999) inferred that the metamorphic cooling was due to extensional unroofing of mid-crustal rocks as metamorphic core complexes. The unroofing was accompanied by ductile deformation in rocks of the Sepik Terrane (Figure 1) later overprinted by brittle low-angle thrust deformation in the Late Miocene and strike-slip deformation in the Pliocene (Crowhurst et al. 1996, 1997). The 20–18 Ma metamorphic cooling ages coincide with a possible temporary cessation of igneous activity as recorded by a gap in the K–Ar ages for that period (Rogerson et al. 1987). Carbonate deposition continued throughout the New Guinea Fold Belt in the Middle Miocene, but widespread igneous activity is reported in the Mobile Belt of Papua New Guinea, the Maramuni Arc (Figure 1) of Dow (1977). Although mainly Middle Miocene in age, Page (1976), Rogerson et al. (1987) and Hill and Raza (1999) reported that the igneous activity ranges from late Early Miocene to Late Miocene (ca 17–10 Ma) with minor activity continuing to the Pliocene. The principal products of the arc were subaerial pyroclastics and lavas, with thick, widespread volcanogenic sediments in addition to granodioritic and dioritic intrusions. The Maramuni Arc extends west only to the Irian Jaya border and there was insignificant volcanism in western New Guinea during the Neogene. The Neogene volcanic rocks from Irian Jaya that have been studied are Late Miocene and have an unusual chemical character that is ‘post-collisional’ quite different from Neogene magmatic 273 rocks elsewhere in New Guinea (Housh & McMahon 2000). The Middle Miocene igneous activity in Papua New Guinea has generally been interpreted as subductionrelated and most authors suggest southwest-dipping subduction beneath Papua (Hamilton 1979; Cullen & Pigott 1989; Francis 1990; Hill & Raza 1999) although some authors are uncertain of the polarity (Cullen 1996) or have suggested north-directed subduction (Abbott 1995). There have also been suggestions that volcanic activity of the Maramuni Arc was not subduction-related (Mason & MacDonald 1978; Johnson & Jaques 1980; Findlay et al. 1997) but was due to melting associated with uplift following Oligo-Miocene arc–continent collision. A doubly dipping subduction zone similar to that of the Molucca Sea region has been inferred to extend westward under New Guinea from the Solomon Sea, based on present seismicity (Cooper & Taylor 1987; Pegler et al. 1995). However, tomography does not show a long south-dipping slab (Hall & Spakman 2003). This implies little or no south-directed subduction or, alternatively, subduction of a slab within a few million years of its formation similar to that on the eastern side of the present-day Woodlark Basin. Beneath western New Guinea, there is a seismically poorly defined zone suggesting south-dipping subduction of about 100 km of ocean crust at the New Guinea Trench, but tomography shows no slab (Spakman & Bijwaard 1998; Hall & Spakman 2003). The widespread deformation of Miocene and older rocks throughout the Fold Belt and Mobile Belt of New Guinea demonstrates the Late Miocene to Pliocene orogenesis. Hill and Raza (1999) and Kendrick (2000) used apatite fission track analyses to infer the onset of compressional deformation in the Mobile Belt at the end of the Middle Miocene, around 14–12 Ma. They demonstrated that the orogeny propagated south into the Fold Belt in the Late Miocene to Pliocene, but suggested a transition to strike-slip deformation in the Pliocene. Kendrick (2000) recognised that in Irian Jaya the structural style of the 5 kmhigh frontal mountains was very different from the lowlying Fold Belt in Papua New Guinea. Hill et al. (in press) suggested that this was due to the orogeny impinging on the strong Australian lithosphere causing inversion of the large extensional fault along the continent edge that had been active from the Neoproterozoic to the Miocene. In contrast, Cloos et al. (1998) suggested that the crustal-scale uplift was due to collisional delamination following locking of a subduction zone to the north at ca 7 Ma, for which there is no evidence. Pliocene igneous activity is mainly manifested as localised stocks in the Fold Belt representing considerably smaller volumes of magma than during the Maramuni event. Davies (1991) pointed out that this igneous activity progressed southwards across the Fold Belt, coincident with the southwards propagation of deformation and uplift, suggesting that it was related to crustal thickening. Pleistocene igneous activity comprises stratovolcanoes in the New Guinea Fold Belt. Hamilton et al. (1983) carried out Sr and Nd analyses on six Pleistocene stratovolcanoes and concluded that they were mantle-derived magmas contaminated by movement through continental crust, but they could not exclude derivation from subducted crust or 274 K. C. Hill and R. Hall metasomatised mantle. In contrast, the magma generation has been attributed to partial melting of the lower crust (Mason & Heaslip 1980) or mantle (Johnson et al. 1978) as a consequence of crustal uplift and reduction of overburden pressure. At both ends of the Fold Belt, in the D’Entrecasteaux Islands in eastern Papua New Guinea and the Wandaman Peninsula in northwest Irian Jaya (Figure 1), medium- to high-grade metamorphic rocks are exposed at surface which have yielded Pliocene–Pleistocene cooling ages (Hill & Baldwin 1993; Baldwin et al. 1993; Bladon 1988). Hill and Baldwin (1993) interpreted the metamorphic rocks of the D’Entrecasteaux Islands to have resulted from rapid extension associated with the westward propagation of the Woodlark spreading ridge (Figure 1) (Taylor et al. 1995, 1999). The Pliocene–Pleistocene cooling ages reflect rapid exhumation of metamorphic core complexes (Baldwin et al. 1993). The Wandaman Peninsula is less thoroughly mapped, but radiometric dating indicates rapid cooling of mid-crustal rocks during the Pliocene–Pleistocene (Bladon 1988). This, too, is here interpreted to result from extension associated with orogenic collapse of the ends of the Fold Belt (Hill et al. 2002). Cenozoic tectonics and plate models It is generally agreed that there were several rapid changes in tectonic setting during the Cenozoic but there is still disagreement about the site of formation of ophiolites, how many arcs there were, the timing of arc–continent collision, polarity of subduction, the importance of extension and the role of strike-slip faulting. Constraining the geological interpretation of the New Guinea margin is only possible by considering its tectonic interactions with the plates to the north. It is of particular importance in interpreting New Guinea geology to consider not only Pacific–Australia plate vectors, but also relative motions between Australia and small plates such as the Philippine Sea and Caroline Plates. Rather than try to summarise the plethora of different plate-tectonic models, the different inferences that have been made, and reconstruct the thoughts of different authors from fragmentary interpretations shown in diagrammatic cross-sections and local plate models, we have tried below to identify the key features of different groups of models. There are few complete plate-tectonic models for this region, largely due to a paucity of data since the region between New Guinea and the Tonga Arc is vast, sparsely populated, difficult to access, covered in rainforest, and remote. The marginal basins of this region are largely undrilled, and there has been little marine investigation of many of them and their ages and spreading histories remain uncertain. Early accounts of New Guinea were based mainly on the work of Dutch geologists that was summarised by van Bemmelen (1949), and exploration studies for oil and minerals (Australian Petroleum Company 1961; Visser & Hermes 1962). These authors recognised the eugeosynclinal character of northern New Guinea to the north of young mountains at the northern edge of the Australian continent which van Bemmelen (1933, 1939) interpreted in a pre-plate-tectonic undation model. In contrast, Carey (1958) was one of the few to have a mobilist view of the region and emphasised the importance of strike-slip faulting: north New Guinea formed part of his Melanesian shear or Tethyan megashear. After plate tectonics became the dominant theory it was quickly recognised (Davies 1971; Curtis 1973; Packham 1973) that northern New Guinea included volcanic-arc rocks and these were interpreted as indicating arc–continent collision followed by subduction polarity reversal (Dewey & Bird 1970; Hamilton 1970). However, it was not long before the polarity reversal model was contested (Johnson 1976; Johnson & Jaques 1980). Plate-tectonic models that have been proposed since the late 1970s fall into three groups. First, there are plate models which consider the regional history of some or all of the major plates of southeast Asia, the western Pacific and Australia in varying detail (Crook & Belbin 1978; Hamilton 1979; Wells 1989; Jolivet et al. 1989; Rangin et al. 1990; Smith 1990; Daly et al. 1991; Yan & Kroenke 1993; Lee & Lawver 1995; Hall 1996, 1997, 1998, 2002). Second, there are local models which have been proposed to account for details of the evolution of sections of the Australian margin, for example in eastern New Guinea or the Melanesian arcs (Jaques & Robinson 1977; Johnson & Jaques 1980; Cooper & Taylor 1987; Pigram & Davies 1987; Hill & Hegarty 1987; Hill et al. 1993; Struckmeyer et al. 1993; Abbott 1995; Musgrave & Firth 1999; Hill & Raza 1999; Charlton 2000). Finally, there are many incomplete plate-tectonic models which are typically in the form of diagrams with cross-sections of subduction zones at different stages, or maps which show key features, but without portraying the entire history of the region (Johnson 1976; Falvey 1978; Falvey & Pritchard 1982; Ridgeway 1987; Richards et al. 1990; Benes et al. 1994; Petterson et al. 1997; Monnier et al. 1999; Weiler & Coe 2000; Findlay 2003). In the New Guinea region almost all models advocate northward-subduction of oceanic crust north of Australia before collision of the Australian margin with a south-facing arc. A few authors have proposed that the northern Australian margin was an active margin and there was southward-subduction beneath this margin before the active north Australia margin collided with an arc (Hill & Hegarty 1987; Hill et al. 1993; Monnier et al. 1999, 2000). Early models, and many authors since, suggested a single arc–continent collision. Variations include opening of marginal basins within the Australian margin with the marginal basins being destroyed by subsequent subduction, and multiple subduction zones. In some cases these subduction zones have a consistent polarity, for example, many models suggest more than one subduction zone dipping northwards north of Australia, but in other models there are subduction zones which have opposing polarities and which appear to have lived for short periods. The more complex models proposed, or implied, multiple arc–continent collision events. Pigram and Davies (1987) influenced many subsequent ideas on New Guinea when they proposed a terrane interpretation which implied a plate model involving accretion of fragments during subduction and collision. Subsequently, Daly et al. (1991), Lee and Lawver (1995) and many others have suggested the addition of arc fragments onto the northern New Guinea margin at several periods during the Cenozoic. However, even in the cases where the same arc is discussed there is considerable Tectonic evolution of New Guinea divergence of views about the timing of arc–continent collision. For example, ophiolites were emplaced in northern New Guinea in the Early Cenozoic but it is not clear if this was a single event or whether there were multiple obduction events at different places on the margin (Davies 1971; Hutchison 1975; Davies & Smith 1971; Pigram & Davies 1987; Davies et al. 1996; Monnier et al. 1999). Monnier et al. (2000) suggested that parts of the Central Ophiolite Belt in Irian Jaya are Jurassic backarc basin ophiolites. This implies that the ophiolites (approximately along-strike from the April, Marum and Papuan ophiolites) were formed within the Australian margin which is interpreted by most other workers as a passive margin. It is also widely reported that there was a Late Oligocene or Early Miocene collision of an arc with the northern New Guinea margin in the region of Papua although many authors have identified this collision as younger than Early Miocene and have suggested collisions which may continue up to the present day (Jaques & Robinson 1977; Abbott 1995; Davies et al. 1996; Hill & Raza 1999; Weiler & Coe 2000). A common theme in many of the models, particularly those which involve multiple subduction zones and multiple collisions, is the diachronous collision of a volcanic arc with the northern New Guinea margin (Hamilton 1979; Cooper & Taylor 1987; Richards et al. 1990; Weiler & Coe 2000). Almost all authors who consider a diachronous collision suggest that it started earliest in the west and progressed eastwards to where collision is apparently occurring between the New Britain – Finisterre arc system and the northern New Guinea margin. To the east of this is the Solomon Sea and some evidence has been interpreted to suggest that beneath the Papuan end of this segment of the collision zone is a doubly vergent subducted slab which dips to the north and south in similar fashion to the wellknown inverted U-shaped configuration of the Molucca Sea (Cooper & Taylor 1987; Pegler et al. 1995). It has been suggested that both north- and south-directed subduction have occurred between a volcanic arc and New Guinea and diachronous collision has progressively added the arc to the northern edge of New Guinea and closed the intervening ocean from west to east leaving the present Solomon Sea as a remnant of a wider ocean. Many of the regional models concerned with New Guinea terminate at Papua but because the relative motion between Australia and the Pacific involves a large westward component it is important to consider the Melanesian region east of Papua to determine where the arcs originated. The term Melanesian Arc is given to the system of arcs extending from New Ireland, through the Solomons, to Tonga (Figure 1). The Melanesian Arc originated in the Eocene either by rifting away the edge of the Australian margin above a subduction zone (Crook & Belbin 1978) or by initiation of subduction in an intra-oceanic setting (Yan & Kroenke 1993). Most models agree that from the Early Cenozoic there was south- or southwest-directed subduction of the Pacific beneath the Melanesian Arc and this continued until the Solomons collided with the Ontong Java Plateau. The arc–plateau collision took place in the Miocene but may have occurred over an extended period, with an early soft collision phase involving shortening of the Ontong Java Plateau, followed by a hard collision phase when the Solomons were transferred to the Pacific Plate 275 (Kroenke 1984; Petterson et al. 1997). Pacific–Australia plate convergence by subduction in the Solomons region then ceased and subduction transferred to other locations in the Early to Middle Miocene. It was at this stage that the plate boundary in the Solomons ceased to be a subduction zone and instead became a strike-slip fault zone. For New Guinea, the importance of this event is that the termination of subduction beneath the Solomons is suggested to have led to initiation of new subduction beneath the Papuan Peninsula to form the Maramuni Arc (Hill & Raza 1999). This southwest-dipping subduction zone was active from the Early Miocene but by the Late Miocene subduction probably ceased and instead new north-dipping subduction began beneath the New Hebrides and New Britain arcs (Hill & Raza 1999), followed by formation of the Woodlark Basin and Bismarck Sea (Taylor 1979; Benes et al. 1994; Taylor et al. 1995, 1999). TECTONIC MODEL FOR THE EVOLUTION OF AUSTRALIA’S NORTHERN MARGIN A regional plate reconstruction which includes the northern Australian margin and synthesises information from the ocean basins with terrestrial geology has been described by Hall (1997, 1998, 2002). Below, we use this reconstruction to provide the context to interpret the Cenozoic history of the New Guinea margin. It has features in common with many of the models outlined above, and proposes several arc–continent collisions, but differs in postulating a dominantly strike-slip plate-boundary zone in northern New Guinea during the Neogene. Before 45 Ma ophiolites were formed in the forearcs of intra-oceanic arcs in the Pacific region and these were emplaced by diachronous arc–continent collision with the north Australian passive margin between New Guinea and New Caledonia, probably in the Eocene. Following a major plate reorganisation at ca 45 Ma, Australia began to move rapidly northwards and the Philippine Sea, Caroline and Solomon Sea Plates formed as backarc basins. At about 25 Ma, collision of the Philippines–Halmahera Arc with the Australian margin and collision of the Ontong Java Plateau with the Melanesian Arc led to the loss of two major subduction zones. The arcs between the Philippines and Melanesia became a single arc system which rotated clockwise at the leading edge of the Pacific Plate, accommodated by intra-plate subduction at the eastern edge of the Philippine Sea Plate. This created a Neogene strike-slip fault system in northern New Guinea with major sinistral displacement of the arc relative to New Guinea. Below is our model for the tectonic evolution of Australia’s northern margin (New Guinea) incorporating all the information reviewed above and the plate-tectonic reconstruction. Palaeozoic Northern Australia was characterised by two crustal provinces, the Tasman Orogen in the east and a stable craton subjected to Gondwana rifting in the west. The Tasman Line, defining the eastern and northern limit of intact Proterozoic continental crust, trended north–south ~100 km west of the current Irian Jaya – Papua New Guinea bor- 276 K. C. Hill and R. Hall Figure 5 3D sketch of the New Guinea margin in the Jurassic, looking west. It is inferred that the strong Proterozoic crust west of the Tasman line results in a narrower zone of rifting and that the northeast structural grain influences the pattern of breakup, resulting in rectilinear promontories and embayments as observed on Australia’s North West Shelf. der and then west-northwest near the present mountain front and through the Bird’s Head (Figure 3). To the east and north were accreted terranes of the Tasman Orogen, including the northern Bird’s Head. The preservation of northeast-trending lineaments through to at least the northern limit of the Fold Belt, contiguous with those in Proterozoic crust of northeast Australia, may indicate that extended Proterozoic and older lithosphere underlies the accreted Tasman terranes. Triassic Along Australia’s North West Shelf, Gondwanan post-rift subsidence had created large Triassic depocentres whilst in southern Irian Jaya there was minimal Triassic deposition and perhaps uplift and erosion. In Papua New Guinea, and probably northern Irian Jaya, Upper Permian to Lower Triassic clastics (and unknown older rocks) of Australian affinity were deformed and metamorphosed in the Early Triassic in a continuation to the northwest of the New England Orogeny in eastern Australia. Orogenesis was probably related to subduction to the southwest beneath the northern and eastern Australia margins, that may have continued along the outer margin of New Guinea through the Bird’s Head and Sula Spur, accounting for the widespread Middle Triassic arc (Figure 4). In the Middle to Late Triassic, the tectonic regime abruptly changed so that the New Guinea margin and the New England Orogen to the southeast underwent rifting, associated with Late Triassic volcanism. Jurassic Mid–Late Triassic and Early–Mid Jurassic rifting in New Guinea was followed by Middle Jurassic breakup, but it is possible that only a relatively narrow Triassic arc and forearc region separated from the continent (Figure 4). We suggest that the underlying crustal architecture strongly influenced the style of rifting and breakup. West of the Tasman Line, rifting was confined to the north of the Mapenduma Fault, the extensional continent-bounding fault active since the Proterozoic (Figure 5). East of the Tasman Line extensional faulting was more widespread. In addition, the strong northeast structural grain influenced the pattern of breakup, creating promontories of extended continental crust separated by embayments of Jurassic oceanic crust. Early Cretaceous In the Late Jurassic and Early Cretaceous New Guinea was a passive margin. Thermal subsidence resulted in deposition of the shelf, shoreline and basin facies which now host significant hydrocarbon reserves, and hence are reasonably well known. In the Aptian, and particularly the Albian, there was widespread volcanism in New Guinea as all along the eastern Australia margin. We suggest that this was associated with renewed subduction beneath the margin, as inferred by Norvick et al. (2001) and Sutherland et al. (2001). The subduction beneath New Guinea would have caused the Jurassic oceanic crust in the embayments along the margin to have been deformed into accretionary prisms. Late Cretaceous and Paleocene In the Late Cretaceous, rifting was initiated along the northern margin of the Australian continent, perhaps associated with subduction (Figure 6). The rifting led to formation of marginal basins in the Late Cretaceous and Paleocene sep- Tectonic evolution of New Guinea 277 Figure 6 (a) Latest Cretaceous palaeogeography (after Hall 2002) and (b) a 3D sketch of the margin. Jurassic–Cretaceous subsidence led to deposition of a thick clastic sequence, including widespread Aptian–Albian volcaniclastics which may indicate subduction beneath the margin (Norvick et al. 2001) and formation of an accretionary prism in the oceanic embayments. Late Cretaceous–Paleocene rifting is inferred to have led to formation of marginal basins, including opening of the Coral Sea (Figure 1). Regional uplift may be associated with rifting (Home et al. 1990) or with dynamic topographic rebound related to movement over a Cretaceous subducted slab (Gurnis et al. 2000). 278 K. C. Hill and R. Hall Figure 7 3D sketch of the New Guinea margin in the Palaeogene (after Hall 2002). There was significant uplift and denudation in southern Papua New Guinea, but subsidence in Irian Jaya. Conceptual marginal basins to the north of New Guinea are unproven, but synchronous with Coral Sea spreading to the southeast (Davies et al. 1996, 1997). The existence of subduction beneath New Guinea is uncertain. arating slivers of extended continental crust from the margin. In the Paleocene, the Papuan Peninsula separated from Queensland during opening of the Coral Sea. Paleocene marginal basins probably formed along the north coast of New Guinea and Paleocene oceanic crust may have formed in Cenderawasih Bay adjacent to the Bird’s Head (Figure 6). In the Late Cretaceous and Paleocene the platform in southern Papua New Guinea was uplifted and eroded, previously suggested (Home et al. 1990) to have been associated with rifting in the Coral Sea (Figure 7). Alternatively, this may have been associated with dynamic topographic rebound associated with the movement of Australia over a Cretaceous subducted slab as suggested for the eastern Australia margin by Gurnis et al. (2000). Widespread Paleocene and Eocene sandstones in western Irian Jaya probably resulted from the uplift and denudation of Papua New Guinea and eastern Australia. Eocene By the Eocene, spreading in the Coral Sea and other marginal basins had ceased. During the Eocene, and possibly the Oligocene, some of the hot buoyant crust of the marginal basins, or intra-oceanic forearcs, was obducted diachronously along the margin, for instance the Papuan, Marum and parts of the April Ultramafics. Southern Papua New Guinea was still mainly emergent, but limestone and fine clastics were deposited in the Mobile Belt and northern part of the Fold Belt and New Guinea Limestone was deposited through northern Irian Jaya. At about 45 Ma there was a major plate reorganisation. Australia began to move rapidly northwards; new northdipping subduction began about 2000 km north of Australia beneath the Philippines–Halmahera Arc; new south-dipping subduction began northeast and east of Australia forming the Melanesian Arc (Figure 8). Between 45 and 25 Ma the Philippines–Halmahera Arc remained in approximately the same position. East of Australia, rollback of the subduction hinge led to the formation of a wide backarc Solomon Sea Basin as the Melanesian Arc rotated north. From about 40 Ma the Caroline Sea opened as a backarc basin by rollback of the Pacific subduction hinge at the eastern margin of the Philippine Sea Plate. The South Caroline Arc was located east of this backarc basin and was the site of formation of arc terranes now found in northern New Guinea. Oligocene As Australia moved north in the Oligocene, the residual rifted terranes and marginal basins to the north reached the Philippines–Halmahera–Caroline subduction zone. They were probably incorporated into the Philippines– Halmahera–Caroline subduction complex and accretionary prism, such that they now occur as basement beneath the northern parts of the Meervlakte, Sepik and Ramu successor basins and in the Weyland Terrane (Figure 1). At about 25 Ma there was an arc–continent collision when the Philippines–Halmahera Arc collided with a continental promontory along the Australian margin (Figure 9). We infer that this event was akin to the Pliocene arc–continent collision in Timor in that a fold and thrust belt was created in an island along the oceanic portion of the margin, with relatively little effect on most of the northern margin, other than Miocene subsidence and the local supply of coarse clastic sediment. Low-energy Late Oligocene to Early Miocene limestones were deposited southwest of the Tectonic evolution of New Guinea 279 Figure 8 (a) Eocene and (b) Oligocene palaeogeography of the New Guinea margin (after Hall 2002). The Eocene onset of convergence as Australia started moving rapidly to the north may have resulted in obduction of the marginal basins. Oceanic subduction over 1000 km to the north led to formation of the Philippine and Caroline Arcs and associated backarc basins. Subduction to the southeast beneath the Papuan Peninsula was associated with formation of the Solomon Sea Plate. G, Gauttier terrane, now in northern Irian Jaya (Figure 1). 280 K. C. Hill and R. Hall Figure 9 (a) Palaeogeography and (b) 3D sketch of the margin in the Late Oligocene. The Philippine–Caroline Arc has collided with a continental promontory along the margin creating an island orogeny as in Timor today. There was a major change in plate motions around this time such that the New Guinea margin changed from being convergent to a strike-slip-dominated margin. G, Gauttier terrane; BT, Bewani–Torricelli Mountains (Figure 1). Tectonic evolution of New Guinea 281 Figure 10 3D sketch of the New Guinea margin in the Early Miocene. There was significant subsidence with regional deposition of up to 1 km of Lower Miocene carbonates to the south and relatively starved basins to the north. The northern margin is inferred to have been a divergent strike-slip system, giving rise to local metamorphic core complexes with 20–18 Ma cooling ages (Crowhurst 1999) adjacent to starved basins. The Solomon Sea Plate was obliquely subducted to the west beneath the Papua New Guinea margin (Figures 9, 11). accreted arc, in the Sepik Basin. The origin of the Late Oligocene to earliest Miocene gabbros and diorites in the Sepik Terrane (Figure 1) is uncertain. 1–2 km of Miocene shallow-water limestone in southern New Guinea and an abrupt transition into deeper water along the northern margin of the Fold Belt. Early Miocene Middle Miocene The Oligocene arc–continent collision, combined with collision of the Ontong Java Plateau with the Melanesian Arc, led to the loss of two subduction zones and a major plate reorganisation. After these collisions, throughout the Miocene, the arcs between the Philippines and Melanesia became broadly a single arc system which rotated clockwise at the leading edge of the Pacific Plate, accommodated by intra-plate subduction at the eastern edge of the Philippine Sea Plate. The Philippine – South Caroline arc terranes, moved along the north Australian margin in a complex, regionally extensive strike-slip zone. There was therefore no significant subduction at the northern Australian margin in Irian Jaya. We infer that the sinistral strike-slip system in northern New Guinea had a significant extensional component and that parts of the extended continental promontories may have been rifted away and transported along the margin with the collided arcs. The extension also led to the formation of metamorphic core complexes that cooled rapidly from ~500°C to near-surface temperatures around 20 Ma (Figure 10). At the same time, the adjacent Sepik Basin was starved of sediment and associated with Early Miocene carbonate deposition. Regionally, the whole New Guinea margin subsided rapidly leading to widespread deposition of With the loss of the Melanesian subduction zone, Pacific–Australia convergence was accommodated by development of new subduction zones: southwest-dipping oblique subduction of the Solomon Sea Plate began first beneath Papua to form the Maramuni Arc (Figure 11). However, we cannot rule out the alternative, that the arc volcanism may have resulted from the ongoing oblique extension that gave rise to the metamorphic core complexes and North New Guinea Basins, accompanied by little Middle Miocene subduction beneath the eastern Papua margin, but by melting of a previously metasomatised mantle. Strike-slip movement continued in the Middle Miocene and the widespread volcanism of the Maramuni Arc filled the North New Guinea basins with volcanogenic sediments. Late Miocene to Pliocene When subduction beneath eastern Papua New Guinea ceased, the New Hebrides Trench propagated west to start new north-directed subduction beneath the Solomons and New Britain Arcs. The inferred cessation of subduction beneath eastern Papua New Guinea coincided with the onset of compression at around 14–12 Ma, causing deformation, uplift and denudation of the Mobile Belt in the Late 282 K. C. Hill and R. Hall Figure 11 (a) Middle Miocene palaeogeography (after Hall 2002) and (b) 3D sketch of the margin. Continued oblique subduction of the Solomon Sea Plate to the west has given rise to the Maramuni Arc in Papua New Guinea, although, alternatively, it could have been related to ongoing transtension. Volcaniclastic deposition was widespread in northern New Guinea, filling the previously starved basins. G, Gauttier terrane; BT, Bewani–Torricelli Mountains; AR, Adelbert Ranges; FR, Finisterre Ranges (Figure 1). Tectonic evolution of New Guinea 283 Figure 12 (a) Pliocene palaeogeography (after Hall 2002) and (b) 3D sketch of the margin. Convergence of the Caroline Arc with New Guinea from the end of the Middle Miocene created the New Guinea Orogen causing thrusting in the Mobile Belt and Fold Belt, but ongoing strike-slip motion between the Mobile Belt and the accreted arc. The Fold Belt was low-lying as it was built on warm, weak and broken lithosphere. Northward subduction of the Paleocene oceanic crust in Cenderawasih Bay beneath a continental fragment from northern Papua New Guinea caused it to collide with the Bird’s Head (Sutriyono 1999). BT, Bewani–Torricelli Mountains; AR, Adelbert Ranges; FR, Finisterre Ranges (Figure 1). 284 K. C. Hill and R. Hall Figure 13 3D sketch of the New Guinea margin in the Pleistocene. The orogeny impinged on the strong Proterozoic lithosphere west of the Tasman Line, which acted as a buttress causing crustal thrusting. This created 5 km-high mountains in Irian Jaya and an adjacent foreland basin. During the Pliocene–Pleistocene, local extensional reactivation of northeast-trending lineaments allowed mantlederived volcanism, including stocks associated with gold mineralisation, e.g. Porgera and Grasberg (Hill et al. 2003). The Fold Belt in Papua New Guinea remained low with no significant foreland basin. Compression abated around 3 Ma and strike-slip faulting became more dominant (Crowhurst et al. 1997). MA, Muller Anticline. Miocene. We suggest that the continued oblique plate convergence was largely partitioned into a sinistral strike-slip fault north of the New Guinea basins and southwestdirected thrusting to the south. Thrust faulting in the Mobile Belt included complex juxtaposition of the ophiolites, accretionary prisms and metamorphic rocks with the Mesozoic–Palaeogene distal sediments and Maramuni Arc intrusive, volcanic and sedimentary rocks. In the Late Miocene to Pliocene, the deformation propagated south to the fold and thrust belt (Figure 12). Throughout New Guinea this fold and thrust belt is interpreted to have been low-lying and lacking a foreland basin because the fold belt overlay weak, hot, broken and extended continental crust that subsided beneath it. When convergence recommenced in west Irian Jaya, we infer that a continental sliver, possibly rifted from a northern Papua New Guinea promontory, was adjacent to the oceanic crust in Cenderawasih Bay. Subduction of the oceanic material is interpreted to have drawn this (Weyland) terrane south to collide with the Bird’s Head forming the Lengguru Fold Belt. Pliocene to Holocene In Irian Jaya the orogenesis impinged on the Mapenduma Fault marking the northern limit of strong, thick, cold, Proterozoic Australian lithosphere. The fault was inverted and a 15 km thick slab of the crust was thrust to the surface building the 5 km-high mountains in the Irian Jaya Fold Belt and creating a substantial foreland basin to the southwest (Figure 13). East of the Tasman Line the orogenesis continued to produce a low-lying fold belt overlying weak lithosphere. As the crust was now in compression, few mag- mas were emplaced, except in local areas of dilation at the intersection of old extensional faults with north-northeasttrending fracture zones. This allowed the emplacement of Cu±Au-bearing magmas from deep in the crust or mantle (Grasberg and Porgera on Figure 13). During the Late Pliocene, compression probably waned, but did not cease, and strike-slip motion became dominant in the Mobile Belt and Irian Jaya Fold Belt. To the east of New Guinea, slab-pull forces at the New Britain Trench led to the formation of the Woodlark spreading centre at the former Solomon Sea transform (Figure 12). Very young lithosphere of the Woodlark Basin is now being subducted at the South Solomon Trench. Throughout this period, west-dipping subduction of the Pacific Plate continued at the Tonga–Kermadec Trench with rapid rollback of the subduction hinge in the last 10 million years. In the Bird’s Head area, we suggest that the Mobile Belt adjacent to the Lengguru Fold Belt collapsed towards the northeast, giving rise to the 2 Ma metamorphic core complex exposed in the Wandamen Peninsula and leaving the present Cenderawasih Bay floored by extended continental crust. The Pleistocene to Holocene stratovolcanoes in the Fold Belt may be related to extension associated with the Woodlark and Cenderawasih orogenic collapse. However, their confinement to the Fold Belt suggests an underlying feature such as the south-dipping Solomon Sea slab suggested by Davies (1991). DISCUSSION Here, we note the rapid changes in tectonic setting associated with the development of New Guinea during the Tectonic evolution of New Guinea Cenozoic. New Guinea cannot be understood simply as the product of an arc–continent collision, or the result of Australia–Pacific plate interaction, although both are part of the story. Some of the conflicts between interpretations of New Guinea geology reflect the difficulties of the terrain, and an inadequate knowledge of this large region, but some probably reflect the complexity of the history of minor plate motions, the differing character of the lithosphere within the orogenic belt, and its unpredictable behaviour during both contraction and extension. It is clear from outcrop data and tomographic studies that strong, cold Proterozoic lithosphere underlies southwestern New Guinea and that this had a significant influence on Phanerozoic tectonics. It is likely that the northeast- and northwest-trending Proterozoic structural grain influenced Jurassic rifting and breakup, resulting in a rectilinear margin of extended continental promontories separated by embayments of Jurassic oceanic crust. Furthermore, the extensional faulting associated with rifting appears to be confined to the area north and east of the strong Proterozoic lithosphere, reactivating a continent-bounding extensional fault that had been active since the Proterozoic. To the east, in the accreted Palaeozoic terranes of the Tasman Orogen, the extension was more widespread. Finally, during orogenesis, the strong Proterozoic litho-sphere had a profound effect when the orogeny impinged on it. This created a foreland basin along the New Guinea margin in the Neogene and created mountains over 5 km high. In contrast, the area overlying the accreted Tasman terranes remained a relatively low-lying fold belt with no adjacent foreland basin. The change in character of the lithosphere at the interpreted position of the Tasman Line is also associated with other tectonic features. From about 45 to 25 Ma this was the position where there was a change in subduction polarity from north-dipping north of New Guinea to south-dipping in the Melanesian Arc. There is a strong positive lower mantle anomaly of uncertain origin seen on tomographic images below Queensland (Hall & Spakman 2003). Together, these observations suggest the Tasman Line is the site of significant changes in the properties of the mantle and crust which influenced Cenozoic tectonic history. An important feature of the Cenozoic plate-tectonic reconstruction is the persistent presence of a sinistral strike-slip fault system in northern New Guinea throughout the Neogene at the same time as the Philippine–Caroline Plate rotated clockwise along the margin. These were the dominant features controlling Neogene orogenesis in New Guinea. We infer that the strike-slip system was transtensional in the Early Miocene resulting in metamorphic core complexes and graben formation and possibly contributing to volcanic activity of the Maramuni Arc. Towards the end of the Middle Miocene, the strike-slip system became transpressional resulting in ~100–200 km of shortening along the New Guinea margin. The fault system was probably partitioned into a strike-slip fault along the northern margin, and fold and thrust structures in the Mobile Belt and Fold Belt to the south. In the Pliocene, as convergence was diminished, strike-slip motion became dominant throughout the Mobile Belt and in the Irian Jaya Fold Belt. At the same time the areas immediately east and west of mainland New Guinea underwent extensional collapse exposing Pleistocene metamorphic core complexes. 285 There is a substantial divergence in the interpretation of Neogene tectonics between this paper and Findlay (2003) which largely reflects the very different approach taken to address the problem. Here, we attempt an analysis on a plate-margin scale incorporating as much regional and local geological data as possible and inevitably there are discrepancies that need testing. In contrast Findlay (2003) and others in the Geological Survey of Papua New Guinea have carried out important new mapping in the area between the Bena Bena – Goroka Terrane and the Finisterre Ranges. Citing observed interfingering relationships and Oligocene K–Ar dates as critical data, Findlay (2003) concludes that the Finisterre Ranges were close to their present location at the end of the Oligocene and that there has been no significant motion between the Bena Bena – Goroka Terrane and the Finisterre Ranges since then, ruling out >1000 km of sinistral offset, as inferred in the model presented here. We acknowledge the interfingering of continental and arc detritus established by Findlay (2003) and the significant thrusting of the Finisterre Range to the south in the last one million years. However, we suggest that prior to that thrusting, there was a major sinistral strike-slip fault between the Finisterre Terrane and New Guinea and that arc detritus could have been derived from any of the Philippine – Caroline Arc fragments adjacent to the margin. It is worth noting that Liu and Crook (2001) mapped the same area as Findlay (2002) and studied sediments in the adjacent Huon Gulf and inferred initial collision of the Finisterre Terrane with New Guinea in the latest Middle Miocene to earliest Late Miocene. Furthermore they estimated propagation of the collision zone to the east at 39–55 km/106 y. Our analysis suggests that the New Guinea Mobile Belt comprises a collision zone between a north-facing Australian indented continental margin and a south-facing Mesozoic and Palaeogene accretionary prism in front of an arc, all subsequently cut by a Neogene strike-slip fault system with more than 1000 km sinistral displacement. This model implies the consumption of most Cretaceous and Palaeogene terranes that may have existed to the north of New Guinea. Although integration with plate-kinematic syntheses greatly increases the reliability of the model, it still relies heavily on the onshore geological record, particularly prior to the Eocene. The dataset is far from complete and considerable new fieldwork and isotopic and palaeontological dating need to be done. Some of the main issues and uncertainties that need to be addressed to test the model are listed below. Major issues and uncertainties to be tested (1) The interpretation of a west-northwest-trending Tasman line through Irian Jaya relies on unpublished dating from Parris (1994). The age and nature of ?Cambrian oceanic crust cropping out in the Irian Jaya Fold Belt should be confirmed. (2) Our model assumes a relatively fixed position of the Bird’s Head relative to Australia. This can in part be tested by studies of the relationship between the Bird’s Head and North West Shelf sediments to the south following the current acquisition of considerable new, high-quality reflection-seismic data due to the very large gas discoveries in Bintuni Bay in the late 1990s. It can also be tested by fur- 286 K. C. Hill and R. Hall ther provenance studies of the Mesozoic sandstones in the Bird’s Head region. (3) Whilst substantial strike-slip motion is inferred in northern New Guinea, our model infers minimal strike-slip motion between the Fold Belt and the Mobile Belt, in part based on northeast-trending lineaments cutting across the boundary, due to inferred continental promontories (Hill et al. 2003). This contrasts strongly with previous models and should be tested using remote sensing images and further isotopic analysis in the Mobile Belt. (4) A considerable portion of our New Guinea tectonic model relies on the dating of the ophiolites, particularly interpretations of Paleocene marginal basins. New dating and geochemical analysis of the ophiolites would improve all models. (5) The arc or rift nature of Triassic and Permian plutons is unknown and could be tested geochemically. (6) Similarly, the chemical character of the Maramuni Arc magmatic activity is almost unknown and a subduction-related interpretation is not supported by tomography. Geochemical studies may be able to differentiate between interpretations of subduction-related, transtension-related, or post-collision magmatism. (7) Substantial Neogene strike-slip motion is here inferred between the Finisterre volcanics and the sediments of the Markham Basin to the south, overprinted by rapid Pleistocene convergence and several kilometres of uplift. This should be further tested by field mapping and dating and structural restoration of the Finisterre Ranges. 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