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AUTHORS
Andrew Quarles van Ufford Depart-ment of Geological Sciences, University ofTexas at Austin, Austin, Texas 78712
Andrew Quarles van Ufford earned a B.A. de-gree in geology from Haverford College, Hav-erford, Pennsylvania, in 1989 and a Ph.D. fromthe University of Texas at Austin in 1996. Heworked as a geologist on the Asia-Pacific explo-ration for ARCO until 2000. He obtained anM.B.A. degree from Northwestern Universityin 2002 and has since been manager of planningat Pioneer Natural Resources, U.S.A. in Irving,Texas.
Mark Cloos Department of GeologicalSciences, University of Texas at Austin, Austin,Texas 78712; [email protected]
Mark Cloos earned a B.S. degree in geologyfrom the University of Illinois, Champaign-Urbana (1976) and a Ph.D. from the Universityof California-Los Angeles (1981). He joined thefaculty at the University of Texas at Austin in 1981as a structural geologist and is now professorand Getty Oil Company Centennial Chair. Hisresearch interests involve all aspects of thegeology of convergent plate margins.
ACKNOWLEDGEMENTS
We thank James R. Moffett of Freeport McMoRan,Inc., whose idea and support made the ErtsbergProject possible. Dave Potter, Steve Van Nort,Dave Mayes, Tom Collinson, Mark Gilliam, GaryOConnor, Kris Hefton, Jay Pennington, KeithParris, Bambang Trisetyo, Peter Sedgwick, ImantsKavalieris, and Art Ona provided discussions anddirect assistance. Special thanks to AmeliusBeanal, Julianus Magal, Etinus Tabuni, Domin-ikus Mom, Tiranus Beanal, Benny Dolame, andthe Tembagapura helicopter operations crewfor assistance during fieldwork. We also thankour University of Texas colleagues Robert E.Boyer, William R. Muehlberger, Sharon Mosher,Rich Weiland, Stefan Boettcher, Paul Warren,Benyamin Sapiie, Eric Beam, Tim McMahon,Eric James, and Stacey Tyburski for discussionsand assistance. Reviews by Eli Silver, W. R.Dickinson, E. A. Mancini, and A. Tripathy aregreatly appreciated. This is Ertsberg ProjectContribution No. 21.
Cenozoic tectonics ofNew GuineaAndrew Quarles van Ufford and Mark Cloos
ABSTRACT
Major hydrocarbon discoveries have been made in eastern and
westernmost New Guinea, and there is great potential for additional
discoveries. Although the island is a type locality for arc-continent
collision during the Cenozoic, the age, number, and plate kine-
matics of the events that formed the island are vigorously argued.
The northern part of the island is underlain by rocks with oceanic
island arc affinities, and the southern part is underlain by the Aus-
tralian continental crust. Based on regional sedimentation patterns,
it is argued herein that the Cenozoic tectonic history of the island
involves two distinct collisional orogenic events.
The first Cenozoic event, the Peninsular orogeny of Oligocene
age (3530 Ma), was restricted to easternmost New Guinea.Emergent uplifts that shed abundant detritus resulted from the
subduction of the northeastern corner of the Australian continent
beneath part of the Inner Melanesian arc. This collision uplifted the
Papuan ophiolite and formed the associated mountainous uplift
that was the primary source of siliciclastic sediments that largely
filled the Aure trough. Between the Oligocene and Miocene, the
paleogeography of the region was similar to present-day New Cale-
donia. The continental crust under central and western New Guinea
remained a passive margin.
The second event, the Central Range orogeny, began in the
latest middle Miocene, when the bulldozing of Australian passive-
margin strata first created emergent uplifts above a north-dipping
subduction zone beneath the western part of the Outer Melanesian
arc. The cessation of carbonate shelf sedimentation and widespread
initiation of siliciclastic sedimentation on top of the Australian con-
tinental basement is dated at about 12 Ma. This collision emplaced
the Irian ophiolite and created the present mountainous topography
forming the spine of the island.
AAPG Bulletin, v. 89, no. 1 (January 2005), pp. 119140 119
Copyright #2005. The American Association of Petroleum Geologists. All rights reserved.
Manuscript received July 10, 2003; provisional acceptance October 8, 2003; revised manuscript receivedAugust 11, 2004; final acceptance August 30, 2004.
DOI:10.1306/08300403073
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INTRODUCTION
New Guinea is a type locality of island arc-continent
collision during the Cenozoic (Dewey and Bird, 1970).
The northern half of the island is underlain by a crys-
talline basement of ocean crust with arc affinities de-
rived from the floor of the Pacific basin (Figure 1). The
southern half of the island is composed of passive-
margin strata overlying the Australian continental base-
ment. Debate exists, however, regarding the number
and timing of the events that created the Central Range,
the approximately 1300-km (800-mi)-long mountain-
ous spine of the island.
Significant oil and gas accumulations have been
discovered in eastern and westernmost New Guinea,
and the region has significant potential (McLennan
et al., 1990). Several significant hydrocarbon accumu-
lations, including the giant Hides gas field (5 tcf ), have
already been developed in the fold and thrust belt of
Papua New Guinea (Figure 1) (Carman and Carman,
1990, 1993). The highlands in the Indonesian half of
the island are much less explored. In westernmost
New Guinea, significant oil production has come from
the Salawati and Bintuni basins (Katili, 1986, 1991).
The early 1990s discovery of the Tangguh gas field
(14+ tcf ) in the Bintuni basin proves that at leastone supergiant gas accumulation is present. All of the
hydrocarbon discoveries known to us are in structures
produced during Cenozoic tectonism.
The outline of the island of New Guinea has been
described as similar to a bird flying westward (Figure 1).
As a result, the island is commonly geographically divided
Figure 1. Tectonic map of New Guinea, modified from Hamilton (1979), Cooper and Taylor (1987), and Sapiie et al. (1999).Spreading centers northwest and southeast of New Guinea are slow (
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into the Birds Head and Tail regions. The central
portion of New Guinea, the Birds Body, can be divided
into four lithotectonic provinces from south to north:
the foreland basin, the Central Range fold and thrust
belt, a metamorphic belt with an overlying ophiolite
complex, and an accreted oceanic arc complex.
Tectonic models for the origin of the Central
Range have been based primarily on field relationships
on the eastern half of the island, the nation of Papua
New Guinea. In this paper, new biostratigraphic infor-
mation is reported for the Cenozoic strata near the
Ertsberg (Gunung Bijih) mining district, which is lo-
cated in western New Guinea near the Puncak Jaya
(labeled with box on Figure 1, 4884 m [16,023 ft]),
the highest part of the Central Range of Papua (for-
merly Irian Jaya), a province of Indonesia. This infor-
mation is integrated with published stratigraphic studies
and other geologic data from across New Guinea to
identify the major Cenozoic orogenic events that formed
the island.
Tectonic Models
The number and timing of Cenozoic orogenic events
in New Guinea has been debated. In part, this is be-
cause the geologic history of the area is complex, be-
cause it is located near the junction of the Pacific,
Australian,and Philippine plates. The primary obser-
vations for which the plate interaction models must
account are the timing and locations of arc volcanism,
patterns of deformation, and the ages of regional meta-
morphism, as well as type and thickness of sedimentation.
Many theories exist for the Cenozoic tectonic
evolution of New Guinea. Based on geologic relation-
ships in eastern New Guinea, Dow et al. (1972) and
Dow (1977) concluded that there was evidence for two
distinct orogenic events. To explain the origin of the
New Guinea trench, the bathymetric depression north
of the island, and volcanism in the highlands, Hamil-
ton (1979) proposed that the island is the result of the
collision of the Australian continent with a south-facing
arc in the early Miocene, followed by subduction re-
versal in the middle Miocene (Figure 2A). In a regional
tectonic synthesis, Kroenke (1984) proposed that there
were two major arc-continent collisions, and that the
New Guinea trench is a recently reactivated relict of
an older oceanic subduction zone (Figure 2B). Milsom
(1985) proposed that an Eocene collision was followed
by subduction reversal in the early Miocene to form the
New Guinea trench, which changed into left-lateral
transform faulting in the late Miocene (Figure 2C).
Cooper and Taylor (1987) proposed a doubly dipping
oceanic plate, separating two active volcanic arcs on
the Australian and Pacific plates, zippered shut from
west to east since the Oligocene (Figure 2D). In this
model, the New Britain trench is a part of the north-
dipping subduction zone, and the Trobriand trough is
a relict of the south-dipping zone. Dow and Sukamto
(1984a, b) and Dow et al. (1988) proposed that New
Guinea is the product of two distinct islandwide arc-
continent collisions (Figure 2E). One is in the Oligo-
cene (Oligocene orogeny), and the other is in the latest
Miocene (Melanesian orogeny).
Most of the recent models for the Cenozoic tec-
tonic history of New Guinea are significantly different
from the ones just mentioned. Pigram and Davies (1987)
proposed that New Guinea formed as the result of ac-
cretion (docking) of at least 32 distinct tectonostrati-
graphic terranes along the northern Australian margin
from the middle Oligocene to the Pliocene. Audley-
Charles (1991) argued for multiple collisions in the
Miocene. Struckmeyer et al. (1993) present paleogeo-
graphic maps that explicitly illustrate a complex his-
tory of accretion. Pigram et al. (1989, p. 199) believe
that the amalgamation of several arc complexes, oce-
anic plateaus, and microcontinents began northeast
of the present-day New Guinea at about 30 Ma and
continues to the present (Figure 2F). They argued that
the northern Australian margin changed from a passive
margin to a foreland basin setting in the Oligocene
(30 Ma, at least as far west as 135j longitude). Ac-cording to their maps (see also Pigram and Symonds,
1991), the front of the south-verging foreland thrust
belt advanced about 100 km (62 mi) in 17 m.y. (at a
rate of 0.5 cm/yr [0.2 in./yr]). It is important torecognize that the approximately 100-km (62-mi)
advancement they show for the thrust front can ac-
count for only a small fraction of the total convergence
(>1000 km [>620 mi] at 12 cm/yr [5 in./yr] alongan azimuth of 245j) between the Pacific and Aus-tralian plates during this time span (see plate recon-
structions of Scotese et al., 1988; Jolivet et al., 1989).
For this tectonic model to be correct, most of the
PacificAustralian plate convergence must have been
accomodated somewhere else.
Hall (1996) presented a kinematic model for the
region that was based on poles of rotation he deduced
for the Philippine plate with respect to the Pacific, Aus-
tralian, and Eurasian plates. He concluded that the island
of New Guinea margin was one of strike-slip tectonism
from about 25 to about 5 Ma. This kinematic model
is not consistent with the pattern of magmatism and
Quarles van Ufford and Cloos 121
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sedimentation across the island, but the inferred Philip-
pine plate movements are compatible with the regional
tectonic model presented at the end of this paper.
In sum, tectonic models of New Guinea range from
one discrete collisional event to prolonged and piece-
meal accretion. None of these workers have evaluated
the details of how the change from steady-state subduc-
tion to collisional orogenesis would manifest itself in
the rock record; this will be discussed in some detail.
Our starting point is a reappraisal of the Cenozoic stra-
tigraphy of western New Guinea, supplemented with
new data from the core of the highlands.
Figure 2. Schematic diagramsillustrating various tectonic mod-els proposed for the Cenozoicplate-tectonic history of NewGuinea. (A) Modified from Ham-ilton (1979); (B) modified fromKroenke (1984); (C) modifiedfrom Milsom (1985); (D) mod-ified from Cooper and Taylor(1987); (E) modified from Dowet al. (1988); (F) modified fromPigram et al. (1989).
122 Cenozoic Tectonics of New Guinea
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REGIONAL SEDIMENTATION
Biostratigraphy
Determining absolute ages from biostratigraphy is some-
what problematic because correlations have changed over
time. In their classic regional report on Irian biostra-
tigraphy, Visser and Hermes (1962, enclosure 7) report
ages in terms of numbered T-stages (e.g., T26T27).
More recent workers use lettered T-stages (e.g., Ta2;
Adams, 1984; Simon Petroleum, 1992, personal com-
munication). This paper uses Adams (1984) to place
the lettered T-stage notation in an absolute timescale.
Dow et al. (1988, their figure 28) proposed a corre-
lation between numbered to lettered T-stage notations
that has most of the boundaries as nearly the same age
as in this paper, but a few differ by as much as 3 m.y.
(for example, base of T27).
Regional Cenozoic Stratigraphy
Stratigraphic columns were compiled from the lit-
erature for Cenozoic strata deposited on the north-
ern Australian margin that is now exposed in the
Lengguru, Irian, Papuan, and Aure fold and thrust
belts and from the stratigraphy of wells in the Sala-
wati, Bintuni, and southern Central Range foreland
basins (Figures 3, 4). Included in the regional strat-
igraphic synthesis is new data (Quarles van Ufford,
1996) based on biostratigraphic analyses of the 1700-m
(5600-ft)-thick section of Cenozoic limestone exposed
in the glaciated areas near Puncak Jaya (Figure 1), the
highest mountain peak in all of southeast Asia (Figure 4,
column R).
The stratigraphic columns show that during the
Cenozoic, carbonate shelf deposition occurred across
most of the southern part of western and central New
Guinea (Figure 4, columns AT). Along the north-
westernmost slope and rise of the Australian margin
(Birds Head and Neck regions), deep-water carbon-
ates of the Imskin Formation accumulated and are now
uplifted and exposed in the Lengguru fold belt (Figure 4,
columns I and N).
A disconformity, locally overlain by sandstone of
Oligocene age, is recognized across much of western
and central New Guinea (Figure 4, columns D, E, H,
and KT). As many workers have placed great em-
phasis on this disconformity, we report in some detail
Figure 3. Tectonic map of New Guinea showing the location of the Cenozoic chronostratigraphic section in Figure 4. See Figure 1for explanation of map symbols.
Quarles van Ufford and Cloos 123
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on the nature of the associated siliciclastic sediments
we found exposed at this stratigraphic level in the core
of the Central Range near Puncak Jaya (Figure 1).
Oligocene Siliciclastic Strata in Western New Guinea:Sirga Formation
The only siliciclastic unit in central and western New
Guinea between the Eocene and middle Miocene that
is of sufficient thickness and continuity to warrant for-
mation status overlies the Oligocene disconformity. This
unit, the Sirga Formation (Visser and Hermes, 1962,
p. 8485), is a 10100-m (33330-ft) or so thick quartz
sand-rich unit (Figure 4, columns D, E, H, and KR).
This unit was called the Adi Member by Pigram and
Panggabean (1983) and Pieters et al. (1983).
The Sirga Formation was deposited during a pe-
riod of subaerial exposure, as indicated by plant fos-
sils and coal films in the type locality in the Birds Head
( Visser and Hermes, 1962, p. 8485). Near Puncak
Figure 4. Cenozoic chronostratigraphic cross section. The location is shown in Figure 3. Shallow-marine limestone is defined ascarbonate rock deposited on a shelf (
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Jaya, Quarles van Ufford (1996) found that the Sirga
Formation is characterized by (1) a lower 1020-m
(3366-ft)-thick, nonfossiliferous, fine-grained to gran-
ule quartz sandstone with cross-beds; (2) a planar, or
a 12-m (3.36.6-ft)-thick lower boundary with re-
worked clasts of the underlying Eocene shallow-marine
carbonate Faumai Formation; (3) a distinctive coal seam,
up to 30 cm (12 in.) thick, which is present within 1 m
(3.3 ft) of the base of the formation at two of the seven
localities where the contact outcrops; (4) a lower con-
tact, which marks a biostratigraphic gap (only fossil
zone Tc is missing); and (5) an upper boundary, which
is gradational more than 20 m (66 ft) and has layers of
quartz sand grading upward into glauconitic quartz
foraminiferal sand to glauconitic marly foraminiferal
packstone and finally into foraminiferal packstone.
Overall, this section is similar to those described by
Rossetter (1978), Pieters et al. (1983), Pigram and Pang-
gabean (1983, p. 78), Brash et al. (1991), and Lunt and
Djaafar (1991) for the Sirga Formation in the Birds
Head region.
The Sirga Formation in the Puncak Jaya region is an
extremely mature quartz sandstone. Dozens of hand
specimens were examined with a hand lens to identify
the best representatives of the different variants of sand-
stone. A dozen samples were thin sectioned and exam-
ined petrographically. Four representative samples were
stained for feldspar and point counted (Figure 5A).
The samples are clean quartz arenites (>95% quartz).
Using the criteria of Dickinson and Suczek (1979), all
of the samples would be classified as derived from a
craton interior or continental block provenance on a
QFL ternary provenance plot. Three of the four point-
counted samples would be classified as from the same
provenance on a QmFLt ternary provenance plot. The
fourth sample is very coarse grained, with 28% granules
of polycrystalline quartz. Using the petrographic cri-
teria of Dickinson (1985), this sample is inappropriate
for provenance discrimination.
Tabular cross-beds are abundant in the Puncak Jaya
region. After correcting for local bedding tilt caused by
folding, 64 cross-bed measurements indicate a north-
easterly flowing depositional current (Figure 5B). This
indicates a source in the direction of the Australian
craton. The presence of coal, the glauconitic character
of the quartz-sand layers, cross-bedding, paleocurrent
directions, and contact relationships led to the inter-
pretation that the Sirga Formation near Puncak Jaya
was deposited in a transgressive fluvial and/or beach
environment (Figure 6A).
Oligocene Disconformity
The disconformity of Oligocene age that is present
across most of the Australian continental shelf that
underlies southern New Guinea (Figure 4, west of col-
umn U) has been given profound tectonic significance.
Dow et al. (1988, p. 199200) argue the unconformity
formed as a result of continental basement uplift
and deep erosional incision during their islandwide
Figure 5. (A) Ternary petrofacies diagrams of the OligoceneSirga Formation (four samples) in the Puncak Jaya area (Figure 4,column R). Provenance interpretation is modified after Dickin-son and Suczek (1979). Q (total quartz) = Qm (monocrystallinequartz) + Qp (polycrystalline quartz); F = total feldspar; L =unstable lithic fragments; and Lt = L + Qp. Three hundred pointcounts per sample (Quarles van Ufford, 1996). CB = continentalblock. (B) Cross-bedding orientation in the Sirga Formation nearPuncak Jaya.
Quarles van Ufford and Cloos 125
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Oligocene orogeny. Pigram et al. (1989, 1990) and
Pigram and Symonds (1991) argue the Oligocene un-
conformity in the carbonate section atop the continen-
tal basement of southern New Guinea resulted from a
flexural forebulge of the basement (after Jacobi, 1981)
migrating southward approximately 100 km (62 mi) in
advance of a landmass emerging from the Pacific basin.
However, a global phenomenon that must ac-
count, at least in part, for the origin of the Oligocene
disconformity is the largest eustatic fall in sea level in
the Cenozoic (Haq et al., 1987). This event, now well
dated as between 33 and 30 Ma (Vakarcs et al., 1998),
involved a sea level fall of about 90-m (300-ft). It has
been detected in cores drilled into mid-Pacific atolls
(Schlanger and Premoli Silva, 1986) and is now well
known to have had a dramatic effect on deep-sea sed-
imentation patterns in the region (Fulthorpe et al.,
1996). An Oligocene disconformity, similar to that re-
corded in the stratigraphy of southern New Guinea is
present in carbonate strata on the northwest (Apthorpe,
1988) and northeast shelves of Australia (Davies et al.,
1989). These nearby areas did not undergo tectonic
movements in the middle Cenozoic.
In New Guinea, the Oligocene transgression and
deposition that followed the sea level low created a
variable pattern of unconformity because of variable
preexisting relief and depth of erosion. In the Birds
Head region, Mesozoic to middle Miocene formations
onlap the Kemum high (Figure 1), indicating that it
was a basement exposure and minor source of silici-
clastic detritus throughout the Cenozoic (Dow et al.,
1988, p. 3134, 194).
In the western highlands, the unconformity typ-
ically spans a range of about 5 m.y. In the Puncak Jaya
section, which must have been in an outer shelf envi-
ronment, only one foraminiferal zone (Tc) has not
been found, and the cessation of deposition may be as
short as 1 m.y. (Figure 4).
In central New Guinea around the Arafura high,
the unconformity is overlain by the lower Miocene
Darai Limestone and juxtaposes rocks as old as Me-
sozoic. It appears that the Arafura high was also a
Figure 6. Interpretedblock diagram for SirgaFormation terrestrial toshallow-marine deposi-tion. (A) Ertsberg miningdistrict depositional set-ting. (B) Regional depo-sitional setting.
126 Cenozoic Tectonics of New Guinea
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pre-Cenozoic feature because no Permian to Jurassic
strata occurred beneath onlapping Cretaceous strata
or Neogene carbonates. In six out of nine wells that
penetrate the basement in the Arafura high, more
than 400 m (1300 ft) of Cretaceous sediments rest
disconformably on top of what is correlated as either
the Silurian to Devonian Modio-Brug Formation or Pre-
cambrian or early Paleozoic Kariem Formation (Visser
and Hermes, 1962, enclosure 8; Simon Petroleum, 1992,
personal communication). Part of the Arafura high
apparently rose hundreds of meters during the Paleo-
cene as the area was gently tilted during movements
associated with the opening of the Coral Sea basin
from 62 to 56 Ma (Weissel and Watts, 1979; Davies,
1990; Davies et al., 1991). In addition, it is likely that
the Arafura high was similarly uplifted, and the flanks
were again gently tilted during the about 3530 Ma
Peninsular orogeny (discussed below).
In sum, we conclude that the widespread Oligo-
cene unconformity in central and western New Guinea
is primarily the result of the short-lived, but large, sea
level fall. The Sirga Formation is simply the basal trans-
gressive unit deposited during sea level rise and does
record evidence of a major, local tectonic event.
Source of the Quartz Sand
The primary source of the quartz grains is uncertain.
Grains may have migrated from Australia as shore-
lines transgressed northward. Probable sources for
quartz are the Kemum and Arafura highs. We empha-
size that some, and perhaps most, quartz must have
been released during exposure and dissolution of early
Cenozoic carbonate formations. Scattered grains and
thin quartz sand beds are present in Eocene limestone
in the western Central Range and in the Birds Neck
region (Figure 4, columns K and P). Quarles van Ufford
(1996) found that in the Puncak Jaya area, nearly 60 m
(200 ft) of the approximately 400-m (1300-ft)-thick
Paleocene to Eocene Waripi Formation dolomites con-
tains more than 10% quartz. The rest of the Waripi
as well as the lower part of the Eocene to Oligocene
Faumai Formation limestones contain scattered grains
of mostly silt-sized quartz.
Imskin Formation: Pelagic Carbonate Sedimentation
Carbonate strata directly indicate that the northern
and eastern edges of the Birds Head region was a pas-
sive margin until the middle Miocene. The Imskin
Formation, a pelagic limestone that grades shelfward
(south and west) into shallow-water carbonate forma-
tions (Figure 4, columns I and N), outcrop in the Leng-
guru fold and thrust belt in the Birds Neck region at
about 135j300E (Visser and Hermes, 1962, p. 7779;Koesoemadinata, 1978; Pieters et al., 1983; Brash et al.,
1991). The Imskin Formation consists of marl, chalk,
chert, and abundant pelagic foraminifera. In the Leng-
guru area, the Imskin Formation lacks siliciclastic beds,
except during the period of Sirga Formation deposition
(Figure 4, column N) (Brash et al., 1991). The regional
setting in the Birds Head region at the time of Sirga
deposition is illustrated in Figure 6B.
Synorogenic Sedimentation on Continental Basement:Central Southern New Guinea
There are two distinct ages for the initiation of imma-
ture siliciclastic sedimentation on top of the conti-
nental basement of New Guinea (Figure 4). The evi-
dence is (1) an older Oligocene event recorded in the
Aure trough of easternmost New Guinea and (2) a
younger late Miocene event recorded across the width
of New Guinea.
The regional synthesis reveals that from the Oligo-
cene to middle Miocene, immature siliciclastic sedi-
mentation atop Australian basement was restricted to
easternmost New Guinea, east of 144jE (Figure 4, col-umn Papua New Guinea (PNG) 2 and eastward). The
oldest synorogenic sediments are found in the Aure
trough near the Papuan Peninsula. Up to 7 km (4 mi) of
sediment has accumulated in the depression since the
middle Oligocene (Te14, 32 Ma; Edwards, 1950;Brown et al., 1975; Slater et al., 1988). They were de-
rived from sedimentary, igneous, and metamorphic ter-
ranes exposed in the Papuan Peninsula. Similar sedi-
mentation occurred northeast of the Papuan Peninsula
in the Cape Vogel depocenter (Davies et al., 1984).
Synorogenic sediments deposited on the continen-
tal basement of western New Guinea are significantly
younger. Kilometer-thick sequences of siliciclastic
strata have accumulated in the Salawati, Bintuni, Aki-
meugah, and Iwur basins (Figure 7) since the latest
middle Miocene (Figure 4) (Visser and Hermes, 1962,
p. 8899; Dow et al., 1988, p. 170178). The propor-
tion of different siliciclastic lithologies varies signifi-
cantly from basin to basin, and different names have
been used for the formations overlying middle Miocene
carbonate strata. The Klasaman, Steenkool, Akimeugah,
Buru, and Iwur formations are reported to contain shale,
siltstone, sandstone, lignite, fossiliferous marl, and lo-
cally, a basal conglomerate containing New Guinea
Quarles van Ufford and Cloos 127
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limestone group clasts (Figures 4, 7) (Visser and Hermes,
1962, p. 9698). The oldest siliciclastic deposits are
found in the bottom of the Iwur basin of central New
Guinea (Tf1 stage, 14 Ma; Dow, 1977), but the Tf2age (12 Ma) is when the siliciclastic deposition be-came widespread on top of the continental basement
of Papua.
The youngest pulse of coarse synorogenic sedi-
mentation on top of continental basement appears to
be slightly younger in Papua New Guinea. The oldest
strata related to the generation of the Papuan fold and
thrust belt that caused the imbrication of Darai Lime-
stone (Hobson, 1986) appear to be Pliocene (Figure 4,
PNG columns 26). In the Aure trough area, silici-
clastic sediments derived from the Papuan Peninsula
region had been continuously accumulating since the
Oligocene. However, a distinct change to conglomeratic
nonmarine deposition indicates that a substantial new
uplift of the source terrane occurred in the Pliocene
(Era beds, Figure 4, column PNG1) (Slater et al., 1988;
Davies, 1990; Klimchuk, 1993; Kugler, 1993).
Synorogenic Sedimentation on Oceanic Basement:Central Northern New Guinea
In northern Papua, coarse clastic material had accumu-
lated in the North Coast basin (also known as the
Meervlakte; Figure 1) since the early middle Miocene.
These deposits bury most of the collided arc complex.
The oldest siliciclastic material, the Makats Forma-
tion, which blankets this oceanic basement, appears
to be dated as early middle Miocene (1614 Ma;
Visser and Hermes, 1962, p. 100111). These silici-
clastic deposits on the north side of the island appear
to predate, by several millions of years, the beginning
of widespread synorogenic sedimentation on top of
Australian continental basement. The Makats Forma-
tion contains clasts of metamorphic rocks, mica schist,
[and] slates (Visser and Hermes, 1962, p. 100106),
indicating deep denudation of the source landmass
emerging to the south.
Interbedded clastic and carbonate sediments in the
Sepik basin of northern Papua New Guinea (Figure 1)
were deposited on top of deformed sediments and meta-
morphic rock that correlates with the Owen Stanley
metamorphic belt that underlies the Papuan Penin-
sula (Dow, 1977; Doust, 1990). Geochronology of the
metamorphic belt indicates recrystallization before
about 25 Ma, in the latest Oligoceneearliest Miocene
(Hill et al.,1993). Clastic strata (including the poly-
mictic Amogu conglomerate) of early Miocene age
accumulated along the southern margin of the Sepik
basin (near the northern edge of the present-day Pap-
uan Peninsula). At the same time, carbonate strata
accumulated in the rest of the basin (Doust, 1990;
Wilson et al., 1993).
Miocene Volcanism: Maramuni ArcEastern New Guineaand the Trobriand Trough
A period of southwest-dipping subduction occurred
along northeastern New Guinea at the Trobriand
trough (Figure 1) between 20 and 10 Ma and gen-
erated the Maramuni arc (Page, 1976; Dow, 1977;
Figure 7. Simplified chronostratigraphic section since the early Miocene located on the Australian shallow-water carbonate platform andnear the modern southern deformational boundary for the Central Range. Based on Figure 4 and modified from Visser and Hermes (1962).NGLG = New Guinea limestone group; cgl = conglomerate; ls = limestone; ss = sandstone; terr = terrigenous; and v = volcaniclastic.
128 Cenozoic Tectonics of New Guinea
-
Davies et al., 1984; Davies, 1990). This magmatic arc
was emplaced into the Australian continental crust east
of 140.5jE (Figure 3). The westernmost occurrence ofprobable Maramuni intrusives, just west of the inter-
national border, are undated (marked with a question
mark in Figure 3) (McMahon, 2000a). Maramuni-age
igneous rocks farther to the west are only found em-
placed into allochthonous Pacific plate terranes, the
Weyland overthrust (Utawa batholith), and the west-
ernmost Irian ophiolite belt next to the Birds Neck
(Figure 3) (Dow et al., 1988, p. 149154; McMahon,
2000b).
Maramuni arc plutons intruded the largely sub-
merged belt of deformed and metamorphosed conti-
nental strata in easternmost New Guinea, but volca-
nism provided detritus to the Sepik and Ramu basins
that locally overwhelmed carbonate sedimentation (Wil-
son et al., 1993). Only in the northwestern Papuan fold
and thrust belt (Star Mountains region, near the In-
donesia and Papua New Guinea border) did Maramuni
arc volcanism become recorded in the middle Miocene
shelf sediments (Pnyang and Lai siltstones) deposited
atop continental crust (Davies, 1990). These partly
volcanogenic strata are located north of the section in
Figure 3 and are not reported in the wells south of the
Papuan fold and thrust belt.
The formation of the Maramuni arc is a distinct
tectonic event that all tectonic models for New Guinea
must account. It is the product of a short-lived period
of subduction at the Trobriand trough. Its northern ex-
tension is actively being overridden by the Finisterre
Huon forearc terrane (Silver et al., 1991).
Middle Cenozoic Tectonism: Westernmost New Guinea(Birds Head)
The Cenozoic history of the Birds Head region of
westernmost New Guinea should now be well under-
stood because of major hydrocarbon discoveries in the
Salawati and Bintuni basins. The Kemun high was a
basement exposure and a minor source of siliciclastic
detritus throughout the Cenozoic. Most importantly,
the extensive Kais Formation reef complexes along
the southern and western margin of the Kemum high
(Figure 3) indicate that rapidly eroding highlands
were not present near the Bintuni and Salawati basins
until the latest Miocene (Visser and Hermes, 1962,
p. 85; Dow et al., 1988).
Based on the literature known to us, it appears
that the deposition of the partially siliciclastic Klasafet
Formation in the Bintuni and Salawati basins of the
Birds Head region began in the early middle Miocene
(early Tf1, 1716 Ma; Figure 4, column E and H)
(Visser and Hermes, 1962, p. 8892; Froidevaux, 1978).
The only direct structural evidence for Oligocene de-
formation of strata deposited on the continental base-
ment of western New Guinea is found near the island
of Misool (Figure 1; columns A and B in Figure 4). In
this area, an angular unconformity with up to about
15j of discordance separates tilted Oligocene ZaagLimestone and older formations from the near-
horizontal Miocene Kasim marl and Openta Limestone
(Pigram et al., 1982; Robinson et al., 1988). We believe
that the Oligocene tilting recorded near Misool, the
renewed uplift of the Sele high, and perhaps the ini-
tiation of Klasafet sedimentation are the easternmost
manifestation of collisional plate interactions occurring
west of the present-day New Guinea in the Sulawesi
area (Silver et al., 1985; Bergman et al., 1996).
TWO OROGENIES: PENINSULAR ANDCENTRAL RANGE
Regional stratigraphic patterns and fieldwork reported
in Quarles van Ufford (1996) and summarized in this
paper indicate that in the vicinity of the Papuan Penin-
sula, two Cenozoic orogenic events affected the Aus-
tralian continental basement, but only the younger
event is evident in central and western New Guinea. In
eastern New Guinea, two collisional orogenic events
were separated by an episode of southwestward sub-
duction along the Trobriand trough, which emplaced
the Maramuni arc into the Australian continental base-
ment. In this paper, the orogenies are named for the
largest mountainous uplift each created: the Peninsular
Range and Central Range.
Several tectonic enigmas near New Guinea must
be considered in any plate model for the history of the
island. These include (1) the northwestern and south-
eastern limit of Trobriand troughMaramuni arc mag-
matism (Figure 3); (2) the origin of the New Guinea
trench (Figure 1); and (3) the western limit of arc
magmatism associated with subduction at the New
Britain trench (near 145jE, Figure 3). The discussionthat follows is a hypothesized sequence of plate-tectonic
adjustments in the southwest Pacific that account for
these three issues as well as the geologic history of
the island as described in this report. The reconstruc-
tions are based on the fact that the overall motions of
the Pacific and Australian plates are well constrained
Quarles van Ufford and Cloos 129
-
(Scotese et al., 1988). They are an attempt to place
the New Guinea region into a more complete plate-
tectonic context than is illustrated by reconstructions
such as those shown in Figure 2. Complete justifica-
tion requires the analysis of the history of magmatism
and deformation in the islands to the east and west of
New Guinea and must be the subject of another paper.
Major Change in Pacific Plate Motion at 43 Ma
A major Eocene change in Pacific plate motion is evi-
dent from the distinct bend in the HawaiianEmperor
seamount chain that is dated at 431 Ma (Clague and
Dalrymple, 1989). This change is either the result or
the cause of the formation of new subduction zones
in the western Pacific basin (Hilde et al., 1977; Kroenke,
1984) and corresponds to a large increase in spread-
ing rate between the Australian and Antarctic plates
(Veevers et al, 1990).
North of the equator, the west-dipping IzuBonin
Mariana subduction system was established. To the
south, two subduction systems were started. One was
northeast-dipping at the PapuanRennellNew Cale-
donian trench system and generated the Inner Mela-
nesian arc. The other, far to the northeast (1500 km;900 mi), was a southwest-dipping subduction zoneand generated the Outer Melanesian arc that is now
the inactive New GuineaManusKilinailauSolomon
trench system (Figure 8A). Whether subduction started
at the same time at both of the Melanesian arc systems
is uncertain, but major events along the inner sub-
duction zone (southern arc) soon caused all the Pacific
Australian plate convergence to become concentrated
at the outer subduction zone (northern outer arc).
Subduction Accretion and Collisional Orogenesis
Before further synthesis is discussed, it is important to
differentiate the generally short-lived collisional oro-
genesis processes from the commonly long-lived, near-
steady-state, subduction processes of offscraping and
underplating. The terms subduction and collision have
been used as synonyms in textbooks and many papers.
From the perspective of describing fundamental con-
vergent margin processes, formal differentiation of
these terms is desirable (Cloos, 1993). We restrict the
term collision for subduction zone events that lead to
some kind of change in plate motions and the uprooting
of crystalline basement or thick-skinned deformation.
Subduction can (but not always) cause the accumula-
tion of an accretionary prism that is the product of
prolonged tectonism (see Cloos and Shreve, 1988a, b
for discussion of the mechanics of subduction accretion
and nonaccretion and erosion). Subduction accretion is
a thin-skinned deformational process that commonly
occurs steadily for many tens of millions of years,
forming an accretionary prism without the detachment
of the underthrusting layer of ocean crust. Offscraping
widens accretionary prisms, whereas underplating
thickens them. By themselves, these processes cause
tectonism but not orogeny (which historically was
defined as the generation of mountains that are subject
to erosion). All active subduction zones with large
accretionary prisms are nearly entirely underwater.
Accretionary complexes can, of course, become parts
of mountainous uplifts when they are involved in
continent-arc and continent-continent collisions. Col-
lisions, in our definition, involve the jamming of a sub-
duction zone with consequent deformation involving
the crystalline top of the descending plate and some
kind of rearrangement of plate motions. Discrimina-
tion of subduction, offscraping, underplating, and
collisional tectonism provides insight into the tectonic
controls on sedimentation and the significance of ter-
rane boundaries in New Guinea and elsewhere.
The basic physics of subduction is obviously related
to but fundamentally different from collision. Steady-
state subduction can continue as long as the bulk den-
sity of the downgoing lithosphere (crystalline crust
and underlying lithospheric mantle) is greater than the
bulk density of the underlying asthenospheric mantle
(Figure 9A). With little modification, it can continue
for many tens of millions of years and result in subduc-
tion erosion truncating a margin or subduction accre-
tion growing an accretionary prism. Collisional oro-
genesis, however, begins when incoming lithosphere
that is less dense than the asthenospheric mantle begins
to turn downward to subduct. Incoming lithosphere
is positively buoyant when it has a sufficiently thick
capping of crystalline continental crust or a large oce-
anic arc complex (Figure 9C). Depending primarily on
the speed of convergence and other factors as well,
collisional orogenesis can be prolonged, but it is com-
monly a short-lived event (a few million years) that
ends when there is a change in plate motions, a re-
arrangement of plate boundaries, or both (Cloos, 1993).
One common manifestation is the contraction of the
overriding plate in the area of the arc because this
is a thermally weakened lithosphere. Where oceanic
arcs are involved in a collision, lithospheric rupture
and subduction reversal, with the line of the old arc
becoming the axis of the new trench, are common.
130 Cenozoic Tectonics of New Guinea
-
Three episodes of subduction reversal occurred in the
New Guinea region in the Cenozoic. Once, it occurred
immediately following the jamming of a subduction
zone, and twice, it was delayed, and the result of later
tectonic reorganizations caused subduction to initiate
in the still warm and weak arc environment.
The Papuan and Irian ophiolite complexes are up-
lifted forearc terranes. They overlay the Owen Stanley
and Ruffaer metamorphic belts that are composed
largely, if not entirely, of the Australian continental
rise, slope, and outer shelf protoliths. Any oceanic ac-
cretionary prism composed of pelagic and trench axis
materials that are scraped off the Pacific plate is very
minor, if any is present. The subduction zones above
which these ophiolites were uplifted because of con-
tinental margin underthrusting and collision was sim-
ilar to the intraoceanic Mariana convergent margin (a
sediment-poor trench with active subduction erosion)
north of New Guinea. Only a small prism could have
accumulated, and thermal thinning of the lithosphere
beneath the coeval arc was limited because the periods
of subduction before the collisions were short.
Along northern Australia, the outer portions of
the continental margin were bulldozed, and the top of
the deforming pile formed small islands prior to colli-
sional orogenesis and the uprooting of crystalline base-
ment. This precollision complex involves the sequen-
tial deformation of first the rise, then slope, and finally
shelf strata (Figure 9B). In the Puncak Jaya region of
western New Guinea, crystalline basement became
involved in the deformation at about 8 Ma (Weiland
and Cloos, 1996). The regional field relations indicate
that collisional tectonism involving crystalline basement
began about 4 m.y. after the approximately 12-Ma
initiation of the Central Range orogeny as marked by
widespread siliciclastic sedimentation.
Collision Event 1: The Peninsular Orogeny
The Peninsular orogeny of New Guinea is an Oligo-
cene event (3530 Ma) restricted to the vicinity ofthe Papuan Peninsula. This event was caused by the
underthrusting of the northeastern edge of the Aus-
tralian continent and marked the end of a 1015-m.y.
episode of northeast-dipping subduction (Figure 8A, B).
Total convergence at the Papuan subduction zone was
probably only a few hundred kilometers because only
a minor volcanic arc was generated in the Trobriand
Sea (Kroenke, 1984; Davies et al., 1984; Davies and
Warren, 1988). By about 30 Ma, the Papuan segment
was fully jammed (Figure 9C). The underthrusting
of bulldozed sediments (Owen Stanley metamorphic
belt) and the Australian continental margin caused the
uplift and exposure of the crystalline oceanic forearc
block, the Papuan ophiolite (Davies, 1971; Davies and
Jaques, 1984). The Oligocene to Pliocene paleogeog-
raphy in the area of the Papuan Peninsula was probably
quite similar to present-day New Caledonia, which
formed at about the same time and in a similar manner.
The Eocene to Oligocene subduction zone form-
ing the Inner Melanesian arc continued eastward from
the Papuan Peninsula to the Rennell trench and arc
complex and from there southward to New Caledonia
Norfolk ridge (Figure 9B) (Parrot and Dugas, 1980) and
to New Zealand, where it is known as the Kaikoura
orogeny (Brothers, 1974; Hayward et al., 1989). The
combination of nearly coeval collisional orogenesis at
the Papuan, New Caledonia, and New Zealand trenches
stopped northeast-dipping subduction along the entire
length of the Inner Melanesian arc. This jamming
caused all Pacific and Australian plate convergence to
become accomodated far offshore at the outer Mela-
nesian trench system, and the oceanic lithosphere be-
tween the arc was effectively welded to the Australian
plate.
The Peninsular deformational belt extended as a
submarine terrane north of the present-day Papuan
Peninsula into the Sepik region and beyond. It was
blanketed by Miocene mudstone and limestone and
then intruded by the magmas of the Maramuni arc
(Dow, 1977; Doust, 1990; Wilson et al., 1993).
Since the Oligocene, the Peninsular uplift has been
the source of immature sediment (Slater et al., 1988;
Davies, 1990) and more than 7 km (4 mi) of siliciclastic-
rich strata, the Aure Group, accumulated in the Aure
trough (Figure 4). This depression, the trench of the
extinct Papuan subduction zone, was a sediment trap
protecting the Australian carbonate shelf from the
influx of siliciclastic detritus from the Peninsular high-
lands. Much of the Aure Group is composed of fine- to
medium-grained turbidite and mass flow deposits that
grade upward from deep to shallow marine (Kugler,
1967; Brown et al., 1975; Francis et al., 1986). The
Miocene part of the Aure Group contains abundant
carbonate clasts of Eocene age (Carman, 1990), as well
as some ophiolite and blueschist detritus. Basement
detritus became much more abundant in Pliocene and
younger strata (Francis et al., 1986; Klimchuk, 1993).
The stratigraphic succession records the progressive
erosional unroofing of the Papuan Peninsula.
Marine geophysical studies in the eastern plateau
region (Figure 1) indicate that the Oligocene jamming
Quarles van Ufford and Cloos 131
-
132 Cenozoic Tectonics of New Guinea
-
of the Papuan subduction zone had some comparative-
ly subtle, regional effects that generated structures that
could be hydrocarbon traps. As much as 1 km (0.6 mi)
of reverse slip during the late Oligocene to early Mio-
cene occurred along preexisting basement normal faults
more than 200 km (120 mi) southwest of the Papuan
Peninsula (Davies et al., 1989). It seems probable that
the collision caused similar movements in the Arafura
Figure 8. Tectonic evolution of New Guinea consistent with regional deformation and magmatic and sedimentation patternssummarized in this paper and from the literature for the surrounding region. AUS = Australian plate; BT = BewaniTorricelli arc;EQTR = equator; FH = FinisterreHuon arc; IMA = Inner Melanesian arc; IOB = Irian ophiolite belt; K = Kilinailau; LHR = Lord Howerise; MB = ManusBismarck arc; NGT = New Guinea trench; NR = Norfolk ridge; NS = north Solomon arc; OMA = Outer Melanesianarc; OJP = Ontong Java Plateau; PAC = Pacific plate; PHS = Philippine plate; POB = Papuan ophiolite belt; SS = South Solomon. T1, T2,T3 = postulated transform faults. Approximate paleolatitudes are modified from Scotese et al. (1988) and Veevers et al. (1991).
Figure 9. Lithospheric-scale cross sections through time of the AustralianPacific, arc-continent collision. Northern Australia was a passivemargin since rifting in the Triassic (Pigram and Panggabean, 1984). See Figure 3 for line of section. (A) Prior to collisional orogenesis, theoceanic portion of the Australian plate is subducted toward the north. The passive continental margin is not involved. (B) Initiation oforogeny at about 12 Ma from contractional thickening of bulldozed passive-margin strata and the underthrusting of Australian continentalbasement. Passive-margin strata on the Australian plate are bulldozed and contracted to such a point that a subaerial high underlain by aprecollision complex is formed. Erosional detritus from this precollision complex accumulates nearby and records the beginning of theCentral Range orogeny in the stratigraphic record. (C) Initiation of collisional orogenesis at about 8 Ma. The point of neutral buoyancy on theAustralian plate has reached the subduction zone. Convergence between the Australia and Pacific plates is no longer accommodated byconvergence. Positive and negative lithospheric buoyancy are with respect to the asthenospheric mantle. NGT = New Guinea trench.
Quarles van Ufford and Cloos 133
-
high (Figure 1). This region was the site of modest
tilting and uplift during the Paleocene opening of the
Coral Sea (Weissel and Watts, 1979; Davies, 1990).
Near-vertical uplift of a few hundred meters in the
late Oligocene could fully account for the Paleocene to
Miocene unconformity centered on the Arafura high
(Figure 4, columns U to PNG2). Exposure of Mesozoic
siliciclastic formations (Kembelangan Group), such as
that detected on the eastern plateau, could have been
a source for the quartz sands in the Sirga Formation.
Except around the uplifted area centered on the
present-day Papuan Peninsula, carbonate sedimenta-
tion occurred elsewhere on the Australian continental
shelf (Figure 4). The Oligocene sea level drop of about
90-m (300-ft) at about 3330 Ma caused large areas
of the Australian shelf to become emergent. This
created the regional disconformity over what is now
western and central New Guinea that was overlain by
the transgressive Sirga Formation during the middle to
late Oligocene (Figure 4, columns D, E, H, and KR).
Pelagic limestone deposition north of the Birds Head
indicates that a complete passive-margin setting (shelf-
slope-rise) existed in western New Guinea until the end
of the middle Miocene.
Subduction at the Trobriand Trough andMaramuni Magmatism about 2010 Ma
A mechanical cause for the short-lived subduction
event forming the Trobriand troughMaramuni arc is
unaccounted for in existing tectonic models. We be-
lieve a simple explanation for the initiation of sub-
duction at the Trobriand trough is the jamming of
southwest-dipping subduction at the KilinailauNorth
Solomon trench and Outer Melanesian arc segment by
the collision of the Ontong Java Plateau. A profound
change in plate interactions was caused by the sub-
duction of the leading edge of the more than 30-km
(18-mi)-thick oceanic crust underlying the Ontong
Java Plateau (Kroenke, 1984). To account for the east-
ern limit of the New Guinea trench and the north-
western and southeastern limits of the Trobriand trough
and Maramuni arc and the western limit of the New
Britain trench, we postulate that three major transform
fault zones (T1T3 in Figure 8C) formed between 25
and 20 Ma. Movements postulated for transform fault
T3 are compatible with the late Cenozoic motions of
the Philippine plate deduced by Hall (1996).
We postulate that southwest-dipping subduction
occurred at the Trobriand trough between two major
transforms (T1 and T2 in Figure 8C). This was con-
current with the initiation of northeast-dipping con-
vergence by the immediate subduction reversal behind
the western part of the Outer Melanesian arc between
the northern transform and another still farther north
(T2 and T3 in Figure 8C). The Trobriand trough is
located where the Inner Melanesian arc would have
been located during the period of subduction at the
Papuan trench from the Eocene to about 30 Ma. Sub-
duction to accommodate the AustralianPacific plate
convergence probably started here because it was still
a belt of thermally weakened lithosphere. This was a
delayed subduction reversal.
Contractile deformation of Aure trough strata
began to form the Aure fold and thrust in the middle
Miocene (Slater et al., 1988). It seems likely that some
of this deformation was the result of the short-lived,
minor reactivation of the old Papuan subduction zone
before subduction was fully established at the Tro-
briand trough. Subsidence analysis of wells in the Gulf
of Papua shows that the rate of sediment accumula-
tion became more rapid at about 25 Ma (Pigram and
Symonds, 1991; Wang and Stein, 1992). Uplift of the
present-day Papuan Peninsula and faster erosion is
another manifestation of crustal movements caused by
the initiation of subduction at the Trobriand trough.
The Cape Vogel basin was a forearc basin that ponded
substantial sediment (Davies and Smith, 1971).
Collision Event 2: Central Range Orogeny
The name Central Range orogeny is proposed for the
event that generated the about 1300-km (800-mi)
mountainous spine of New Guinea that stretches from
the Birds Neck (135jE) up to the Papuan Peninsula(146jE). This chain includes the Sneeuw Mountains(Hamilton, 1979, his figure 119), also known as the
Pegunungan Maoke or Central Range (Allison and Pe-
terson, 1989) of Papua, as well as the New Guinea
Highlands and Papuan foothills (Dow, 1977), also
known as the Central Cordillera of Papua New Guinea
(Davies, 1990).
Fast, north-dipping subduction along an approx-
imately 500-km (300-mi) length of the Outer Mela-
nesian arc (Figure 8D) began between 30 and 25 Ma
(Figure 9A). This soon caused submarine tectonism and
metamorphism of the Australian continental margin
deposits. The first evidence of land emergence caused
by the approach of this subduction zone comes from
the early middle Miocene (early Tf1, 1614 Ma)Makats Formation in the North Coast basin (Visser and
Hermes, 1962, p. 103105). This basin is underlain by
134 Cenozoic Tectonics of New Guinea
-
the Irian ophiolite (at the time a forearc terrane) and
the oceanic island arc complex along the Irian north
coast. The oldest Makats Formation appears to be re-
stricted to the eastern part of the basin and seemingly
was sourced from small early middle Miocene islands
near the modern international border. The Makats
Formation reflects the development of isolated islands
formed by the progressive tectonic bulldozing of the
sedimentary sequence deposited on oceanic basement
(continental rise) and then transitional Australian crust
(continental slope and outer shelf; Figure 9B). The
paleogeography of the eroding bathymetric high(s) was
probably similar to present-day subduction zones off
Sumatra forming Nias Island and in the Lesser Antilles
forming Barbados Island.
The lower Iwur Formation on the south side of
the Central Range contains the oldest siliciclastic de-
posits known to have been deposited on continental
basement outside of the Aure trough. This occurrence
probably marks the filling of the Aure trough to the
east, but it could mark the first debris shed southward
from an emerging landmass to the north. In any case,
the Central Range orogeny did not begin until the
latest middle Miocene because carbonate shelf sedi-
mentation continued across the Australian continental
shelf until about 12 Ma. At that time, the regional
change in depositional patterns records that a sub-
stantial landmass underlain by bulldozed continental
margin deposits was elongated east-west. The Klasa-
man, Akimeugah, Iwur, and lower Buru formations
contain voluminous shale, siltstone, and sandstone,
which indicates a siliciclastic-rich landmass extended
more than 500 km (300 mi) (Figure 4, columns CU).
A pronounced sedimentological change occurred
in the PliocenePleistocene (5 Ma), when the coarse-ness of deposits along the flanks of the Central Range
increased dramatically for hundreds of kilometers along
strike (Figure 7). The Sele, Steenkool, Dakebo, upper
Buru, Birim, and Era formations contain boulder beds
that appear to date when the Central Range attained a
topography similar to that of today. Shortly before,
crystalline Australian basement first became involved
in the deformation. Field relations and fission-track
analysis studies in the Puncak Jaya region indicate that a
30-km (18-mi)-wide basement block, the Mapenduma
anticline, was pushed southward with unroofing be-
ginning at about 8 Ma (Weiland and Cloos, 1996).
Collisional orogenesis in western New Guinea must
have begun at that time (Figure 9C).
The collisional jamming that formed the Central
Range changed the force balance on the plates, and we
believe this explains the cause of major tectonic events
to the west and east. Subduction began along trans-
form T3 (Figure 8c), extending the Java subduction
system eastward forming the Banda trench. This seg-
ment is now undergoing collisional tectonism resulting
from subduction of the northwestern part of the Aus-
tralian continent near Timor. Northward-dipping sub-
duction also began by delayed subduction reversal along
the easterly segment of the still warm Outer Melane-
sian arc. This created the present New BritainSolomon
arc system. As PacificAustralian convergence became
accommodated at this zone, subduction at the Tro-
briand trough ceased. We emphasize that this geom-
etry and sequence of events accounts for the location
of the western end of the New Britain arc (Figure 1).
The most northwestern segment of the Outer
Melanesian arc complex, the product of southwest-
dipping subduction between about 40 and 25 Ma and
the northeast-dipping subduction from 25 to about
10 Ma, is built on Mesozoic oceanic basement. With
the initiation of northeast-dipping subduction between
transform faults T2 and T3 at about 25 Ma, the older
arc complex became a forearc basement terrane. A
large piece of this terrane, the Irian ophiolite belt, was
uplifted during the Central Range orogeny, as Austra-
lian continental margin sediments (Ruffaer metamor-
phic belt) and the crystalline basement were underthrust.
The associated arc (Maramuni time equivalent) was
parked near the present north coast of New Guinea at
about 10 Ma.
Ongoing Collisional Orogeny in Eastern New Guinea
The Central Range orogeny is a slightly younger and
ongoing event in eastern New Guinea. From about 20
to 10 Ma, PacificAustralian convergence was accom-
modated by the southwest-dipping subduction at the
Trobriand trough (Figure 8C, D). The jamming of
the subduction zone in western New Guinea corre-
sponds to the slowing and eventual cessation of fast
subduction at the Trobriand trough. The initiation
of northeast-dipping subduction began beneath the
extinct but still warm eastern segment of the Outer
Melanesian arc. Since about 10 Ma, nearly all Pacific
Australian convergence has been accommodated at the
New BritainSolomon Arc system (Figure 8D). The
forearc terrane for this system is now exposed in the
Adelbert and Finisterre Ranges and Huon Peninsula
and eastward-forming New Britain Island. As with the
Irian ophiolite, these terranes contain remnants of the
Quarles van Ufford and Cloos 135
-
Outer Melanesian arc volcanism that was the product
of southwest-dipping subduction between about 40
and 25 Ma.
In eastern New Guinea, progressive jamming of
the north-dipping subduction zone is underway. This
uplift of the Adelbert Range and FinisterreHuon
Ranges has formed a mountain belt along the north
coast of eastern New Guinea. In the Sepik basin re-
gion, collisional tectonism deforms the shallow-marine
Miocene strata (Dow, 1977; Doust, 1990) that blan-
keted rocks deformed in the Peninsular orogeny.
The change in force balance from the collision
that formed the Central Range orogeny caused pro-
found plate-tectonic changes in the immediate area
and probably the entire Pacific basin. Cox and Engeb-
retson (1985) and Pollitz (1986) delineate a change in
Pacific plate motion at that time. It appears that the
prong of the Pacific plate directly north of New Guinea
was temporarily detached and moved as its own kine-
matic entity from about 5 to 3 Ma (Figure 1) (the
Caroline plate of Weissel and Anderson, 1978). The
back-arc area of the New Britain arc ruptured, forming
the Bismarck microplate with strike-slip faulting and
sea-floor spreading along its northern boundary (Figure 1).
This piece of lithosphere has been a distinct kinematic
entity since about 3.5 Ma (Taylor, 1979). The trans-
form zone extends westward, emerging onland as the
BewaniTorricelli fault zone (BTFZ in Figure 1),
which links to the Yapen and Sorong faults (SYFZ in
Figure 1) in western New Guinea (Sapiie et al., 1999).
Transform faulting was localized in this region be-
cause the lithosphere was locally thin beneath the
recently extinct arc. The relict arc, largely buried be-
neath the deposits of the North Coast basin, has been
dismembered as a result of about 300 km (190 mi) of
left-lateral, postcollision, transform offset.
Another tectonic result of the change in force
balance is found to the east of New Guinea. The north-
ern corner of the Australian plate ruptured, forming
the Woodlark spreading center. The Solomon Sea mi-
croplate is just a tear in the Australian plate that has
been moving northward into the New Britain trench
since about 3.5 Ma (Weissel et al., 1982).
Major uplift in eastern New Guinea is well dated
from a change in provenance from continental to
volcanic-rich materials near the FinisterreHuon
forearc terrane at about 4 Ma ( Abbott et al., 1994a, b;
Abbott, 1995). Directly south, collisional deformation
caused the initial unroofing of basement-cored up-
lifts in the Papuan fold and thrust belt at about 4 Ma
according to the apatite fission-track analysis of Hill
and Gleadow (1989). Farther south, renewed uplift in
the Papuan Peninsula is also evident by the appearance
of conglomeratic, molasse-type, deposits (Era beds) in
the Aure trough region (Brown et al., 1975; Pigram
et al., 1989; Davies, 1990), the renewed faulting and
folding in the Aure trough (Kugler, 1993), and a five-
fold increase in sediment accumulation rates in the
Gulf of Papua (Pigram and Symonds, 1991; Wang and
Stein, 1992).
Collisional deformation in eastern New Guinea is
ongoing. Thrust-type earthquakes are found beneath
the FinisterreHuon terranes and along the southern
flank of the Papuan highlands (Abers and McCaffrey,
1988; Sapiie et al., 1999). The juncture between the
active deformation front at the New Britain trench and
the inactive front at the Trobriand trough (Figure 1)
is propagating eastward at a rate between 110 and
210 km/m.y. (68 and 130 mi/m.y.) (Silver et al., 1991;
Abbott et al., 1994a). The active involvement of the
Australian continental basement in collisional orogen-
esis is dated at about 10 Ma in western New Guinea
(dated at 8 Ma near Puncak Jaya) and at about 5 Manear the international border. The Central Range col-
lisional orogenesis has propagated along the length
of the entire island at a rate of about 150 km/m.y.
(93 mi/m.y.).
Implications for Hydrocarbon Exploration in New Guinea
The implications of the model for the Cenozoic tec-
tonics of New Guinea presented in this paper center
on the timing of uplift and erosion, thick sedimenta-
tion, deep burial and heating, and structural trap for-
mation in the different parts of the island. On the
Papua New Guinea side of the island, there have been
two distinct collisional orogenic events since about
35 Ma. In the western half of the Birds Body, one col-
lisional orogenic event since about 12 Ma considers the
regional tectonic relationships.
Collisional tectonism causes complex folding and
thrusting in the mountainous core zone and common-
ly causes modest movements in the basement of the
foreland, where sediments eroded from the rising moun-
tains are deposited. Movements in the foreland base-
ment can be located 100 km (62 mi) or more from the
collision-generated mountain front. The Oligocene tec-
tonic event that caused minor folding in the western
edge of the Birds Head block is probably the east-
ernmost manifestation of collisional tectonism in the
Sulawesi region.
136 Cenozoic Tectonics of New Guinea
-
In central and western New Guinea, the Oligo-
cene sea level fall led to widespread exposure of the
shelf. The Sirga Formation, a well-sorted transgressive
quartz sandstone unit should have, at least locally,
good reservoir characteristics. This unit is highly var-
iable in thickness but has widespread distribution in
both the highly deformed highlands and beneath the
southern foreland basin.
CONCLUSIONS
The Cenozoic tectonic history of New Guinea records
two major orogenic events (emergent mountain build-
ing and erosion) related to arc-continent collision at
northward-dipping subduction zones. The first event,
the Peninsular orogeny, caused the cessation of con-
vergence that began in the Eocene, when the northern
corner of the Australian continental crust jammed the
subduction zone in the Oligocene. The effects are re-
stricted to eastern New Guinea. The Papuan ophiolite
was emplaced above the Owen Stanley metamorphic
belt, and the orogeny generated the Papuan Peninsula.
The uplifted area, similar to present-day New Cale-
donia, was the source of abundant siliciclastic sedi-
ment deposited in the Aure trough.
The Oligocene disconformity and the transgressive
deposition of the quartzose Sirga Formation in western
and central Papua is the result of the about 90-m (300-ft)
drop in sea level between 33 and 30 Ma. Regional
stratigraphic relationships indicate that the Australian
margin in central and western New Guinea remained
a carbonate shelf until the late middle Miocene.
The Central Range orogeny was the event that
formed the present-day shape and topography of New
Guinea. Before the orogeny began, the Australian
margin rise and slope sediments were bulldozed and
metamorphosed. The top of the deformed sediment
pile was locally emergent and eroded in the middle
Miocene (1614 Ma). Uplifts causing widespreadsiliciclastic sedimentation above the Australian conti-
nental basement formed in the latest middle Miocene
at about 12 Ma. Collisional orogenesis involving crys-
talline basement probably began beneath the west-
ernmost Central Range at about 10 Ma, propagated
eastward reaching central New Guinea at about 5 Ma,
and is ongoing in eastern New Guinea. The eastern
edge of the collisional orogenesis has propagated
along the length of the entire island at a rate of about
150 km/m.y. (93 mi/m.y.). Cenozoic tectonism form-
ing the island of New Guinea has generated a network
of folds, faults, and stratigraphic complexities that host
a major hydrocarbon province in the eastern highlands
but is still largely untested in the western highlands.
REFERENCES CITED
Abbott, L. D., 1995, Neogene tectonic reconstruction of theAdelbertFinisterreNew Britain collision, northern PapuaNew Guinea: Journal of Southeast Asian Earth Sciences, v. 11,p. 3351.
Abbott, L. D., E. A. Silver, and J. Galewsky, 1994a, Structuralevolution of a modern arc-continent collision in Papua NewGuinea: Tectonics, v. 13, p. 10071034.
Abbott, L. D., E. A. Silver, P. R. Thompson, M. V. Filewicz, C.Schneider, and Abdoerrias, 1994b, Stratigraphic constraintson the development and timing of arc-continent collision innorthern Papua New Guinea: Journal of Sedimentary Research,v. 64, p. 169183.
Abers, G., and R. McCaffrey, 1988, Active deformation in the NewGuinea fold and thrust belt: Seismological evidence for strike-slip faulting and basement involved thrusting: Journal ofGeophysical Research, v. 93, p. 13,32213,354.
Adams, C. G., 1984, Neogene larger foraminifera, evolutionary andgeological events in the context of datum lanes, in N. Ikebeand R. Tsuchi, eds., Pacific Neogene datum planes: Tokyo, Japan,University of Tokyo Press, p. 4767.
Allison, I., and J. A. Peterson, 1989, Glaciers of Irian Jaya, In-donesia: U.S. Geological Survey Professional Paper 1386-H-1,p. H1H20.
Apthorpe, M., 1988, Cainozoic depositional history of the NorthWest Shelf, in P. G. Purcell and R. R. Purcell, eds., The NorthWest Shelf, Australia: Petroleum Exploration Society of Aus-tralia, p. 5584.
Audley-Charles, M. G., 1991, Tectonics of the New Guinea area:Annual Reviews of Earth and Planetary Sciences, v. 19, p. 1741.
Australasian Petroleum Company, 1961, Geological results ofpetroleum exploration in western Papua, 19371961: Geo-logical Society of Australia Journal, v. 8, p. 1133.
Bar, C. B., H. J. Cortel, and A. E. Escher, 1961, Geological resultsof the Star Mountains (Sterrengebergte) expedition: NovaGuinea: Geology, v. 4, p. 4099.
Belford, D. J., D. Burger, S. K. Skwarko, and B. Kummel, 1974,Foraminifera from the Ilaga Valley, Nassau Range, Irian Jaya:Bureau of Mineral Resources Geology and GeophysicsBulletin, v. 150, p. 126.
Bergman, S. C., D. Q. Coffield, J. P. Talbot, and R. A. Garrard,1996, Tertiary tectonic and magmatic evolution of westernSulawesi and the Makassar Strait, Indonesia: Evidence for a Mio-cene continent-continent collision: Geological Society (London)Special Publication 106, p. 391429.
Brash, R. A., L. F. Henage, B. H. Harahap, D. T. Moffat, and R. W.Tauer, 1991, Stratigraphy and depositional history of the NewGuinea limestone group, Lengguru, Irian Jaya: Proceedings ofthe Indonesian Petroleum Association, v. 20, p. 6884.
Brothers, R. N., 1974, Kaikoura orogeny in Northland, NewZealand: New Zealand Journal of Geology and Geophysics,v. 17, p. 118.
Brown, C. M., P. E. Pieters, and G. P. Robinson, 1975, Stratigraphicand structural development of the Aure trough and adjacentshelf and slope areas: The Australian Petroleum ExplorationAssociation Journal, v. 15, p. 6171.
Carman, G. J., 1990, Occurrence and nature of Eocene strata in theeastern Papuan Basin, in G. J. Carman and Z. Carman, eds.,
Quarles van Ufford and Cloos 137
-
Petroleum exploration in Papua New Guinea: Proceedings ofthe First PNG Petroleum Convention, Port Moresby, February1214, 1990, p. 169183.
Carman, G. J., and Z. Carman, eds., 1990, Petroleum exploration inPapua New Guinea: Proceedings of the First PNG PetroleumConvention, Port Moresby, February 1214, 1990, 597 p.
Carman, G. J., and Z. Carman, eds., 1993, Petroleum exploration inPapua New Guinea: Proceedings of the Second PNG Petro-leum Convention, Port Moresby, May 31June 2, 1993, 687 p.
Clague, D. A., and G. B. Dalrymple, 1989, Tectonics, geochronol-ogy, and origin of the HawaiianEmperor volcanic chain:Geological Society of America, The Geology of North America,v. N, p. 188217.
Cloos, M., 1993, Lithospheric buoyancy and collisional orogenesis:Subduction of oceanic plateaus, continental margins, islandarcs, spreading ridges, and seamounts: Geological Society ofAmerica Bulletin, v. 105, p. 715737.
Cloos, M., and R. L. Shreve, 1988a, Subduction-channel model ofprism accretion, melange formation, sediment subduction, andsubduction erosion at convergent plate margins: 1. Backgroundand description: Pure and Applied Geophysics, v. 128, p. 455500.
Cloos, M., and R. L. Shreve, 1988b, Subduction-channel model ofprism accretion, melange formation, sediment subduction, andsubduction erosion at convergent plate margins: 2. Implicationsand discussions: Pure Applied Geophysics., v. 128, p. 501545.
Cooper, P., and B. Taylor, 1987, Seismotectonics of New Guinea: Amodel for arc reversal following arc-continent collision: Tec-tonics, v. 6, p. 5368.
Cox, A., and D. Engebretson, 1985, Change in motion of Pacificplate at 5 Myr BP: Nature, v. 313, p. 472474.
Davies, H. L., 1971, Peridotite-gabbro-basalt complex in easternPapua: An overthrust plate of oceanic mantle and crust: Bu-reau of Mineral Resources Geology and Geophysics Bulletin,v. 128, 48 p.
Davies, H. L., 1990, Structure and evolution of the border region ofNew Guinea, in G. J.Carman and Z. Carman, eds., Petroleumexploration in Papua New Guinea: Proceedings of the FirstPNG Petroleum Convention, Port Moresby, February 1214,1990, p. 245269.
Davies, H. L., and A. L. Jaques, 1984, Emplacement of ophiolite inPapua New Guinea: Geological Society (London) Special Pub-lication 13, p. 341349.
Davies, H. L., and I. E. Smith, 1971, Geology of eastern Papua:Geological Society of America Bulletin, v. 82, p. 32993312.
Davies, H. L., and R. G. Warren, 1988, Origin of eclogite-bearing,domed, layered, metamorphic complexes (core complexes)in the Islands, Papua New Guinea: Tectonics, v. 7, p. 121.
Davies, H. L., P. A. Symonds, and I. D. Ripper, 1984, Structure andevolution of the southern Solomon Sea region: Bureau of Min-eral Resources Journal of Australian Geology and Geophysics,v. 9, p. 4968.
Davies, P. J., P. A. Symonds, D. A. Feary, and C. J. Pigram, 1989,The evolution of the carbonate platforms of northeast Aus-tralia: SEPM Special Publication 44, p. 233258.
Davies, P. J., P. A. Symonds, D. A. Feary, and C. J. Pigram, 1991,The evolution of the carbonate platforms of northeast Aus-tralia: Geological Society of Australia Special Publication 18,p. 4478.
DeMets, C., R. G. Gordon, D. F. Argus, and S. Stein, 1994, Effectof recent revisions to the geomagnetic reversal time scale onestimates of current plate motions: Geophysical Research Let-ters, v. 21, p. 21912194.
Dewey, J. F., and J. M. Bird, 1970, Mountain belts and the new globaltectonics: Journal of Geophysical Research, v. 75, p. 26252647.
Dickinson, W. R., 1985, Interpreting provenance relations fromdetrital modes of sandstones, in G. G. Zuffa, ed., Provenanceof arenites: Dordrecht, Netherlands, Reidel, p. 333361.
Dickinson, W. R., and C. A. Suczek, 1979, Plate tectonics andsandstone compositions: AAPG Bulletin, v. 63, p. 21642182.
Doust, H., 1990, Geology of the Sepik basin, Papua New Guinea, inG. J. Carman and Z. Carman, eds., Petroleum exploration inPapua New Guinea: Proceedings of the First PNG PetroleumConvention, Port Moresby, February 1214, 1990, p. 461478.
Dow, D. B., 1977, A geological synthesis of Papua New Guinea:Bureau of Mineral Resources Geology and Geophysics Bul-letin, v. 201, 41 p.
Dow, D. B., and R. Sukamto, 1984a, Late Tertiary to Quaternarytectonics of Irian Jaya: Episodes, v. 7, p. 39.
Dow, D. B., and R. Sukamto, 1984b, Western Irian Jaya: The end-product of oblique plate convergence in the late Tertiary:Tectonophysics, v. 106, p. 109139.
Dow, D. B., J. A. J. Smit, J. H. C. Bain, and R. J. Ryburn, 1972,Geology of the south Sepik region, New Guinea: Bureau ofMineral Resources Geology and Geophysics Bulletin, v. 133,88 p.
Dow, D. B., G. P. Robinson, U. Hartono, and N. Ratman, 1988,Geology of Irian Jaya: Irian Jaya Geological Mapping Project,Geological Research and Development Center, Indonesia, incooperation with the Bureau of Mineral Resources, Australiaon behalf of the Department of Mines and Energy, Indonesiaand the Australian Development Assistance Bureau, 298 p.
Edwards, A. B., 1950, The petrology of the Miocene sediments ofAure trough, Papua: Proceedings of the Royal Society of Vic-toria, v. 60, p. 123148.
Francis, G., R. Rogerson, D. W. Haig, and J. Sari, 1986, Neogenestratigraphy, sedimentation and petroleum potential of theOiapuYule IslandOroi Region, Papua New Guinea: Geo-logical Society of Malaysia Bulletin, v. 19, p. 123152.
Froidevaux, C. M., 1978, Tertiary tectonic history of the Salawatiarea, Irian Jaya, Indonesia: AAPG Bulletin, v. 62, p. 11271150.
Fulthorpe, C. S., R. M. Carter, K. G. Miller, and J. Wilson, 1996,Marshall paraconformity: A mid-Oligocene record of inceptionof the Antarctic circumpolar current and coeval glacio-eustaticlowstand: Marine and Petroleum Geology, v. 13, p. 6177.
Gibson-Robinson, C., and H. Soedirdja, 1986, Transgressive devel-opment of Miocene reefs, Salawati basin, Irian Jaya: Proceed-ings of the Indonesian Petroleum Association, v. 15, p. 377403.
Hall, R., 1996, Reconstructing Cenozoic SE Asia: Geological So-ciety (London) Special Publication 106, p. 153184.
Hamilton, W., 1979, Tectonics of the Indonesian region: U.S.Geological Survey Professional Paper 1078, 345 p.
Haq, B. U., J. Hardenbol, and P. R. Vail, 1987, Chronology of fluc-tuating sea levels since the Triassic: Science, v. 235, p. 11561167.
Hayward, B. W., F. J. Brook, and M. J. Isaac, 1989, Cretaceous tomiddle Tertiary stratigraphy, petrology and tectonic history ofNorthland, New Zealand: Bulletin of the Royal Society of NewZealand, v. 26, p. 4764.
Hilde, T. W., C. S. Uyeda, and L. Kroenke, 1977, Evolution of thewestern Pacific and its margin: Tectonophysics, v. 38, p. 145165.
Hill, K. C., and A. J. W. Gleadow, 1989, Uplift and thermal historyof the Papuan fold belt, Papua New Guinea: Apatite fission-track analysis: Australian Journal of Earth Sciences, v. 36,p. 515539.
Hill, K. C., A. Grey, D. Foster, and R. Barrett, 1993, An alternativemodel for the Oligo-Miocene evolution of northern PNG and
138 Cenozoic Tectonics of New Guinea
-
the Sepik-Ramu basins, in G. J. Carman and Z. Carman, eds.,Petroleum exploration in Papua New Guinea: Proceedings ofthe Second PNG Petroleum Convention, Port Moresby, May31June 2, 1993, p. 241259.
Hobson, D. M., 1986, A thin skinned model for the Papuan thrustbelt and some implications for hydrocarbon exploration:Australian Petroleum Exploration Association Journal, v. 26,p. 214224.
Home, P. C., D. G. Dalton, and J. Brannan, 1990, Geologicalevolution of the western Papuan basin, in G. J. Carman andZ. Carman, eds., Petroleum exploration in Papua New Guinea:Proceedings of the First PNG Petroleum Convention, PortMoresby, February 1214, 1990, p. 107117.
Jacobi, R. D., 1981, Peripheral bulge-causal mechanism for theLower/Middle Ordovician unconformity along the western mar-gin of the northern Appalachians: Earth and Planetary ScienceLetters, v. 56, p. 245251.
Jolivet, L., P. Huchon, and C. Rangin, 1989, Tectonic setting ofwestern Pacific marginal basins: Tectonophysics, v. 160, p. 2347.
Katili, J. A., 1986, Geology and hydrocarbon potential of theArafura Sea: Future petroleum provinces of the world: AAPGMemoir 40, p. 487501.
Katili, J. A., 1991, Tectonic evolution of eastern Indonesia and itsbearing on the occurrence of hydrocarbons: Marine and Petro-leum Geology, v. 8, p. 7083.
Klimchuk, G. A., 1993, Provenance and depositional setting of thePliocene Era Formation, Aure fold and thrust belt, Papua NewGuinea: M.A. thesis, University of Texas at Austin, Austin,130 p.
Koesoemadinata, R. P., 1978, Tertiary carbonate sedimentation inIrian Jaya with special reference to the northern part of theBintuni basin, in Proceedings of the Carbonate Seminar, Ja-karta, September 1219, 1976: Indonesian Petroleum Asso-ciation, p. 7992.
Kroenke, L. W., 1984, Cenozoic tectonic development of thesouthwest Pacific: United Nations Economic and Social Com-mission for Asia and the Pacific, Coordinating Committee forGeoscience Programmes in East and Southeast AsiaSouthPacific Applied Geoscience Commission Technical Bulletin,v. 6, 126 p.
Kugler, A., 1967, The stratigraphy, structure and tectonics of theKukukuku Lobe, Permit 22, Papua: Ph.D. dissertation, Uni-versity of Tasmania, Hobart, 365 p.
Kugler, K. A., 1993, Detailed analysis from seismic data of thestructure within the Aure fold and thrust belt, Gulf of Papua,Papua New Guinea, in G. J. Carman and Z. Carman, eds., Pe-troleum exploration and development in Papua New Guinea:Proceedings of the Second PNG Petroleum Convention, PortMoresby, May 31June 2, 1993, p. 399411.
Lunt, P., and R. Djaafar, 1991, Aspects of the stratigraphy ofwestern Irian Jaya and implications for the development ofsandy facies: Proceedings of the Indonesian Petroleum Asso-ciation, v. 20, p. 107124.
McLennan, J. M., J. S. Rasidi, R. L. Holmes, and G. C. Smith, 1990,The geology and petroleum potential of the western ArafuraSea: The Australian Petroleum Exploration Association Jour-nal, v. 30, p. 91106.
McMahon, T. P., 2000a, Origin of syn- to post-collisional magmatismin New Guinea: Buletin Geologi, Jurusan Teknik Geologi-Institute Teknologi Bandung, v. 32, no. 2, p. 89104.
McMahon, T. P., 2000b, Magmatism in an arc-continent collisionzone: An example from Irian Jaya (western New Guinea),Indonesia: Buletin Geologi, Jurusan Teknik Geologi-InstituteTeknologi Bandung, v. 32, no. 1, p. 122.
Milsom, J., 1985, New Guinea and the western Melanesian arcs, in
A. E. M. Nairn, F. G. Stehli, and S. Uyeda, eds., The oceanbasins and margins: The Pacific Ocean 7A: New York, PlenumPress, p. 551605.
Page, R. W., 1976, Geochronology of igneous and metamorphicrocks in the New Guinea Highlands: Bureau of Mineral Re-sources Geology and Geophysics Bulletin, v. 162, 117 p.
Parrot, J. F., and F. Dugas, 1980, The disrupted ophiolitic belt ofthe southwest Pacific: Evidence of an Eocene subduction zone:Tectonophysics, v. 66, p. 349372.
Pieters, P. E., C. J. Pigram, D. S. Trail, D. B. Dow, N. Ratman, andR. Sukamto, 1983, The stratigraphy of western Irian Jaya: Bul-letin Geologic Research and Development Center (Bandung,Indonesia), v. 8, p. 1448.
Pigram, C. J., and H. L. Davies, 1987, Terranes and the accretionhistory of the New Guinea orogen: Bureau of Mineral Re-sources Journal of Australian Geology and Geophysics, v. 10,p. 193212.
Pigram, C., and H. Panggabean, 1983, Geological data recordWaghete ( Yapekopra), scale 1:250,000 sheet area, Irian Jaya:Irian Jaya Geological Mapping Project, Geological Researchand Development Center, Indonesia, in cooperation with theBureau of Mineral Resources, Australia, on behalf of the De-partment of Mines and Energy, Indonesia, and the AustralianDevelopment Assistance Bureau, 126 p.
Pigram, C., and H. Panggabean, 1984, Rifting of the northernmargin of the Australian continent and the origins of somemicrocontinents in eastern Indonesia: Tectonophysics, v. 107,p. 331353.
Pigram, C., and P. A. Symonds, 1991, A review of the timing of themajor tectonic events in the New Guinea orogen: Journal ofSoutheast Asian Earth Sciences, v. 6, p. 307318.
Pigram, C., A. B. Challinor, F. Hasibuan, E. Rusmana, and U. Hartono,1982, Lithostratigraphy of the Misool archipelago, Irian Jaya,Indonesia: Geologie en Mijnbouw, v. 61, p. 265279.
Pigram, C., P. J. Davies, D. A. Feary, and P. A. Symonds, 1989,Tectonic controls on carbonate platform evolution in southernPapua New Guinea: Passive margin to foreland basin: Geology,v. 17, p. 199202.
Pigram, C., P. J. Davies, D. A. Feary, P. A. Symonds, and G. C. H.Chaproniere, 1990, Controls on Tertiary carbonate platformevolution in the Papuan Basin: New play concepts, in G. J.Carman and Z. Carman, eds., Petroleum exploration in PapuaNew Guinea: Proceedings of the First PNG Petroleum Con-vention, Port Moresby, February 1214, 1990, p. 185195.
Pollitz, F. F., 1986, Pliocene change in Pacific-plate motion: Nature,v. 320, p. 738741.
Quarles van Ufford, A. I., 1996, Stratigraphy, structural geology,and tectonics of a young forearc-continent collision, westernCentral Range, Irian Jaya (western New Guinea), Indonesia:Ph.D. Dissertation, University of Texas at Austin, Austin, Texas,420 p., 8 enclosures.
Robinson, G. P., R. J. Ryburn, S. L. Tobing, and A. Achdan, 1988,Steenkool (Wasior)Kaimana 1:250,000 sheet area geologicaldata record: Irian Jaya Geological Mapping Project, GeologicalResearch and Development Center, Indonesia, in cooperationwith the Bureau of Mineral Resources, Australia, on behalf ofthe Department of Mines and Energy, Indonesia, and the Aus-tralian Development Assistance Bureau, 153 p.
Rossetter, R. J., 1978, New Guinea limestone group BomberaiPeninsula, Irian Jaya, in Proceedings of the Carbonate Seminar,Jakarta, September 1219, 1976: Indonesian Petroleum Asso-ciation, p. 7992.
Sapiie, B., D. H. Natawidjaya, and M. Cloos, 1999, Strike-sliptectonics of New Guinea: Transform motion between theCaroline and Australian p