tectonica del cenozoico de nueva guinea

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

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 (

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

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

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.


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 (

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.


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.


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


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.


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


(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.


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


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.


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


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


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.


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.

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