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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]
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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
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(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.
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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.
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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.
(8) New models need to be developed to explain
episodes of rifting of the north and east Australian margins
to produce thin elongate slivers of crust several times during hundreds of millions of years.
ACKNOWLEDGEMENTS
This paper greatly benefited from an early informal review
by Martin Norvick and helpful and constructive refereeing
by Heike Struckmeyer, Bob Findlay and Carmen Gaina
with additional comments from Dietmar Müller. The
research over many years that has contributed to this
review has been funded by many oil and mining companies
who are gratefully acknowledged, particularly in their support of PhD theses by Crowhurst, Kendrick and Sutriyono.
Hill was in part supported by the Australian Geodynamics
Cooperative Research Centre. Financial support to R. Hall
has also been provided by NERC, the Royal Society, the
London University Central Research Fund, and the Royal
Holloway SE Asia Research Group.
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Received 27 August 2001; accepted 11 June 2002